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

===== 3.2.1.2.4 Carbon and ocean acidification =====

Various elements of marine biogeochemistry and geochemistry in the polar regions are of global importance. Here we focus on aspects relevant to carbon and ocean acidification; others (e.g., changes in dissolved oxygen) are assessed in Section 5.2.2. Compiled datasets on observed trends in ocean acidification from different observational platforms can be found in Table SM5.3.

About a quarter of carbon dioxide (CO 2 ) released by human activities is taken up by the ocean (WGI AR5, their Section 3.8). This dissolves in surface water to form carbonic acid, which, upon dissociation, causes a decrease in pH (acidification) and carbonate ion (CO 3 2– ) concentration. This can affect organisms that form shells and skeletons using calcium carbonate (CaCO 3 , aragonite and calcite as dominant mineral forms). Since AR5, new observations have demonstrated the spatial and temporal variability of ocean acidification and controlling mechanisms of carbon systems in different regions (Bellerby et al., 2018 <sup>[[#fn:r276|276]]</sup> ).

Robbins et al. (2013) <sup>[[#fn:r277|277]]</sup> showed aragonite undersaturation for about 20% of surface waters in the Canada and Makarov Basins, where substantial sea ice melt occurred. Qi et al. (2017) <sup>[[#fn:r278|278]]</sup> reported that aragonite undersaturation has expanded northward by at least 5° of latitude, and deepened by ~100 m between the 1990s and 2010 primarily due to increased Pacific Winter Water transport. In the East Siberian Arctic Shelf, extreme aragonite undersaturation was driven by the degradation of terrestrial organic matter and runoff of Arctic river water with elevated CO 2 concentrations, reflecting pH changes in excess of those projected in this region for 2100 (Semiletov et al., 2016 <sup>[[#fn:r279|279]]</sup> ) ( ''high confidence'' ); this was also observed along the continental margin and traced in the deep Makarov and Canada Basins (Anderson et al., 2017a <sup>[[#fn:r280|280]]</sup> ). The variable buffering capacities of rivers flowing through watersheds with different bedrock geology also influenced the state of ocean acidification in coastal regions (Tank et al., 2012 <sup>[[#fn:r281|281]]</sup> ; Azetsu-Scott et al., 2014 <sup>[[#fn:r282|282]]</sup> ).

The dissolved inorganic carbon (DIC) concentration increased in subsurface waters (150–1400 m) in the central Arctic between 1991 and 2011 (Ericson et al., 2014 <sup>[[#fn:r283|283]]</sup> ). The rate of increase was 0.6–0.9 µmol kg –1 yr –1 in the Arctic Atlantic Water and 0.4–0.6 µmol kg –1 yr –1 in the upper Polar Deep Water due to anthropogenic CO 2 , while no trend was observed in nutrient concentrations. In waters below 2000 m, no significant trend was observed for DIC and nutrient concentrations. Observation-based estimates (MacGilchrist et al., 2014 <sup>[[#fn:r284|284]]</sup> ) revealed a net summertime pan-Arctic export of 231 ± 49 TgC yr –1 of DIC across the Arctic Ocean gateways to the North Atlantic; at least 166 ± 60 TgC yr –1 of this was sequestered from the atmosphere ( ''medium confidence)'' . Similar to other regions (Table SM5.3), observed changes in the carbonate chemistry of the Arctic are indicative of ongoing ocean acidification ( ''high confidence'' ).

Studies covering seasonal-to-decadal variability in the Arctic are limited, with most conducted in ice-free or low ice periods during summer to autumn. However, it has been demonstrated that biological processes, respiration and photosynthesis, control the CaCO 3 saturation states in Chukchi Sea bottom water (Yamamoto-Kawai et al., 2016 <sup>[[#fn:r285|285]]</sup> ). Sea ice formation and melt influence the dynamics of ikaite (CaCO 3 precipitation trapped in sea ice during brine rejection), and therefore local carbonate chemistry (Rysgaard et al., 2013 <sup>[[#fn:r286|286]]</sup> ; Bates et al., 2014 <sup>[[#fn:r287|287]]</sup> ; Geilfus et al., 2016 <sup>[[#fn:r288|288]]</sup> ; Fransson et al., 2017 <sup>[[#fn:r289|289]]</sup> ). Although the increase of pH and saturation states by biological carbon fixation that consumes DIC in surface water is well documented (Azetsu-Scott et al., 2014 <sup>[[#fn:r290|290]]</sup> ; Yamamoto-Kawai et al., 2016 <sup>[[#fn:r291|291]]</sup> ) ( ''high confidence)'' , it has been shown that long photoperiods in Arctic summers sustain high pH in kelp forests, slowing ocean acidification (Krause-Jensen et al., 2016 <sup>[[#fn:r292|292]]</sup> ).

