Editing IPCC:AR6/WGI/Chapter-7 (section)

==== 7.4.4.2 Tropical Pacific Sea Surface Temperature Gradients ====

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Research published since AR5 identifies changes in the tropical Pacific Ocean zonal SST gradient over time as a key factor affecting how radiative feedbacks may evolve in the future ( [[#7.4.4.3|Section 7.4.4.3]] ). There is now a much-improved understanding of the processes that govern the tropical Pacific SST gradient ( [[#7.4.4.2.1|Section 7.4.4.2.1]] ) and the paleoclimate record provides evidence for its equilibrium changes from time periods associated with changes in CO <sub>2</sub> [[#7.4.4.2.2|Section 7.4.4.2.2]] ).

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===== 7.4.4.2.1 Critical processes determining changes in tropical Pacific sea surface temperature gradients =====

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A weakening of the equatorial Pacific Ocean east–west SST gradient, with greater warming in the east than the west, is a common feature of the climate response to greenhouse gas forcing as projected by ESMs on centennial and longer time scales (e.g., Figure 7.14b; see ( [[IPCC:Wg1:Chapter:Chapter-4#4.5.1|Section 4.5.1]] ). There are thought to be several factors contributing to this pattern. In the absence of any changes in atmospheric or oceanic circulations, the east–west surface temperature difference is theorized to decrease owing to weaker evaporative damping, and thus greater warming in response to forcing, where climatological temperatures are lower in the eastern Pacific cold tongue ( [[#Xie--2010|Xie et al., 2010]] ; [[#Luo--2015|Luo et al., 2015]] ). Within atmospheric ESMs coupled to a mixed-layer ocean, this gradient in damping has been linked to the rate of change with warming of the saturation specific humidity, which is set by the Clausius–Clapeyron relation ( [[#Merlis--2011|Merlis and Schneider, 2011]] ). Gradients in low-cloud feedbacks may also favour eastern equatorial Pacific warming ( [[#DiNezio--2009|DiNezio et al., 2009]] ).

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[[File:60bc04f65baf68bf30b88df95e74c1aa IPCC_AR6_WGI_Figure_7_14.png]]

'''Figure''' '''7.14 |''' '''Illustration of tropospheric temperature and low-cloud response to observed and projected Pacific Ocean sea surface temperature trends. (a)''' Atmospheric response to linear sea surface temperature trend observed over 1870–2019 (HadISST1 dataset; [[#Rayner--2003|Rayner et al., 2003]] ). '''(b)''' Atmospheric response to linear sea-surface temperature trend over 150 years following ''abrupt 4xCO2'' forcing as projected by CMIP6 ESMs ( [[#Dong--2020|Dong et al., 2020]] ). Relatively large historical warming in the western tropical Pacific has been communicated aloft (a shift from grey to red atmospheric temperature profile), remotely warming the tropical free troposphere and increasing the strength of the inversion in regions of the tropics where warming has been slower, such as the eastern equatorial Pacific. In turn, an increased inversion strength has increased the low-cloud cover ( [[#Zhou--2016|Zhou et al., 2016]] ) causing an anomalously negative cloud and lapse-rate feedbacks over the historical record ( [[#Andrews--2018|Andrews et al., 2018]] ; [[#Marvel--2018|Marvel et al., 2018]] ). Relatively large projected warming in the eastern tropical Pacific is trapped near the surface (shift from grey to red atmospheric temperature profile), decreasing the strength of the inversion locally. In turn, a decreased inversion strength combined with surface warming is projected to decrease the low-cloud cover, causing the cloud and lapse-rate feedbacks to become less negative in the future. Figure adapted from [[#Mauritsen--2016|Mauritsen (2016)]] . Further details on data sources and processing are available in the chapter data table (Table 7.SM.14).

In the coupled climate system, changes in atmospheric and oceanic circulations will influence the east-west temperature gradient as well. It is expected that as global temperature increases and as the east–west temperature gradient weakens, east–west sea level pressure gradients and easterly trade winds (characterizing the Walker circulation) will weaken as well (Sections 4.5.3, 8.2.2.2 and 8.4.2.3, and Figure 7.14b; [[#Vecchi--2006|Vecchi et al., 2006]] , 2008). This would, in turn, weaken the east–west temperature gradient through a reduction of equatorial upwelling of cold water in the east Pacific and a reduction in the transport of warmer water to the western equatorial Pacific and Indian Ocean ( [[#England--2014|England et al., 2014]] ; [[#Dong--2017|Dong and McPhaden, 2017]] ; [[#Li--2017|Li et al., 2017]] ; [[#Maher--2018|Maher et al., 2018]] ).

