Editing IPCC:AR6/WGI/Chapter-4 (section)

===== 4.5.1.1.2 Polar amplification =====

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It is ''very likely'' that under all SSPs the warming in the Arctic will be more pronounced than in the global average over the 21st century. Since AR5 the understanding of the physical mechanisms driving polar amplification has improved.

The Arctic surface is projected to warm more than the global average over the 21st century, with annual-average Arctic warming of about 3°C (SSP1-2.6), 10°C (SSP3-7.0) and 12°C in (SSP5-8.5) by 2081–2100 relative to 1995–2014 (Figure 4.19). This phenomenon, known as polar or Arctic amplification, is a ubiquitous feature of the response to GHG forcing simulated by climate models ( [[#Manabe--1975|Manabe and Wetherald, 1975]] , 1980; [[#Manabe--1980|Manabe and Stouffer, 1980]] ; [[#Robock--1983|Robock, 1983]] ; [[#Hansen--1984|Hansen et al., 1984]] ; [[#Manabe--1991|Manabe et al., 1991]] ; [[#Holland--2003|Holland and Bitz, 2003]] ; [[#Winton--2006|Winton, 2006]] ; [[#Pithan--2014|Pithan and Mauritsen, 2014]] ) and has been observed over recent decades concurrent with Arctic sea ice loss ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.2.1|Section 2.3.2.1]] ; [[#Serreze--2011|Serreze and Barry, 2011]] ). Based on robust scientific understanding and agreement across multiple lines of evidence (Section 7.4.4.1), there is ''high confidence'' that the rate of Arctic surface warming will continue to exceed the global average over the 21st century.

A variety of mechanisms contribute to Arctic amplification (Section 7.4.4.1.1). While surface-albedo feedbacks associated with the loss of sea ice and snow have long been known to play important roles ( [[#Arrhenius--1896|Arrhenius, 1896]] ; [[#Manabe--1980|Manabe and Stouffer, 1980]] ; [[#Robock--1983|Robock, 1983]] ; [[#Hall--2004|Hall, 2004]] ), it is now recognized that temperature (lapse-rate and Planck) feedbacks also contribute to Arctic amplification through a less efficient longwave radiative damping to space with warming at high latitudes ( [[#Winton--2006|Winton, 2006]] ; [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Goosse--2018|Goosse et al., 2018]] ; [[#Stuecker--2018|Stuecker et al., 2018]] ). Increases in poleward atmospheric latent heat transport and oceanic heat transport also contribute to Arctic warming ( [[#Holland--2003|Holland and Bitz, 2003]] ; [[#Bitz--2006|Bitz et al., 2006]] ; [[#Lee--2011|Lee et al., 2011]] , [[#Lee--2017|Lee et al., 2017]] ; [[#Alexeev--2013|Alexeev and Jackson, 2013]] ; [[#Marshall--2014|Marshall et al., 2014]] , 2015; [[#Woods--2016|Woods and Caballero, 2016]] ; [[#Nummelin--2017|Nummelin et al., 2017]] ; [[#Singh--2017|Singh et al., 2017]] ; [[#Merlis--2018|Merlis and Henry, 2018]] ; [[#Oldenburg--2018|Oldenburg et al., 2018]] ; [[#Armour--2019|Armour et al., 2019]] ; [[#Beer--2020|Beer et al., 2020]] ). Projected reduction in the strength of the AMOC over the 21st century is expected to reduce Arctic warming, but even a strong AMOC reduction would not eliminate Arctic amplification entirely ( ''medium confidence'' ) ( [[#Liu--2017|Liu et al., 2017]] ; [[#Liu--2018|Liu et al., 2018]] ; [[#Wen--2018|Wen et al., 2018]] ).

