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

==== 7.4.4.1 Polar Amplification ====

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Polar amplification describes the phenomenon where surface temperature change at high latitudes exceeds the global average surface temperature change in response to radiative forcing of the climate system. Arctic amplification, often defined as the ratio of Arctic to global surface warming, is a ubiquitous emergent feature of climate model simulations ( [[IPCC:Wg1:Chapter:Chapter-4#4.5.1|Section 4.5.1]] and Figure 7.12a; [[#Holland--2003|Holland and Bitz, 2003]] ; [[#Pithan--2014|Pithan and Mauritsen, 2014]] ) and is also seen in observations ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1|Section 2.3.1]] ). However, both climate models and observations show relatively less warming of the SH high latitudes compared to the Northern Hemisphere (NH) high latitudes over the historical record ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1|Section 2.3.1]] ), a characteristic that is projected to continue over the 21st century ( [[IPCC:Wg1:Chapter:Chapter-4#4.5.1|Section 4.5.1]] ). Since AR5 there is a much-improved understanding of the processes that drive polar amplification in the NH and delay its emergence in the Sh ( [[#7.4.4.1.1|Section 7.4.4.1.1]] ). Furthermore, the paleoclimate record provides evidence for polar amplification from multiple time periods associated with changes in CO <sub>2</sub> ( [[#Hollis--2019|Hollis et al., 2019]] ; [[#Cleator--2020|Cleator et al., 2020]] ; [[#McClymont--2020|McClymont et al., 2020]] ; [[#Tierney--2020b|Tierney et al., 2020b]] ), and allows an evaluation of polar amplification in model simulations of these periods ( [[#7.4.4.1.2|Section 7.4.4.1.2]] ). Research since AR5 identifies changes in the degree of polar amplification over time, particularly in the SH, as a key factor affecting how radiative feedbacks may evolve in the future ( [[#7.4.4.3|Section 7.4.4.3]] ).

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[[File:06a8b9db625dc756a4847f3a68f254f1 IPCC_AR6_WGI_Figure_7_12.png]]

'''Figure 7.12''' '''|''' '''Contributions of effective radiative forcing, ocean heat uptake, atmospheric heat transport, and radiative feedbacks to regional surface temperature changes at year 100 of''' ''abrupt 4xCO2'' '''simulations of CMIP6 Earth system models (ESMs).''' '''Figure 7.12: (a)''' Pattern of near-surface air temperature change. '''(b–d)''' Contributions to net Arctic (&gt;60°N), tropical (30°S–30°N), and Antarctic (&lt;60°S) warming calculated by dividing regional-average energy inputs by the magnitude of the regional-average Planck response. The contributions from radiative forcing, changes in moist, dry-static, and total atmospheric energy transport, ocean heat uptake, and radiative feedbacks (orange bars) all sum to the value of net warming (grey bar). Inset shows regional warming contributions associated with individual feedbacks, all summing to the total feedback contribution. Uncertainties (represented by black whiskers) show the interquartile range (25th and 75th percentiles) across models. The warming contributions (units of °C) for each process are diagnosed by calculating the energy flux (units of W m <sup>–2</sup> ) that each process contributes to the atmosphere over a given region, either at the top-of-atmosphere or surface, then dividing that energy flux by the magnitude of the regional Planck response (around 3.2 W m <sup>–2</sup> °C <sup>–1</sup> but varying with region). By construction, the individual warming contributions sum to the total warming in each region. Radiative kernel methods ( [[#7.4.1|Section 7.4.1]] ) are used to decompose the net energy input from radiative feedbacks into contributions from changes in atmospheric water vapour, lapse rate, clouds, and surface albedo ( [[#Zelinka--2020|Zelinka et al. (2020)]] using the [[#Huang--2017|Huang et al. (2017)]] radiative kernel). The CMIP6 models included are those analysed by [[#Zelinka--2020|Zelinka et al. (2020)]] and the warming contribution analysis is based on that of [[#Goosse--2018|Goosse et al. (2018)]] . Further details on data sources and processing are available in the chapter data table (Table 7.SM.14).

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<span id="critical-processes-driving-polar-amplification"></span>
===== 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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<span id="polar-amplification-from-proxies-and-models-during-past-climates-associated-with-co-2-change"></span>
===== 7.4.4.1.2 Polar amplification from proxies and models during past climates associated with CO 2 change =====

