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

== 5.3 Ocean Acidification and Deoxygenation ==

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The surface ocean has absorbed a quarter of all anthropogenic CO <sub>2</sub> emissions, mainly through physical–chemical processes ( [[#McKinley--2016|McKinley et al., 2016]] ; [[#Gruber--2019b|Gruber et al., 2019b]] ; [[#Friedlingstein--2020|Friedlingstein et al., 2020]] ). Once dissolved in seawater, CO <sub>2</sub> reacts with water and forms carbonic acid, which in turn dissociates, leading to a decrease in the concentration of carbonate (CO <sub>3</sub> <sup>–2</sup> ) ions, and increasing both bicarbonate (HCO <sub>3</sub> <sup>–</sup> ) and hydrogen (H <sup>+</sup> ) ion concentration. This process has caused a shift in the carbonate chemistry towards a less basic state, commonly referred to as ‘ocean acidification’ ( [[#Caldeira--2003|Caldeira and Wickett, 2003]] ; [[#Orr--2005|Orr et al., 2005]] ; [[#Doney--2009|Doney et al., 2009]] ). Although the societal concern regarding ocean acidification is relatively recent (about the last 20 years), the physical–chemical basis for the ocean absorption (sink) of atmospheric CO <sub>2</sub> has been discussed much earlier by [[#Revelle--1957|Revelle and Suess (1957)]] . The AR5 and SROCC assessments were of ''robust evidence'' that the H <sup>+</sup> ion concentration is increasing in the surface ocean, thereby reducing seawater pH (= -log [H <sup>+</sup> ]) ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.4.1|Section 2.3.4.1]] ; [[#Orr--2005|Orr et al., 2005]] ; [[#Feely--2009|Feely et al., 2009]] ; [[#Ciais--2013|Ciais et al., 2013]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ), and there is ''high confidence'' that ocean acidification is impacting marine organisms ( [[#Bindoff--2019|Bindoff et al., 2019]] ).

Ocean oxygen decline, or deoxygenation, is driven by changes in ocean ventilation and solubility ( [[#Bindoff--2019|Bindoff et al., 2019]] ). It is ''virtually certain'' that anthropogenic forcing has made a substantial contribution to the ocean heat content increase over the historical period ( [[#Bindoff--2019|Bindoff et al., 2019]] ; [[#IPCC--2019c|IPCC, 2019c]] ; Chapter 9, [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.1|Section 2.3.3.1]] ), strengthening upper water column stratification. Ocean warming decreases the solubility of dissolved oxygen in seawater, and it contributes to about 15% of the dissolved oxygen decrease in the oceans according to estimates based on solubility and the recent SROCC assessment ( ''medium confidence'' ), especially in sub-surface waters, between 100–600 m depth ( [[#Helm--2011|Helm et al., 2011]] ; [[#Schmidtko--2017|Schmidtko et al., 2017]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ; [[#Oschlies--2018|Oschlies et al., 2018]] ; SROCC, [[#5.3.1|Section 5.3.1]] ). Stratification reduces the ventilation flux into the ocean interior, contributing to most of the remaining ocean deoxygenation ( [[#Schmidtko--2017|Schmidtko et al., 2017]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ; [[IPCC:Wg1:Chapter:Chapter-3#3.6.2|Section 3.6.2]] ). Deoxygenation may enhance emissions of nitrous oxide, especially from oxygen minimum zones (OMZs) or hypoxic coastal areas ( [[#Breitburg--2018|Breitburg et al., 2018]] ; [[#Oschlies--2018|Oschlies et al., 2018]] ). Since SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ), CMIP6 model simulation results agree with the reported 2% loss (4.8 ± 2.1 Pmoles O <sub>2</sub> ) in total dissolved oxygen in the upper ocean layer (100–600 m) for the 1970–2010 period ( [[#Helm--2011|Helm et al., 2011]] ; [[#Ito--2017|Ito et al., 2017]] ; [[#Schmidtko--2017|Schmidtko et al., 2017]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ; [[IPCC:Wg1:Chapter:Chapter-2#2.3.4.2|Section 2.3.4.2]] ). The response of marine organisms to the coupled effects of ocean warming, acidification and deoxygenation occur at different metabolic levels on different groups, and include respiratory stress and reduction of thermal tolerance ( [[#Gruber--2011|Gruber, 2011]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ; [[#IPCC--2019c|IPCC, 2019c]] ; [[#Kawahata--2019|Kawahata et al., 2019]] ). An assessment of these effects on marine biota is found in WGII AR6 Chapter 2.

This section assesses past events of ocean acidification and deoxygenation ( [[#5.3.1|Section 5.3.1]] ), the historical trends and spatial variability for the upper ocean ( [[#5.3.2|Section 5.3.2]] ) and the ocean interior ( [[#5.3.3|Section 5.3.3]] ). Future projections for ocean acidification and the drivers in the coastal ocean are assessed in Sections 5.3.4 and 5.3.5, respectively.

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=== 5.3.1 Paleoclimate Context ===

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==== 5.3.1.1 Paleocene–Eocene Thermal Maximum ====

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The Paleocene–Eocene Thermal Maximum (PETM) was an episode of global warming exceeding pre-industrial temperatures by 4°C–8°C ( [[#McInerney--2011|McInerney and Wing, 2011]] ; [[#Dunkley%20Jones--2013|Dunkley]] [[#Jones--2013|]] [[#Jones--2013|Jones et al., 2013]] ) that occurred 55.9–55.7 Ma. The PETM involved a large pulse of geologic CO <sub>2</sub> released into the ocean–atmosphere system in 3–20 kyr ( [[#Zeebe--2016|Zeebe et al., 2016]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Kirtland%20Turner--2017|Kirtland]] [[#Turner--2017|Turner et al., 2017]] ; [[#Kirtland%20Turner--2018|Kirtland Turner, 2018]] ; [[#Gingerich--2019|Gingerich, 2019]] ; [[#5.2.1.1|Section 5.2.1.1]] ). In response, observationally constrained model simulations report an increase in atmospheric CO <sub>2</sub> concentrations ranging from about 900 ppm to &gt;2000 ppm (Chapter 2; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Cui--2018|Cui and Schubert, 2018]] ; [[#Anagnostou--2020|Anagnostou et al., 2020]] ). The PETM thus provides a test for our understanding of the ocean’s response to the increase in carbon (and heat) emissions over geologically short time scales.

