Editing IPCC:AR6/WGI/Chapter-2 (section)

=== 2.3.3 Ocean ===

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This section focuses on large-scale changes in a subset of physical components of the ocean (Cross-Chapter Box 2.2). [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] assesses the role of the ocean in Earth system heating and evaluates the Earth’s energy budget. [[IPCC:Wg1:Chapter:Chapter-9|Chapter 9]] undertakes a holistic assessment of changes in the ocean integrating observations, modelling and theoretical understanding. [[IPCC:Wg1:Chapter:Chapter-11|Chapter 11]] assesses extremes such as marine heat waves and storm surges. SSTs are assessed in [[#2.3.1.1|Section 2.3.1.1]] as they constitute a critical component of GMST estimation.

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<span id="ocean-temperature-heat-content-and-thermal-expansion"></span>
==== 2.3.3.1 Ocean Temperature, Heat Content and Thermal Expansion ====

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AR5 assessed that since 1971, global ocean warming was ''virtually certain'' for the upper 700 m and ''likely'' for the 700–2000 m layer. The SROCC reported linear warming trends for the 0–700 m and 700–2000 m layers of the ocean, respectively, of 4.35 ± 0.8 and 2.25 ± 0.64 ZJ yr <sup>–1</sup> over 1970–2017; 6.28 ± 0.48 and 3.86 ± 2.09 ZJ yr <sup>–1</sup> over 1993–2017; and 5.31 ± 0.48 and 4.02 ± 0.97 ZJ yr <sup>–1</sup> over 2005–2017. Both AR5 and SROCC assessed that the ocean below 2000 m had ''likely'' warmed since 1992. The SROCC reported global mean thermosteric sea level (ThSL) rise, associated with thermal expansion of the ocean, with a trend of 0.89 ± 0.05 mm yr <sup>–1</sup> for 1970–2015; 1.36 ± 0.40 mm yr <sup>–1</sup> for 1993–2015; and 1.40 ± 0.40 mm yr <sup>–1</sup> for 2006–2015, and also reported that the rate of ocean warming over 1993–2017 had ''likely'' more than doubled since 1969–1992.

New ocean heat content (OHC) reconstructions derived from paleo proxies ( [[#Bereiter--2018|Bereiter et al., 2018]] ; [[#Baggenstos--2019|Baggenstos et al., 2019]] ; [[#Shackleton--2019|Shackleton et al., 2019]] ; [[#Gebbie--2021|Gebbie, 2021]] ) indicate that the global ocean warmed by 2.57°C ± 0.24°C, at an average rate of about 0.3°C ka <sup>–1</sup> (equivalent to an OHC change rate of 1.3 ZJ yr <sup>–1</sup> ) from the LGM (about 20 ka) to the early Holocene (about 10 ka; Section 9.2.2.1 and Figure 9.9). Over the LDT, ocean warming occurred in two stages, offset by some heat loss during the Antarctic Cold Reversal (14.58–12.75 ka). Only during a short period of rapid warming at the end of the Younger Dryas (12.75–11.55 ka) were rates comparable to those observed since the 1970s ( [[#Bereiter--2018|Bereiter et al., 2018]] ; [[#Shackleton--2019|Shackleton et al., 2019]] ). Ice cores imply a small decrease in the global mean ocean temperature during the early Holocene (&lt;0.4°C) ( [[#Bereiter--2018|Bereiter et al., 2018]] ; [[#Baggenstos--2019|Baggenstos et al., 2019]] ). Sediment cores from the equatorial Pacific and Atlantic Ocean (0–1000 m) indicate a stronger regional cooling (compared to mean ocean temperature) of 1.0°C ± 0.7°C to 1.8°C ± 0.4°C from the early/mid-Holocene to ca.1750 CE ( [[#Rosenthal--2013|Rosenthal et al., 2013]] , 2017; [[#Morley--2014|Morley et al., 2014]] ; [[#Kalansky--2015|Kalansky et al., 2015]] ). Sediment cores from the western equatorial Pacific suggest 0.8°C ± 0.1°C higher temperatures in the upper 700 m of the ocean during 950–1100 CE compared to 1400–1750 CE. These changes are consistent with a global estimate derived from combined surface and subsurface ocean temperature proxy records ( [[#PAGES%202k%20Consortium--2013|PAGES 2k Consortium, 2013]] ; [[#McGregor--2015|McGregor et al., 2015]] ). A combined study of model and observational data further confirmed these results, treating temperature as a passive tracer ( [[#Gebbie--2019|Gebbie and Huybers, 2019]] ) and addressing the role of circulation dynamics ( [[#Scheen--2020|Scheen and Stocker, 2020]] ). Collectively, the proxy records indicate a global OHC decrease of about 400 ± 70 ZJ (about 170 ± 100 ZJ in the Pacific) in the upper 700 m between 950–1100 CE and 1400–1750 CE, and also suggest that the deep Pacific is still adjusting to this cooling ( [[#Rosenthal--2013|Rosenthal et al., 2013]] ), partially offsetting the global increase since 1750 CE ( [[#Gebbie--2019|Gebbie and Huybers, 2019]] ; [[#Gebbie--2021|Gebbie, 2021]] ).

For the instrumental era, since AR5 and SROCC, new and updated OHC and ThSL observation-based analyses ( [[#Johnson--2020|Johnson et al., 2020]] ; [[#von%20Schuckmann--2020|von Schuckmann et al., 2020]] ) enhance an existing large ensemble of direct and indirect OHC estimates (Figure 2.26), although some rely to varying degrees upon information from ocean-climate models. Direct estimates benefit from improved: bias adjustments (e.g., [[#Cheng--2018|Cheng et al., 2018]] ; [[#Leahy--2018|Leahy et al., 2018]] ; [[#Palmer--2018|Palmer et al., 2018]] ; [[#Ribeiro--2018|Ribeiro et al., 2018]] ; [[#Wang--2018|]] [[#Wang--2018|B. Wang et al., 2018]] ; [[#Bagnell--2020|Bagnell and DeVries, 2020]] ; [[#Gouretski--2020|Gouretski and Cheng, 2020]] ); interpolation methods ( [[#Kuusela--2018|Kuusela and Stein, 2018]] ; [[#Su--2020|Su et al., 2020]] ); and characterization of sources of uncertainty (e.g., [[#Good--2017|Good, 2017]] ; [[#Wunsch--2018|Wunsch, 2018]] ; [[#Allison--2019|Allison et al., 2019]] ; [[#Garry--2019|Garry et al., 2019]] ; [[#Meyssignac--2019|Meyssignac et al., 2019]] ; [[#Palmer--2021|Palmer et al., 2021]] ), including those originating from forced and intrinsic ocean variability ( [[#Penduff--2018|Penduff et al., 2018]] ). After 2006 direct OHC estimates for the upper 2000 m layer benefit from the near-global ARGO array with its superior coverage over 60°S–60°N ( [[#Roemmich--2019|Roemmich et al., 2019]] ). Indirect estimates include OHC and ThSL series inferred from satellite altimetry and gravimetry since 2003 ( [[#Meyssignac--2019|Meyssignac et al., 2019]] ), the passive uptake of OHC (ThSL) at centennial timescales inferred from observed SST anomalies, and time-invariant circulation processes from an ocean state estimation (e.g., [[#Zanna--2019|Zanna et al., 2019]] ). [[#Resplandy--2019|Resplandy et al. (2019)]] estimate the rate of global OHC uptake over 1991–2016 from changes in atmospheric composition and physical relationships based on CMIP5 model simulations. The uncertainties are broader than from direct estimates but the estimate is qualitatively consistent.

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[[File:b5c47168906cc523362848773535b3bc IPCC_AR6_WGI_Figure_2_26.png]]

'''Figure 2.2''' '''6 |''' '''Changes in ocean heat content (OHC).''' Changes are shown over '''(a)''' full depth of the ocean from 1871–2019 from a selection of indirect and direct measurement methods. The series from Table 2.7 is shown in solid black in both (a) and (b) (see Table 2.7 caption for details). '''(b)''' as (a) but for 0–2000 m depths only and reflecting the broad range of available estimates over this period. For further details see chapter data table (Table 2.SM.1).

Collectively, the new and updated analyses strengthen AR5 and SROCC findings of a sustained increase in global OHC (Figure 2.26 and Table 2.7) and associated ThSL rise. Larger warming rates are observed in the upper 700 m compared to deeper layers, with more areas exhibiting significant warming than significant cooling ( [[#Johnson--2020|Johnson and Lyman, 2020]] ). There is an improved consistency among available estimates of OHC rates in the upper 2000 m since 2006. [[#Cheng--2020|Cheng et al. (2020)]] , [[#von%20Schuckmann--2020|von Schuckmann et al. (2020)]] and [[#Johnson--2020|Johnson et al. (2020)]] have further confirmed that the central estimates of rates of OHC change in the upper 2000 m depths have increased after 1993 and particularly since 2010 ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.1.3|Section 3.5.1.3]] and Figures 2.26 and 3.26), although uncertainties are large (Table 2.7). Ocean reanalyses support findings of continued upper ocean warming ( [[#Balmaseda--2013|Balmaseda et al., 2013]] ; [[#von%20Schuckmann--2018|von Schuckmann et al., 2018]] ; [[#Meyssignac--2019|Meyssignac et al., 2019]] ), albeit with higher spread than solely observational estimates, particularly in the poorly sampled deep ocean below 2000 m ( [[#Storto--2017|Storto et al., 2017]] ; [[#Palmer--2018|Palmer et al., 2018]] ).