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Paleoclimate proxy data provide observational evidence of large-scale patterns of surface warming in response to past forcings, and allow an evaluation of the modelled response to these forcings (Sections 3.3.1.1 and 3.8.2.1). In particular, paleoclimate data provide evidence for long-term changes in polar amplification during time periods in which the primary forcing was a change in atmospheric CO <sub>2</sub> , although data sparsity means that for some time periods this evidence may be limited to a single hemisphere or ocean basin, or the evidence may come primarily from the mid-latitudes as opposed to the polar regions. In this context, there has been a modelling and data focus on the Last Glacial Maximum (LGM) in the context of PMIP4 ( [[#Cleator--2020|Cleator et al., 2020]] ; [[#Tierney--2020b|Tierney et al., 2020b]] ; [[#Kageyama--2021|Kageyama et al., 2021]] ), the mid-Pliocene Warm Period (MPWP) in the context of PlioMIP2 (Cross-Chapter Box 2.4; [[#Salzmann--2013|Salzmann et al., 2013]] ; [[#Haywood--2020|Haywood et al., 2020]] ; [[#McClymont--2020|McClymont et al., 2020]] ), the Early Eocene Climatic Optimum (EECO) in the context of DeepMIP ( [[#Hollis--2019|Hollis et al., 2019]] ; [[#Lunt--2021|Lunt et al., 2021]] ), and there is growing interest in the Miocene ( [[#Goldner--2014b|Goldner et al., 2014b]] ; [[#Steinthorsdottir--2021|Steinthorsdottir et al., 2021]] ; for definitions of time periods see Cross-Chapter Box 2.1). For all these time periods, in addition to the CO <sub>2</sub> forcing there are long-term feedbacks associated with ice sheets ( [[#7.4.2.6|Section 7.4.2.6]] ), and in particular for the Early Eocene there is a forcing associated with paleogeographic change ( [[#Farnsworth--2019|Farnsworth et al., 2019]] ). However, because these non-CO <sub>2</sub> effects can all be included as boundary conditions in model simulations, these time periods allow an assessment of the patterns of modelled response to known forcing (although uncertainty in the forcing increases further back in time). Because these changes to boundary conditions can be complex to implement in models, and because long simulations (typically longer than 500 years) are required to approach equilibrium, these simulations have been carried out mostly by pre-CMIP6 models, with relatively few (or none for the Early Eocene) fully coupled CMIP6 models in the ensembles.

At the time of AR5, polar amplification was evident in proxy reconstructions of paleoclimate sea surface temperature (SST) and surface air temperature (SAT) from the LGM, MPWP and the Early Eocene, but uncertainties associated with proxy calibrations ( [[#Waelbroeck--2009|Waelbroeck et al., 2009]] ; [[#Dowsett--2012|Dowsett et al., 2012]] ; [[#Lunt--2012a|Lunt et al., 2012a]] ) and the role of orbital forcing (for the MPWP; [[#Lisiecki--2005|Lisiecki and Raymo, 2005]] ) meant that the degree of polar amplification during these time periods was not accurately known. Furthermore, although some models (CCSM3; [[#Winguth--2010|Winguth et al., 2010]] ; [[#Huber--2011|Huber and Caballero, 2011]] ) at that time were able to reproduce the strong polar amplification implied by temperature proxies of the Early Eocene, this was achieved at higher CO <sub>2</sub> concentrations (&gt;2000 ppm) than those indicated by CO <sub>2</sub> proxies (&lt;1500 ppm; [[#Beerling--2011|Beerling and Royer, 2011]] ).

Since AR5 there has been progress in improving the accuracy of proxy temperature reconstructions of the LGM ( [[#Cleator--2020|Cleator et al., 2020]] ; [[#Tierney--2020b|Tierney et al., 2020b]] ), the MPWP ( [[#McClymont--2020|McClymont et al., 2020]] ), and the Early Eocene ( [[#Hollis--2019|Hollis et al., 2019]] ) time periods. In addition, reconstructions of the MPWP have been focused on a short time slice with an orbit similar to modern-day (isotopic stage KM5C; [[#Haywood--2013|Haywood et al., 2013]] , 2016b). Furthermore, there are more robust constraints on CO <sub>2</sub> concentrations from the MPWP ( [[#Martínez-Botí--2015|Martínez-Botí et al., 2015]] ; [[#de%20la%20Vega--2020|de la Vega et al., 2020]] ) and the Early Eocene ( [[#Anagnostou--2016|Anagnostou et al., 2016]] , 2020). As such, polar amplification during the LGM, MPWP, and Early Eocene time periods can now be better quantified than at the time of AR5, and the ability of climate models to reproduce this pattern can be better assessed; model-data comparisons for SAT and SST for these three time periods are shown in Figure 7.13.

Since AR5, there has been progress in the simulation of polar amplification by paleoclimate models of the Early Eocene. Initial work indicated that changes to model parameters associated with aerosols and/or clouds could increase simulated polar amplification and improve agreement between models and paleoclimate data ( [[#Kiehl--2013|Kiehl and Shields, 2013]] ; [[#Sagoo--2013|Sagoo et al., 2013]] ), but such parameter changes were not physically based. In support of these initial findings, a more recent (CMIP5) climate model, that includes a process-based representation of cloud microphysics, exhibits polar amplification in better agreement with proxies when compared to the models assessed in AR5 ( [[#Zhu--2019a|Zhu et al., 2019a]] ). Since then, some other CMIP3 and CMIP5 models in the DeepMIP multi-model ensemble ( [[#Lunt--2021|Lunt et al., 2021]] ) have obtained polar amplification for the EECO that is consistent with proxy indications of both polar amplification and CO <sub>2</sub> . Although there is a lack of tropical proxy SAT estimates, both proxies and DeepMIP models show greater terrestrial warming in the high latitudes than the mid-latitudes in both hemispheres (Figure 7.13a,d). SST proxies also exhibit polar amplification in both hemispheres, but the magnitude of this polar amplification is too low in the models, in particular in the south-west Pacific (Figure 7.13g,j).

