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

===== 7.4.4.1.1 Critical processes driving polar amplification =====

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Several processes contribute to polar amplification under greenhouse gas forcing, including the loss of sea ice and snow (an amplifying surface-albedo feedback), the confinement of warming to near the surface in the polar atmosphere (an amplifying lapse-rate feedback), and increases in poleward atmospheric and oceanic heat transport ( [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Goosse--2018|Goosse et al., 2018]] ; [[#Dai--2019|Dai et al., 2019]] ; [[#Feldl--2020|Feldl et al., 2020]] ). Modelling and process studies since AR5 have led to an improved understanding of the combined effect of these different processes in driving polar amplification and how they differ between the hemispheres.

Idealized modelling studies suggest that polar amplification would occur even in the absence of any amplifying polar surface-albedo or lapse-rate feedbacks owing to changes in poleward atmospheric heat transport under global warming ( [[#Hall--2004|Hall, 2004]] ; [[#Alexeev--2005|Alexeev et al., 2005]] ; [[#Graversen--2009|Graversen and Wang, 2009]] ; [[#Alexeev--2013|Alexeev and Jackson, 2013]] ; [[#Graversen--2014|Graversen et al., 2014]] ; [[#Roe--2015|Roe et al., 2015]] ; [[#Merlis--2018|Merlis and Henry, 2018]] ; [[#Armour--2019|Armour et al., 2019]] ). Poleward heat transport changes reflect compensating changes in the transport of latent energy (moisture) and dry-static energy (sum of sensible and potential energy) by atmospheric circulations ( [[#Alexeev--2005|Alexeev et al., 2005]] ; [[#Held--2006|Held and Soden, 2006]] ; [[#Hwang--2010|Hwang and Frierson, 2010]] ; [[#Hwang--2011|Hwang et al., 2011]] ; [[#Kay--2012|Kay et al., 2012]] ; [[#Huang--2014|Huang and Zhang, 2014]] ; [[#Feldl--2017a|Feldl et al., 2017a]] ; [[#Donohoe--2020|Donohoe et al., 2020]] ). ESMs project that within the mid-latitudes, where eddies dominate the heat transport, an increase in poleward latent energy transport arises from an increase in the equator-to-pole gradient in atmospheric moisture with global warming, with moisture in the tropics increasing more than at the poles as described by the Clausius–Clapeyron relation ( [[IPCC:Wg1:Chapter:Chapter-8#8.2|Section 8.2]] ). This change is partially compensated by a decrease in dry-static energy transport arising from a weakening of the equator-to-pole temperature gradient as the polar regions warm more than the tropics.

Energy balance models that approximate atmospheric heat transport in terms of a diffusive flux down the meridional gradient of near-surface moist static energy (sum of dry-static and latent energy) are able to reproduce the atmospheric heat transport changes seen within ESMs ( [[#Flannery--1984|Flannery, 1984]] ; [[#Hwang--2010|Hwang and Frierson, 2010]] ; [[#Hwang--2011|Hwang et al., 2011]] ; [[#Rose--2014|Rose et al., 2014]] ; [[#Roe--2015|Roe et al., 2015]] ; [[#Merlis--2018|Merlis and Henry, 2018]] ), including the partitioning of latent and dry-static energy transports ( [[#Siler--2018b|Siler et al., 2018b]] ; [[#Armour--2019|Armour et al., 2019]] ). These models suggest that polar amplification is driven by enhanced poleward latent heat transport and that the magnitude of polar amplification can be enhanced or diminished by the latitudinal structure of radiative feedbacks. Amplifying polar feedbacks enhance polar warming and in turn cause a decrease in the dry-static energy transport to high latitudes ( [[#Alexeev--2013|Alexeev and Jackson, 2013]] ; [[#Rose--2014|Rose et al., 2014]] ; [[#Roe--2015|Roe et al., 2015]] ; [[#Bonan--2018|Bonan et al., 2018]] ; [[#Merlis--2018|Merlis and Henry, 2018]] ; [[#Armour--2019|Armour et al., 2019]] ; [[#Russotto--2020|Russotto and Biasutti, 2020]] ). Poleward latent heat transport changes act to favour polar amplification and inhibit tropical amplification ( [[#Armour--2019|Armour et al., 2019]] ), resulting in a strongly polar-amplified warming response to polar forcing and a more latitudinally uniform warming response to tropical forcing within ESMs ( [[#Alexeev--2005|Alexeev et al., 2005]] ; [[#Rose--2014|Rose et al., 2014]] ; [[#Stuecker--2018|Stuecker et al., 2018]] ). The important role for poleward latent energy transport in polar amplification is supported by studies of atmospheric reanalyses and ESMs showing that episodic increases in latent heat transport into the Arctic can enhance surface downwelling radiation and drive sea ice loss on sub-seasonal time scales ( [[#Woods--2016|Woods and Caballero, 2016]] ; [[#Gong--2017|Gong et al., 2017]] ; [[#Lee--2017|Lee et al., 2017]] ; B. [[#Luo--2017|]] [[#Luo--2017|Luo et al., 2017]] ), however this may be a smaller driver of sea ice variability than atmospheric temperature fluctuations ( [[#Olonscheck--2019|Olonscheck et al., 2019]] ).