A limited number of independent proxy records indicate that the PETM was associated with a surface ocean pH decline ranging from 0.15 to 0.30 units ( [[#Penman--2014|Penman et al., 2014]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Babila--2018|Babila et al., 2018]] ) ''.'' It was also accompanied by a rapid (&lt;10 ka) shallowing of the carbonate saturation horizon, resulting in the widespread dissolution of sedimentary carbonate, followed by a gradual (100 kyr) recovery ( [[#Zachos--2005|Zachos et al., 2005]] ; [[#Bralower--2018|Bralower et al., 2018]] ). The remarkable similarity among sedimentary records spanning a wide range of ecosystems suggests with ''medium confidence'' that the perturbation in the ocean carbonate saturation was global ( [[#Babila--2018|Babila et al., 2018]] ) and directly resulted from elevated atmospheric CO <sub>2</sub> levels. The degree of acidification is similar to the 0.4 pH unit decrease projected for the end of the 21st century under RCP8.5 ( [[#Gattuso--2015|Gattuso et al., 2015]] ) and is estimated to have occurred at a rate about one order of magnitude slower than the current rate of ocean acidification ( [[#Zeebe--2016|Zeebe et al., 2016]] ). There is ''low confidence'' in the inferred rates of ocean acidification inherent to the range of uncertainties affecting rates estimates based on marine sediments ( [[#5.1.2.1|Section 5.1.2.1]] ).

Recent model outputs and globally distributed geochemical data reveal with ''medium confidence'' widespread ocean deoxygenation during the PETM ( [[#Dickson--2012|Dickson et al., 2012]] , 2014; [[#Winguth--2012|Winguth et al., 2012]] ; [[#Chang--2018|Chang et al., 2018]] ; [[#Remmelzwaal--2019|Remmelzwaal et al., 2019]] ), with parts of the ocean potentially becoming drastically oxygen-depleted (anoxic; [[#Yao--2018|Yao et al., 2018]] ; [[#Clarkson--2021|Clarkson et al., 2021]] ). Deoxygenation affected the surface ocean globally (including the Arctic Ocean; [[#Sluijs--2006|Sluijs et al., 2006]] ), due to vertical and lateral expansion of OMZs ( [[#Zhou--2014|Zhou et al., 2014]] ) that resulted from warming and related changes in ocean stratification ''.'' Expansion of OMZs may have stimulated N <sub>2</sub> O production through water-column (de)nitrification ( [[#Junium--2018|Junium et al., 2018]] ). The degree to which N <sub>2</sub> O production impacted PETM warming, however, has not yet been established.

The feedbacks associated with recovery from the PETM are uncertain, yet could include drawdown associated with silicate weathering ( [[#Zachos--2005|Zachos et al., 2005]] ) and regrowth of terrestrial and marine organic carbon stocks ( [[#Bowen--2010|Bowen and Zachos, 2010]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ).

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==== 5.3.1.2 Last Deglacial Transition ====

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The Last Deglacial Transition (LDT) is the best documented climatic transition in the past associated with a substantial atmospheric CO <sub>2</sub> rise ranging from 190 to 265 ppm between 18–11 ka ( [[#Marcott--2014|Marcott et al., 2014]] ). The amplitude of the deglacial CO <sub>2</sub> rise is thus on the same order of magnitude as the increase since the Industrial Revolution.

Boron isotope ( δ <sup>11</sup> B) data suggest a 0.15–0.05 unit decrease in sea surface pH ( [[#Hönisch--2005|Hönisch and Hemming, 2005]] ; [[#Henehan--2013|Henehan et al., 2013]] ) across the LDT, an average rate of decline of about 0.002 units per century compared with the current rate of more than 0.1 units per century ( [[#Bopp--2013|Bopp et al., 2013]] ; [[#Gattuso--2015|Gattuso et al., 2015]] ). Planktonic foraminiferal shell weights decreased by 40% to 50% ( [[#Barker--2002|Barker and Elderfield, 2002]] ), and coccolith mass decreased by about 25% ( [[#Beaufort--2011|Beaufort et al., 2011]] ) across the LDT. Independent proxy reconstructions thus highlight with ''high confidence'' that pH values decreased as atmospheric CO <sub>2</sub> concentrations increased across the LDT. There is, however, ''low confidence'' in the inferred rate of ocean acidification owing to multiple sources of uncertainties affecting rates estimates based on marine sediments ( [[#5.1.2.1|Section 5.1.2.1]] ).

Geochemical and micropaleontological evidence suggest that intermediate-depth OMZs almost vanished during the Last Glacial Maximum (LGM) ( [[#Jaccard--2014|Jaccard et al., 2014]] ). However, multiple lines of evidence suggest with ''medium confidence'' that the deep (&gt;1500 m) ocean became depleted in O <sub>2</sub> (concentrations were possibly lower than 50 μmol kg <sup>–1</sup> ) globally (Jaccard and Galbraith, 2012; [[#Hoogakker--2015|Hoogakker et al., 2015]] , 2018; [[#Gottschalk--2016|Gottschalk et al., 2016]] , 2020a; [[#Anderson--2019|Anderson et al., 2019]] ) as a combined result of sluggish ventilation of the ocean subsurface ( [[#Gottschalk--2016|Gottschalk et al., 2016]] , 2020a; [[#Skinner--2017|Skinner et al., 2017]] ) and a generally more efficient marine biological carbon pump ( [[#Buchanan--2016|Buchanan et al., 2016]] ; [[#Yamamoto--2019|Yamamoto et al., 2019]] ; [[#Galbraith--2020|Galbraith and Skinner, 2020]] ).