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'''Table 2.7 |''' '''Rates of global ocean heat content (OHC) and global mean thermosteric sea level (ThSL) change for four depth integrations over different periods.''' For the period up to 1971, the assessment for all depth layers is based on [[#Zanna--2019|Zanna et al. (2019)]] . From 1971 onwards, consistent with AR5, Domingues et al. (2008, updated) is the central estimate for 0–700 m along with uncertainty from a five-member ensemble ( [[#Domingues--2008|Domingues et al., 2008]] , updated; [[#Levitus--2012|Levitus et al., 2012]] ; [[#Good--2013|Good et al., 2013]] ; [[#Cheng--2017|Cheng et al., 2017]] ; [[#Ishii--2017|Ishii et al., 2017]] ), following the approach of [[#Palmer--2021|Palmer et al. (2021)]] . Similarly, [[#Ishii--2017|Ishii et al. (2017)]] is the central estimate for 700–2000 m with uncertainty based on a 3-member ensemble ( [[#Levitus--2012|Levitus et al., 2012]] ; [[#Cheng--2017|Cheng et al., 2017]] ; [[#Ishii--2017|Ishii et al., 2017]] ). For depths below 2000 m, both central estimate and uncertainty are from Purkey and Johnson (2010, updated). In cases when OHC estimates do not have a ThSL counterpart (e.g., [[#Good--2013|Good et al., 2013]] ; [[#Cheng--2017|Cheng et al., 2017]] ), OHC was converted into ThSL using the average linear regression coefficients for 0–700 m and 700–2000 m from all available ensemble members. For consistency with the energy and sea-level budgets presented in Chapters 7 and 9, reported rates are based on the difference between the first and last annual mean value in each period ( [[#Palmer--2021|Palmer et al., 2021]] , Box 7.2, Cross-Chapter Box 9.1). N/A indicates not applicable. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

{| class="wikitable"
|-
| rowspan="2"| Depth

| rowspan="2"| Period

| rowspan="2"| OHC Rate (ZJ y <sup>–1</sup> )

| rowspan="2"| ThSL Rate (mm yr <sup>–1</sup> )

| colspan="2"| Relative Full Ocean Depth Contribution

|-
| OHC

| ThSL

|-
| rowspan="5"| '''0–700 m'''

| 1901–1990

| 2.50

[1.16 to 3.85]

| 0.31

[0.16 to 0.45]

| 81%

| 86%

|-
| 1901–2018

| 3.11

[2.18 to 4.04]

| 0.40

[0.30 to 0.50]

| 66%

| 74%

|-
| 1971–2018

| 5.14

[3.46 to 6.82]

| 0.71

[0.51 to 0.90]

| 61%

| 70%

|-
| 1993–2018

| 6.06

[4.56 to 7.55]

| 0.89

[0.69 to 1.10]

| 58%

| 68%

|-
| 2006–2018

| 6.28

[4.06 to 8.50]

| 0.91

[0.51 to 1.31]

| 54%

| 65%

|-
| rowspan="5"| '''700–2000 m'''

| 1901–1990

| 0.50

[–0.59 to 1.60]

| 0.04

[–0.07 to 0.16]

| 16%

| 11%

|-
| 1901–2018

| 1.26

[0.43 to 2.09]

| 0.11

[0.02 to 0.19]

| 27%

| 20%

|-
| 1971–2018

| 2.62

[2.04 to 3.20]

| 0.23

[0.16 to 0.31]

| 31%

| 23%

|-
| 1993–2018

| 3.31

[2.40 to 4.22]

| 0.30

[0.19 to 0.41]

| 32%

| 23%

|-
| 2006–2018

| 4.14

[2.41 to 5.86]

| 0.36

[0.15 to 0.58]

| 36%

| 26%

|-
| rowspan="5"| '''&gt;2000 m'''

| 1901–1990

| 0.07

[0.02 to 0.12]

| 0.01

[0.00 to 0.01]

| 2%

| 3%

|-
| 1901–2018

| 0.32

[0.18 to 0.46]

| 0.03

[0.02 to 0.05]

| 7%

| 6%

|-
| 1971–2018

| 0.66

[0.33 to 0.99]

| 0.07

[0.03 to 0.10]

| 8%

| 7%

|-
| 1993–2018

| 1.15

[0.58 to 1.72]

| 0.12

[0.06 to 0.18]

| 11%

| 9%

|-
| 2006–2018

| 1.15

[0.58 to 1.72]

| 0.12

[0.06 to 0.18]

| 10%

| 9%

|-
| rowspan="5"| '''Full-depth'''

| 1901–1990

| 3.08

[1.36 to 4.79]

| 0.36

[0.17 to 0.54]

| N/A

| N/A

|-
| 1901–2018

| 4.68

[3.45 to 5.92]

| 0.54

[0.40 to 0.68]

| N/A

| N/A

|-
| 1971–2018

| 8.42

[6.08 to 10.77]

| 1.01

[0.73 to 1.29]

| N/A

| N/A

|-
| 1993–2018

| 10.52

[7.76 to 13.28]

| 1.31

[0.95 to 1.66]

| N/A

| N/A

|-
| 2006–2018

| 11.57

[7.20 to 15.94]

| 1.39

[0.74 to 2.05]

| N/A

| N/A

|}

In summary, current multi-decadal to centennial rates of OHC gain are greater than at any point since the last deglaciation ( ''medium confidence'' ). At multi-centennial timescales, changes in OHC based upon proxy indicators demonstrate a tight link with surface temperature changes during the last deglaciation ( ''high confidence'' ), as well as during the Holocene and CE ( ''low confidence'' ). It is ''likely'' the global ocean has warmed since 1871, consistent with the observed increase in sea surface temperature. It is ''virtually certain'' that OHC increased between 1971 and 2018 in the upper 700 m and ''very likely'' in the 700–2000 m layer, with ''high confidence'' since 2006. It is ''likely'' the OHC below 2000 m has increased since 1992. Confidence in the assessment of multi-decadal OHC increase is further strengthened by consistent closure of both global sea level and energy budgets (Section 7.2.2.2, Box 7.2, Cross-Chapter Box 9.1).

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==== 2.3.3.2 Ocean Salinity ====

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The AR5 concluded that subtropical regions of high salinity (where evaporation dominates over precipitation) had become more saline, while regions of low salinity (mostly in the tropics and high latitudes) had ''very likely'' become fresher since the 1950s, both at the near-surface, and in the ocean interior along ventilation pathways. From 1950 to 2008, the mean surface contrast between high- and low-salinity regions increased by 0.13 [0.08 to 0.17] (PSS-78, [[#UNESCO/ICES/SCOR/IAPSO--1981|UNESCO/ICES/SCOR/IAPSO, 1981]] ). Across basins, the Atlantic Ocean had become saltier and the Pacific and Southern Oceans had freshened ( ''very likely'' ) .

Prior to the instrumental record, reconstructions of near-surface salinity change are accomplished by combining isotopic and elemental proxy data from microfossil plankton shells and skeletons preserved in deep-sea sediments. These data highlight changes in the salinity contrast between the Pacific and Atlantic oceans during past glacials ( [[#Broecker--1989|Broecker, 1989]] ; [[#Keigwin--2007|Keigwin and Cook, 2007]] ; [[#Costa--2018|Costa et al., 2018]] ) and for repeated episodes of increased subtropical salinity ( [[#Schmidt--2004|Schmidt et al., 2004]] , 2006) and subpolar freshening ( [[#Cortijo--1997|Cortijo et al., 1997]] ; [[#Thornalley--2011|Thornalley et al., 2011]] ) in the North Atlantic ocean. These episodes were associated with disruptions to the large-scale deep ocean circulation ( [[#Buizert--2015|Buizert et al., 2015]] ; [[#Henry--2016|Henry et al., 2016]] ; [[#Lynch-Stieglitz--2017|Lynch-Stieglitz, 2017]] ). Further quantification of paleo salinity changes is complicated by incomplete understanding of proxy-salinity relationships and the relative influence of atmospheric and ocean processes across regions and paleo periods ( [[#Rohling--2007|Rohling, 2007]] ; [[#LeGrande--2011|LeGrande and Schmidt, 2011]] ; [[#Holloway--2016|Holloway et al., 2016]] ; [[#Conroy--2017|Conroy et al., 2017]] ).

