Editing IPCC:AR6/SROCC/Chapter-3 (section)

==== 3.2.2.2 Physical Oceanography ====

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Consistent with the projected sea ice decline, there is ''high confidence'' that the Arctic Ocean will warm significantly towards the end of this century at the surface and in the deeper layers. Most CMIP5 models capture the seasonal changes in surface heat and freshwater fluxes for the present day climate, and show that the excess summer solar heating is used to melt sea ice, in a positive ice albedo feedback (Ding et al., 2016 <sup>[[#fn:r484|484]]</sup> ). Using RCP8.5, Vavrus et al. (2012) found that the Atlantic layer is projected to warm by 2.5°C at around 400 m depth at the end of the century, but only by 0.5°C in the surface mixed layer. Consistent results for lower Atlantic Water layer warming were found by Koenigk and Brodeau (2014) <sup>[[#fn:r485|485]]</sup> for RCP2.5 (0.5°C), RCP4.5 (1.0°C) and RCP8.5 (2.0°C).

Poleward ocean heat transport contributes to Arctic Ocean warming ( ''medium confidence'' ). Comparing 20 CMIP5 simulations for RCP8.5, Nummelin et al. (2017) found a 2°C–6°C range in Arctic amplification of surface air temperature north of 70°N, consistent with increased ocean heat transport. Comparing 26 different CMIP5 simulations for RCP4.5, Burgard and Notz (2017) found that ocean heat transport changes explain the Arctic Ocean multi-model mean warming, but that differences between models are compensated by changes in surface fluxes. Increased ocean heat transport into the Barents Sea beyond 2020 appears as a probable mechanism with continued warming (Koenigk and Brodeau, 2014 <sup>[[#fn:r486|486]]</sup> ; Årthun et al., 2019 <sup>[[#fn:r487|487]]</sup> ). Based on four CMIP5 models, the Barents Sea is projected to become ice-free during winter beyond 2050 under RCP8.5 (Onarheim and Årthun, 2017 <sup>[[#fn:r488|488]]</sup> ), to which the main response is an increased ocean-to-atmosphere heat flux and related surface warming (Smedsrud et al., 2013 <sup>[[#fn:r489|489]]</sup> ). The ocean heat transport increases in all Arctic gateways, but is dominated by the Barents Sea, and when winter sea ice disappears here the heat loss cannot increase further and the excess ocean heat continues into the Arctic Basin (Koenigk and Brodeau, 2014 <sup>[[#fn:r490|490]]</sup> ).

The surface mixed layer of the Arctic Ocean is expected to freshen in future because an intensified hydrological cycle will increase river runoff (Haine et al., 2015 <sup>[[#fn:r491|491]]</sup> ) ( ''medium confidence'' ). The related increase in stratification has the potential to contribute to the warming of the deep Atlantic Water layer, as upward vertical mixing will be reduced (Nummelin et al., 2016 <sup>[[#fn:r492|492]]</sup> ). There are, however, biases in salinity of ~1 across the Arctic Basin for the present day climate (Ilicak et al., 2016 <sup>[[#fn:r493|493]]</sup> ) in forced global ice-ocean models with configurations comparable to CMIP5, suggesting limited predictive skill for the Arctic freshwater cycle.

CMIP5 projections (Figure 3.3) indicate that observed Southern Ocean warming trends will continue under RCP4.5 and RCP8.5 scenarios, leading to 1°C–3°C warming by 2100 mostly in the upper ocean (Sallée et al., 2013a <sup>[[#fn:r494|494]]</sup> ). Projections demonstrate a similar distribution of heat storage to historical observations, notably focused in deep pools north of the Subantarctic Front (e.g., Armour et al., 2016 <sup>[[#fn:r495|495]]</sup> ). Antarctic Bottom Water becomes coherently warmer by up to 0.3°C by 2100 across the model ensemble under RCP8.5 (Heuzé et al., 2015 <sup>[[#fn:r496|496]]</sup> ). The upper ocean also becomes considerably fresher (salinity decrease of approximately 0.1) (Sallée et al., 2013b <sup>[[#fn:r497|497]]</sup> ) with an overall increase in stratification and a shallowing of mixed layers (Sallée et al., 2013a <sup>[[#fn:r498|498]]</sup> ). Although the sign of model changes appear mostly robust, there is ''low confidence'' in magnitude due to the large inter-model spread in projections and significant warm biases in historical water mass properties (Sallée et al., 2013a <sup>[[#fn:r499|499]]</sup> ) and sea surface temperature, which may be up to 3°C too high in the historical runs (Wang et al., 2014 <sup>[[#fn:r500|500]]</sup> ). Projections of changes in Southern Ocean circulation are discussed in Cross-Chapter Box 7 in Chapter 3.

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