During the LDT, deep ocean ventilation increased as Antarctic Bottom Water (AABW) ( [[#Skinner--2010|Skinner et al., 2010]] ; [[#Gottschalk--2016|Gottschalk et al., 2016]] ; [[#Jaccard--2016|Jaccard et al., 2016]] ) and subsequently the Atlantic meridional overturning circulation ( [[#McManus--2004|McManus et al., 2004]] ; [[#Lippold--2016|Lippold et al., 2016]] ) resumed, transferring previously sequestered remineralized carbon from the ocean interior to the upper ocean, and eventually the atmosphere ( [[#Skinner--2010|Skinner et al., 2010]] ; [[#Galbraith--2015|Galbraith and Jaccard, 2015]] ; [[#Gottschalk--2016|Gottschalk et al., 2016]] ; [[#Ronge--2016|Ronge et al., 2016]] , 2020; [[#Sikes--2016|Sikes et al., 2016]] ; [[#Rae--2018|Rae et al., 2018]] ), contributing to the deglacial CO <sub>2</sub> rise. Intermediate depths lost oxygen as a result of sluggish ventilation and increasing temperatures (decreasing saturation). As the world emerged from the last Glacial period, OMZs underwent a large volumetric increase at the beginning of the Bølling-Allerød (B/A), a northern-hemisphere wide warming event, 14.7 ka ( [[#Jaccard--2012|Jaccard and Galbraith, 2012]] ; [[#Praetorius--2015|Praetorius et al., 2015]] ) with deleterious consequences for benthic ecosystems (e.g., [[#Moffitt--2015|Moffitt et al., 2015]] ). These observations indicate with ''high confidence'' that the rate of warming, affecting the solubility of oxygen and upper water column stratification, coupled with changes in subsurface ocean ventilation, impose a direct control on the degree of ocean deoxygenation, implying a high sensitivity of ocean oxygen loss to warming. The expansion of OMZs contributed to a widespread increase in water column (de)nitrification (Galbraith and Kienast, 2013), which contributed substantially to enhanced marine N <sub>2</sub> O emissions. Nitrogen stable isotope measurements on N <sub>2</sub> O extracted from ice cores suggest that approximately one-third (of the order of 0.7 ± 0.3TgN yr <sup>–1</sup> ) of thedeglacial increase in N <sub>2</sub> O emissions relates to oceanic sources (Schilt et al., 2014; H. [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ).

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=== 5.3.2 Historical Trends and Spatial Characteristics in the Upper Ocean ===

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==== 5.3.2.1 Reconstructed Centennial Ocean Acidification Trends ====

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Ocean pH time series are based on the reconstruction of coral boron isotope ratios ( δ <sup>11</sup> B). A majority of coral δ <sup>11</sup> B data have been generated from the western Pacific region with a few records from the Atlantic Ocean. Biweekly resolution paleo-pH records show monsoonal variation of about 0.5 pH unit in the South China Sea ( [[#Liu--2014|Liu et al., 2014]] ). Interannual ocean pH variability in the range of 0.07–0.16 pH unit characterizes southwest Pacific corals that are attributed to El Niño–Southern Oscillation (ENSO) ( [[#Wu--2018|]] [[#Wu--2018|H.C. Wu et al., 2018]] ) and river runoff ( [[#D’Olivo--2015|D’Olivo et al., 2015]] ). Decadal (10-, 22- and 48-year) ocean pH variations in the south-west Pacific have been linked to the Inter-decadal Pacific Oscillation, causing variations of up to 0.30 pH unit in the Great Barrier Reef ( [[#Pelejero--2005|Pelejero et al., 2005]] ; [[#Wei--2009|Wei et al., 2009]] ) but weaker (about 0.08 pH unit) in the open ocean ( [[#Wu--2018|]] [[#Wu--2018|H.C. Wu et al., 2018]] ). Decadal variations in the South China Sea pH changes of 0.10–0.20 have also been associated with the variation in the East Asian monsoon ( [[#Liu--2014|Liu et al., 2014]] ; [[#Wei--2015|Wei et al., 2015]] ), as a weakening of the Asian winter monsoon leads to sluggish water circulation within the reefs, building up localised CO <sub>2</sub> concentration in the water due to calcification and respiration.

Since the beginning of the industrial period in the mid-19th century, coral δ <sup>11</sup> B-derived ocean pH has decreased by 0.06–0.24 pH unit in the South China Sea ( [[#Liu--2014|Liu et al., 2014]] ; [[#Wei--2015|Wei et al., 2015]] ) and 0.12 pH unit in the south-west Pacific ( [[#Wu--2018|]] [[#Wu--2018|H.C. Wu et al., 2018]] ). Since the mid-20th century, a distinct feature of coral δ <sup>11</sup> B records relates to ocean acidification trends, albeit having a wide range of values: 0.12–0.40 pH unit in the Great Barrier Reef ( [[#Wei--2009|Wei et al., 2009]] ; [[#D’Olivo--2015|D’Olivo et al., 2015]] ), 0.05–0.08 pH unit in the north-west Pacific ( [[#Shinjo--2013|Shinjo et al., 2013]] ) and 0.04–0.09 pH unit in the Atlantic Ocean ( [[#Goodkin--2015|Goodkin et al., 2015]] ; [[#Fowell--2018|Fowell et al., 2018]] ). Concurrent coral carbon isotopic ( δ <sup>13</sup> C) measurements infer ocean uptake of anthropogenic CO <sub>2</sub> from the combustion of fossil fuel, based on the lower abundance of <sup>13</sup> C in fossil fuel carbon. Western Pacific coral records show depleted δ <sup>13</sup> C trends since the late 19th century that are more prominent since the mid-20th century ( ''high confidence'' ) (Pelejeroet al., 2005; [[#Wei--2009|Wei et al., 2009]] ; [[#Shinjo--2013|Shinjo et al., 2013]] ; [[#Liu--2014|Liu et al., 2014]] ; [[#Kubota--2017|Kubota et al., 2017]] ; [[#Wu--2018|]] [[#Wu--2018|H.C. Wu et al., 2018]] ).

Overall, many of the records show a highly variable seawater pH underlaying strong imprints of internal climate variability ( ''high confidence'' ) and, in most instances, superimposed on a decreasing δ <sup>11</sup> B trend that is indicative of anthropogenic ocean acidification in recent decades ( ''medium confidence'' ). The robustness of seawater pH reconstructions is currently limited by the uncertainty on the calibration of The δ <sup>11</sup> B proxy in different tropical coral species.

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==== 5.3.2.2 Observations of Ocean Acidification over Recent Decades ====

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The SROCC ( [[#5.2.2.3|Section 5.2.2.3]] ) indicated that it is ''virtually certain'' that the ocean has undergone acidification globally in response to ocean CO <sub>2</sub> uptake, and concluded that pH in open ocean surface water has changed by a ''virtually certain'' range of –0.017 to –0.027 pH units per decade since the late 1980s. Since SROCC, evidence of the progress of acidification across all regions of the oceans has been further strengthened by continued observations of seawater carbonate chemistry at ocean time series stations, and compiled shipboard studies providing temporally resolved and methodologically consistent datasets ( [[#Jiang--2019|Jiang et al., 2019]] ) (Figure 5.20; Supplementary Material Table 5.SM.3; [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.5|Section 2.3.3.5]] ).

