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

==== 5.5.1.4 Combined Assessment of TCRE ====

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Studies differ in how they define TCRE, in the methods they use, and their assumptions, such as the assumed climate sensitivity distribution or the choice of metrics of global temperature change (e.g., GMST or GSAT, see Table 5.7). This makes TCRE estimates from individual studies difficult to compare. The combined assessment of TCRE therefore takes advantage of the well-established decomposition of TCRE in two factors: the TCR and the AF ( [[#5.5.1.3|Section 5.5.1.3]] ). This provides a TCRE assessment range for CO <sub>2</sub> -induced warming at the time of doubling CO <sub>2</sub> concentrations that builds on the broader Working Group 1 assessment. Expert judgement based on the airborne fraction range found in CMIP6 models ( [[#Arora--2020|Arora et al., 2020]] ; [[#Jones--2020|Jones and Friedlingstein, 2020]] ) suggest a value of 53% with a 1-sigma range of ±6%, which is double the sigma range based on the spread of CMIP6 models only. Combining this range with the AR6 TCR assessment ( [[IPCC:Wg1:Chapter:Chapter-7#7.5|Section 7.5]] , best estimate 1.8°C, 1.4°C–2.2°C ''likely'' and 1.2°C–2.4°C ''very likely'' range) results in a 5–95% range of 1.0–2.3°C per 1000 PgC (0.27°C–0.63°C per 1000 GtCO <sub>2</sub> ). Based on expert judgement that accounts for the incomplete coverage of all Earth system components, this results in a consolidated assessment that TCRE would fall ''likely'' in the range of 1.0–2.3°C per 1000 PgC, with a best estimate of 1.65°C per 1000 PgC (0.45°C per 1000 GtCO <sub>2</sub> ). Warming here reflects the human-induced GSAT increase and assumes a normal distribution. Some studies using observational constraints support a lognormal shape for the TCRE distribution ( [[#Spafford--2020|Spafford and Macdougall, 2020]] ), but such a distribution is currently not supported by the combined assessment of TCR and airborne fraction. Finally, this assessed TCRE range needs to be considered in combination with the ZEC ( [[IPCC:Wg1:Chapter:Chapter-4#4.7.1.1|Section 4.7.1.1]] ) when estimating the CO <sub>2</sub> -induced warming of low-emissions scenarios.

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'''Cross-Chapter Box 5.3 | The Ocean Carbon–Heat Nexus and Climate Change Commitment'''

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'''Contributors:''' Pedro M.S. Monteiro (South Africa), Jean-Baptiste Sallée (France), Piers Foster (United Kingdom), Baylor Fox-Kemper (United States of America), Helen T. Hewitt (United Kingdom), Masao Ishii (Japan), Joeri Rogelj (United Kingdom/Belgium), Kirsten Zickfeld (Canada/Germany)

'''Context'''
In the past 60 years, the ocean has taken up and stored 23 ± 5% of anthropogenic carbon emissions ( ''medium confidence'' ) ( [[#5.2.1.3|Section 5.2.1.3]] ) as well as more than 90% of the heat that has accumulated in the Earth system (referred to as excess heat) since the 1970s (Sections 7.2.2, 9.2.2 and 9.2.3, and Box 7.2; [[#Frölicher--2015|Frölicher et al., 2015]] ; [[#Talley--2016|Talley et al., 2016]] ; [[#Gruber--2019b|Gruber et al., 2019b]] ; [[#Hauck--2020|Hauck et al., 2020]] ). The interplay between heat and CO <sub>2</sub> uptake by the ocean has played a major role in slowing the rate of global warming, and also provides a first order influence in determining the unique properties of a metric of the coupled climate–carbon cycle response – transient climate response to cumulative CO <sub>2</sub> emissions (TCRE) – which is critical to setting the future remaining carbon emissions budget (Sections 5.5.1.3 and 5.5.4). This role of the ocean in the uptake of heat and anthropogenic CO <sub>2</sub> and related feedbacks is what we term the ‘ocean carbon–heat nexus’. The ocean processes behind this nexus are important in shaping and understanding the near-linear relationship between cumulative CO <sub>2</sub> emissions and global warming (TCRE) as well as the uncertainties in future projections of TCRE properties ( [[#Zickfeld--2016|Zickfeld et al., 2016]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ; [[#Jones--2020|Jones and Friedlingstein, 2020]] ), its path independence ( [[#MacDougall--2017|MacDougall, 2017]] ), and the warming commitment after cessation of greenhouse gas emissions – the zero emissions commitment (ZEC; [[#5.5.2|Section 5.5.2]] ; [[#Zickfeld--2016|Zickfeld et al., 2016]] ; [[#Ehlert--2017|Ehlert and Zickfeld, 2017]] ). In this box, we assess the role of the ocean and its physical and chemical thermodynamic processes that shape these striking characteristics.