Since AR5, new and extended multi-decadal analyses have strengthened the observational support for increased contrast between high and low near-surface salinity regions and inter-basin contrast since the mid-20th century (Section 9.2.2.2; [[#Durack--2010|Durack and Wijffels, 2010]] ; [[#Good--2013|Good et al., 2013]] ; [[#Skliris--2014|Skliris et al., 2014]] ; [[#Aretxabaleta--2017|Aretxabaleta et al., 2017]] ; [[#Cheng--2020|Cheng et al., 2020]] ). These analyses employ different statistical algorithms for interpolation, and only [[#Cheng--2020|Cheng et al. (2020)]] use CMIP5 model simulations to constrain observation-based signals in data-sparse regions.

The 1950–2019 trends reveal near-surface freshening of the northern and western Warm (and fresh) Pool of the Pacific and increased salinity maxima in the subtropical Atlantic, strengthening the inter-basin contrast (Figure 2.27a). There are indications that the subpolar freshening and subtropical salinification of the Atlantic ocean may extend back to at least 1896 ( [[#Friedman--2017|Friedman et al., 2017]] ). Over recent decades, new observations from Argo floats and ocean reanalyses provide general support that changes in the global patterns of near-surface salinity contrast are broadly associated with an intensification of the hydrological cycle (Sections 2.3.1.3.5 and 8.3.1.1). However, this assessment is complicated by changing observational techniques ( [[IPCC:Wg1:Chapter:Chapter-1#1.5.1|Section 1.5.1]] ), temporally and spatially inhomogeneous sampling and uncertainties in interpolation algorithms and the substantial influence of modes of natural variabiltity and ocean circulation processes over interannual timescales ( [[#Skliris--2014|Skliris et al., 2014]] ; [[#Durack--2015|Durack, 2015]] ; [[#Grist--2016|Grist et al., 2016]] ; [[#Aretxabaleta--2017|Aretxabaleta et al., 2017]] ; [[#Vinogradova--2017|Vinogradova and Ponte, 2017]] ; [[#Liu--2020|Liu et al., 2020]] ). Following AR5, based on the updated analysis from [[#Durack--2010|Durack and Wijffels (2010)]] which infills in situ gaps to recover large-scale patterns the mean salinity contrast between high- and low- near-surface salinity regions increased by 0.14 [0.07 to 0.20] from 1950 to 2019.

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[[File:f2b2b1d6eb71e00065f0e9c6df0f57a7 IPCC_AR6_WGI_Figure_2_27.png]]

'''Figure 2.2''' '''7 |''' '''Changes in ocean salinity.''' Estimates of salinity trends using a total least absolute differences fitting method for '''(a)''' global near-surface salinity (SSS) changes and '''(b)''' global zonal mean subsurface salinity changes. Black contours show the associated climatological mean salinity (either near-surface (a) or subsurface (b)) for the analysis period (1950–2019). Both panels represent changes in Practical Salinity Scale 1978 [PSS-78], per decade. In both panels green denotes freshening regions and orange/brown denotes regions with enhanced salinities (‘×’ marks denote non-significant changes). Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Changes in the global patterns of near-surface salinity contrast are transferred to the ocean interior via ventilation pathways (Figure 2.27b). Large scale similarities in subsurface salinity changes across observational estimates point to decreasing (increasing) salinity in regions where salinity is lower (higher) than the global average, with freshening in subpolar regions and salinification in the subtropical gyres ( [[#Durack--2010|Durack and Wijffels, 2010]] ; [[#Good--2013|Good et al., 2013]] ; [[#Skliris--2014|Skliris et al., 2014]] ; [[#Durack--2015|Durack, 2015]] ; [[#Aretxabaleta--2017|Aretxabaleta et al., 2017]] ; [[#Cheng--2020|Cheng et al., 2020]] ). Regional changes in salinity are assessed in Section 9.2.2.2.

In summary, it is ''virtually certain'' that since 1950 near-surface high salinity regions have become more saline, while low salinity regions have become fresher, and it is ''very likely'' that this extends to the ocean interior along ventilation pathways. Across basins, it is ''very likely'' that the Atlantic has become saltier and the Pacific and Southern oceans have freshened. The differences between high-salinity and low-salinity regions are linked to an intensification of the hydrological cycle ( ''medium confidence'' ).

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<span id="sea-level"></span>
==== 2.3.3.3 Sea Level ====

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The AR5 concluded based on proxy and instrumental data that the rate of global mean sea level (GMSL) rise since the mid-19th century was larger than the mean rate during the previous two millennia ( ''high confidence'' ). The SROCC reported with ''high confidence'' that GMSL increases were 1.5 [1.1 to 1.9] mm yr <sup>–1</sup> for 1902–2010 (with an acceleration rate between –0.002 and +0.019 mm yr <sup>–2</sup> ), 2.1 [1.8 to 2.3] mm yr <sup>–1</sup> for 1970–2015, 3.2 [2.8 to 3.5] mm yr <sup>–1</sup> for 1993–2015 and 3.6 [3.1 to 4.1] mm yr <sup>–1</sup> for 2006–2015. AR5 reported that GMSL during the LIG was, over several thousand years, between 5 and 10 m higher than 1985–2004 ( ''medium confidence'' ) whereas SROCC concluded it was ''virtually certain'' that GMSL exceeded current levels ( ''high confidence'' ), and reached a peak that was ''likely'' 6–9 m higher than today, but did not exceed 10 m ( ''medium confidence'' ). The AR5 concluded with ''high confidence'' that there were two intra-LIG GMSL peaks and that the millennial-scale rate during these periods exceeded 2 mm yr <sup>–1</sup> . The AR5 had ''high confidence'' that GMSL during the MPWP did not exceed 20 m above present. Based on new understanding, SROCC placed the upper bound at 25 m but with ''low confidence'' .

The Earth was largely ice free during the EECO ( [[#Cramer--2011|Cramer et al., 2011]] ; [[#Miller--2020|Miller et al., 2020]] , Section 9.6.2), and complete loss of current land ice reservoirs would raise GMSL by 65.6 ± 1.8 m ( [[#Morlighem--2017|Morlighem et al., 2017]] , 2020; [[#Farinotti--2019|Farinotti et al., 2019]] ). Given that GMSL change must be due to some combination of transient land ice growth and changes in terrestrial water storage, additional global mean thermosteric sea-level increase of 7 ± 2 m ( [[#Fischer--2018|Fischer et al., 2018]] ) implies a peak EECO GMSL of 70–76 m ( ''low confidence'' ). Changes in ocean basin size driven by plate tectonics contributed a comparable amount to global mean geocentric sea level in the Eocene, but are definitionally excluded from GMSL assessment ( [[#Wright--2020|Wright et al., 2020]] ).

For the MPWP, several studies of coastal features have provided additional quantitative sea-level estimates of: 5.6–19.2 m from Spain ( [[#Dumitru--2019|Dumitru et al., 2019]] ), approximately 14 m from South Africa ( [[#Hearty--2020|Hearty et al., 2020]] ), 15 m from the United States ( [[#Moucha--2017|Moucha and Ruetenik, 2017]] ), and 25 m from New Zealand ( [[#Grant--2019|Grant et al., 2019]] ). Thus, consistent with SROCC, GMSL during the MPWP was higher than present by 5–25 m ( ''medium confidence'' ).

Reconstructions of GMSL from marine oxygen isotopes in foraminifera shells show variations of more than 100 m over intervals of 10–100 kyr during glacial-interglacial cycles of the Quaternary ( [[#Shackleton--1987|Shackleton, 1987]] ; [[#McManus--1999|McManus et al., 1999]] ; [[#Waelbroeck--2002|Waelbroeck et al., 2002]] ; [[#Miller--2020|Miller et al., 2020]] ). Correction for past temperatures and a calibration for ice-volume changes implies uncertainty estimates of ± 10–13 m (1 SD) ( [[#Grant--2014|Grant et al., 2014]] ; [[#Shakun--2015|Shakun et al., 2015]] ; [[#Spratt--2016|Spratt and Lisiecki, 2016]] ). A recent marine oxygen-isotope-based GMSL reconstruction ( [[#Spratt--2016|Spratt and Lisiecki, 2016]] ) agrees with previous reconstructions, while focusing on the past 800 kyr (Figure 2.28). It shows that GMSL during the Holocene was among the highest over this entire interval, and was surpassed only during the LIG (Marine Isotope Stage (MIS 5e)) and MIS 11 ( ''medium confidence'' ); however, relatively brief (about 2 kyr) highstands during other interglacial periods might be obscured by dating limitations.