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[[File:f8c1048a44db02ea97346655a5cec7e7 IPCC_AR6_WGI_Figure_5_20.png]]

'''Figure 5.20 |''' '''Multi-decadal trends of pH (Total Scale) in surface layer at various sites of the oceans and a global distribution of annual mean pH adjusted to the year 2000''' . Time-series data of pH are from [[#Dore--2009|Dore et al. (2009)]] , [[#Olafsson--2009|Olafsson et al. (2009)]] , [[#González-Dávila--2010|González-Dávila et al. (2010)]] , [[#Bates--2014|Bates et al. (2014)]] , [[#Takahashi--2014|Takahashi et al. (2014)]] , [[#Wakita--2017|Wakita et al. (2017)]] , [[#Merlivat--2018|Merlivat et al. (2018)]] , [[#Ono--2019|Ono et al. (2019)]] , and [[#Bates--2020|Bates and Johnson (2020)]] . Global distribution of annual mean pH have been evaluated from data of surface ocean ''p'' CO <sub>2</sub> <sup></sup> measurements ( [[#Bakker--2016|Bakker et al., 2016]] ; [[#Jiang--2019|Jiang et al., 2019]] ). Acronyms in panels: KNOT and K2 – Western Pacific subarctic gyre time series; HOT – Hawaii Ocean Time-series; BATS – Bermuda Atlantic Time-series Study; DYFAMED – Dynamics of Atmospheric Fluxes in the Mediterranean Sea; ESTOC – European Station for Time-series in the Ocean Canary Islands; CARIACO – Carbon Retention in a Colored Ocean Time-series. Further details on data sources and processing are available in the chapter data table (Table 5.SM.6).

In the subtropical open oceans, decreases in pH have been reported with a ''very'' ''likely'' rate range from –0.016 to –0.019 pH units per decade since 1980s, which equates to approximately 4 % increase in hydrogen ion concentration ([H <sup>+</sup> ]) per decade. Accordingly, the saturation state Ω (=[Ca <sup>2+</sup> ][CO <sub>3</sub> <sup>2-</sup> ]/ ''K'' <sub>sp</sub> ) of seawater with respect to calcium carbonate mineral aragonite has been declining at rates ranging from –0.07 to –0.12 per decade ( [[#González-Dávila--2010|González-Dávila et al., 2010]] ; [[#Feely--2012|Feely et al., 2012]] ; [[#Bates--2014|Bates et al., 2014]] ; [[#Takahashi--2014|Takahashi et al., 2014]] ; [[#Ono--2019|Ono et al., 2019]] ; [[#Bates--2020|Bates and Johnson, 2020]] ; Supplementary Material Table 5.SM.3). These rates are consistent with the rates expected from the transient equilibration with increasing atmospheric CO <sub>2</sub> concentrations, but the variability of rate in decadal time scale has also been detected with ''robust evidence'' (Ono et al., 2019; [[#Bates--2020|Bates and Johnson, 2020]] ). In the tropical Pacific, its central and eastern upwelling zones exhibited a faster pH decline of –0.022 to –0.026 pH unit per decade due to increased upwelling of CO <sub>2</sub> -rich sub-surface waters in addition to anthropogenic CO <sub>2</sub> uptake ( [[#Sutton--2014|Sutton et al., 2014]] ; [[#Lauvset--2015|Lauvset et al., 2015]] ). By contrast, warm pools in the western tropical Pacific exhibited slower pH decline of –0.010 to –0.013 pH unit per decade (Supplementary Material Table 5.SM.3; Lauvsetet al., 2015; [[#Ishii--2020|Ishii et al., 2020]] ). Observational and modelling studies ( [[#Nakano--2015|Nakano et al., 2015]] ; [[#Ishii--2020|Ishii et al., 2020]] ) consistently suggest that slower acidification in this region is attributable to the anthropogenic CO <sub>2</sub> taken up in the extratropics around a decade ago and transported to the tropics via shallow meridional overturning circulations.

In open subpolar and polar zones, the ''very likely'' range (–0.003 to –0.026 pH unit per decade) and uncertainty (up to 0.010) observed in pH decline are larger than in the subtropics, reflecting the complex interplay between physical and biological forcing mechanisms ( [[#Olafsson--2009|Olafsson et al., 2009]] ; [[#Midorikawa--2012|Midorikawa et al., 2012]] ; [[#Bates--2014|Bates et al., 2014]] ; [[#Takahashi--2014|Takahashi et al., 2014]] ; [[#Lauvset--2015|Lauvset et al., 2015]] ; [[#Wakita--2017|Wakita et al., 2017]] ; [[#Merlivat--2018|Merlivat et al., 2018]] ). Nevertheless, the ''high agreement'' of pH decline among these available time-series studies leads to ''high confidence'' in the trend of acidification in these zones. In the Arctic Ocean, a temporally limited time series of carbonate chemistry measurements prevents drawing robust conclusions on ocean acidification trends. However, the carbonate saturation state (Ω) is generally low, and observational studies show with ''robust evidence'' that the recent extensive melting of sea ice leading to enhanced air–sea CO <sub>2</sub> exchange, large freshwater inputs, together with river discharge and glacial drainage, as well as the degradation of terrestrial organic matter in seawater, result in the decline of Ω of aragonite to undersaturation ( [[#Bates--2009|Bates et al., 2009]] ; [[#Chierici--2009|Chierici and Fransson, 2009]] ; [[#Yamamoto-Kawai--2009|Yamamoto-Kawai et al., 2009]] ; [[#Azetsu-Scott--2010|Azetsu-Scott et al., 2010]] ; [[#Robbins--2013|Robbins et al., 2013]] ; [[#Fransson--2015|Fransson et al., 2015]] ; [[#Semiletov--2016|Semiletov et al., 2016]] ; [[#Anderson--2017|Anderson et al., 2017]] ; [[#Qi--2017|Qi et al., 2017]] ; [[#Beaupré-Laperrière--2020|Beaupré-Laperrière et al., 2020]] ; Y. [[#Zhang--2020|]] [[#Zhang--2020|]] [[#Zhang--2020|Zhang et al., 2020]] ; SROCC [[IPCC:Wg1:Chapter:Chapter-3#3.2.1.2.4|Section 3.2.1.2.4]] , [[#IPCC--2019b|IPCC, 2019b]] ). The low saturation state of aragonite (Ω about 1) has also been observed in surface waters of the Antarctic coastal zone associated with freshwater input from glaciers ( [[#Mattsdotter%20Björk--2014|Mattsdotter Björk et al., 2014]] ) and upwelling of deep water ( [[#Hauri--2015|Hauri et al., 2015]] ) as well as along eastern boundary upwelling systems ( [[#Feely--2016|Feely et al., 2016]] ).