Regional energy budget analyses are commonly used to diagnose the relative contributions of radiative feedbacks and energy fluxes to polar amplification as projected by ESMs under increased CO <sub>2</sub> concentrations (Figure 7.12; [[#Feldl--2013|Feldl and Roe, 2013]] ; [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Goosse--2018|Goosse et al., 2018]] ; [[#Stuecker--2018|Stuecker et al., 2018]] ). These analyses suggest that a primary cause of amplified Arctic warming in ESMs is the latitudinal structure of radiative feedbacks, which warm the Arctic more than the tropics (Figure 7.12b), and enhanced latent energy transport into the Arctic. That net atmospheric heat transport into the Arctic does not change substantially within ESMs, on average, under CO <sub>2</sub> forcing (Figure 7.12b) reflects a compensating decrease in poleward dry-static energy transport as a response to polar amplified warming ( [[#Hwang--2011|Hwang et al., 2011]] ; [[#Armour--2019|Armour et al., 2019]] ; [[#Donohoe--2020|Donohoe et al., 2020]] ). The latitudinal structure of radiative feedbacks primarily reflects that of the surface-albedo and lapse-rate feedbacks, which preferentially warm the Arctic ( [[#Graversen--2014|Graversen et al., 2014]] ; [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Goosse--2018|Goosse et al., 2018]] ). Latitudinal structure in the lapse-rate feedback reflects weak radiative damping to space with surface warming in polar regions, where atmospheric warming is constrained to the lower troposphere owing to stably stratified conditions, and strong radiative damping in the tropics, where warming is enhanced in the upper troposphere owing to moist convective processes. This is only partially compensated by latitudinal structure in the water-vapour feedback ( [[#Taylor--2013|Taylor et al., 2013]] ), which favours tropical warming ( [[#Pithan--2014|Pithan and Mauritsen, 2014]] ). While cloud feedbacks have been found to play little role in Arctic amplification in CMIP5 models ( [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Goosse--2018|Goosse et al., 2018]] ; Figure 7.12b), less-negative cloud feedbacks at high latitude, as seen within some CMIP6 models ( [[#Zelinka--2020|Zelinka et al., 2020]] ), tend to favour stronger polar amplification ( [[#Dong--2020|Dong et al., 2020]] ). A weaker Planck response at high latitudes, owing to less efficient radiative damping where surface and atmospheric temperatures are lower, also contributes to polar amplification ( [[#Pithan--2014|Pithan and Mauritsen, 2014]] ). The effective radiative forcing of CO <sub>2</sub> is larger in the tropics than at high latitudes, suggesting that warming would be tropically amplified if not for radiative feedbacks and poleward latent heat transport changes (Figure 7.12b–d; [[#Stuecker--2018|Stuecker et al., 2018]] ).

