Editing IPCC:AR6/WGI/TS (section)

== TS.3 Understanding the Climate System Response and Implications for Limiting Global Warming ==

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This section summarizes advances in our knowledge of Earth’s energy budget, including the time evolution of forcings and climate feedbacks that lead to the climate system responses summarized in Section TS.2. It assesses advances since AR5 and SR1.5 in the estimation of remaining carbon budgets, the Earth system response to carbon dioxide removal, and the quantification of metrics that allow comparisons of the relative effects of different forcing agents. The section also highlights: future climate and air pollution responses due to projected changes in short-lived climate forcers (SLCFs); the state of understanding of the climate response to potential interventions related to solar radiation modification (SRM); and irreversibility, tipping points and abrupt changes in the climate system.

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=== TS.3.1 Radiative Forcing and Energy Budget ===

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'''Since AR5, the accumulation of energy in the Earth system, quantified by observations of warming of the ocean, atmosphere, and land and melting of ice, has become established as a robust measure of the rate of global climate change on interannual-to-decadal time scales. Compared to changes in global surface temperature, the increase in the global energy inventory exhibits less variability, and thus better indicates underlying climate trends.''' '''The global energy inventory increased by 282 [177 to 387] zettajoules (ZJ, equal to 10 <sup>21</sup> joules) for the period 1971–2006 <sup></sup> and 152 [100 to 205] ZJ for the period 2006–2018 (Figure TS.13), with more than 90% accounted for by ocean warming. To put these numbers in context, the 2006–2018 average Earth energy imbalance is equivalent to approximately 20 times the annual rate of global energy consumption in 2018. The accumulation of energy is driven by a positive total anthropogenic effective radiative forcing (ERF) relative to 1750.''' '''The best estimate ERF of 2.72 W m <sup>−2</sup> has increased by 0.43 W m <sup>–2</sup> relative to that given in AR5 (for 1750–2014) due to an increase in the greenhouse gas ERF that is partly compensated by a more negative aerosol ERF compared to AR5. The greenhouse gas ERF has been revised due to changes in atmospheric concentrations and updates to forcing efficiencies, while the revision to aerosol ERF is due to increased understanding of aerosol–cloud interactions and is supported by improved agreement between different lines of evidence. Improved quantifications of ERF, the climate system radiative response, and the observed energy increase in the Earth system for the period 1971–2018 demonstrate improved closure of the global energy budget (i.e., the extent to which the sum of the integrated forcing and the integrated radiative response equals the energy gain of the Earth system) compared to AR5 ( ''high confidence'' ). (See FAQ 7.1). Links to chapters 7.2.2, 7.3.5, 7.5.2, Box 7.2, Table 7.1'''

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'''Figure TS.13 |''' '''Estimates of the net cumulative energy change (ZJ = 10''' '''21''' '''Joules) for the period 1971–2018 associated with (a) observations of changes in the global energy inventory, (b) integrated radiative forcing, and (c) integrated radiative response.''' ''The intent is to show assessed changes in energy budget and effective radiative forcings (ERFs).'' Black dotted lines indicate the central estimate with ''likely'' and ''very likely'' ranges as indicated in the legend. The grey dotted lines indicate the energy change associated with an estimated 1850–1900 Earth energy imbalance of 0.2 W m <sup>–2</sup> (panel a) and an illustration of an assumed pattern effect of –0.5 W m <sup>–2</sup> °C <sup>–1</sup> (panel c). Background grey lines indicate equivalent heating rates in W m <sup>–2</sup> per unit area of Earth’s surface. Panels (d) and (e) show the breakdown of components, as indicated in the legend, for the global energy inventory and integrated radiative forcing, respectively. Panel (f) shows the global energy budget assessed for the period 1971–2018, that is, the consistency between the change in the global energy inventory relative to 1850–1900 and the implied energy change from integrated radiative forcing plus integrated radiative response under a number of different assumptions, as indicated in the figure legend, including assumptions of correlated and uncorrelated uncertainties in forcing plus response. Shading represents the ''very likely'' range for observed energy change relative to 1850–1900 and ''likely'' range for all other quantities. Forcing and response time series are expressed relative to a baseline period of 1850–1900. Links to chapters Box 7.2, Figure 1