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[[File:afaeaca739981232e0771a71f50accc9 IPCC_AR6_WGI_Figure_2_28.png]]

'''Figure 2.2''' '''8 |''' '''Changes in global mean sea level. (a)''' Reconstruction of sea-level from ice core oxygen isotope analysis for the last 800 kyr. For target paleo periods (CCB2.1) and MIS11 the estimates based upon a broader range of sources are given as box whiskers. Note the much broader axis range (200 m) than for later panels (tenths of metres). '''(b)''' Reconstructions for the last 2500 years based upon a range of proxy sources with direct instrumental records superposed since the late 19th century. '''(c)''' Tide-gauge and, more latterly, altimeter-based estimates since 1850. The consensus estimate used in various calculations in Chapters 7 and 9 is shown in black. '''(d)''' The most recent period of record from tide-gauge and altimeter-based records. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Few sites globally have well-preserved MIS 11 sea-level indicators ( [[#Dutton--2015|Dutton et al., 2015]] ). As reported in AR5, [[#Raymo--2012|Raymo and Mitrovica (2012)]] used glacial isostatic adjustment models to correct the elevation of MIS 11 sea-level proxies from Bermuda and Bahamas to estimate a peak MIS 11 GMSL between 6 and 13 m above present-day. This agrees with the elevation of 13 ± 2 m for the MIS 11 subtidal–intertidal transition in South Africa ( [[#Roberts--2012|Roberts et al., 2012]] ). A revised glacial isostatic adjustment at this location resulted in a peak GMSL estimate of 8–11.5 m ( [[#Chen--2014|Chen et al., 2014]] ). In light of these data, and the review by [[#Dutton--2015|Dutton et al. (2015)]] , the AR5 estimate of 6–13 m for MIS 11 remains the best available ( ''medium confidence'' ).

Recent studies have highlighted uncertainties in estimates of GMSL during the LIG, including the extent of GMSL variability ( [[#Capron--2019|Capron et al., 2019]] ). Vertical land motions ( [[#Austermann--2017|Austermann et al., 2017]] ) are starting to be considered quantitatively (e.g., [[#Stephenson--2019|Stephenson et al., 2019]] ), but are still bounded by large uncertainties. The distribution and thickness of pre-LIG ice sheets ( [[#Dendy--2017|Dendy et al., 2017]] ; [[#Rohling--2017|Rohling et al., 2017]] ) and isostasy driven by sediment loading since the LIG ( [[#Pico--2020|Pico, 2020]] ) add further uncertainty. In light of these recent studies and previous assessments, there is ''medium confidence'' that peak GMSL during the LIG was ''likely'' between 5 and 10 m higher than modern. Relative sea-level estimates from some sites (e.g., Bahamas and Seychelles) report ephemeral, metre-scale fluctuations ( [[#Vyverberg--2018|Vyverberg et al., 2018]] ). Different generations of LIG reef growth at other sites (e.g., Yucatan Peninsula, Western Australia) suggest the occurrence of sudden accelerations in GMSL change ( [[#Blanchon--2009|Blanchon et al., 2009]] ; [[#O’Leary--2013|O’Leary et al., 2013]] ). However, other sites (e.g., South Australia, Mediterranean), indicate that LIG sea level was substantially stable (T.-Y. [[#Pan--2018|]] [[#Pan--2018|Pan et al., 2018]] ; [[#Polyak--2018|Polyak et al., 2018]] ). In addition, there are uncertainties in the interpretation of local relative sea level from some GMSL reconstructions ( [[#Barlow--2018|Barlow et al., 2018]] ). Therefore, ''low confidence'' is assigned to any GMSL rate of change estimated within the LIG.

New geological proxies and glacial isostatic adjustment (GIA) modelling studies confirm that, at the LGM, GMSL was 125–134 m below present ( [[#Lambeck--2014|Lambeck et al., 2014]] ; [[#Yokoyama--2018|Yokoyama et al., 2018]] ). During the LDT, GMSL rose from approximately –120 m to –50 m, implying an average rate of about 10 mm yr <sup>–1</sup> ( [[#Lambeck--2014|Lambeck et al., 2014]] ). The fastest rise occurred during Meltwater Pulse 1A, at about 14.6–14.3 ka ( [[#Deschamps--2012|Deschamps et al., 2012]] ; [[#Sanborn--2017|Sanborn et al., 2017]] ), when GMSL rose by between 8 m and 15 m ( ''medium confidence'' ) (J. [[#Liu--2016|]] [[#Liu--2016|Liu et al., 2016]] ) at an average rate of 24–44 mm yr <sup>–1</sup> .

Recent GIA modelling studies tuned to both near- and far-field relative sea level (RSL) data yield MH GMSL estimates of –3.8 to –1.0 m ( [[#Lambeck--2014|Lambeck et al., 2014]] ; [[#Peltier--2015|Peltier et al., 2015]] ; [[#Bradley--2016|Bradley et al., 2016]] ; [[#Roy--2017|Roy and Peltier, 2017]] ). Estimates from relatively stable locations where the effects of GIA are small and relatively insensitive to parameters defining Earth rheology, and where RSL is expected to approximate GMSL to within about 1 m (e.g., [[#Milne--2008|Milne and Mitrovica, 2008]] ), suggest that RSL was between about –6 to +1.5 m at around 6 ka at multiple locations ( [[#Camoin--1997|Camoin et al., 1997]] ; [[#Braithwaite--2000|Braithwaite et al., 2000]] ; [[#Frank--2006|Frank et al., 2006]] ; [[#Montaggioni--2008|Montaggioni and Faure, 2008]] ; [[#Vacchi--2016|Vacchi et al., 2016]] ; [[#Khan--2017|Khan et al., 2017]] ; [[#Hibbert--2018|Hibbert et al., 2018]] ). The assessment of GMSL change at 6 kyr is challenging considering the proportionately large GIA effect ( [[#Kopp--2016|Kopp et al., 2016]] ), insufficient resolution of marine geochemical proxies ( δ <sup>18</sup> O, Mg/Ca) and uncertainties in the contribution of the Antarctic Ice Sheet during the MH ( [[#2.3.2.4|Section 2.3.2.4]] ). The possibility that GMSL was at least somewhat higher than present cannot be excluded.

For the last 3 kyr, GMSL has been estimated from global databases of sea-level proxies, including numerous densely-sampled high-resolution salt-marsh records with decimetre scale vertical resolution and sub-centennial temporal resolution ( [[#Kopp--2016|Kopp et al., 2016]] ; [[#Kemp--2018|Kemp et al., 2018]] ). Over the last about 1.5 kyr, the most prominent century-scale GMSL trends include average maximum rates of lowering and rising of –0.7 ± 0.5 mm yr <sup>–1</sup> (2 SD) over 1020–1120 CE, and 0.3 ± 0.5 (2 SD) over 1460–1560, respectively. Between 1000 and 1750 CE, GMSL is estimated to have been within the range of about –0.11 to +0.09 m relative to 1900 ( [[#Kemp--2018|Kemp et al., 2018]] ). This was followed by a sustained increase of GMSL that began between 1820 and 1860 and has continued to the present day. New analyses demonstrate that it is ''very likely'' that GMSL rise over the 20th century was faster than over any preceding century in at least 3 kyr ( [[#Kopp--2016|Kopp et al., 2016]] ; [[#Kemp--2018|Kemp et al., 2018]] ) (Figure 2.28).

Since SROCC, two new tide gauge reconstructions of 20th century GMSL change have been published, although both rely upon CMIP models to varying degrees (Figure 2.28). [[#Frederikse--2020|Frederikse et al. (2020)]] used a ‘virtual station’ method and a probabilistic framework to estimate GMSL change and its uncertainties since 1900. [[#Dangendorf--2019|Dangendorf et al. (2019)]] combined a Kalman Smoother ( [[#Hay--2015|Hay et al., 2015]] ) with Reduced Space Optimal Interpolation ( [[#Church--2011|Church and White, 2011]] ; [[#Ray--2011|Ray and Douglas, 2011]] ) in an effort to better represent both the long-term GMSL change while preserving information on sea-level variability. In addition, new ensemble-based methods for quantifying GMSL change have been presented that account for both structural and parametric uncertainty ( [[#Palmer--2021|Palmer et al., 2021]] ). Altimeter time series of GMSL change (Figure 2.28) have been extended to 2019/2020 but bias adjustments ( [[#Watson--2015|Watson et al., 2015]] ; [[#Beckley--2017|Beckley et al., 2017]] ; [[#Dieng--2017|Dieng et al., 2017]] ; [[#Ablain--2019|Ablain et al., 2019]] ; [[#Legeais--2020|Legeais et al., 2020]] ) did not change since SROCC.

Based on the ensemble approach of [[#Palmer--2021|Palmer et al. (2021)]] and an updated [[#WCRP%20Global%20Sea%20Level%20Budget%20Group--2018|WCRP Global Sea Level Budget Group (2018)]] assessment (Figure 2.28) GMSL rose at a rate of 1.32 [0.58 to 2.06] mm yr <sup>–1</sup> for the period 1901–1971, increasing to 1.87 [0.82 to 2.92] mm yr <sup>–1</sup> between 1971 and 2006, and further increasing to 3.69 [3.21 to 4.17] mm yr <sup>–1</sup> for 2006–2018 ( ''high confidence'' ). The average rate for 1901–2018 was 1.73 [1.28 to 2.17] mm yr <sup>–1</sup> with a total rise of 0.20 [0.15 to 0.25] m (Table 9.5). The acceleration rate ''very likely'' is 0.094 [0.082 to 0.115] mm yr <sup>–2</sup> for 1993–2018 ( [[#WCRP%20Global%20Sea%20Level%20Budget%20Group--2018|WCRP Global Sea Level Budget Group, 2018]] , updated), consistent with other estimates ( [[#Watson--2015|Watson et al., 2015]] ; X. [[#Chen--2017|]] [[#Chen--2017|Chen et al., 2017]] ; [[#Nerem--2018|Nerem et al., 2018]] ; [[#WCRP%20Global%20Sea%20Level%20Budget%20Group--2018|WCRP Global Sea Level Budget Group, 2018]] ; [[#Ablain--2019|Ablain et al., 2019]] ; [[#Legeais--2020|Legeais et al., 2020]] ). For the period 1902–2010 the updated tide gauge reconstructions published since SROCC also show a robust acceleration over the 20th century and the ensemble estimate of [[#Palmer--2021|Palmer et al. (2021)]] gives a value of 0.0053 [0.0042 to 0.0073] mm yr <sup>–2</sup> , based on an unweighted quadratic fit.