Research published since AR5 ( [[#Burls--2014b|Burls and Fedorov, 2014b]] ; [[#Fedorov--2015|Fedorov et al., 2015]] ; [[#Erfani--2019|Erfani and Burls, 2019]] ) has built on an earlier theory ( [[#Liu--1997|Liu and Huang, 1997]] ; [[#Barreiro--2008|Barreiro and Philander, 2008]] ) linking the east–west temperature gradient to the north–south temperature gradient. In particular, model simulations suggest that a reduction in the equator-to-pole temperature gradient (polar amplification) increases the temperature of water subducted in the extra-tropics, which in turn is upwelled in the eastern Pacific. Thus, polar amplified warming, with greater warming in the mid-latitudes and subtropics than in the deep tropics, is expected to contribute to the weakening of the east–west equatorial Pacific SST gradient on decadal to centennial time scales.

The transient adjustment of the equatorial Pacific SST gradient is influenced by upwelling waters which delay surface warming in the east since they have not been at the surface for years-to-decades to experience the greenhouse gas forcing. This ‘thermostat mechanism’ ( [[#Clement--1996|Clement et al., 1996]] ; [[#Cane--1997|Cane et al., 1997]] ) is not thought to persist to equilibrium since it does not account for the eventual increase in temperatures of upwelled waters ( [[#Liu--2005|Liu et al., 2005]] ; [[#Xie--2010|Xie et al., 2010]] ; Y. [[#Luo--2017|]] [[#Luo--2017|Luo et al., 2017]] ) which will occur as the subducting waters in mid-latitudes warm by more than the tropics on average as polar amplification emerges. An individual CMIP5 ESM (GFDL’s ESM2M) has been found to exhibit a La Niña-like pattern of Pacific temperature change through the 21st century, similar to the SST trends seen over the historical record ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1|Section 9.2.1]] and Figure 7.14a), owing to a weakening asymmetry between El Niño and La Niña events ( [[#Kohyama--2017|Kohyama et al., 2017]] ), but this pattern of warming may not persist to equilibrium ( [[#Paynter--2018|Paynter et al., 2018]] ).