The global energy inventory change for the period 1971–2006 corresponds to an Earth energy imbalance (Box TS.1) of 0.50 [0.32 to 0.69] W m <sup>–2</sup> , increasing to 0.79 [0.52 to 1.06] W m <sup>–2</sup> for the period 2006 '''–''' 2018. Ocean heat uptake is by far the largest contribution and accounts for 91% of the total energy change. Land warming, melting of ice and warming of the atmosphere account for about 5%, 3% and 1% of the total change, respectively. More comprehensive analysis of inventory components, cross-validation of satellite and in situ-based estimates of the global energy imbalance, and closure of the global sea level budget have led to a strengthened assessment relative to AR5. ( ''high confidence'' ) Links to chapters 7.2.2, 7.5.2.3, Box 7.2, Table 7.1, 9.6.1, Cross-Chapter Box 9.1, Table 9.5

As in AR5, the perturbations to Earth’s top-of-atmosphere energy budget are quantified using ERFs (see also Section TS.2.2). These include any consequent adjustments to the climate system (e.g., from changes in atmospheric temperatures, clouds and water vapour as shown in Figure TS.14), but exclude any surface temperature response. Since AR5, ERFs have been estimated for a larger number of forcing agents and shown to be more closely related to the temperature response than the stratospheric-temperature-adjusted radiative forcing. ( ''high confidence'' ) Links to chapters 7.3.1

Improved quantifications of ERF, the climate system radiative response, and the observed energy increase in the Earth system for the period 1971–2018 demonstrate improved closure of the global energy budget relative to AR5 (Figure TS.13). Combining the ''likely'' range of ERF over this period with the central estimate of radiative response gives an expected energy gain of 340 [47 to 662] ZJ. Both estimates are consistent with an independent observation-based assessment of the global energy increase of 284 [96 to 471] ZJ ( ''very likely'' ''range'' ), expressed relative to the estimated 1850–1900 Earth energy imbalance. ( ''high confidence'' ) Links to chapters 7.2.2, 7.3.5, Box 7.2

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'''Figure TS.14 |''' '''Schematic representation of changes in the top-of-atmosphere (TOA) radiation budget following a perturbation.''' ''The intent of this figure is to illustrate the concept of adjustments in the climate system following a perturbation in the radiation budget.'' The baseline TOA energy budget (a) responds instantaneously to perturbations (b) , leading to adjustments in the atmospheric meteorology and composition and land surface that are independent of changes in surface temperature (c) . Surface temperature changes (here using an increase as an example) lead to physical, biogeophysical and biogeochemical feedback processes (d) . Long-term feedback processes, such as those involving ice sheets, are not shown here. Links to chapters adapted from Figure 7.2; FAQ 7.2, Figure 1; and Figure 8.3

The assessed greenhouse gas ERF over the 1750–2019 period (Section TS.2.2) has increased by +0.59 W m <sup>−2</sup> over AR5 estimates for 1750–2011. This increase includes +0.34 W m <sup>–2</sup> from increases in atmospheric concentrations of well-mixed greenhouse gases (including halogenated species) since 2011, +0.15 W m <sup>–2</sup> from upwards revisions of their radiative efficiencies and +0.10 W m <sup>–2</sup> from re-evaluation of the ozone and stratospheric water vapour ERF. Links to chapters 7.3.2, 7.3.4, 7.3.5

For CO <sub>2</sub> , CH <sub>4</sub> , N <sub>2</sub> O, and chlorofluorocarbons, there is now evidence to quantify the effect on ERF of tropospheric adjustments. The assessed ERF for a doubling of CO <sub>2</sub> compared to 1750 levels (3.9 ± 0.5 Wm <sup>–2</sup> ) is larger than in AR5. For CO <sub>2</sub> , the adjustments include the physiological effects on vegetation. The reactive well-mixed greenhouse gases (CH <sub>4</sub> , N <sub>2</sub> O, and halocarbons) cause additional chemical adjustments to the atmosphere through changes in ozone and aerosols (Figure TS.15a). The ERF due to CH <sub>4</sub> emissions is 1.19 [0.81 to 1.58] W m <sup>–2</sup> , of which 0.35 [0.16 to 0.54] W m <sup>–2</sup> is attributed to chemical adjustments mainly via ozone. These chemical adjustments also affect the emissions metrics (Section TS.3.3.3). Changes in sulphur dioxide (SO <sub>2</sub> <sub>)</sub> emissions make the dominant contribution to the ERF from aerosol–cloud interactions ( ''high confidence'' ). Over the 1750–2019 period, the contributions from the emitted compounds to global surface temperature changes broadly match their contributions to the ERF ( ''high confidence'' ) (Figure TS.15b). Since a peak in emissions-induced SO <sub>2</sub> eRF has already occurred recently (Section TS.2.2) and since there is a delay in the full global surface temperature response owing to the thermal inertia in the climate system, changes in SO <sub>2</sub> emissions have a slightly larger contribution to global surface temperature change compared with changes in CO <sub>2</sub> emissions, relative to their respective contributions to ERF. Links to chapters 6.4.2, 7.3.2