The role of the ocean in setting the coupled climate–carbon cycle response is threefold. First, the ocean and land carbon sinks together set the airborne fraction (AF) of CO <sub>2</sub> in the atmosphere, which sets the radiative forcing that drives the additional heat in the atmosphere, most of which is taken up by the ocean (Sections 7.2 and 9.2; [[#Katavouta--2019|Katavouta et al., 2019]] ; [[#Williams--2019|Williams et al., 2019]] ). the land carbon sink does not appear to play an important role in determining the linearity and path-independence of TCRE ( [[#5.5.1.1|Section 5.5.1.1]] ; [[#Goodwin--2015|Goodwin et al., 2015]] ; [[#MacDougall--2015|MacDougall and Friedlingstein, 2015]] ; [[#Ehlert--2017|Ehlert et al., 2017]] ). Second, the ocean sets the thermal response through ocean heat uptake ( [[IPCC:Wg1:Chapter:Chapter-9#9.2|Section 9.2]] ; [[#Frölicher--2015|Frölicher et al., 2015]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). Third, there is a feedback within the ocean carbon–heat nexus as ocean warming, particularly under low or no mitigation scenarios, weakens the ocean sink of CO <sub>2</sub> , which influences the AF, and hence the radiative forcing (Box 7.1; [[#Williams--2019|Williams et al., 2019]] ). The near-linear relationship between cumulative CO <sub>2</sub> emissions and global warming (TCRE) is thought to arise, to a large extent, from the compensation between the decreasing ability of the ocean to take up heat and CO <sub>2</sub> at higher cumulative CO <sub>2</sub> emissions, pointing to similar processes that determine ocean uptake of heat and carbon ( [[#5.5.1.1|Section 5.5.1.1]] ; [[#Goodwin--2015|Goodwin et al., 2015]] ; [[#MacDougall--2015|MacDougall and Friedlingstein, 2015]] ; [[#Williams--2016|Williams et al., 2016]] ; [[#Zickfeld--2016|Zickfeld et al., 2016]] ; [[#Ehlert--2017|Ehlert et al., 2017]] ).

'''Processes that drive the ocean carbon–heat nexus and its change'''
The air–sea flux of heat and all gases across the ocean interface is driven by a common set of complex and turbulent diffusion and mixing processes that are difficult to observe (Sections 5.2.1.3 and 9.2.1.2; [[#Wanninkhof--2009|Wanninkhof et al., 2009]] ; [[#Wanninkhof--2014|Wanninkhof, 2014]] ; [[#Cronin--2019|Cronin et al., 2019]] ; [[#Watson--2020|Watson et al., 2020]] ). These processes are typically simplified into widely verified expressions that link the flux to wind stress, the solubility and the gradient across the air–sea interface ( ''medium confidence'' ). Because the ocean has a higher heat capacity than the atmosphere (the heat capacity of the upper 100 m of the ocean is about 30 times larger than the heat capacity of the atmosphere), the partitioning of heat between the atmosphere and the ocean is primarily influenced by the temperature differences between air and seawater. Similarly, the unique seawater carbonate buffering capacity enables CO <sub>2</sub> to be stored in the ocean as dissolved salts, rather than just as dissolved gas; this increases the capacity of seawater to store CO <sub>2</sub> by two orders of magnitude beyond the solubility of CO <sub>2</sub> gas and approximates the partitioning ratio of heat between the atmosphere and the ocean ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.2.1|Section 9.2.2.1]] ; [[#Zeebe--2009|Zeebe and Wolf-Gladrow, 2009]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). The role of the biological carbon pump in influencing the ocean sink of anthropogenic carbon into the ocean interior is assessed to be minimal during the historical period, but this may change, particularly in regional contexts, by 2100 ( ''medium confidence'' ) ( [[#Laufkötter--2015|Laufkötter et al., 2015]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Its role is important in the natural or pre-industrial carbon cycle ( ''medium confidence'' ) ( [[#Henson--2016|Henson et al., 2016]] ).