In summary, GMSL is rising, and the rate of GMSL rise since the 20th century is faster than over any preceding century in at least the last three millennia ( ''high confidence'' ). Since 1901, GMSL has risen by 0.20 [0.15 to 0.25] m at an accelerating rate. Further back in time, there is ''medium confidence'' that GMSL was within –3.5 to +0.5 m ( ''very likely'' range) of present during the MH, 5–10 m higher ( ''likely range'' ) during the LIG, and 5–25 m higher ( ''very likely'' range) during the MPWP.

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<span id="ocean-circulation"></span>
==== 2.3.3.4 Ocean Circulation ====

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<div id="2.3.3.4.1" class="h4-container"></div>

<span id="atlantic-meridional-overturning-circulation-amoc"></span>
===== 2.3.3.4.1 Atlantic Meridional Overturning Circulation (AMOC) =====

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The AR5 concluded that there was no evidence of a trend in the AMOC during the period of instrumental observations. However, AR5 also stressed insufficient evidence to support a finding of change in the heat transport of the AMOC. SROCC assessed that there was emerging evidence in sustained observations, both in situ (2004–2017) and revealed from SST-based reconstructions, that the AMOC had weakened during the instrumental era relative to 1850–1900 ( ''medium confidence'' ), although there were insufficient data to quantify the magnitude of the weakening. SROCC also concluded with ''low confidence'' an increase of the Southern Ocean upper cell overturning circulation. SROCC also reported with ''medium confidence'' that the production of Antarctic Bottom Water had decreased since the 1950s consistent with a decreased lower cell overturning circulation, and potentially modulating the strength of the AMOC.

On multi-millennial timescales, proxy evidence indicates that the AMOC varied repeatedly in strength and vertical structure. During the last glacial period, particularly around the LGM, AMOC was estimated to be shallower than present, although there is continued debate about the magnitude of the shoaling ( [[#Lynch-Stieglitz--2007|Lynch-Stieglitz et al., 2007]] ; [[#Gebbie--2014|Gebbie, 2014]] ), and whether this change was associated with a weaker overturning ( [[#Ritz--2013|Ritz et al., 2013]] ; [[#Menviel--2017|Menviel et al., 2017]] ; [[#Muglia--2018|Muglia et al., 2018]] ). There are indications that substantial variations in AMOC were associated with abrupt climate changes during the glacial intervals, including Dansgaard-Oeschger and Heinrich events (14–70 ka) ( [[#McManus--2004|McManus et al., 2004]] ; [[#Böhm--2015|Böhm et al., 2015]] ; [[#Henry--2016|Henry et al., 2016]] ; [[#Lynch-Stieglitz--2017|Lynch-Stieglitz, 2017]] ). During these millennial-scale oscillations, weakened AMOC was associated with dramatic cooling in the NH and warming in the SH ( [[#Buizert--2015|Buizert et al., 2015]] ; [[#Henry--2016|Henry et al., 2016]] ), while hemispheric changes of opposite sign accompanied strengthened AMOC. After the final demise of the Laurentide ice sheet about 8 ka, the mean overall strength of AMOC has been relatively stable throughout the rest of the Holocene compared to the preceding 100 kyr ( [[#Hoffmann--2018|Hoffmann et al., 2018]] ; [[#Lippold--2019|Lippold et al., 2019]] ). There are however indications of episodic variations in AMOC during the Holocene ( [[#Bianchi--1999|Bianchi and McCave, 1999]] ; [[#Oppo--2003|Oppo et al., 2003]] ; [[#Thornalley--2013|Thornalley et al., 2013]] ; [[#Ayache--2018|Ayache et al., 2018]] ), and past interglacial intervals ( [[#Galaasen--2014|Galaasen et al., 2014]] , 2020; [[#Hayes--2014|Hayes et al., 2014]] ; [[#Mokeddem--2014|Mokeddem et al., 2014]] ; H. [[#Huang--2020|]] [[#Huang--2020|Huang et al., 2020]] ). Over the last 3 kyr, there are indications that AMOC variability was potentially linked to decreasing production of Labrador Sea Water (LSW), one of the water masses contributing to AMOC ( [[#Alonso-Garcia--2017|Alonso-Garcia et al., 2017]] ; [[#Moffa-Sánchez--2017|Moffa-Sánchez and Hall, 2017]] ; [[#Moffa-Sánchez--2019|Moffa-Sánchez et al., 2019]] ).

Numerous proxy records collectively imply that AMOC is currently at its weakest point in the past 1.6 ka ( [[#Rahmstorf--2015|Rahmstorf et al., 2015]] ; [[#Caesar--2018|Caesar et al., 2018]] , 2021; [[#Thibodeau--2018|Thibodeau et al., 2018]] ; [[#Thornalley--2018|Thornalley et al., 2018]] ). [[#Caesar--2021|Caesar et al. (2021)]] analyse a compilation of various available indirect AMOC proxies from marine sediments, in situ-based reconstructions and terrestrial proxies, which show a decline beginning in the late 19th century and over the 20th century superimposed by large decadal variability in the second half of the 20th century. Indirect reconstructions of AMOC components based on coastal sea level records in the western North Atlantic ( [[#Ezer--2013|Ezer, 2013]] ; [[#McCarthy--2015|McCarthy et al., 2015]] ; [[#Piecuch--2020|Piecuch, 2020]] ) show an AMOC decline since the late 1950s, with only a short period of recovery during the 1990s.

However, other studies highlight that proxy records do not show such clear signals ( [[#Moffa-Sánchez--2019|Moffa-Sánchez et al., 2019]] ), and the use of SST- and coastal sea level-based proxies of AMOC places uncertainties on these results ( [[#Little--2019|Little et al., 2019]] ; [[#Jackson--2020|Jackson and Wood, 2020]] ; [[#Menary--2020|Menary et al., 2020]] ). For instance, SSTs are additionally influenced by atmospheric and non-AMOC related ocean variability ( [[#Josey--2018|Josey et al., 2018]] ; [[#Keil--2020|Keil et al., 2020]] ; [[#Menary--2020|Menary et al., 2020]] ), while sea level responds to a variety of factors (e.g., atmospheric pressure and local winds) independent of the AMOC ( [[#Woodworth--2014|Woodworth et al., 2014]] ; [[#Piecuch--2015|Piecuch and Ponte, 2015]] ; [[#Piecuch--2016|Piecuch et al., 2016]] ). Finally, large decadal variability is present in many reconstructions and obscures estimation of the long-term trend over the 20th century ( [[#Ezer--2013|Ezer, 2013]] ; [[#McCarthy--2015|McCarthy et al., 2015]] ; [[#Yashayaev--2016|Yashayaev and Loder, 2016]] ; [[#Thornalley--2018|Thornalley et al., 2018]] ; [[#Caesar--2021|Caesar et al., 2021]] ). It is also noted that the proxy reported AMOC decline, beginning in the late 19th century, is not supported by model-based evidence (Sections 3.5.4.1 and 9.2.3.1).

Since the 1980s, multiple lines of observational evidence for AMOC change exist. Ship-based hydrographic estimates of AMOC as far back as the 1980s show no overall decline in AMOC strength ( [[#Fu--2020|Fu et al., 2020]] ; [[#Worthington--2021|Worthington et al., 2021]] ). Direct indications from in-situ observations report a –2.5 ± 1.4 Sv change between 1993 and 2010 across the OVIDE section, superimposed on large interannual to decadal variability ( [[#Mercier--2015|Mercier et al., 2015]] ). At 41°N and 26°N, a decline of –3.1 ± 3.2 Sv per decade and –2.5 ± 2.1 Sv per decade respectively has been reported over 2004–2016 ( [[#Baringer--2018|Baringer et al., 2018]] ; [[#Smeed--2018|Smeed et al., 2018]] ). However, [[#Moat--2020|Moat et al. (2020)]] report an increase in AMOC strength at 26°N over 2009–2018. Recent time series of moored observations at 11ºS ( [[#Hummels--2015|Hummels et al., 2015]] ), 34°S ( [[#Meinen--2018|Meinen et al., 2018]] ; [[#Kersalé--2020|Kersalé et al., 2020]] ), and between 57 and 60ºN ( [[#Lozier--2019|Lozier et al., 2019]] ) are currently too short to permit robust conclusions about changes. The directly observed AMOC weakening since 2004, while significant, is over too short a period to assess whether it is part of a longer term trend or dominated by decadal‐scale internal variability ( [[#Smeed--2014|Smeed et al., 2014]] ; [[#Collins--2019|Collins et al., 2019]] ; [[#Moat--2020|Moat et al., 2020]] ). Notably an increase and subsequent decline in the 1990s is present in estimates of AMOC and associated heat transport constructed from reanalyses or auxiliary data (Section 9.2.3.1; [[#Frajka-Williams--2015|Frajka-Williams, 2015]] ; [[#Jackson--2016|Jackson et al., 2016]] ; [[#Trenberth--2017|Trenberth and Fasullo, 2017]] ; [[#Jackson--2020|Jackson and Wood, 2020]] ).