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'''Figure TS.15 |''' '''Contribution to (a) effective radiative forcing (ERF) and (b) global surface temperature change from component emissions for''' '''1750–2019''' '''based on Coupled Model Intercomparison Project Phase 6 (CMIP6) models and (c) net aerosol ERF for 1750–2014 from different lines of evidence.''' ''The intent of this figure is to show advances since AR5 in the understanding of (a) emissions-based ERF, (b) global surface temperature response for'' ''short-lived'' ''climate forcers as estimated in Chapter 6, and (c) aerosol ERF from different lines of evidence as assessed in Chapter 7.'' In panel (a), ERFs for well-mixed greenhouse gases (WMGHGs) are from the analytical formulae. ERFs for other components are multi-model means based on Earth system model simulations that quantify the effect of individual components. The derived emissions-based ERFs are rescaled to match the concentration-based ERFs in Figure 7.6. Error bars are 5–95% and for the ERF account for uncertainty in radiative efficiencies and multi-model error in the means. In panel (b), the global mean temperature response is calculated from the ERF time series using an impulse response function. In panel (c), the AR6 assessment is based on energy balance constraints, observational evidence from satellite retrievals, and climate model-based evidence. For each line of evidence, the assessed best-estimate contributions from ERF due to aerosol–radiation interactions (ERFari) and aerosol–cloud interactions (ERFaci) are shown with darker and paler shading, respectively. Estimates from individual CMIP Phase 5 (CMIP5) and CMIP6 models are depicted by blue and red crosses, respectively. The observational assessment for ERFari is taken from the instantaneous forcing due to aerosol–radiation interactions (IRFari). Uncertainty ranges are given in black bars for the total aerosol ERF and depict ''very likely'' ranges. Links to chapters 6.4.2, Figure 6.12, 7.3.3, Cross-Chapter Box 7.1, Table 7.8, Figure 7.5

Aerosols contributed an ERF of –1.3 [–2.0 to –0.6] W m <sup>–2</sup> over the period 1750 to 2014 ( ''medium confidence'' ). The ERF due to aerosol–cloud interactions (ERFaci) contributes most to the magnitude of the total aerosol ERF ( ''high confidence'' ) and is assessed to be –1.0 [–1.7 to –0.3] W m <sup>–2</sup> ( ''medium confidence'' ), with the remainder due to aerosol–radiation interactions (ERFari), assessed to be –0.3 [–0.6 to 0.0] W m <sup>–2</sup> ( ''medium confidence'' ). There has been an increase in the estimated magnitude – but a reduction in the uncertainty – of the total aerosol ERF relative to AR5, supported by a combination of increased process-understanding and progress in modelling and observational analyses (Figure TS.15c). Effective radiative forcing estimates from these separate lines of evidence are now consistent with each other, in contrast to AR5, and support the assessment that it is ''virtually certain'' that the total aerosol ERF is negative. Compared to AR5, the assessed magnitude of ERFaci has increased, while that of ERFari has decreased ''.'' Links to chapters 7.3.3, 7.3.5

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=== TS.3.2 Climate Sensitivity and Earth System Feedbacks ===

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==== TS.3.2.1 Equilibrium Climate Sensitivity, Transient Climate Response, and Transient Climate Response to Cumulative Carbon-dioxide Emissions ====