Since 1870, observed SSTs in the tropical western Pacific Ocean have increased while those in the tropical eastern Pacific Ocean have changed less (Figure 7.14a and ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1|Section 9.2.1]] ). Much of the resultant strengthening of the equatorial Pacific temperature gradient has occurred since about 1980 due to strong warming in the west and cooling in the east (Figure 2.11b) concurrent with an intensification of the surface equatorial easterly trade winds and Walker circulation (Sections 3.3.3.1, 3.7.6, 8.3.2.3 and 9.2, and Figures 3.16f and 3.39f; [[#England--2014|England et al., 2014]] ). This temperature pattern is also reflected in regional ocean heat content trends and sea level changes observed from satellite altimetry since 1993 ( [[#Bilbao--2015|Bilbao et al., 2015]] ; [[#Richter--2020|Richter et al., 2020]] ). The observed changes may have been influenced by one or a combination of temporary factors including sulphate aerosol forcing ( [[#Smith--2016|Smith et al., 2016]] ; [[#Takahashi--2016|Takahashi and Watanabe, 2016]] ; [[#Hua--2018|Hua et al., 2018]] ), internal variability within the Indo-Pacific Ocean ( [[#Luo--2012|Luo et al., 2012]] ; [[#Chung--2019|Chung et al., 2019]] ), teleconnections from multi-decadal tropical Atlantic SST trends ( [[#Kucharski--2011|Kucharski et al., 2011]] , 2014, 2015; [[#McGregor--2014|McGregor et al., 2014]] ; [[#Chafik--2016|Chafik et al., 2016]] ; X. [[#Li--2016|]] [[#Li--2016|Li et al., 2016]] ; [[#Kajtar--2017|Kajtar et al., 2017]] ; [[#Sun--2017|Sun et al., 2017]] ), teleconnections from multi-decadal Southern Ocean SST trends ( [[#Hwang--2017|Hwang et al., 2017]] ), and coupled ocean–atmosphere dynamics which slow warming in the equatorial eastern Pacific ( [[#Clement--1996|Clement et al., 1996]] ; [[#Cane--1997|Cane et al., 1997]] ; [[#Seager--2019|Seager et al., 2019]] ). CMIP3 and CMIP5 ESMs have difficulties replicating the observed trends in the Walker circulation and Pacific Ocean SSTs over the historical record ( [[#Sohn--2013|Sohn et al., 2013]] ; [[#Zhou--2016|Zhou et al., 2016]] ; [[#Coats--2017|Coats and Karnauskas, 2017]] ), possibly due to model deficiencies including insufficient multi-decadal Pacific Ocean SST variability ( [[#Laepple--2014|Laepple and Huybers, 2014]] ; [[#Bilbao--2015|Bilbao et al., 2015]] ; [[#Chung--2019|Chung et al., 2019]] ), mean state biases affecting the forced response or the connection between Atlantic and Pacific basins ( [[#Kucharski--2014|Kucharski et al., 2014]] ; [[#Kajtar--2018|Kajtar et al., 2018]] ; [[#Luo--2018|Luo et al., 2018]] ; [[#McGregor--2018|McGregor et al., 2018]] ; [[#Seager--2019|Seager et al., 2019]] ), and/or a misrepresentation of radiative forcing (Sections 9.2.1 and 3.7.6). However, the observed trends in the Pacific Ocean SSTs are still within the range of internal variability as simulated by large initial condition ensembles of CMIP5 and CMIP6 models ( [[#Olonscheck--2020|Olonscheck et al., 2020]] ; Watanabe et al., 2021). Because the causes of observed equatorial Pacific temperature gradient and Walker circulation trends are not well understood ( [[IPCC:Wg1:Chapter:Chapter-3#3.3.3.1|Section 3.3.3.1]] ), there is ''low confidence'' in their attribution to anthropogenic influences ( [[IPCC:Wg1:Chapter:Chapter-8#8.3.2.3|Section 8.3.2.3]] ), while there is ''medium confidence'' that the observed changes have resulted from internal variability (Sections 3.7.6 and 8.2.2.2).

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===== 7.4.4.2.2 Tropical Pacific temperature gradients in past high-CO 2 climates =====

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The AR5 stated that paleoclimate proxies indicate a reduction in the longitudinal SST gradient across the equatorial Pacific during the Mid-Pliocene Warm Period (MPWP; [[#Masson-Delmotte--2013|Masson-Delmotte et al., 2013]] ; see Cross-Chapter Box 2.1 and Cross-Chapter Box 2.4 in this Report). This assessment was based on SST reconstructions between two sites situated very close to the equator in the heart of the western Pacific warm pool and eastern Pacific cold tongue, respectively. Multiple SST reconstructions based on independent paleoclimate proxies generally agreed that during the Pliocene the SST gradient between these two sites was reduced compared with the modern long-term mean ( [[#Wara--2005|Wara et al., 2005]] ; [[#Dekens--2008|Dekens et al., 2008]] ; [[#Fedorov--2013|Fedorov et al., 2013]] ).

Since AR5, the generation of new SST records has led to a variety of revised gradient estimates, specifically the generation of a new record for the warm pool ( [[#Zhang--2014|Zhang et al., 2014]] ), the inclusion of SST reconstructions from sites in the South China Sea as warm pool estimates ( [[#O’Brien--2014|O’Brien et al., 2014]] ; [[#Zhang--2014|Zhang et al., 2014]] ), and the inclusion of several new sites from the eastern Pacific as cold tongue estimates ( [[#Zhang--2014|Zhang et al., 2014]] ; [[#Fedorov--2015|Fedorov et al., 2015]] ). Published estimates of the reduction in the longitudinal SST difference for the Late Pliocene, relative to either Late Quaternary (0–0.5 million years ago) or pre-industrial values, include 1°C to 1.5°C ( [[#Zhang--2014|Zhang et al., 2014]] ), 0.1°C to 1.9°C ( [[#Tierney--2019|Tierney et al., 2019]] ), and about 3°C ( [[#Ravelo--2014|Ravelo et al., 2014]] ; [[#Fedorov--2015|Fedorov et al., 2015]] ; [[#Wycech--2020|Wycech et al., 2020]] ). All of these studies report a further weakening of the longitudinal gradient based on records extending into the Early Pliocene. While these revised estimates differ in magnitude due to differences in the sites and SST proxies used, they all agree that the longitudinal gradient was weaker, and this is supported by the probabilistic approach of [[#Tierney--2019|Tierney et al. (2019)]] . However, given that there are currently relatively few western equatorial Pacific records from independent site locations, and due to uncertainties associated with the proxy calibrations ( [[#Haywood--2016a|Haywood et al., 2016a]] ), there is only ''medium confidence'' that the average longitudinal gradient in the tropical Pacific was weaker during the Pliocene than during the Late Quaternary.