For the MPWP, model simulations are now in better agreement with proxies than at the time of AR5 ( [[#Haywood--2020|Haywood et al., 2020]] ; [[#McClymont--2020|McClymont et al., 2020]] ). In particular, in the tropics new proxy reconstructions of SSTs are warmer and in better agreement with the models, due in part to the narrower time window in the proxy reconstructions. There is also better agreement at higher latitudes (primarily in the North Atlantic), due in part to the absence of some very warm proxy SSTs due to the narrower time window ( [[#McClymont--2020|McClymont et al., 2020]] ), and in part to a modified representation of Arctic gateways in the most recent Pliocene model simulations ( [[#Otto-Bliesner--2017|Otto-Bliesner et al., 2017]] ), which have resulted in warmer modelled SSTs in the North Atlantic ( [[#Haywood--2020|Haywood et al., 2020]] ). Furthermore, as for the Eocene, improvements in the representation of aerosol–cloud interactions have also led to improved model-data consistency at high latitudes ( [[#Feng--2019|Feng et al., 2019]] ). Although all PlioMIP2 models exhibit polar amplification of SAT, due to the relatively narrow time window there are insufficient terrestrial proxies to assess this (Figure 7.13b,e). However, polar SST amplification in the PlioMIP2 ensemble mean is in reasonably good agreement with that from SST proxies in the Northern Hemisphere (Figure 7.13h,k).

The Last Glacial Maximum (LGM) also gives an opportunity to evaluate model simulation of polar amplification under CO <sub>2</sub> forcing, albeit under colder conditions than today ( [[#Kageyama--2021|Kageyama et al., 2021]] ). Terrestrial SAT and marine SST proxies exhibit clear polar amplification in the Northern Hemisphere, and the PMIP4 models capture this well (Figure 7.13c,f,i,l), particularly for SAT. There is less proxy data in the mid- to high latitudes of the Southern Hemisphere, but here the models exhibit polar amplification of both SST and SAT. LGM regional model-data agreement is also assessed in ( [[IPCC:Wg1:Chapter:Chapter-3|Chapter 3]] [[IPCC:Wg1:Chapter:Chapter-3#3.8.2|Section 3.8.2]] ).

Overall, the proxy reconstructions give ''high confidence'' that there was polar amplification in the LGM, MPWP and EECO, and this is further supported by model simulations of these time periods (Figure 7.13; [[#Zhu--2019a|Zhu et al., 2019a]] ; [[#Haywood--2020|Haywood et al., 2020]] ; [[#Kageyama--2021|Kageyama et al., 2021]] ; [[#Lunt--2021|Lunt et al., 2021]] ). For both the MPWP and EECO, models are more consistent with the temperature and CO <sub>2</sub> proxies than at the time of AR5 ( ''high confidence'' ). For the LGM Northern Hemisphere, which is the region with the most data and the time period with the least uncertainty in model boundary conditions, polar amplification in the PMIP4 ensemble mean is in good agreement with the proxies, especially for SAT ( ''medium confidence'' ). Overall, the confidence in the ability of models to accurately simulate polar amplification is higher than at the time of AR5, but a more complete model evaluation could be carried out if there were more CMIP6 paleoclimate simulations included in the assessment.

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===== 7.4.4.1.3 Overall assessment of polar amplification =====

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Based on mature process understanding of the roles of poleward latent heat transport and radiative feedbacks in polar warming, a high degree of agreement across a hierarchy of climate models, observational evidence, paleoclimate proxy records of past climates associated with CO <sub>2</sub> change, and ESM simulations of those past climates, there is ''high confidence'' that polar amplification is a robust feature of the long-term response to greenhouse gas forcing in both hemispheres. Stronger warming in the Arctic than the global average has already been observed ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1|Section 2.3.1]] ) and its causes are well understood. It is ''very likely'' that the warming in the Arctic will be more pronounced than the global average over the 21st century ( ''high confidence'' ) [[IPCC:Wg1:Chapter:Chapter-4#4.5.1.1|Section 4.5.1.1]] ). This is supported by models’ improved ability to simulate polar amplification during past time periods, compared with at the time of AR5 ( ''high confidence'' ); although this is based on an assessment of mostly non-CMIP6 models.

Southern Ocean SSTs have been slow to warm over the instrumental period, with cooling since about 1980 owing to a combination of upper-ocean freshening from ice-shelf melt, intensification of surface westerly winds from ozone depletion, and variability in ocean convection ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1|Section 9.2.1]] ). This stands in contrast to the equilibrium warming pattern either inferred from the proxy record or simulated by ESMs under CO <sub>2</sub> forcing. There is ''high confidence'' that the SH high latitudes will warm more than the tropics on centennial time scales as the climate equilibrates with radiative forcing and Southern Ocean heat uptake is reduced. However, there is only ''low confidence'' that this feature will emerge this century.

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