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

==== 3.2.2.3 Carbon and Ocean Acidification ====

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The Arctic and Southern Ocean have a systemic vulnerability to aragonite undersaturation (Orr et al., 2005 <sup>[[#fn:r501|501]]</sup> ). For the RCP8.5 scenario, the entire Arctic and Southern Ocean surface waters will ''very likely'' be typified by year-around conditions corrosive for aragonite minerals for 2090–2100 (Figure 3.4) (Hauri et al., 2015 <sup>[[#fn:r502|502]]</sup> ; Sasse et al., 2015 <sup>[[#fn:r503|503]]</sup> ), whilst under RCP2.6 the extent of undersaturated waters are reduced markedly. At a basin/circumpolar scale, there is ''high confidence'' in these projections due to our robust understanding of the driving mechanisms. However, there is ''medium confidence'' for the response of specific locations, due to the need for improved resolution of the local circulation, interactions with sea ice, and other processes that modulate the rate of acidification.

Under RCP8.5, melting ice causes the greatest declining rate of pH and CaCO 3 saturation state in the Central Arctic, Canadian Arctic Archipelago and Baffin Bay (Popova et al., 2014 <sup>[[#fn:r504|504]]</sup> ). In the Canada Basin, projections using RCP8.5 show reductions in mean surface pH from approximately 8.1 in 1986–2005 to 7.7 by 2066–2085, and aragonite saturation from 1.52–0.74 during the same period (Steiner et al., 2014 <sup>[[#fn:r505|505]]</sup> ). A shoaling of the aragonite saturation horizon of approximately 1200 m, a large increase in area extent of undersaturated surface waters, and a pH change in the surface water of –0.19 are projected using the SRES A1B scenario (broadly comparable to RCP6.0) in the Nordic Sea from 2000 to 2065 (Skogen et al., 2014 <sup>[[#fn:r506|506]]</sup> ). Under the same scenario, aragonite undersaturation is projected to occur in the bottom waters over the entire Kara Sea shelf by 2040 and over most of the Barents and East Greenland shelves by 2070 due to the accumulation of anthropogenic CO 2 (Wallhead et al., 2017 <sup>[[#fn:r507|507]]</sup> ).

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'''Figure 3.4'''

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'''The upper ocean (0–10 m) at end of this century (2081–2100), characterised by undersaturated conditions for aragonite across 90% confidence intervals (dark red to light red) for the Representative Concentration Pathway (RCP)8.5 (a) and RCP2.6 (b) scenarios in the Coupled Model Intercomparison Project Phase 5 (CMIP5). Saturation states are averaged and confidence intervals calculated at […]'''
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The upper ocean (0–10 m) at end of this century (2081–2100), characterised by undersaturated conditions for aragonite across 90% confidence intervals (dark red to light red) for the Representative Concentration Pathway (RCP)8.5 (a) and RCP2.6 (b) scenarios in the Coupled Model Intercomparison Project Phase 5 (CMIP5). Saturation states are averaged and confidence intervals calculated at each geographic location across the CNRM-CM5, HadGEM2-ES, GFDL-ESM2G, GFDL-ESM2G, IPSL-CM5-LR, IPSL-CM5-MR, MPI-LR, MPI-MR and NCAR-CESM1 models.

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Under RCP8.5, the rate of CO 2 uptake by the Southern Ocean is projected to increase from the contemporary 0.91 Pg C yr ''–'' 1 to 2.38(1.65–2.55) Pg C yr ''–'' 1 by 2100, but the growth in uptake rate will slow and likely stop around 2070 ± 10 corresponding to cumulative CO 2 emissions of 1600 Gt C (Kessler and Tjiputra, 2016 <sup>[[#fn:r508|508]]</sup> ; Wang et al., 2016b <sup>[[#fn:r509|509]]</sup> ). This halt in the increase in the uptake rate of CO 2 is linked to the combined feedbacks from well-understood reductions in buffering capacity and warming, as well as the increased upwelling rate of carbon-rich Circumpolar Deep Water (Hauck and Volker, 2015 <sup>[[#fn:r509|509]]</sup> ) (Cross-Chapter Box 7 in Chapter 3). Although there is ''high agreement'' amongst models, contemporary biases in the fluxes of CO 2 in CMIP5 models in the Southern Ocean (Mongwe et al., 2018 <sup>[[#fn:r510|510]]</sup> ) suggest ''medium confidence'' levels for these projections.

Alongside the mean state changes, Southern Ocean aragonite saturation is also affected by the seasonal cycle of carbonate as well as by the impact of reduced buffering capacity (SM3.2.4) on the seasonal cycle of CO 2 (Sasse et al., 201511v; McNeil and Sasse, 2016 <sup>[[#fn:r512|512]]</sup> ). This leads to an amplification of the seasonal variability of pCO 2 (Hauck and Volker, 2015 <sup>[[#fn:r513|513]]</sup> ; McNeil and Sasse, 2016 <sup>[[#fn:r514|514]]</sup> ; Landschützer et al., 2018 <sup>[[#fn:r515|515]]</sup> ) and the hydrogen ion concentration that accelerates the onset of hypercapnia (i.e., high pCO 2 levels; pCO2 &gt; 1000 μ atm) to nearly 2 decades (~2085) ahead of anthropogenic CO 2 forcing (McNeil and Sasse, 2016 <sup>[[#fn:r516|516]]</sup> ). The seasonal cycles of pH and aragonite saturation will be attenuated (Kwiatkowski and Orr, 2018 <sup>[[#fn:r517|517]]</sup> ) (Section 5.2.2.3), however when the mean state changes are combined with the changes in seasonality, the onset of undersaturation is brought forward by 10–20 years (Table SM3.5). Model projections remain uncertain and affected by the resolution of local ocean physics, which leads to overall ''medium confidence'' in the timing of undersaturation and hypercapnia.

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