To avoid the influence of local biases, changes in the longitudinal temperature difference within Pliocene model simulations are typically evaluated using domain-averaged SSTs within chosen east and west Pacific regions and as such there is sensitivity to methodology. Unlike the reconstructed estimates, longitudinal gradient changes simulated by the Pliocene Model Intercomparison Project Phase 1 (PlioMIP1) models do not agree on the change in sign and are reported as spanning approximately –0.5°C to +0.5°C by [[#Brierley--2015|Brierley et al. (2015)]] and approximately –1°C to +1°C by [[#Tierney--2019|Tierney et al. (2019)]] . Initial PlioMIP Phase 2 (PlioMIP2) analysis suggests responses similar to PlioMIP1 ( [[#Feng--2019|Feng et al., 2019]] ; [[#Haywood--2020|Haywood et al., 2020]] ). Models that include hypothetical modifications to cloud albedo or ocean mixing are required to simulate the substantially weaker longitudinal differences seen in reconstructions of the Early Pliocene ( [[#Fedorov--2013|Fedorov et al., 2013]] ; [[#Burls--2014a|Burls and Fedorov, 2014a]] ).

While more western Pacific warm pool temperature reconstructions are needed to refine estimates of the longitudinal gradient, several Pliocene SST reconstructions from the east Pacific indicate enhanced warming in the centre of the eastern equatorial cold tongue upwelling region ( [[#Liu--2019|Liu et al., 2019]] ). This enhanced warming in the east Pacific cold tongue appears to be dynamically consistent with reconstruction of enhanced subsurface warming ( [[#Ford--2015|Ford et al., 2015]] ) and enhanced warming in coastal upwelling regions, suggesting that the tropical thermocline was deeper and/or less stratified during the Pliocene. The Pliocene data therefore suggest that the observed cooling trend over the last 60 years in parts of the eastern equatorial Pacific ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1.1|Section 9.2.1.1]] and Figure 9.3; [[#Seager--2019|Seager et al., 2019]] ), whether forced or due to internal variability, involves transient processes that are probably distinct from the longer-time scale process ( [[#Burls--2014a|Burls and Fedorov, 2014a]] , b; [[#Luo--2015|Luo et al., 2015]] ; [[#Heede--2020|Heede et al., 2020]] ) that maintained warmer eastern Pacific SST during the Pliocene.

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===== 7.4.4.2.3 Overall assessment of tropical Pacific sea surface temperature gradients under CO 2 forcing =====

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The paleoclimate proxy record and ESM simulations of the MPWP, process understanding, and ESM projections of climate response to CO <sub>2</sub> forcing provide ''medium evidence'' and a ''medium agreement'' and thus ''medium confidence'' that equilibrium warming in response to elevated CO <sub>2</sub> will be characterized by a weakening of the east–west tropical Pacific SST gradient.

Overall the observed pattern of warming over the instrumental period, with a warming minimum in the eastern tropical Pacific Ocean (Figure 7.14a), stands in contrast to the equilibrium warming pattern either inferred from the MPWP proxy record or simulated by ESMs under CO <sub>2</sub> forcing. There is ''medium confidence'' that the observed strengthening of the east–west SST gradient is temporary and will transition to a weakening of the SST gradient on centennial time scales. However, there is only ''low confidence'' that this transition will emerge this century owing to a low degree of agreement across studies about the factors driving the observed strengthening of the east–west SST gradient and how those factors will evolve in the future. These trends in tropical Pacific SST gradients reflect changes in the climatology, rather than changes in ENSO amplitude or variability, which are assessed in ( [[IPCC:Wg1:Chapter:Chapter-4|Chapter 4]] [[IPCC:Wg1:Chapter:Chapter-4#4.3.3|Section 4.3.3]] ).

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