Editing IPCC:AR6/WGI/Chapter-5 (section)

=== 5.3.3 Ocean Interior Change ===

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==== 5.3.3.1 Ocean Memory: Acidification in the Ocean Interior ====

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Advances in observations and modelling for ocean physics and biogeochemistry and established knowledge of ocean carbonate chemistry show with ''very high confidence'' that anthropogenic CO <sub>2</sub> taken up into the ocean surface layer is further spreading into the ocean interior through ventilation processes, including vertical mixing, diffusion, subduction and meridional overturning circulations (Sections 2.3.3.5, 5.2.1.3 and 9.2.2.3; [[#Sallée--2012|Sallée et al., 2012]] ; [[#Bopp--2015|Bopp et al., 2015]] ; [[#Nakano--2015|Nakano et al., 2015]] ; [[#Iudicone--2016|Iudicone et al., 2016]] ; [[#Toyama--2017|Toyama et al., 2017]] ; [[#Pérez--2018|Pérez et al., 2018]] ; [[#Gruber--2019b|Gruber et al., 2019b]] ) and is causing acidification in the ocean interior. The net change in oxygen consumption by aerobic respiration of marine organisms further influences acidification by releasing CO <sub>2</sub> ( [[#5.3.3.2|Section 5.3.3.2]] ; [[#Chen--2017|Chen et al., 2017]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ; [[#Robinson--2019|Robinson, 2019]] ).

Observations over past decades of basin-wide and global syntheses of ocean interior carbon show that the extent of acidification due to anthropogenic CO <sub>2</sub> invasion tends to diminish with depth ( ''very high confidence'' ) ( [[#5.2.1.3.3|Section 5.2.1.3.3]] and Figure 5.21; [[#Woosley--2016|Woosley et al., 2016]] ; [[#Carter--2017|Carter et al., 2017]] ; [[#Lauvset--2020|Lauvset et al., 2020]] ). The regions of deep convection such as subpolar North Atlantic and Southern Ocean present the deepest acidification detections below 2000 m ( ''medium confidence'' ). Mid-latitudinal zones within the subtropical cells and tropical regions present a relatively deep and shallow detection, respectively. A pH decrease has also been observed on the Antarctic continental shelf ( [[#Hauck--2010|Hauck et al., 2010]] ; [[#Williams--2015|Williams et al., 2015]] ). Acidification is also underway in the subsurface to intermediate layers of the Arctic Ocean due to the inflow of ventilated waters from the North Atlantic and the North Pacific ( [[#Qi--2017|Qi et al., 2017]] ; [[#Ulfsbo--2018|Ulfsbo et al., 2018]] ).

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[[File:003c6023a9968316104dcfa15c261c17 IPCC_AR6_WGI_Figure_5_21.png]]

'''Figure 5.21 |''' '''Spread of ocean acidification from the surface into the interior of ocean since pre-industrial times''' . '''(a)''' Map showing the three transects used to create the cross sections shown in (b) and (c); vertical sections of the changes in '''(b)''' pH and '''(c)''' saturation state of aragonite (Ω <sub>arag</sub> ) between 1800–2002 due to anthropogenic CO <sub>2</sub> invasion (colour). Contour lines are their contemporary values in 2002. The red transect begins in the Nordic Seas and then follows the GO-SHIP lines A16 southward in the Atlantic Ocean, SR04 and S04P westward in the Southern Ocean, and P16 northward in the Pacific Ocean. The purple line follows the GO-SHIP line I09 southward in the Indian Ocean. The green line on the smaller inset crosses the Arctic Ocean from the Bering Strait to North Pole along 175°W and from the North Pole to the Fram Strait along 5°E (Lauvset et al., 2020). Further details on data sources and processing are available in the chapter data table (Table 5.SM.6).

A significant increase in acidification resulting from net metabolic CO <sub>2</sub> release coupled with ocean circulation changes has been shown with ''high confidence'' in large swathes of intermediate waters in the Pacific and Atlantic oceans ( [[#Dore--2009|Dore et al., 2009]] ; [[#Byrne--2010|Byrne et al., 2010]] ; [[#Ríos--2015|Ríos et al., 2015]] ; [[#Chu--2016|Chu et al., 2016]] ; [[#Carter--2017|Carter et al., 2017]] ; [[#Lauvset--2020|Lauvset et al., 2020]] ). For example, ocean circulation contributes a pH change of –0.013 ± 0.013 to the overall observed change of –0.029 ± 0.014 for 1993–2013 at depths around 1000 m at 30°S–40°S in the South Atlantic ocean ( [[#Ríos--2015|Ríos et al., 2015]] ). Long-term repeated observations in the North Pacific show a decline in dissolved oxygen (–4.0 μmol kg <sup>−1</sup> per decade at maximum) being sustained in the intermediate water since the 1980s (Takatani et al., 2012; [[#Sasano--2015|Sasano et al., 2015]] ). The amplification of acidification associated with the weakening ventilation is thought to have been occurring persistently. In contrast, for the North Pacific subtropical mode water, large decadal variability in pH and aragonite saturation state with amplitudes of about 0.02 and about 0.1, respectively, are superimposed on secular declining trends due to anthropogenic CO <sub>2</sub> invasion ( [[#Oka--2019|Oka et al., 2019]] ). This is associated with the variability in ventilation due to the approximately 50% variation in the formation volume of the mode water that is forced remotely by the Pacific Decadal Oscillation ( [[#Qiu--2013|Qiu et al., 2013]] ; [[#Oka--2015|Oka et al., 2015]] ).

