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

=== 5.4.4 Climate Effects on Future Ocean Carbon Uptake ===

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==== 5.4.4.1 Physical Drivers of Future Ocean Carbon Uptake and Storage ====

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The principal contribution to increasing global ocean carbon is the air–sea flux of CO <sub>2</sub> , which changes the dissolved inorganic carbon (DIC) inventory ( [[#5.4.2|Section 5.4.2]] ; [[#Arora--2020|Arora et al., 2020]] ). The processes that influence the variability and trends of the ocean carbon–heat nexus are assessed in Cross-Chapter Box 5.3. Climate has three important impacts on the ocean uptake of anthropogenic CO <sub>2</sub> : (i) ocean warming reduces the solubility of CO <sub>2</sub> , which increases ''p'' CO <sub>2</sub> and increases the stratification of the mixed layer, both acting as positive feedbacks weakening the ocean sink ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1|Section 9.2.1]] and Cross-Chapter Box 5.3; [[#Arora--2020|Arora et al., 2020]] ); (ii) changing the temporal and spatial characteristics of wind stress and storms alters mixing – entrainment in, and across the bottom of, the mixed layer ( [[#Bronselaer--2018|Bronselaer et al., 2018]] ); and (iii) warming and wind stress influence the large-scale meridional overturning circulation (MOC) circulation, which modifies the rate of ventilation, storage or outgassing of ocean carbon in the ocean interior ( [[#5.2.3.1|Section 5.2.3.1]] ; [[#Gruber--2019b|Gruber et al., 2019b]] ; [[#Arora--2020|Arora et al., 2020]] ). The land-to-ocean riverine flux and the carbon burial in ocean sediments play a minor role ( ''low confidence'' ) ( [[#Arora--2020|Arora et al., 2020]] ). Based on ''high agreement'' of projections by coupled climate models, there is ''high confidence'' that the resultant climate–carbon cycle feedbacks are positive, but the extent of the ocean sink weakening is scenario dependent ( [[#Arora--2020|Arora et al., 2020]] ).

Regionally, the Southern Ocean is a major sink of anthropogenic CO <sub>2</sub> (Figure 5.8a), although challenges in modelling its circulation and Antarctic sea ice transport (Sections 3.4.1.2, 9.2.3.2 and 9.3.2) generate uncertainty in the response of its sink to future carbon–climate feedbacks. Increased freshwater input may cause a slowdown of the lower overturning circulation, leading to increased Southern Ocean biological carbon storage ( [[#Ito--2015|Ito et al., 2015]] ); alternatively, increased winds may intensify the overturning circulation, reducing the net CO <sub>2</sub> sink in the Southern Ocean ( [[#Bronselaer--2018|Bronselaer et al., 2018]] ; [[#Saunders--2018|Saunders et al., 2018]] ). On centennial time scales, there is thus ''low confidence'' in the overall effect of intensifying winds in the Southern Ocean on CO <sub>2</sub> uptake.

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==== 5.4.4.2 Biological Drivers of Future Ocean Carbon Uptake ====

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While physical drivers control the present-day anthropogenic carbon sink, biological processes are responsible for the majority of the vertical gradient in DIC (natural carbon storage). A small fraction of the organic carbon fixed by primary production (PP) reaches the sea floor, where it can be stored in sediments on geological time scales, making the biological carbon pump (BCP) an important mechanism for very long-term CO <sub>2</sub> storage. Projected reductions in ocean ventilation ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1.4|Section 9.2.1.4]] ) would lengthen residence time and lead to DIC accumulating in the deep ocean due to organic carbon remineralization.

Since AR5 (Section 6.3.2.5.6), progress has been made in understanding the biological drivers of ocean carbon uptake in both coupled climate models and observations (SROCC, [[#5.2.2.6|Section 5.2.2.6]] ). Here we focus on potential feedbacks between biological processes and climate. In CMIP5 models, the direction of modelled PP in response to increased atmospheric CO <sub>2</sub> concentration and climate warming wasunclear ( [[#Taucher--2011|Taucher and Oschlies, 2011]] ; [[#Laufkötter--2015|Laufkötter et al., 2015]] ). This remains the case in the CMIP6 models; inter-model uncertainty has increased in CMIP6 models, compared to CMIP5. The projected global multi-model mean change in PP in 13 models run under the SSP5−8.5 scenario is −3 ± 9% (2080–2099 mean values relative to 1870–1899 ± the inter-model standard deviation; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Under the low-emissions, high-mitigation scenario SSP1−2.6, the global change in PP is −0.56 ± 4%. Observations in the contemporary period provide little direct constraint on the modelled responses of PP to climate change, partly due to insufficiently long records ( [[#Henson--2016|Henson et al., 2016]] ). However, there is some indication of an emergent constraint on changes in tropical PP based on interannual variability derived from remote sensing ( [[#5.4.6|Section 5.4.6]] ; [[#Kwiatkowski--2017|Kwiatkowski et al., 2017]] ).