Under climate change, the buffering capacity of the ocean decreases (increasing Revelle Factor), which reflects a decreasing capacity for the ocean to take up additional anthropogenic CO <sub>2</sub> and store it in the dissolved inorganic carbon reservoir ( [[#Egleston--2010|Egleston et al., 2010]] ). In contrast to CO <sub>2</sub> , there is no physical limitation that would reduce the ability of surface ocean temperature to equilibrate with the atmospheric temperature. However, both carbon and heat fluxes depend on air–sea heat fluxes that in turn depend on gradients of characteristics at the air–sea interface. These gradients at the air–sea interface respond to ocean dynamics, such as the volume of the surface mixed-layer that equilibrates with the atmosphere, and ocean circulation that can flush the surface layer with water masses that have not equilibrated with the atmosphere for a long time. Limited recent evidence suggests that the effect of small-scale dynamics absent in climate and Earth system models might be locally important ( [[#Bachman--2020|Bachman and Klocker, 2020]] ). In summary, changes in heat and carbon uptake by the ocean rely on a combination of unique chemical and shared physical processes, any of which have the potential to disrupt the coherence of heat and CO <sub>2</sub> change in the ocean.

'''Spatial pattern of air–sea fluxes and storage'''
Large-scale regional and global ocean circulation shape the spatial pattern of the uptake and storage of both CO <sub>2</sub> and heat (see Figure 5.8 for carbon; Figure 9.6 for heat observations; [[IPCC:Wg1:Chapter:Chapter-9#9.2|Section 9.2]] ; [[#Frölicher--2015|Frölicher et al., 2015]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). This coherence of spatial patterns driven by the large-scale ocean circulation has three aspects. First, notwithstanding interannual-decadal variability in heat and CO <sub>2</sub> uptake, there is a spatial coherence of the temporally integrated uptake at the air–sea boundary, particularly in the Southern Ocean (Cross-Chapter Box 5.3, Figure 1; [[#Talley--2016|Talley et al., 2016]] ; [[#Keppler--2019|Keppler and Landschützer, 2019]] ; [[#Auger--2021|Auger et al., 2021]] ). Second, the importance of the meridional overturning circulation in the subsequent storage of both heat and CO <sub>2</sub> in mode, intermediate and deep waters of the ocean interior ( [[IPCC:Wg1:Chapter:Chapter-9#9.2|Section 9.2]] ). Third, of particular note, the roles of the North Atlantic Ocean ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.3.1|Section 9.2.3.1]] ) and the Southern Ocean ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.3.2|Section 9.2.3.2]] ) in linking the spatial pattern of air–sea fluxes, the storage of heat and carbon, and ultimately in understanding and predicting the sensitivity of the carbon-heat nexus to climate change ( [[#Frölicher--2015|Frölicher et al., 2015]] ; [[#Thomas--2018|Thomas et al., 2018]] ; [[#Wu--2019|Wu et al., 2019]] ).

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[[File:ef5d8d921b5a54a76f265843b8749cce IPCC_AR6_WGI_CCBox_5_3_Figure_1.png]]

'''Cross-Chapter Box 5.3, Figure 1 |'''

'''CMIP6 multi-model mean of changes in zonally integrated (a) heat and (b) carbon storage in the ocean''' '''between the pre-industrial and the modern period''' . Carbon corresponds to dissolved in organic carbon. Data are shown for the upper 2000 m. The modern period is 1995–2014. Adapted from [[#Frölicher--2015|Frölicher et al. (2015)]] .