These trends of acidification in the ocean interior lead to ''high confidence'' in shoaling of the saturation horizons of calcium carbonate minerals where Ω = 1. In the Pacific Ocean where the aragonite saturation horizon is shallower (a few hundred metres to 1200 m; Figure 5.21c), the rate of its shoaling is in the order of 1–2 m yr <sup>–1</sup> ( [[#Feely--2012|Feely et al., 2012]] ; [[#Ross--2020|Ross et al., 2020]] ). In contrast, shoaling rates of 4 m yr <sup>–1</sup> to 1710 m for 1984–2008 and of 10–15 m yr <sup>–1</sup> to 2250 m for 1991–2016 have been observed in the Iceland sea and the Irminger sea, respectively ( [[#Olafsson--2009|Olafsson et al., 2009]] ; [[#Pérez--2018|Pérez et al., 2018]] ).

In summary, ocean acidification is spreading into the ocean interior. Its rates at depths are controlled by the ventilation of the ocean interior as well as anthropogenic CO <sub>2</sub> uptake at the surface, thereby diminishing with depth ( ''very high confidence'' ) (Figure 5.21). Variability in ocean circulation modulates the trend of ocean acidification at depths through the changes in ventilation and their impacts on metabolic CO <sub>2</sub> content. However, the large knowledge gap around ventilation changes leads to ''low confidence'' in their impacts in many ocean regions (Sections 5.3.3.2; 9.2.2.3 and 9.3.2).

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==== 5.3.3.2 Ocean Deoxygenation and its Implications for Greenhouse Gases ====

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As summarized in SROCC ( [[#5.2.2.4|Section 5.2.2.4]] ), there is a growing consensus that between 1970 and 2010 the open ocean has ''very likely'' lost 0.5–3.3% of its dissolved oxygen in the upper 1000 m depth ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.6|Section 2.3.3.6]] ; [[#Helm--2011|Helm et al., 2011]] ; [[#Ito--2017|Ito et al., 2017]] ; [[#Schmidtko--2017|Schmidtko et al., 2017]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ). Regionally, the equatorial and North Pacific, the Southern Ocean and the South Atlantic have shown the greatest oxygen loss of up to 30 mol m <sup>–2</sup> per decade ( [[#Schmidtko--2017|Schmidtko et al., 2017]] ). Warming – via solubility reduction and circulation changes – mixing and respiration are considered the major drivers, with 50% of the oxygen loss for the upper 1000 m of the global oceans attributable to the solubility reduction ( [[#Schmidtko--2017|Schmidtko et al., 2017]] ). Climate variability also modifies the oxygen loss on interannual and decadal time scales especially for the tropical ocean OMZs ( [[#Deutsch--2011|Deutsch et al., 2011]] , 2014; [[#Llanillo--2013|Llanillo et al., 2013]] ) and the North Pacific subarctic zone ( [[#Whitney--2007|Whitney et al., 2007]] ; [[#Sasano--2018|Sasano et al., 2018]] ; [[#Cummins--2020|Cummins and Ross, 2020]] ). However, quantifying the oxygen decline and variability and attributing them to processes in different regions remains challenging (Levin, 2018; [[#Oschlies--2018|Oschlies et al., 2018]] ). Earth system models (ESMs) in CMIP5 and CMIP6 corroborate the decline in ocean oxygen, and project a continuing and accelerating decline with a strong impact of natural climate variability under high-emissions scenarios (Bopp et al., 2013; [[#Long--2016|Long et al., 2016]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). However, CMIP5 models did not reproduce observed patterns for oxygen changes in the tropical thermocline, and generally simulated only about half the oxygen loss inferred from observations ( [[#Oschlies--2018|Oschlies et al., 2018]] ). CMIP6 models have a more realistic simulated mean state of ocean biogeochemistry than CMIP5 models due to improved ocean physical processes and better representation of biogeochemical processes ( [[#Séférian--2020|Séférian et al., 2020]] ). Theyalso exhibit enhanced ocean warming as a result of an increase in the equilibrium climate sensitivity (ECS) of CMIP6 relative to CMIP5 models, which contributes to increased stratification and reduced subsurface ventilation (Sections 4.3.1, 4.3.4, 5.3.3.2, 7.4.2, 7.5.6, 9.2.1, and TS2.4). Consequently, CMIP6 model ensembles reproduce the ocean deoxygenation trend of −0.30 to −1.52 mmol m <sup>−3</sup> per decade between 1970–2010 reported in SROCC ( [[#5.2.2.4|Section 5.2.2.4]] ) with a very ''likely'' range, and also project 32–71% greater subsurface (100–600 m) oxygen decline relative to their Representative Concentration Pathway (RCP) analogues in CMIP5, reaching to the ''likely'' range of decline of 6.4 ± 2.9 mmol m <sup>–3</sup> under SSP1–2.6 and 13.3 ± 5.3 mmol m <sup>–3</sup> under SSP5–8.5, from 1870–1899 to 2080–2099 ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). It is concluded that the oxygen content of subsurface ocean is projected to transition to historically unprecedented condition with decline over the 21st century ( ''medi'' ''um confidence'' ).