In CMIP5 models run under RCP8.5, particulate organic carbon (POC) export flux is projected to decline by 1–12% by 2100 ( [[#Taucher--2011|Taucher and Oschlies, 2011]] ; [[#Laufkoetter--2015|Laufkoetter et al., 2015]] ). Similar values are predicted in 18 CMIP6 models, with declines of 2.5–21.5% (median –14%) or 0.2–2 GtC (median –0.8 GtC) between 1900 and 2100 under the SSP5-8.5 scenario. The mechanisms driving these changes vary widely between models due to differences in parametrization of particle formation, remineralization and plankton community structure.

Ocean warming reduces the vertical supply of nutrients to the upper ocean due to increasing stratification ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1.4|Section 9.2.1.4]] ) but may also act to alleviate seasonal light limitation. The projected effect is to decrease PP at low latitudes and increase PP at high latitudes ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Future changes to dust deposition due to desertification ( [[#Mahowald--2017|Mahowald et al., 2017]] ), alterations to the nitrogen cycle ( [[#5.3.3.2|Section 5.3.3.2]] ; SROCC, [[#5.2.3.1.2|Section 5.2.3.1.2]] ), and reducing sea ice cover (Ardyna and Arrigo, 2020) all have the potential to alter PP regionally. Higher ocean temperatures tend to result in higher metabolic rates, although respiration may increase more rapidly than PP ( [[#Boscolo-Galazzo--2018|Boscolo-Galazzo et al., 2018]] ; [[#Brewer--2019|Brewer, 2019]] ; [[#Cavan--2019|Cavan et al., 2019]] ). Ocean warming and reduced PP are expected to result in lower zooplankton abundance, and the expansion of oxygen minimum zones (OMZs) may reduce the ability of zooplankton to remineralize POC, thus increasing the efficiency of the BCP and forming a negative climate feedback ( [[#Cavan--2017|Cavan et al., 2017]] ). Increased microbial respiration due to warming may result in greater quantities of organic carbon transferred into the dissolved organic carbon pool ( [[#Jiao--2014|Jiao et al., 2014]] ; [[#Legendre--2015|Legendre et al., 2015]] ; [[#Roshan--2017|Roshan and DeVries, 2017]] ) which, while increasing the residence time of carbon in the ocean, would ultimately reduce the sedimentary burial, and hence sequestration on geologic time scales ( [[#Olivarez%20Lyle--2006|Olivarez Lyle and Lyle, 2006]] ).

Most models project that smaller phytoplankton are favoured in future ocean conditions ( ''medium confidence'' ; [[#Cabré--2015|Cabré et al., 2015]] ; [[#Fu--2016|Fu et al., 2016]] ; [[#Flombaum--2020|Flombaum et al., 2020]] ) driven by warming water and/or changing nutrient availability, which would alter the magnitude and efficiency of the BCP by altering the sinking speed, respiration rate and aggregation/fragmentation of sinking particles. There is ''low confidence'' in the sign of the resulting feedback: regions in which small phytoplankton dominate may have a more efficient pump, although the total amount of organic carbon reaching the sea floor is lower ( [[#Herndl--2013|Herndl and Reinthaler, 2013]] ; [[#Bach--2016|Bach et al., 2016]] ; [[#Richardson--2019|Richardson, 2019]] ). Alternatively, an increase in small phytoplankton could result in a less efficient pump, due either to a greater fraction of PP being processed through the upper ocean microbial loop ( [[#Jiao--2014|Jiao et al., 2014]] ) or generation of slower sinking particles ( [[#Guidi--2009|Guidi et al., 2009]] ; [[#Leung--2021|Leung et al., 2021]] ). Variable phytoplankton stoichiometry is predicted to increase the amount of carbon stored via the BCP relative to the amount of PP, so that fixed stoichiometry models (as in CMIP5) may underestimate cumulative ocean carbon uptake to 2100 by 0.5–3.5% (2–15 PgC; RCP8.5 scenario; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Other climate effects such as deoxygenation or ocean acidification could also result in alterations to the magnitude and efficiency of the BCP ( [[#Krumhardt--2019|Krumhardt et al., 2019]] ; [[#Raven--2021|Raven et al., 2021]] ; [[#Taucher--2021|Taucher et al., 2021]] ).

Based on ''high agreement'' across multiple lines of evidence and physical understanding there is ''high confidence'' that feedbacks to climate will arise from alterations to the magnitude and efficiency of the BCP changing PP, and the depth of remineralization. However, the complexity of the mechanisms involved in the export and remineralization of POC result in ''low confidence'' in the magnitude and sign of biological feedbacks to climate. Nevertheless, improved model representation of PP and the BCP is required (which requires better observational constraints), as the contribution of biological processes to CO <sub>2</sub> uptake is expected to become more significant with continued climate change ( [[#Hauck--2015|Hauck et al., 2015]] ).

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