Overall, in agreement with SROCC, it is ''virtually certain'' from these observational studies that ocean surface waters undergo acidification globally with the CO <sub>2</sub> increase in the atmosphere. These sustained measurements over the past decades, and campaign studies of ocean carbonate chemistry, also highlight with ''robust evidence'' that trends of acidification have been modulated by the variability and changes in physical and chemical states of ocean, including those affected by the warming of the cryosphere, and need to be better understood.

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=== 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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<span id="ocean-deoxygenation-and-its-implications-for-greenhouse-gases"></span>
==== 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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<span id="future-projections-for-ocean-acidification"></span>
=== 5.3.4 Future Projections for Ocean Acidification ===

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<div id="5.3.4.1" class="h3-container"></div>

<span id="future-projections-with-earth-system-models-esms"></span>
==== 5.3.4.1 Future Projections with Earth System Models (ESMs) ====

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Projections with CMIP5 ESMs, reported in AR5 (Section 6.4.4) and SROCC ( [[#5.2.2.3|Section 5.2.2.3]] ; [[#IPCC--2019b|IPCC, 2019b]] ), showed changes in global mean surface ocean pH from 1870–1899 to 2080–2099 of –0.14 ± 0.001 (inter-model standard deviation) under RCP2.6 and –0.38 ± 0.005 under RCP8.5 with pronounced regional variability ( [[#Bopp--2013|Bopp et al., 2013]] ; [[#Hurd--2018|Hurd et al., 2018]] ). They also projected faster pH declines in mode waters below seasonal mixed layers ( [[#Resplandy--2013|Resplandy et al., 2013]] ; [[#Watanabe--2017|Watanabe and Kawamiya, 2017]] ) as has been observed in the Atlantic ( [[#Salt--2015|Salt et al., 2015]] ) and in the Pacific ( [[#Carter--2019|Carter et al., 2019]] ), because of the net CO <sub>2</sub> release by respiration and lowering CO <sub>2</sub> buffering capacity of seawater. In these CO <sub>2</sub> concentration-driven simulations, the level of acidification in the surface ocean is primarily determined by atmospheric CO <sub>2</sub> concentration and regional seawater carbonate chemistry, thereby providing consistent projections across models. New projections with CMIP6 ESMs show greater surface pH decline of –0.16 ± 0.002 under the SSP1-2.6 and –0.44 ± 0.005 under SSP5-8.5 from 1870–1899 to 2080–2099 ( [[IPCC:Wg1:Chapter:Chapter-4#4.3.2.5|Section 4.3.2.5]] and Cross-Chapter Box 5.3; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). The greater pH declines in CMIP6 are primarily a consequence of higher atmospheric CO <sub>2</sub> concentrations in SSPs than their CMIP5-RCP analogues ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Ocean acidification is also projected to occur with ''high confidence'' in the Abyssal Bottom Waters in regions such as the northern North Atlantic and the Southern Ocean ( [[#Sulpis--2019|Sulpis et al., 2019]] ), with the rates of global mean pH decline of –0.018 ± 0.001 under SSP1-2.6 and –0.030 ± 0.002 under SSP5-8.5 from 1870–1899 to 2080–2099 in CMIP6 ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ).

In surface ocean, changes in the amplitude of seasonal variations in pH are also projected to occur with ''high confidence.'' ESMs in CMIP6 show +73 ± 12% increase in the amplitude of seasonal variation in hydrogen ion concentration ([H <sup>+</sup> ]) but 10 ± 5% decrease in the seasonal variation in pH (= -log [H <sup>+</sup> ]) from 1995–2014 to 2080–2099 under SSP5-8.5. The simultaneous amplification of [H <sup>+</sup> ] and attenuation of pH seasonal cycles is counterintuitive but is the consequence of a greater increase in the annual mean [H <sup>+</sup> ] due to anthropogenic CO <sub>2</sub> invasion than the corresponding increase in its seasonal amplitude. These changes are consistent with the amplification/attenuation of the seasonal variation of +81 ±16% for [H <sup>+</sup> ] and –16 ± 7% for pH from 1990–1999 to 2090–2099 under RCP8.5 in CMIP5 ( [[#Kwiatkowski--2018|Kwiatkowski and Orr, 2018]] ).

The signal of ocean acidification in surface ocean is large and is projected to emerge beyond the range of natural variability within the time scale of a decade in all ocean basins ( [[#Schlunegger--2019|Schlunegger et al., 2019]] ). There is ''high agreement'' among modelling studies that the largest pH decline and large-scale undersaturation of aragonite in surface seawater start to occur first in polar oceans ( [[#Orr--2005|Orr et al., 2005]] ; [[#Steinacher--2009|Steinacher et al., 2009]] ; [[#Hurd--2018|Hurd et al., 2018]] ; [[#Jiang--2019|Jiang et al., 2019]] ). Under SSP5-8.5, the largest surface pH decline, exceeding 0.45 between 1995–2014 and 2080–2099, occurs in the Arctic Ocean ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). The freshwater input from sea ice melt is an additional factor leading to a faster decline of aragonite saturation level than expected from the anthropogenic CO <sub>2</sub> uptake ( [[#Yamamoto--2012|Yamamoto et al., 2012]] ). The increase in riverine and glacial discharges that provide terrigenous carbon, nutrients and alkalinity as well as freshwater are the other factors modifying the rate of acidification in the Arctic Ocean. However, their impacts have been projected in a limited number of studies with extensive knowledge gaps and model simplifications leading to ''low confidence'' in their impacts ( [[#Terhaar--2019|Terhaar et al., 2019]] ; [[#Hopwood--2020|Hopwood et al., 2020]] ). In the Southern Ocean, the aragonite undersaturation starts in the 2030s in RCP8.5, and the area that experiences aragonite undersaturation for at least one month per year by 2100 is projected to be more than 95%. Under RCP2.6, short periods (less than one month) of aragonite undersaturation are expected to be found in less than 2% of the area during this century ( [[#Sasse--2015|Sasse et al., 2015]] ; [[#Hauri--2016|Hauri et al., 2016]] ; [[#Negrete-García--2019|Negrete-García et al., 2019]] ). These long term projections are modified at interannual time scales by large-scale climate modes ( [[#Ríos--2015|Ríos et al., 2015]] ) such as the ENSO and the Southern Annular Mode ( [[#Conrad--2015|Conrad and Lovenduski, 2015]] ). In other regions, acidification trends are influenced by a range of processes such as changes in ocean circulation, temperature, salinity, carbon cycling, and the structure of the marine ecosystem. As, at present, models do not resolve fine-scale variability of these processes, current projections do not fully capture the changes that the marine environment will experience in the future ( [[#Takeshita--2015|Takeshita et al., 2015]] ; [[#Turi--2016|Turi et al., 2016]] ).