The role of the large-scale circulation in shaping these fluxes is to: (i) flush the ocean surface layer with deep waters that are relatively cold and with weak or no anthropogenic CO <sub>2</sub> and heat content because they have been isolated from the atmosphere for centuries; and (ii) transport the anthropogenic CO <sub>2</sub> and heat at depth, away from the atmosphere ( [[#Frölicher--2015|Frölicher et al., 2015]] ; [[#Marshall--2015|Marshall et al., 2015]] ; [[#Armour--2016|Armour et al., 2016]] ). For instance, in the Southern Ocean, upwelled water masses take up a large amount of anthropogenic CO <sub>2</sub> and heat (Cross-Chapter Box 5.3, Figure 1), which are then exported northward by the circulation to be stored at depth in the Southern Hemisphere subtropical gyres (Cross-Chapter Box 5.3, Figure 1; Figure 9.7). In the North Atlantic, the signature of the Atlantic meridional overturning circulation (AMOC) is also clearly visible, with large amounts of heat and carbon being stored beneath the North Atlantic subtropical gyre at 1 km depth (Cross-Chapter Box 5.3, Figure 1). In summary, the net air–sea fluxes of anthropogenic CO <sub>2</sub> and heat depend on large-scale circulation, which is associated with upper ocean stratification, mixed-layer depth, and water-mass formation, transport and mixing (Sections 9.1–9.3).

'''Changes in ocean processes and impact on the ocean carbon–heat nexus'''
Future projections of the ocean carbon–heat nexus in the second half of the 21st century, particularly those under weak or no mitigation scenarios, are characterized by the strengthening of the two largest positive feedbacks: weakening surface ocean CO <sub>2</sub> buffering capacity (increasing Revelle Factor) and warming that further reduces CO <sub>2</sub> solubility and strengthens ocean stratification, which reduces exchange between the ocean surface and interior ( [[#Jiang--2019|Jiang et al., 2019]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). These are offset by a growing but scenario-dependent negative feedback from increasing carbon and heat air–sea fluxes towards the ocean, due to increased atmospheric temperature and CO <sub>2</sub> concentrations ( [[#Talley--2016|Talley et al., 2016]] ; [[#Jiang--2019|Jiang et al., 2019]] ; [[#McKinley--2020|McKinley et al., 2020]] ). The Southern Ocean in particular is one of the regions where the projected feedback can be largest and where inter-model differences are strongest ( [[#Roy--2011|Roy et al., 2011]] ; [[#Frölicher--2015|Frölicher et al., 2015]] ; [[#Hewitt--2016|Hewitt et al., 2016]] ; [[#Mongwe--2018|Mongwe et al., 2018]] ). These projected trends in ocean carbonate chemistry ( [[#5.4.2|Section 5.4.2]] ), together with surface ocean warming ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.1.1|Section 9.2.1.1]] ), explain the slow down and long-term reduction of the ocean sink for anthropogenic CO <sub>2</sub> even as emissions continue to rise beyond 2050 under weak-to-no-mitigation scenarios (Figures 2.7.1 and 5.25, and Technical Summary TS Box 7). Projected change in the North Atlantic and Southern Ocean overturning circulation also impact air–sea fluxes of heat and carbon. The ''very likely'' decline in AMOC in the 21st century for all shared socio-economic pathways (SSP) scenarios ( [[IPCC:Wg1:Chapter:Chapter-9#9.2.3.1|Section 9.2.3.1]] ) tends to reduce heat and carbon uptake, resulting in a positive feedback. In contrast, in the Southern Ocean, the future 21st century projected increase in upper ocean overturning circulation ( ''low confidence'' ) – due to increasing wind forcing projected for all scenarios, except those with large mitigation (SSP1-2.6) – produces a negative feedback, with increasing heat and carbon uptake and storage despite the increasing stratification and outgassing of natural CO <sub>2</sub> in the upwelling zone (Sections 9.2.3.2 and 5.2.1.3).

In summary, a combination of unique chemical properties of seawater carbonate combined with shared physical ocean processes explain the coherence and scaling in the uptake and storage of both CO <sub>2</sub> and heat in the ocean, which is the basis for the carbon–heat nexus ( ''high confidence'' ). In this way, the processes of the ocean carbon-heat nexus help understand the quasi-linear and path independence of properties of TCRE, which forms the basis for the zero emissions commitment (ZEC; [[#5.5|Section 5.5]] ) ( ''medium confidence'' ). Future projections under low or no mitigation indicate with ''high confidence'' that carbon chemistry and warming will strengthen the positive feedback to climate change by reducing ocean carbon uptake, and ''medium confidence'' that ocean circulation may partially compensate that positive feedback by slightly increasing anthropogenic carbon storage. Increasing ocean warming and stratification may decrease exchanges between the surface and subsurface ocean, which could reduce the path independence of TCRE, though this effect can be partially counterbalanced regionally by increasing circulation associated with increasing winds ( ''l'' ''ow confidence'' ).

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