Repeated full depth in situ measurements report that deep convection – a major driver for AMOC – has recently returned to the Labrador Sea, particularly in 2015 ( [[#Yashayaev--2016|Yashayaev and Loder, 2016]] ; [[#Rhein--2017|Rhein et al., 2017]] ), and to the Irminger Sea ( [[#de%20Jong--2016|de Jong and de Steur, 2016]] ; [[#Gladyshev--2016|Gladyshev et al., 2016]] ; [[#de%20Jong--2018|de Jong et al., 2018]] ) following an extended period with weak convection since 2000. An associated strengthening of the outflow from the Labrador Sea has not been observed ( [[#Zantopp--2017|Zantopp et al., 2017]] ; [[#Lozier--2019|Lozier et al., 2019]] ), while strengthening of the AMOC is tentative ( [[#Desbruyères--2019|Desbruyères et al., 2019]] ; [[#Moat--2020|Moat et al., 2020]] ). A long-term increase of the upper overturning cell in the Southern Ocean since the 1990s can be assessed with ''low confidence'' , and there is ''medium confidence'' of a decrease in Antarctic bottom water (AABW) volume and circulation, which has potential implications for the strength of the AMOC (Section 9.2.3.2).

In summary, proxy-based reconstructions suggest that the AMOC was relatively stable during the past 8 kyr ( ''medium confidence'' ), with a weakening beginning since the late 19th century ( ''medium confidence'' ), but due to a lack of direct observations, ''confidence'' in an overall decline of AMOC during the 20th century ''is low'' . From the mid-2000s to mid-2010s, the directly observed weakening in AMOC ( ''high confidence'' ) cannot be distinguished between decadal-scale variability or a long-term trend ( ''high confidence'' ).

<div id="2.3.3.4.2" class="h4-container"></div>

<span id="western-boundary-currents-and-inter-basin-exchanges"></span>
===== 2.3.3.4.2 Western boundary currents and inter-basin exchanges =====

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Both AR5 and SROCC reported that western boundary currents (WBCs) have undergone an intensification, warming and poleward expansion, except for the Gulf Stream and the Kuroshio, but did not provide confidence statements. The AR5 reported with ''medium'' to ''high confidence'' intensification of the North Pacific subpolar gyre, the South Pacific subtropical gyre, and the subtropical cells, along with an expansion of the North Pacific subtropical gyre since the 1990s. It was pointed out that these changes are ''likely'' predominantly due to interannual-to-decadal variability, and in the case of the subtropical cells represent a reversal of earlier multi-decadal changes. SROCC concluded that it was ''unlikely'' that there has been a statistically significant net southward movement of the mean Antarctic Circumpolar Current (ACC) position over the past 20 years, in contrast to AR5, where this change had been assessed with ''medium confidenc'' e.

The intensity of the Kuroshio current system in the north-west Pacific varied in conjunction with the glaciation cycles over the last 1 Myr, with some limited glacial-interglacial variability in position ( [[#Jian--2000|Jian et al., 2000]] ; [[#Gallagher--2015|Gallagher et al., 2015]] ). The Agulhas current has strengthened substantially during the warming associated with deglaciations of the past 1 Myr ( [[#Peeters--2004|Peeters et al., 2004]] ; [[#Bard--2009|Bard and Rickaby, 2009]] ; [[#Martínez-Méndez--2010|Martínez-Méndez et al., 2010]] ; [[#Marino--2013|Marino et al., 2013]] ; [[#Ballalai--2019|Ballalai et al., 2019]] ). According to sediment core analyses, the Agulhas leakage varied by about 10 Sv during major climatic transitions over the past 640 kyr ( [[#Caley--2014|Caley et al., 2014]] ). Available data suggests that there was relatively little change in the net flow of the ACC during the LGM, with no consensus on the sign of changes ( [[#McCave--2013|McCave et al., 2013]] ; [[#Lamy--2015|Lamy et al., 2015]] ; [[#Lynch-Stieglitz--2016|Lynch-Stieglitz et al., 2016]] ), except at one location at the northern edge of the Drake Passage where a 40% decrease of transport had been reported ( [[#Lamy--2015|Lamy et al., 2015]] ). Longer time series from the northern entrance to Drake Passage suggest a consistent transport variability of 6–16% through glacial climate cycles, with higher current speeds during interglacial times and reduced current speeds during glacial intervals ( [[#Toyos--2020|Toyos et al., 2020]] ). Inferred variability in the size and strength of the North Atlantic subpolar gyre was substantial, and included rapid changes on millennial time scales during both interglacial and glacial intervals over the last 150 kyr ( [[#Born--2010|Born and Levermann, 2010]] ; [[#Mokeddem--2014|Mokeddem et al., 2014]] ; [[#Irvalı--2016|Irvalı et al., 2016]] ; [[#Mokeddem--2016|Mokeddem and McManus, 2016]] ). North Atlantic – Arctic exchange has also varied in the past, with indications of an increasing inflow of Atlantic waters into the Arctic during the late Holocene ( [[#Ślubowska--2005|Ślubowska et al., 2005]] ) with an acceleration to the recent inflow that is now the largest of the past 2 kyr ( [[#Spielhagen--2011|Spielhagen et al., 2011]] ).

A latitudinal shift of subtropical/subpolar gyres on the order of 0.1 ± 0.04° per decade is derived by an indirect method using remote sensing data during 1993–2018 ( [[#Yang--2020|Yang et al., 2020]] ). Direct observations show a systematic poleward migration of WBCs ( [[#Wu--2012|Wu et al., 2012]] ; [[#Yang--2016|Yang et al., 2016]] , 2020; [[#Bisagni--2017|Bisagni et al., 2017]] ). However, they do not support an intensification of WBCs, with a weakening, broadening, or little change reported for the Kuroshio (Y.-L. [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|Wang et al., 2016]] ; [[#Wang--2018|Wang and Wu, 2018]] ; [[#Collins--2019|Collins et al., 2019]] ), Gulf Stream ( [[#McCarthy--2018|McCarthy et al., 2018]] ; [[#Collins--2019|Collins et al., 2019]] ; [[#Dong--2019|Dong et al., 2019]] ; [[#Andres--2020|Andres et al., 2020]] ), Agulhas ( [[#Beal--2016|Beal and Elipot, 2016]] ; [[#Elipot--2018|Elipot and Beal, 2018]] ) and East Australian ( [[#Sloyan--2015|Sloyan and O’Kane, 2015]] ) currents. The Gulf Stream has recently reversed a long-term poleward migration ( [[#Bisagni--2017|Bisagni et al., 2017]] ). Multidecadal variability of the strength and position of WBCs ( [[#Hsin--2015|Hsin, 2015]] ; [[#Bisagni--2017|Bisagni et al., 2017]] ; [[#McCarthy--2018|McCarthy et al., 2018]] ) and short records from direct observations obscure the detection of any long-term trends ( [[#Yang--2020|Yang et al., 2020]] ).

The Pacific to Arctic exchange at the Bering Strait plays a minor role in the total Arctic exchange with the global ocean, which has increased from 0.8 Sv to 1.0 Sv over 1990–2015 ( [[#Woodgate--2018|Woodgate, 2018]] ). For Atlantic-Arctic exchange, major branches of Atlantic Water inflow from the North Atlantic into the Arctic across the Greenland–Scotland Ridge have remained stable since the mid-1990s ( [[#Berx--2013|Berx et al., 2013]] ; [[#Hansen--2015|Hansen et al., 2015]] ; [[#Jochumsen--2017|Jochumsen et al., 2017]] ; [[#Østerhus--2019|Østerhus et al., 2019]] ), with only the smaller pathway of Atlantic Water north of Iceland showing a strengthening trend during 1993–2018 ( [[#Casanova-Masjoan--2020|Casanova-Masjoan et al., 2020]] ), but with associated heat transport strengthening through the 1990s ( [[#Rossby--2020|Rossby et al., 2020]] ; [[#Tsubouchi--2021|Tsubouchi et al., 2021]] ). The Arctic outflow remained broadly stable from the mid-1990s to the mid 2010s ( [[#Østerhus--2019|Østerhus et al., 2019]] ).