Overall, with the rise of atmospheric CO <sub>2</sub> , the physics of CO <sub>2</sub> transfer across the air–sea interface, the carbonate chemistry in seawater, the trends of ocean acidification being observed in the past decades ( [[#5.3.3.2|Section 5.3.3.2]] ) and modelling studies described in this section, it is ''virtually certain'' that ocean acidification will continue to grow. However, the magnitude and sign (direction) of many of ocean carbon–climate feedbacks are still poorly constrained (Matear and Lenton, 2014, 2018), leading to ''low confidence'' in their significant and long-lasting impacts on ocean acidification.

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<span id="reversal-of-ocean-acidification-by-carbon-dioxide-removal"></span>
==== 5.3.4.2 Reversal of Ocean Acidification by Carbon Dioxide Removal ====

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Reversing the increase in atmospheric CO <sub>2</sub> concentrations through negative emissions ( [[#5.6|Section 5.6]] ) will reverse ocean acidification at the sea surface ( ''high confidence'' ) but will not result in rapid amelioration of ocean acidification in the deeper ocean ( [[#5.3.3.2|Section 5.3.3.2]] ). The ocean’s uptake of atmospheric CO <sub>2</sub> will start to decrease as atmospheric CO <sub>2</sub> decreases (Sections 5.4.5, 5.4.10 and 5.6.2.1; [[#Mathesius--2015|Mathesius et al., 2015]] ; [[#Tokarska--2015|Tokarska and Zickfeld, 2015]] ). However, because of the long time scales of the ocean turnover that transfers CO <sub>2</sub> from the upper to the deep ocean, excess carbon will continue to accumulate in the deep ocean even after a decrease in atmospheric CO <sub>2</sub> ( [[#Cao--2014|Cao et al., 2014]] ; [[#Mathesius--2015|Mathesius et al., 2015]] ; [[#Tokarska--2015|Tokarska and Zickfeld, 2015]] ; T. [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ). There is thus ''high confidence'' that CO <sub>2</sub> emissions leave a long-term legacy in ocean acidification, and are therefore irreversible at multi-human generational scales, even with aggressive atmospheric CO <sub>2</sub> removal.

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<span id="coastal-ocean-acidification-and-deoxygenation"></span>
=== 5.3.5 Coastal Ocean Acidification and Deoxygenation ===

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The coastal ocean, from the shoreline to the isobath of 200 m, is highly heterogeneous due to the complex interplay between physical, biogeochemical and anthropogenic factors ( [[#Gattuso--1998|Gattuso et al., 1998]] ; [[#Chen--2009|Chen and Borges, 2009]] ; [[#Dürr--2011|Dürr et al., 2011]] ; [[#Laruelle--2014|Laruelle et al., 2014]] ; [[#McCormack--2016|McCormack et al., 2016]] ). These areas, according to SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) are, with ''high confidence'' , already affected by ocean acidification and deoxygenation. This section assesses the drivers and spatial variability of acidification and deoxygenation based on new observations and data products.

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<span id="drivers"></span>
==== 5.3.5.1 Drivers ====

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Observations and data products including models ( [[#Astor--2013|Astor et al., 2013]] ; [[#Bakker--2016|Bakker et al., 2016]] ; [[#Kosugi--2016|Kosugi et al., 2016]] ; [[#Vargas--2016|Vargas et al., 2016]] ; [[#Laruelle--2017|Laruelle et al., 2017]] , 2018; [[#Orselli--2018|Orselli et al., 2018]] ; [[#Roobaert--2019|Roobaert et al., 2019]] ; [[#Cai--2020|Cai et al., 2020]] ; H. [[#Sun--2020|]] [[#Sun--2020|Sun et al., 2020]] ) confirm the strong spatial and temporal variability in the coastal ocean surface carbonate chemistry and sea-air CO <sub>2</sub> fluxes ( ''high agreement'' , ''robust evidence'' ). The anthropogenic CO <sub>2</sub> -induced acidification is either mitigated or enhanced through biological processes; primary production removes dissolved CO <sub>2</sub> from the surface, and respiration adds CO <sub>2</sub> and consumes oxygen in the subsurface layers. The relative intensity of these processes is controlled by natural or anthropogenic eutrophication. Other drivers of variability include biological community composition, freshwater input from rivers or melting ice, sea ice cover and calcium carbonate precipitation/dissolution dynamics, coastal upwelling and regional circulation, and seasonal surface cooling ( [[#Fransson--2015|Fransson et al., 2015]] , 2017; [[#Feely--2018|Feely et al., 2018]] ; [[#Roobaert--2019|Roobaert et al., 2019]] ; [[#Cai--2020|Cai et al., 2020]] ; [[#Hauri--2020|Hauri et al., 2020]] ; [[#Monteiro--2020b|Monteiro et al., 2020b]] ; H. [[#Sun--2020|]] [[#Sun--2020|Sun et al., 2020]] ). Near-shore surface waters are often supersaturated with CO <sub>2</sub> , regardless of the latitude, especially in highly populated areas receiving substantial amounts of domestic and industrial sewage ( [[#Chen--2009|Chen and Borges, 2009]] ). Nevertheless, thermal or haline-stratified eutrophic coastal areas may act as net atmospheric CO <sub>2</sub> sinks ( [[#Chou--2013|Chou et al., 2013]] ; [[#Cotovicz%20Jr.--2015|Cotovicz Jr. et al., 2015]] ). Continental shelves, excluding near-shore areas, act as CO <sub>2</sub> sinks at a rate of 0.2 ± 0.02 PgC yr <sup>–1</sup> ( [[#Laruelle--2014|Laruelle et al., 2014]] ; [[#Roobaert--2019|Roobaert et al., 2019]] ), considering ice-free areas only. Under increasing atmospheric CO <sub>2</sub> and eutrophication, such ecosystems would be more vulnerable to ecological and seawater chemistry changes, impacting the local economy.