Since AR5, there are new constraints on the seasonal-to-decadal variability in the Southern Ocean CO 2 flux (McNeil and Matear, 2013 <sup>[[#fn:r293|293]]</sup> ; Landschützer et al., 2014 <sup>[[#fn:r294|294]]</sup> ; Landschützer et al., 2015 <sup>[[#fn:r295|295]]</sup> ; Gregor et al., 2017 <sup>[[#fn:r296|296]]</sup> ; Ritter et al., 2017 <sup>[[#fn:r297|297]]</sup> ; Keppler and Landschutzer, 2019 <sup>[[#fn:r298|298]]</sup> ) (Figure SM3.4), with mean annual flux anomalies varying from 0.3 ± 0.1 Pg C yr –1 in 2001–2002 to –0.4 Pg C yr –1 in 2012 (Landschützer et al., 2015 <sup>[[#fn:r299|299]]</sup> ); this can affect the magnitude of the global CO 2 sink (Section 5.2.2). A weakening CO 2 sink during the 1990s (Le Quéré et al., 2007) reversed in the 2000s as part of a decadal cycle (Landschützer et al., 2015 <sup>[[#fn:r300|300]]</sup> ; Munro et al., 2015 <sup>[[#fn:r3+1|3+1]]</sup> ; Williams et al., 2017 <sup>[[#fn:r302|302]]</sup> ) (SM3.2.3; Figure SM3.4), with a weakening again since 2011 (Keppler and Landschutzer, 2019 <sup>[[#fn:r303|303]]</sup> ). While the weakening sink during the 1990s was explained as a response to changes in the circumpolar winds over the Southern Ocean enhancing the outgassing of natural CO 2 , the subsequent changes appear due to a combination of changes in regional winds, temperature and circulation (Landschützer et al., 2015 <sup>[[#fn:r304|304]]</sup> ; Gregor et al., 2017 <sup>[[#fn:r305|305]]</sup> ; Keppler and Landschutzer, 2019 <sup>[[#fn:r306|306]]</sup> ). Data scarcity, especially in winter, remains a challenge (Ritter et al., 2017 <sup>[[#fn:r307|307]]</sup> ; Fay et al., 2018 <sup>[[#fn:r308|308]]</sup> ; Gruber et al., 2019b <sup>[[#fn:r309|309]]</sup> ); recent data from pH-enabled floats highlighted the potential role for winter outgassing south of the Polar Front (Williams et al., 2017 <sup>[[#fn:r310|310]]</sup> ; Gray et al., 2018 <sup>[[#fn:r311|311]]</sup> ). Overall, there is ''medium confidence'' that the Southern Ocean CO 2 sink has experienced significant decadal variations since the 1980s.

Southern Ocean carbon storage is affected by changes in overturning circulation (Cross-Chapter Box 7 in Chapter 3), with the storage of anthropogenic and natural carbon being both variable and out of phase on decadal timescales (DeVries et al., 2017 <sup>[[#fn:r312|312]]</sup> ; Tanhua et al., 2017 <sup>[[#fn:r313|313]]</sup> ) (Table SM3.4). Mode and intermediate waters are strongly involved in changing storage, also showing high sensitivity to shifts in winds (Swart et al., 2014 <sup>[[#fn:r314|314]]</sup> ; Swart et al., 2015a <sup>[[#fn:r315|315]]</sup> ; Tanhua et al., 2017 <sup>[[#fn:r316|316]]</sup> ; Gruber et al., 2019a <sup>[[#fn:r317|317]]</sup> ). Zonal basin differences in the uptake and storage of anthropogenic carbon are not well resolved and there is weak agreement between reanalysis products and Coupled Model Intercomparison Project Phase 5 (CMIP5) models (Swart et al., 2014 <sup>[[#fn:r318|318]]</sup> ). The presence of subduction hotspots suggest that basin-wide studies may be underestimating the importance of mode water subduction as a principal storage mechanism (Langlais et al., 2017 <sup>[[#fn:r319|319]]</sup> ).

Strengthening impacts of Southern Ocean acidification are illustrated by the 3.9 ± 1.3% decrease in derived calcification rates (1998–2014) (Freeman and Lovenduski, 2015 <sup>[[#fn:r320|320]]</sup> ). These have strong regional character, with decreases in the Indian and Pacific sectors (7.5–11.6%) and increases in the Atlantic (14.3 ± 5.1%). There have also been changes in the seasonality of pCO 2 linked to decreasing buffer capacity (McNeil and Sasse, 2016 <sup>[[#fn:r321|321]]</sup> ) (SM3.2.4) or adjustments to primary production (Conrad and Lovenduski, 2015 <sup>[[#fn:r322|322]]</sup> ); seasonal changes are discussed further in Section 5.2.2.

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