The heat and mass transport of the Indonesian throughflow (ITF) shows substantial variability at seasonal to decadal time scales ( [[#Zhuang--2013|Zhuang et al., 2013]] ; Q.-Y. [[#Liu--2015|]] [[#Liu--2015|Liu et al., 2015]] ; [[#Susanto--2015|Susanto and Song, 2015]] ; [[#Feng--2017|Feng et al., 2017]] , 2018; M. [[#Li--2018|]] [[#Li--2018|Li et al., 2018]] ; [[#Sprintall--2019|Sprintall et al., 2019]] ; [[#Xie--2019|Xie et al., 2019]] ). Q.-Y. [[#Liu--2015|]] [[#Liu--2015|Liu et al. (2015)]] reported an increasing trend in the ITF geostrophic transport of 1 Sv per decade over 1984–2013, consistent with direct estimates ( [[#Sprintall--2014|Sprintall et al., 2014]] ), and results from reanalyses (M. [[#Li--2018|]] [[#Li--2018|Li et al., 2018]] ), and this appears to be linked to multi-decadal scale variability rather than a long-term trend ( [[#Kosaka--2013|Kosaka and Xie, 2013]] ; [[#England--2014|England et al., 2014]] ; [[#Lee--2015|Lee et al., 2015]] ).

Southern Ocean circulation changes are assessed in SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ), and are confirmed and synthesized in Section 9.2.3.2 which shows that there is no indication of ACC transport change, and that it is ''unlikely'' that the mean meridional position of the ACC has moved southward in recent decades.

In summary, over the past 3–4 decades, the WBC strength is highly variable ( ''high confidence'' ), and WBCs and subtropical gyres have shifted poleward since 1993 ( ''medium confidence'' ). Net Arctic Ocean volume exchanges with the other ocean basins remained stable over the mid-1990s to the mid-2010s ( ''high confidence'' ). There is ''high confidence'' that the ITF shows strong multi-decadal scale variability since the 1980s.

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<span id="ocean-ph"></span>
==== 2.3.3.5 Ocean pH ====

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The AR5 assessed with ''high confidence'' that the pH of the ocean surface had decreased since preindustrial times, primarily as a result of ocean uptake of CO <sub>2</sub> . SROCC concluded that the global ocean absorbed 20–30% of total CO <sub>2</sub> emissions since the 1980s, with ''virtually certain'' ocean surface pH decline. The SROCC assessed a rate of surface pH decline of 0.017–0.027 pH units per decade across a range of time series of pH observations longer than 15 years. The decline in surface open ocean pH was assessed by SROCC as having ''very likely'' already emerged from background natural variability for more than 95% of the global surface open ocean.

Understanding of changes in surface pH at paleo time-scales has increased since AR5 ( [[#Clarkson--2015|Clarkson et al., 2015]] ; [[#Foster--2016|Foster and Rae, 2016]] ; [[#Zeebe--2016|Zeebe et al., 2016]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Sosdian--2018|Sosdian et al., 2018]] ; [[#Henehan--2019|Henehan et al., 2019]] ; [[#Anagnostou--2020|Anagnostou et al., 2020]] ; [[#Harper--2020|Harper et al., 2020]] ; [[#Müller--2020|Müller et al., 2020]] ). Over the last 65 million years there have been several intervals when the pH of surface waters varied concurrently with climate change such as during the PETM, EECO, and MCO (Figure 2.29a and Section 5.3.1.1). However, only during the PETM is the change sufficiently well-constrained to allow for a direct comparison with recent and current trends ( [[#Kirtland%20Turner--2018|Kirtland]] [[#Turner--2018|Turner, 2018]] ). This event was associated with profound perturbations of the global carbon cycle, ocean warming, deoxygenation and a surface ocean pH decrease ''likely'' ranging from 0.15 to 0.30 units ( [[#Penman--2014|Penman et al., 2014]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Babila--2018|Babila et al., 2018]] ) – a rate that was ''likely'' at least an order of magnitude slower than today ( [[#Cui--2011|Cui et al., 2011]] ; [[#Bowen--2015|Bowen et al., 2015]] ; [[#Frieling--2016|Frieling et al., 2016]] ; [[#Zeebe--2016|Zeebe et al., 2016]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Kirtland%20Turner--2018|Kirtland]] [[#Turner--2018|Turner, 2018]] ; [[#Gingerich--2019|Gingerich, 2019]] ).

Paleo evidence suggests that surface ocean pH has gradually increased over the last 50 Myr ( [[#Anagnostou--2016|Anagnostou et al., 2016]] , 2020; [[#Sosdian--2018|Sosdian et al., 2018]] ) (Figure 2.29a). Global mean surface pH values as low as observed during recent decades are uncommon in the last 2 Myr (Figure 2.29b) ( [[#Martínez-Botí--2015a|Martínez-Botí et al., 2015a]] ; [[#Chalk--2017|Chalk et al., 2017]] ; [[#Dyez--2018|Dyez et al., 2018]] ; [[#Sosdian--2018|Sosdian et al., 2018]] ), and have not been experienced in at least the last 25 kyr (Figure 2.29c; [[#Palmer--2003|Palmer and Pearson, 2003]] ; [[#Foster--2008|Foster, 2008]] ; [[#Palmer--2010|Palmer et al., 2010]] ; [[#Henehan--2013|Henehan et al., 2013]] ; [[#Kirschke--2013|Kirschke et al., 2013]] ; [[#Martínez-Botí--2015a|Martínez-Botí et al., 2015a]] ; [[#Naik--2015|Naik et al., 2015]] ; [[#Ezat--2017|Ezat et al., 2017]] ; [[#Gray--2018|Gray et al., 2018]] ; [[#Shao--2019|Shao et al., 2019]] ). The magnitude of pH change during the Pleistocene glacial–interglacial cycles was 0.1–0.15 pH units – similar to recent changes in the modern era (Figure 2.29c and Section 5.3.1.2; [[#Hönisch--2009|Hönisch et al., 2009]] ; [[#Chalk--2017|Chalk et al., 2017]] ; [[#Shao--2019|Shao et al., 2019]] ). Maximum rates of pH change during the LDT, inferred from changes in atmospheric CO <sub>2</sub> recorded in ice cores ( [[#Marcott--2014|Marcott et al., 2014]] ) and the established relationships between pH and CO <sub>2</sub> changes and the boron isotope proxy ( [[#Hain--2018|Hain et al., 2018]] ), reached –0.02 pH units per century at about 11.7 ka, about 14.8 ka and about 16.3 ka, as previously sequestered CO <sub>2</sub> was transferred from the ocean interior to the subsurface ocean ( [[#Martínez-Botí--2015a|Martínez-Botí et al., 2015a]] ; [[#Jaccard--2016|Jaccard et al., 2016]] ; [[#Rae--2018|Rae et al., 2018]] ).

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[[File:ff55a94c2abf2085ce29b2e2e0c13e49 IPCC_AR6_WGI_Figure_2_29.png]]

'''Figure 2.29 | Low-latitude surface ocean pH over the last 65 million years (65 Myr). (a)''' Low-latitude (30°N–30°S) surface ocean pH over the last 65 Myr, reconstructed using boron isotopes in foraminifera. '''(b)''' as (a) but for the last 3.5 Myr. Double headed arrow shows the approximate magnitude of glacial-interglacial pH changes. '''(c)''' Multisite composite of surface pH. In (a, b, c) uncertainty is shown at 95% confidence as a shaded band. Relevant paleoclimate reference periods (CCB2.1) have been labelled. Period windows for succeeding panels are shown as horizontal black lines in (a) and (b). '''(d)''' Estimated low-latitude surface pH from direct observations (BATS, HOT) and global mean pH (65°S–65°N) from two indirect estimates (CMEMS, OCEAN-SODA). Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Since the 1980s, the global ocean has experienced a decline in surface pH of 0.016 ± 0.006 pH units per decade based on indirect pH products (Figure 2.29d; [[#Lauvset--2015|Lauvset et al., 2015]] ; [[#Hurd--2018|Hurd et al., 2018]] ; [[#IPCC--2019|IPCC, 2019]] ; [[#Gehlen--2020|Gehlen et al., 2020]] ) that agrees with the decline of 0.017–0.025 pH units per decade assessed in SROCC from direct time-series measurements of pH. Section 5.3.2.2 assesses a decline that ranges from 0.01 to 0.026 pH units per decade for the tropical and subtropical open ocean areas, and 0.003–0.026 pH units per decade for the polar and subpolar open ocean regions by using time series and ship-based datasets from the surface ocean CO <sub>2</sub> measurement network ( [[#Bakker--2016|Bakker et al., 2016]] ; [[#Gehlen--2020|Gehlen et al., 2020]] ; [[#Gregor--2021|Gregor and Gruber, 2021]] ). There is general consensus that global surface ocean pH trends over the past two decades have exceeded the natural background variability ( [[#Lauvset--2015|Lauvset et al., 2015]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ; [[#Gehlen--2020|Gehlen et al., 2020]] ). However, for some areas sparse data coverage, and large year-to-year variations hinders the detection of long-term surface ocean pH trends; for example in the Southern Ocean ( [[#Lauvset--2015|Lauvset et al., 2015]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ) and in the Arctic Ocean ( [[#Lauvset--2015|Lauvset et al., 2015]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ; [[#Meredith--2019|Meredith et al., 2019]] ).