Since AR5, ( [[#Ciais--2013|Ciais et al., 2013]] ) and in agreement with SROCC ( [[#IPCC--2019b|IPCC, 2019b]] ), there is now ''high agreement (robust evidence)'' that coastal ocean acidification, whether induced only by increasing atmospheric CO <sub>2</sub> or exacerbated by eutrophication or upwelling, has negative effects on specific groups of marine organisms such as reef-building corals, crabs, pteropods, and sessile fauna (AR6 WGII, Chapter 3; [[#Dupont--2010|Dupont et al., 2010]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ; [[#Bednaršek--2020|Bednaršek et al., 2020]] ; [[#Osborne--2020|Osborne et al., 2020]] ), especially when combined with stressors such as temperature and deoxygenation, and potentially increased bioavailability of toxic elements such as arsenic and copper ( [[#Millero--2009|Millero et al., 2009]] ; [[#Boyd--2015|Boyd et al., 2015]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ).

Since SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ), there is further evidence that anthropogenic eutrophication via continental runoff and atmospheric nutrient deposition, and ocean warming are ''very likely'' the main drivers of deoxygenation in coastal areas ( [[#Levin--2015|Levin and Breitburg, 2015]] ; [[#Levin--2015|Levin et al., 2015]] ; [[#Royer--2016|Royer et al., 2016]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ; [[#Cocquempot--2019|Cocquempot et al., 2019]] ; [[#Fagundes--2020|Fagundes et al., 2020]] ; [[#Limburg--2020|Limburg et al., 2020]] ). Increasing intensity and frequency of wind-driven upwelling is responsible for longer and more intense coastal hypoxia, fuelled by organic matter degradation from primary production ''(medium to high agreement, medium evidence)'' ( [[#Rabalais--2010|Rabalais et al., 2010]] ; [[#Bakun--2015|Bakun et al., 2015]] ; [[#Varela--2015|Varela et al., 2015]] ; [[#Fennel--2019|Fennel and Testa, 2019]] ; [[#Limburg--2020|Limburg et al., 2020]] ). Locally, submarine groundwater discharge may enhance the eutrophication state ( ''low agreement'' , ''limited evidence'' , [[#Luijendijk--2020|Luijendijk et al., 2020]] ). Since AR5 ( [[#Ciais--2013|Ciais et al., 2013]] ) and SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) new observations and model studies confirm the trends in increasing coastal hypoxia caused by eutrophication, ocean warming and changes in circulation ( [[#Claret--2018|Claret et al., 2018]] ; [[#Dussin--2019|Dussin et al., 2019]] ; [[#Limburg--2020|Limburg et al., 2020]] ), as well as the ubiquitous impacts on marine organisms and fisheries (AR6 WGII Chapter 3; [[#Carstensen--2019|Carstensen and Conley, 2019]] ; [[#Fennel--2019|Fennel and Testa, 2019]] ; [[#Osma--2020|Osma et al., 2020]] ). Following open ocean deoxygenation trends since the 1950s, more than 700 coastal regions are being reported as hypoxic (dissolved oxygen concentration &lt;2 mg O <sub>2</sub> L <sup>–1</sup> ) ( [[#Limburg--2020|Limburg et al., 2020]] ). Additionally, deoxygenation or increasing severe hypoxic periods in coastal areas may enhance the sea-to-air fluxes of N <sub>2</sub> O and CH <sub>4</sub> especially through microbial-mediated processes in the water column–sediment interface ( ''medium agreement'' ) ( [[#Middelburg--2009|Middelburg and Levin, 2009]] ; [[#Naqvi--2010|Naqvi et al., 2010]] ; [[#Farías--2015|Farías et al., 2015]] ; [[#Limburg--2020|Limburg et al., 2020]] ).

<div id="5.3.5.2" class="h3-container"></div>

<span id="spatial-characteristics"></span>
==== 5.3.5.2 Spatial Characteristics ====

<div id="h3-28-siblings" class="h3-siblings"></div>

There is ''high agreement'' ( ''robust evidence'' ) that heterogeneity implies different responses of coastal regions to increasing atmospheric CO <sub>2</sub> , decreasing seawater pH and calcium carbonate saturation state, and deoxygenation ( [[#Duarte--2013|Duarte et al., 2013]] ; [[#Regnier--2013|Regnier et al., 2013]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ; [[#Laruelle--2018|Laruelle et al., 2018]] ; [[#Carstensen--2019|Carstensen and Duarte, 2019]] ).

There is ''high agreement'' that long-time series of observations utilizing standard methods are needed to distinguish the climate change signal in seawater carbonate chemistry from the natural variability of coastal sites ( [[#Duarte--2013|Duarte et al., 2013]] ; [[#Salisbury--2018|Salisbury and Jönsson, 2018]] ; [[#IOC--2019|IOC, 2019]] ; [[#Sutton--2019|Sutton et al., 2019]] ; [[#Tilbrook--2019|Tilbrook et al., 2019]] ; [[#Turk--2019|Turk et al., 2019]] ). Despite the increasing availability of data and sea–air CO <sub>2</sub> flux budgets for the coastal ocean (Sections 5.3.5.1 and 5.2.3.1), additional long-term observations are required to constrain the global time of emergence of coastal acidification. There is ''high agreement'' ( ''medium evidence'' ) that, for the coastal subtropical to temperate north-east Pacific and north-west Atlantic, the mean time of emergence for acidification is above two decades ( [[#Sutton--2019|Sutton et al., 2019]] ; [[#Turk--2019|Turk et al., 2019]] ).

Observations and models predict an expansion and intensification of low-pH deep water intrusions for the north-east Pacific coastal upwelling area ( ''high agreement'' , ''robust evidence'' ) ( [[#Hauri--2013|Hauri et al., 2013]] ; [[#Feely--2016|Feely et al., 2016]] ; [[#Cai--2020|Cai et al., 2020]] ). Areas such as the California Current System are naturally exposed to intrusions of low‐pH, high pCO <sub>2sea</sub> deep waters from remineralization processes and anthropogenic CO <sub>2</sub> intrusion ( [[#Feely--2008|Feely et al., 2008]] , 2010, 2018; [[#Chan--2019|Chan et al., 2019]] ; [[#Lilly--2019|Lilly et al., 2019]] ; [[#Cai--2020|Cai et al., 2020]] ).The eastern Pacific coastal upwelling displays seasonality in subsurface aragonite undersaturation as a consequence of the interplay between anthropogenic CO <sub>2</sub> , respiration and intrusion of upwelling waters ( [[#Feely--2008|Feely et al., 2008]] , 2010, 2016, 2018; [[#Hauri--2013|Hauri et al., 2013]] ; [[#Vargas--2016|Vargas et al., 2016]] ; [[#Chan--2019|Chan et al., 2019]] ; [[#Lilly--2019|Lilly et al., 2019]] ). The coastal south-east Pacific upwelling combined with low-pH, low-alkalinity, organic matter-rich river inputs display extreme temporal variability in surface seawater ''p'' CO <sub>2</sub> and low aragonite saturation ( [[#Vargas--2016|Vargas et al., 2016]] ; [[#Osma--2020|Osma et al., 2020]] ).