There remains substantial uncertainty in the magnitude of projected Arctic amplification ( [[#Smith--2020|Smith et al., 2020]] ), with the Arctic warming ranging from two to four times the global average in models ( [[#Holland--2003|Holland and Bitz, 2003]] ; [[#Nummelin--2017|Nummelin et al., 2017]] ). This uncertainty primarily stems from different representations of polar surface-albedo, lapse-rate, and cloud feedbacks, and from different projected poleward energy transport changes ( [[#Holland--2003|Holland and Bitz, 2003]] ; [[#Crook--2011|Crook et al., 2011]] ; [[#Mahlstein--2011|Mahlstein and Knutti, 2011]] ; [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Bonan--2018|Bonan et al., 2018]] ). The magnitude of Arctic amplification may also depend on the mix of radiative forcing agents ( [[#Najafi--2015|Najafi et al., 2015]] ; [[#Sand--2016|Sand et al., 2016]] ; [[#Stjern--2019|Stjern et al., 2019]] ) such as the contribution of ozone depleting substances ( [[#Polvani--2020|Polvani et al., 2020]] ). Tropospheric aerosol emissions tend to reduce simulated Arctic warming over the middle of the 20th century ( [[#Gagné--2017b|Gagné et al., 2017b]] ) and consequently aerosol emission reductions in observations and SSP scenarios enhance simulated Arctic warming over recent and future decades (Section 6.4.3; [[#Gagné--2015|Gagné et al., 2015]] ; [[#Acosta%20Navarro--2016|Acosta Navarro et al., 2016]] ; [[#Wobus--2016|Wobus et al., 2016]] ; [[#Wang--2018|Wang et al., 2018]] ).

Climate models project a weaker polar amplification in the SH than in the NH under transient warming (Figure 4.19). Model simulations ( [[#Hall--2004|Hall, 2004]] ; [[#Danabasoglu--2009|Danabasoglu and Gent, 2009]] ; [[#Li--2013|Li et al., 2013]] ) and paleoclimate proxies indicate polar amplification in both hemispheres near equilibrium, but generally with less warming in the Antarctic than the Arctic (Section 7.4.4.1.2). The primary driver of delayed warming of the southern high latitudes is the upwelling in the Southern Ocean and associated ocean heat uptake that is then transported away from Antarctica by northward flowing surface waters ( [[#Frölicher--2015|Frölicher et al., 2015]] ; [[#Marshall--2015|Marshall et al., 2015]] ; [[#Armour--2016|Armour et al., 2016]] ; [[#Liu--2018|Liu et al., 2018]] ), although asymmetries in feedbacks between the poles also play a role (Section 7.4.4.1.1). Changes in westerly surface winds over the Southern Ocean have the potential to affect the rate of sea-surface warming, but there is currently ''low confidence'' in even the sign of the effect based on a diverse range of climate model responses to wind changes ( [[#Marshall--2014|Marshall et al., 2014]] ; [[#Ferreira--2015|Ferreira et al., 2015]] ; [[#Kostov--2017|Kostov et al., 2017]] ; [[#Seviour--2019|Seviour et al., 2019]] ). A substantial increase in freshwater input to the ocean from the Antarctic ice sheet could further slow the emergence of SH polar amplification by cooling the Southern Ocean surface ( [[#Bronselaer--2018|Bronselaer et al., 2018]] ; [[#Golledge--2019|Golledge et al., 2019]] ; [[#Schloesser--2019|Schloesser et al., 2019]] ), but this process is not represented in current climate models which lack dynamic ice sheets. Thus, while there is ''high confidence'' that the SH high latitudes will warm by more than the tropics on centennial time scales, there is ''low confidence'' that such a feature will emerge this century (Section 7.4.4.1).

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====== Seasonal warming patterns ======

The warming pattern shows distinct seasonal characteristics. The majority of models show a stronger hemispheric winter than summer warming over land poleward of about 55°N and 55°S (Figure 4.20) and thereby a reduced amplitude of the temperature cycle ( [[#Dwyer--2012|Dwyer et al., 2012]] ; [[#Donohoe--2013|Donohoe and Battisti, 2013]] ). On the other hand, over most of the subtropics and mid-latitudinal land regions except for parts of Asia, models project stronger warming in hemispheric summer than winter ( [[#Donohoe--2013|Donohoe and Battisti, 2013]] ; [[#Santer--2018|Santer et al., 2018]] ), leading to an amplification of the seasonal cycle. This phenomenon has been studied particularly in the case of the amplified summer warming over the Mediterranean region ( [[#Seager--2014a|Seager et al., 2014a]] ; [[#Kröner--2017|Kröner et al., 2017]] ; [[#Brogli--2019|Brogli et al., 2019]] ).

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'''Figure 4.20''' '''|''' '''Difference of surface temperature change between June–July–August (JJA) and December–January–February (DJF).''' Displayed are spatial patterns of multi-model mean difference in projected warming in JJA minus warming in DJF in 2081–2100 relative to 1995–2014 for '''(left)''' SSP1-2.6 and '''(right)''' SSP3-7.0. Diagonal lines mark areas where fewer than 80% of the models agree on the sign of change, and no overlay where at least 80% of the models agree. Further details on data sources and processing are available in the chapter data table (Table 4.SM.1).

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