For subsurface pH changes, estimates arise from direct ship measurements from repeated hydrography programs ( [[#Carter--2019|Carter et al., 2019]] ), indirect estimates of pH through calcite and aragonite saturation horizons ( [[#Osborne--2020|Osborne et al., 2020]] ; [[#Ross--2020|Ross et al., 2020]] ), and the very recent biogeochemical Argo floats equipped with pH sensors ( [[#Claustre--2020|Claustre et al., 2020]] ). Global subsurface pH has decreased over the past 20 to 30 years, with signals observed to at least 1000 m depths ( [[#Lauvset--2020|Lauvset et al., 2020]] ). Global findings are supplemented by regional findings from the Pacific Ocean ( [[#Carter--2019|Carter et al., 2019]] ; [[#Ross--2020|Ross et al., 2020]] ); the South Atlantic ( [[#Salt--2015|Salt et al., 2015]] ) and Southern Ocean ( [[#Jones--2017|Jones et al., 2017]] ); the North Atlantic Ocean and along the AMOC ( [[#Woosley--2016|Woosley et al., 2016]] ; [[#Perez--2018|Perez et al., 2018]] ), the Arctic Ocean ( [[#Qi--2017|Qi et al., 2017]] ) and marginal seas ( [[#Chen--2017|]] [[#Chen--2017|C.-T.A. Chen et al., 2017]] ). Further details are given in Section 5.3.3.1.

To conclude, it is ''virtually certain'' that surface open ocean pH has declined globally over the last 40 years by 0.003–0.026 pH per decade, and a decline in the ocean interior has been observed in all ocean basins over the past 2–3 decades ( ''high confidence'' ). A long-term increase in surface open ocean pH occurred over the past 50 Myr ( ''high confidence'' ), and surface open ocean pH as low as recent times is uncommon in the last 2 Myr ( ''medium confidence'' ). There is ''very high confidence'' that open ocean surface pH is now the lowest it has been for at least 26 kyr and current rates of pH change are unprecedented since at least that time.

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<span id="ocean-deoxygenation"></span>
==== 2.3.3.6 Ocean Deoxygenation ====

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SROCC concluded that there has ''very likely'' been a net loss of oxygen over all ocean depths since the 1960s linked to global ocean deoxygenation at a range of 0.3–2.0%, and that the oxygen levels in the global upper 1000 m of the ocean had decreased by 0.5–3.3% during 1970–2010 ( ''medium confidence'' ), alongside an expansion of oxygen minimum zones (OMZ) by 3–8%. For the surface ocean (0–100 m) and the thermocline at 100–600 m, the ''very likely'' range of oxygen decline was assessed to be 0.2–2.1% and 0.7–3.5%, respectively. Multidecadal rates of deoxygenation showed variability throughout the water column and across ocean basins ( ''high confidence'' ).

Since AR5 evidence for changes in oxygen content based on new proxy reconstructions has increased ( [[#Hoogakker--2015|Hoogakker et al., 2015]] , 2018; [[#Gottschalk--2016|Gottschalk et al., 2016]] , 2020; [[#Anderson--2019|Anderson et al., 2019]] ). Paleo records point to past periods of reduced oceanic oxygen levels during the late Permian (about 250 Ma), the Jurassic and the Cretaceous (180–100 Ma), alongside associated global scale disturbances of the global carbon cycle. Emerging studies for the Cenozoic Era suggest more stable ocean oxygenation conditions throughout the interval on million-year time scales (X. [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|Wang et al., 2016]] ). Sedimentary proxy data indicate, however, that the seemingly stable Cenozoic was punctuated by transient, widespread deoxygenation during the PETM ( [[#Dickson--2012|Dickson et al., 2012]] ; [[#Winguth--2012|Winguth et al., 2012]] ; [[#Remmelzwaal--2019|Remmelzwaal et al., 2019]] ), with parts of the ocean reaching anoxic levels (Section 5.3.1.1; [[#Yao--2018|Yao et al., 2018]] ). Since the LGM, there was an overall emergence and expansion of low-oxygen waters into the ocean intermediate depths as a result of rapid warming and a reorganization of the global overturning circulation ( [[#Galbraith--2015|Galbraith and Jaccard, 2015]] ). The maximum expansion of oxygen-depleted waters during the LDT occurred coincidently with rapid warming in the NH at 14.7–12.9 ka ( [[#Jaccard--2012|Jaccard and Galbraith, 2012]] ; [[#Moffitt--2015|Moffitt et al., 2015]] ; [[#Hoogakker--2018|Hoogakker et al., 2018]] ). Deep (&gt;1500 m) ocean oxygen levels increased by 100–150 µmol kg <sup>–1</sup> since the LGM, reaching modern oxygen levels at about 10 ka (Section 5.3.1.2; [[#Hoogakker--2015|Hoogakker et al., 2015]] ; [[#Gottschalk--2016|Gottschalk et al., 2016]] ; [[#Anderson--2019|Anderson et al., 2019]] ).

New findings for ocean oxygen content since SROCC are limited to regional scale assessments. The magnitude of change is difficult to compare across regions to arrive at a global assessment due to differences in depth range, time period, baseline climatology, methodology, and particularly the use of different units. To facilitate comparisons, data are presented as change per decade, and conversions of the SROCC global mean percentage of oxygen decline estimates are provided as a loss of 3.2 μmol kg <sup>−1</sup> in the upper 1000 m of the global ocean (1.93%), 2.0 μmol kg <sup>−1</sup> (0.8 μmol kg <sup>−1</sup> per decade) in the upper 1000 m of the global ocean (1.93%), and 2.0 μmol kg <sup>−1</sup> (0.5 μmol kg <sup>−1</sup> per decade) in the entire water column (1.15%) between 1970 and 2010 ( [[#Bindoff--2019|Bindoff et al., 2019]] ). Oxygen change also shows decadal variability ( [[#Ito--2016|Ito et al., 2016]] ; [[#Stramma--2020|Stramma et al., 2020]] ) that can influence estimates of trends.

Subsurface (100–400 m) oxygen in the California Current system is estimated to have declined by 24 ± 2 μmol kg <sup>−1</sup> (1.0 μmol kg <sup>−1</sup> per decade) between 1993 and 2018, a rate similar to the global upper 1000 m average ( [[#Bindoff--2019|Bindoff et al., 2019]] ). In some locations, however, the magnitude of oxygen loss substantially exceeds global averages ( [[#Queste--2018|Queste et al., 2018]] ; [[#Bronselaer--2020|Bronselaer et al., 2020]] ; [[#Cummins--2020|Cummins and Ross, 2020]] ; [[#Stramma--2020|Stramma et al., 2020]] ). For example, a decline in oxygen content of 11.7 ± 3.5% in the upper 4000 m, including a decline of 20.4 ± 7.2% in the upper 1550 m, is reported in the North Pacific over the period 1958–2018 ( [[#Cummins--2020|Cummins and Ross, 2020]] ) (equivalent to 2.3 μmol kg <sup>−1</sup> per decade in the upper 1550 m and 2.0 μmol kg <sup>−1</sup> decade <sup>− 1</sup> throughout the 4000 m water column). In some regions of the Southern Ocean south of 65°S oxygen in the upper 2000 m has declined by 60 μmol kg <sup>−1</sup> (~52 μmol kg <sup>−1</sup> per decade) based on comparisons of 2014–2019 and 1985–2005 observations ( [[#Bronselaer--2020|Bronselaer et al., 2020]] ). Within some OMZs regions of the Indian ocean, oxygen has declined from 6–12 to &lt;2 μmol kg <sup>−1</sup> between the 1960s and 2015–2016 ( [[#Bristow--2017|Bristow et al., 2017]] ; [[#Al-Said--2018|Al-Said et al., 2018]] ; [[#Queste--2018|Queste et al., 2018]] ).

Findings since SROCC provide further support that the volume of severely oxygen-depleted water has expanded in some locations of the global ocean (Section 5.3.3.2). For example, vertical expansion of low oxygen zones is reported in the North Pacific at a rate of 3.1 ± 0.5 m yr <sup>–1</sup> ( [[#Ross--2020|Ross et al., 2020]] ), and suboxic waters have increased by 20% at a rate of about 19 m per decade from 1982–2010 in the Arabian Sea ( [[#Al-Said--2018|Al-Said et al., 2018]] ; [[#Lachkar--2019|Lachkar et al., 2019]] ), and expanded off the coast of Mexico ( [[#Sánchez-Velasco--2019|Sánchez-Velasco et al., 2019]] ).

In summary, episodes of widespread and long-lasting (100 ka scales) open-ocean deoxygenation were related to warm climate intervals of the Permian-Cretaceous, with conditions becoming generally better oxygenated as the climate cooled over the course of the Cenozoic ( ''high confidence'' ). The largest expansions of oxygen depleted waters over the past 25 ka were strongly linked to rapid warming rates ( ''medium confidence'' ). Open-ocean deoxygenation has occurred in most regions of the open ocean during the mid-20th to early 21st centuries ( ''high confidence'' ), and shows decadal variability ( ''medium confidence'' ). Evidence further confirms SROCC that OMZs are expanding at many locations ( ''high confidence'' ).

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