While the contributions to regional warming can be diagnosed within ESM simulations (Figure 7.12), assessment of the underlying role of individual factors is limited by interactions inherent to the coupled climate system. For example, polar feedback processes are coupled and influenced by warming at lower latitudes ( [[#Screen--2012|Screen et al., 2012]] ; [[#Alexeev--2013|Alexeev and Jackson, 2013]] ; [[#Graversen--2014|Graversen et al., 2014]] ; [[#Graversen--2016|Graversen and Burtu, 2016]] ; [[#Rose--2016|Rose and Rencurrel, 2016]] ; [[#Feldl--2017a|Feldl et al., 2017a]] , 2020; [[#Yoshimori--2017|Yoshimori et al., 2017]] ; [[#Garuba--2018|Garuba et al., 2018]] ; [[#Po-Chedley--2018b|Po-Chedley et al., 2018b]] ; [[#Stuecker--2018|Stuecker et al., 2018]] ; [[#Dai--2019|Dai et al., 2019]] ), while atmospheric heat transport changes are in turn influenced by the latitudinal structure of regional feedbacks, radiative forcing, and ocean heat uptake ( [[#Hwang--2011|Hwang et al., 2011]] ; [[#Zelinka--2012|Zelinka and Hartmann, 2012]] ; [[#Feldl--2013|Feldl and Roe, 2013]] ; [[#Huang--2014|Huang and Zhang, 2014]] ; [[#Merlis--2014|Merlis, 2014]] ; [[#Rose--2014|Rose et al., 2014]] ; [[#Roe--2015|Roe et al., 2015]] ; [[#Feldl--2017b|Feldl et al., 2017b]] ; [[#Stuecker--2018|Stuecker et al., 2018]] ; [[#Armour--2019|Armour et al., 2019]] ). The use of different feedback definitions, such as a lapse-rate feedback partitioned into upper and lower tropospheric components ( [[#Feldl--2020|Feldl et al., 2020]] ) or including the influence of water vapour at constant relative humidity ( [[#Held--2012|Held and Shell, 2012]] ; [[#7.4.2|Section 7.4.2]] ), would also change the interpretation of which feedbacks contribute most to polar amplification.

The energy budget analyses (Figure 7.12) suggest that greater surface warming in the Arctic than the Antarctic under greenhouse gas forcing arises from two main processes. The first is large surface heat uptake in the Southern Ocean (Figure 7.12c) driven by the upwelling of deep waters that have not yet felt the effects of the radiative forcing; the heat taken up is predominantly transported away from Antarctica by northward-flowing surface waters ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1|Section 9.2.1]] ; [[#Marshall--2015|Marshall et al., 2015]] ; [[#Armour--2016|Armour et al., 2016]] ). Strong surface heat uptake also occurs in the subpolar North Atlantic Ocean under global warming ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1|Section 9.2.1]] ). However, this heat is partially transported northward into the Arctic, which leads to increased heat fluxes into the Arctic atmosphere (Figure 7.12b; [[#Rugenstein--2013|Rugenstein et al., 2013]] ; [[#Jungclaus--2014|Jungclaus et al., 2014]] ; [[#Koenigk--2014|Koenigk and Brodeau, 2014]] ; [[#Marshall--2015|Marshall et al., 2015]] ; [[#Nummelin--2017|Nummelin et al., 2017]] ; [[#Singh--2017|Singh et al., 2017]] ; [[#Oldenburg--2018|Oldenburg et al., 2018]] ). The second main process contributing to differences in Arctic and Antarctic warming is the asymmetry in radiative feedbacks between the poles ( [[#Yoshimori--2017|Yoshimori et al., 2017]] ; [[#Goosse--2018|Goosse et al., 2018]] ). This primarily reflects the weaker lapse-rate and surface-albedo feedbacks and more-negative cloud feedbacks in the SH high latitudes (Figure 7.12). However, note the SH cloud feedbacks are uncertain due to possible biases in the treatment of mixed phase clouds ( [[#Hyder--2018|Hyder et al., 2018]] ). Idealized modelling suggests that the asymmetry in the polar lapse-rate feedback arises from the height of the Antarctic Ice Sheet precluding the formation of deep atmospheric inversions that are necessary to produce the stronger positive lapse-rate feedbacks seen in the Arctic ( [[#Salzmann--2017|Salzmann, 2017]] ; [[#Hahn--2020|Hahn et al., 2020]] ). ESM projections of the equilibrium response to CO <sub>2</sub> forcing show polar amplification in both hemispheres, but generally with less warming in the Antarctic than the Arctic (C. [[#Li--2013|]] [[#Li--2013|Li et al., 2013]] ; [[#Yoshimori--2017|Yoshimori et al., 2017]] ).