Temperate, non-upwelling coastal areas along the north-west Atlantic display a trend of decreasing seawater pH, mainly attributed to the combined effects of eutrophication and decreasing seawater buffering capacity ( ''high agreement'' , ''robust evidence'' ). Observations show an increasing north to south gradient of aragonite saturation state ( [[#Sutton--2016|Sutton et al., 2016]] ; [[#Fennel--2019|Fennel et al., 2019]] ; [[#Cai--2020|Cai et al., 2020]] ). Low alkalinity and total inorganic carbon concentration, combined with an ocean signal of acidification, diminishes the buffering capacity along the decreasing salinity gradient from the ocean to the coast. Models suggest that, in this area, the aragonite saturation is seasonally controlled by nutrient availability and primary production, supporting the finding that eutrophication is the main driver for exacerbating acidification ( [[#Cai--2017|Cai et al., 2017]] , 2020). The coastal Gulf of Mexico is facing a parallel increase in bottom water acidification and deoxygenation off the Mississippi Delta driven by eutrophication ( [[#Cai--2011|Cai et al., 2011]] ; [[#Laurent--2017|Laurent et al., 2017]] ; [[#Fennel--2019|Fennel et al., 2019]] ).

Many coastal tropical areas are under heavy anthropogenic eutrophication induced by the effluents from large cities, or receive large riverine inputs of freshwater, nutrients, and organic matter (such as Amazon, Mississippi, Orinoco, Congo, Mekong, or Changjiang rivers). Under strong eutrophication, often sub-surface and bottom waters present pH values lower than average surface open ocean (about 8.0) because increased respiration decreases pH ( ''high agreement'' , ''robust evidence'' ), despite a net atmospheric CO <sub>2</sub> sink in shallow and vertically stratified coastal areas ( [[#Koné--2009|Koné et al., 2009]] ; [[#Wallace--2014|Wallace et al., 2014]] ; [[#Cotovicz%20Jr.--2015|Cotovicz Jr. et al., 2015]] , 2018; [[#Fennel--2019|Fennel and Testa, 2019]] ; [[#Lowe--2019|Lowe et al., 2019]] ; [[#5.3.5.1|Section 5.3.5.1]] ).

There is ''medium evidence'' from observations and models that the coastal north-western Antarctic Peninsula (Southern Ocean) will experience calcium carbonate undersaturation by 2060, considering that anthropogenic emissions reach an atmospheric CO <sub>2</sub> concentration of about 500 pm at that date ( [[#Lencina-Avila--2018|Lencina-Avila et al., 2018]] ; [[#Monteiro--2020a|Monteiro et al., 2020a]] ). The synergies among warming, melt water, sea-air CO <sub>2</sub> equilibrium and circulation may, to some extent, offset the coastal ocean acidification trends in Antarctica ( [[#Henley--2020|Henley et al., 2020]] ). In the coastal western Arctic Ocean, there is increasing ''robust evidence'' that ocean acidification is driven by sea-air CO <sub>2</sub> fluxes and sea-ice melt, and increasing intrusions since the 1990s of low-alkalinity Pacific water, lowering aragonite saturation state ( [[#Qi--2017|Qi et al., 2017]] , 2020; [[#Cross--2018|Cross et al., 2018]] ). The Bering Sea (north-eastern Pacific) shows decreasing trends in calcium carbonate saturation, associated to the increasing atmospheric CO <sub>2</sub> uptake combined with riverine freshwater and carbon inputs ( ''high agreement'' , ''robust evidence'' ) ( [[#Pilcher--2019|Pilcher et al., 2019]] ; H. [[#Sun--2020|]] [[#Sun--2020|Sun et al., 2020]] ).

The spatial distribution of hypoxic areas is highly heterogeneous in the coastal ocean, and there is ''high agreement, robust evidence'' that more severe hypoxia or anoxia is often associated with highly populated coastal areas,or local circulation and upwelling, and seasonal stratification leading to an accumulation of organic matter in subsurface waters (Ciais et al., 2013; [[#Rabalais--2014|Rabalais et al., 2014]] ; M. [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|Li et al., 2016]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ; SROCC Chapter 5). The causes and trends of coastal deoxygenation can only be assessed by making available long-term time series combined with regional modelling ( [[#Fennel--2019|Fennel and Testa, 2019]] ), as in the California current system ( [[#Wang--2017|Wang et al., 2017]] ), the East China Sea ( [[#Chen--2007|Chen et al., 2007]] ; [[#Qian--2017|Qian et al., 2017]] ), the Namibian or along the north-western Atlantic shelves ( [[#Claret--2018|Claret et al., 2018]] ). Other coastal upwelling sites such as the Arabian Sea display seasonal hypoxia but no worsening trends ( [[#Gupta--2016|Gupta et al., 2016]] ).

The Baltic Sea is the largest semi-enclosed sea where hypoxia is reported to have happened before the 1950s ( [[#Carstensen--2014|Carstensen et al., 2014]] ; [[#Rabalais--2014|Rabalais et al., 2014]] ; [[#Łukawska-Matuszewska--2019|Łukawska-Matuszewska et al., 2019]] ). The frequency and volume of seawater inflow from the North Sea decreased after 1950, leading to an expansion of hypoxic areas from 40,000 to 60,000 km² in combination with increasing eutrophication ( [[#Carstensen--2014|Carstensen et al., 2014]] ). From the available observations, there is ''robust evidence'' that many areas in the Baltic Sea are experiencing deoxygenation despite efforts to reduce nutrient loads ( [[#Lennartz--2014|Lennartz et al., 2014]] ; [[#Jokinen--2018|Jokinen et al., 2018]] ).

There is ''medium agreement'' ( ''medium evidence'' ) that simply reducing anthropogenic nutrient inputs may lead to less severe coastal hypoxic conditions, as observed in the coastal north-western Adriatic Sea ( [[#Djakovac--2015|Djakovac et al., 2015]] ). However, low-oxygen sediments may remain a long-term source of phosphorus and ammonium to the water column, and in this way fuelling primary production ( [[#Jokinen--2018|Jokinen et al., 2018]] ; [[#Fennel--2019|Fennel and Testa, 2019]] ; [[#Limburg--2020|Limburg et al., 2020]] ).

<div id="5.4" class="h1-container"></div>

<span id="biogeochemical-feedbacks-on-climate-change"></span>