In oxygen-depleted waters, microbial processes (denitrification and anammox, i.e., anaerobic ammonium oxidation; [[#Kuypers--2005|Kuypers et al., 2005]] ; [[#Codispoti--2007|Codispoti, 2007]] ; [[#Gruber--2008|Gruber and Galloway, 2008]] ) remove fixed nitrogen, and when upwelled waters reach the photic zone, primary production becomes nitrogen-limited ( [[#Tyrrell--2002|Tyrrell and Lucas, 2002]] ). However, in other oceanic regions, increased water-column stratification due to warming may reduce the amount of N <sub>2</sub> O reaching the surface and thereby decrease N <sub>2</sub> O flux to the atmosphere. [[#Landolfi--2017|Landolfi et al. (2017)]] suggest that, by 2100, under the RCP8.5 scenario, total N <sub>2</sub> O production in the ocean may decline by 5% and N <sub>2</sub> O emissions be reduced by 24% relative to the pre-industrial era due to decreased organic matter export and anthropogenic-driven changes in ocean circulation and atmospheric N <sub>2</sub> O concentrations. Projected oxygen loss in the ocean is thought to result in an ocean-climate feedback through changes in the natural emissions of GHGs ( ''l'' ''ow confidence'' ).

The areas with relatively rapid oxygen decrease include OMZs in the tropical oceans, where oxygen content has been decreasing at a rate of 0.9–3.4 µmol kg <sup>–1</sup> per decade in the thermocline for the past five decades (Stramma et al., 2008). Low oxygen, low pH and shallow aragonite saturation horizons in the OMZs of the eastern boundary upwelling regions co-occur, affecting ecosystem structure ( [[#Chavez--2008|Chavez et al., 2008]] ) and function in the water column, including the presently unbalanced nitrogen cycle ( [[#Paulmier--2009|Paulmier and Ruiz-Pino, 2009]] ). The coupling between upwelling, productivity, and oxygen depletion feeds back to biological productivity and the role of these regions as sinks or sources of climate active gases. When OMZ waters upwell and impinge on the euphotic zone, they release significant quantities of GHGs, including N <sub>2</sub> O (0.81–1.35 TgN yr <sup>–1</sup> ), CH <sub>4</sub> (0.27–0.38 TgCH <sub>4</sub> yr <sup>–1</sup> ), and CO <sub>2</sub> (yet to be quantified) to the atmosphere, exacerbating global warming ( [[#Paulmier--2008|Paulmier et al., 2008]] ; [[#Naqvi--2010|Naqvi et al., 2010]] ; [[#Kock--2012|Kock et al., 2012]] ; [[#Arévalo-Martínez--2015|Arévalo-Martínez et al., 2015]] ; [[#Babbin--2015|Babbin et al., 2015]] ; [[#Farías--2015|Farías et al., 2015]] ). Modelling projectionssuggest a global decrease of 4–12% in oceanic N <sub>2</sub> O emissions (from 3.71–4.03 TgN yr <sup>–1</sup> <sup></sup> to 3.54–3.56 TgN yr <sup>–1</sup> ) from 2005 to 2100 under RCP8.5, despite a tendency to increased N <sub>2</sub> O production in the OMZs, associated primarily with denitrification ( [[#Martinez-Rey--2015|Martinez-Rey et al., 2015]] ). It is difficult to single out the contribution of nitrification and denitrification, which can occur simultaneously. A rigorous separation of these two processes would require more mechanistic parametrization, which has been hindered by the still large conceptual and parametric uncertainties ( [[#Babbin--2015|Babbin et al., 2015]] ; [[#Trimmer--2016|Trimmer et al., 2016]] ; [[#Landolfi--2017|Landolfi et al., 2017]] ). Furthermore, the correlation between N <sub>2</sub> O and oxygen varies with microorganisms present, nutrient concentrations, and other environmental variables ( [[#Voss--2013|Voss et al., 2013]] ).

In summary, total oceanic N <sub>2</sub> O emissions were projected to decline by 4–12% from 2005–2100 ( [[#Martinez-Rey--2015|Martinez-Rey et al., 2015]] ) and by 24% from the pre-industrial era to 2100 ( [[#Landolfi--2017|Landolfi et al., 2017]] ) under RCP8.5. However, there is ''low confidence'' in the reduction in N <sub>2</sub> O emissions to the atmosphere, because of large conceptual and parametric uncertainties, a limited number of modelling studies that explored this process, and greater oxygen losses simulated in CMIP6 models than in CMIP5 models ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ).

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