Because multiple processes contribute to polar amplification, it is a robust feature of the projected long-term response to greenhouse gas forcing in both hemispheres. At the same time, contributions from multiple processes make projections of the magnitude of polar warming inherently more uncertain than global mean warming ( [[#Holland--2003|Holland and Bitz, 2003]] ; [[#Roe--2015|Roe et al., 2015]] ; [[#Bonan--2018|Bonan et al., 2018]] ; [[#Stuecker--2018|Stuecker et al., 2018]] ). The magnitude of Arctic amplification ranges from a factor of two to four in ESM projections of 21st-century warming ( [[IPCC:Wg1:Chapter:Chapter-4#4.5.1|Section 4.5.1]] ). While uncertainty in both global and tropical warming under greenhouse gas forcing is dominated by cloud feedbacks ( [[#7.5.7|Section 7.5.7]] ; [[#Vial--2013|Vial et al., 2013]] ), uncertainty in polar warming arises from polar surface-albedo, lapse-rate, and cloud feedbacks, changes in atmospheric and oceanic poleward heat transport, and ocean heat uptake ( [[#Hwang--2011|Hwang et al., 2011]] ; [[#Mahlstein--2011|Mahlstein and Knutti, 2011]] ; [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Bonan--2018|Bonan et al., 2018]] ).

The magnitude of polar amplification also depends on the type of radiative forcing applied ( [[IPCC:Wg1:Chapter:Chapter-4#4.5.1.1|Section 4.5.1.1]] ; [[#Stjern--2019|Stjern et al., 2019]] ), with ( [[IPCC:Wg1:Chapter:Chapter-6|Chapter 6]] (Section 6.4.3) discussing changes in sulphate aerosol emissions and the deposition of black carbon aerosols on ice and snow as potential drivers of amplified Arctic warming. The timing of the emergence of SH polar amplification remains uncertain due to insufficient knowledge of the time scales associated with Southern Ocean warming and the response to surface wind and freshwater forcing ( [[#Bintanja--2013|Bintanja et al., 2013]] ; [[#Kostov--2017|Kostov et al., 2017]] , 2018; [[#Pauling--2017|Pauling et al., 2017]] ; [[#Purich--2018|Purich et al., 2018]] ). ESM simulations indicate that freshwater input from melting ice shelves could reduce Southern Ocean warming by up to several tenths of a °C over the 21st century by increasing stratification of the surface ocean around Antarctica ( ''low confidence'' due to ''medium agreement'' but ''limited evidence'' ) (Sections 7.4.2.6 and 9.2.1, and Box 9.3; [[#Bronselaer--2018|Bronselaer et al., 2018]] ; [[#Golledge--2019|Golledge et al., 2019]] ; [[#Lago--2019|Lago and England, 2019]] ). However, even a large reduction in the Atlantic Meridional Overturning Circulation (AMOC) and associated northward heat transport due, for instance, to greatly increased freshwater runoff from Greenland would be insufficient to eliminate Arctic amplification ( ''medium confidence'' based on ''medium agreement'' and ''medium evidence'' ) ( [[#Liu--2017|Liu et al., 2017]] ; Y. [[#Liu--2018|]] [[#Liu--2018|Liu et al., 2018]] ; [[#Wen--2018|Wen et al., 2018]] ).

Arctic amplification has a distinct seasonality with a peak in early winter (November to January) owing to sea ice loss and associated increases in heat fluxes from the ocean to the atmosphere resulting in strong near-surface warming ( [[#Pithan--2014|Pithan and Mauritsen, 2014]] ; [[#Dai--2019|Dai et al., 2019]] ). Surface warming may be further amplified by positive cloud and lapse-rate feedbacks in autumn and winter ( [[#Burt--2016|Burt et al., 2016]] ; [[#Morrison--2019|Morrison et al., 2019]] ; [[#Hahn--2020|Hahn et al., 2020]] ). Arctic amplification is weak in summer owing to surface temperatures remaining stable as excess energy goes into thinning the summertime sea ice cover, which remains at the melting point, or into the ocean mixed layer. Arctic amplification can also be interpreted through changes in the surface energy budget ( [[#Burt--2016|Burt et al., 2016]] ; [[#Woods--2016|Woods and Caballero, 2016]] ; [[#Boeke--2018|Boeke and Taylor, 2018]] ; [[#Kim--2019|Kim et al., 2019]] ), however such analyses are complicated by the finding that a large portion of the changes in downward longwave radiation can be attributed to the lower troposphere warming along with the surface itself ( [[#Vargas%20Zeppetello--2019|Vargas Zeppetello et al., 2019]] ).

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