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'''Since AR5, substantial quantitative progress has been made in combining new evidence of Earth’s climate sensitivity with improvements in the understanding and quantification of Earth’s energy imbalance, the instrumental record of global surface temperature change, paleoclimate change from proxy records, climate feedbacks and their dependence on time scale and climate state. A key advance is the broad agreement across these multiple lines of evidence, supporting a best estimate of equilibrium climate sensitivity of 3°C, with a ''very likely'' range of 2°C to 5°C. The ''likely'' range of 2.5°C to 4°C is narrower than the AR5 ''likely'' range of 1.5°C to 4.5°C. Links to chapters 7.4, 7.5'''
Constraints on equilibrium climate sensitivity (ECS) and transient climate response (TCR) (see Glossary) are based on four main lines of evidence: feedback process understanding, climate change and variability seen within the instrumental record, paleoclimate evidence, and so-called ‘emergent constraints’, whereby a relationship between an observable quantity and either ECS or TCR established within an ensemble of models is combined with observations to derive a constraint on ECS or TCR. In reports up to and including the IPCC Third Assessment Report, ECS and TCR derived directly from ESMs were the primary line of evidence. However, since AR4, historical warming and paleoclimates provided useful additional evidence (Figure TS.16a). This Report differs from previous reports in not directly using climate model estimates of ECS and TCR in the assessed ranges of climate sensitivity. Links to chapters 1.5, 7.5

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'''Figure TS.16 |''' '''(a) Evolution of equilibrium climate sensitivity (ECS) assessments from the Charney Report through a succession of IPCC Assessment Reports to AR6, and lines of evidence and combined assessment for (b) ECS and (c) transient climate response (TCR) in AR6.''' ''The intent of this figure is to show the progression in estimates of ECS, including uncertainty and the lines of evidence used for assessment, and to show the lines of assessment used to assess ECS and TCR in AR6.'' In panel (a), the lines of evidence considered are listed below each assessment. Best estimates are marked by horizontal bars, ''likely'' ranges by vertical bars, and ''very likely'' ranges by dotted vertical bars. In panel (b) and (c), assessed ranges are taken from Tables 7.13 and 7.14 for ECS and TCR respectively. Note that for the ECS assessment based on both the instrumental record and paleoclimates, limits (i.e., one-sided distributions) are given, which have twice the probability of being outside the maximum/minimum value at a given end, compared to ranges (i.e., two tailed distributions) which are given for the other lines of evidence. For example, the ''extremely likely'' limit of greater than 95% probability corresponds to one side of the ''very likely'' (5% to 95%) range. Best estimates are given as either a single number or by a range represented by grey box. Coupled Model Intercomparison Project Phase 6 (CMIP6) Earth system model (ESM) values are not directly used as a line of evidence but are presented on the figure for comparison. Links to chapters 1.5, 7.5; Tables 7.13 and 7.14; Figure 7.18

It is now clear that when estimating ECS and TCR, the dependence of feedbacks on time scales and the climate state must be accounted for. Feedback processes are expected to become more positive overall (more amplifying of global surface temperature changes) on multi-decadal time scales as the spatial pattern of surface warming evolves and global surface temperature increases, leading to an ECS that is higher than was inferred in AR5 based on warming over the instrumental record ( ''high confidence'' ). Historical surface temperature change since 1870 has shown relatively little warming in several key regions of positive feedbacks, including the eastern equatorial Pacific Ocean and the Southern Ocean, while showing greater warming in key regions of negative feedbacks, including the western Pacific warm pool. Based on process understanding, climate modelling, and paleoclimate reconstructions of past warm periods, it is expected that future warming will become enhanced over the eastern Pacific Ocean ( ''medium confidence'' ) and Southern Ocean ( ''high confidence'' ) on centennial time scales. This new understanding, along with updated estimates of historical temperature change, ERF, and energy imbalance, reconciles previously disparate ECS estimates. Links to chapters 7.4.4, 7.5.2, 7.5.3

The AR6 best estimate of ECS is 3°C, the ''likely'' range is 2.5°C to 4°C and the ''very likely'' range is 2°C to 5°C. There is a high level of agreement among the four main lines of evidence listed above (Figure TS.16b), and altogether it is ''virtually certain'' that ECS is larger than 1.5°C, but currently it is not possible to rule out ECS values above 5°C. Therefore, the 5°C upper end of the ''very likely'' range is assessed with ''medium confidence'' and the other bounds with ''high confidence'' . Links to chapters 7.5.5

Based on process understanding, warming over the instrumental record, and emergent constraints, the best estimate of TCR is 1.8°C, the ''likely'' range is 1.4°C to 2.2°C and the ''very likely'' range is 1.2°C to 2.4°C. There is a high level of agreement among the different lines of evidence (Figure TS.16c) ( ''high confidence'' ). Links to chapters 7.5.5

On average, CMIP6 models have higher mean ECS and TCR values than the CMIP5 generation of models and also have higher mean values and wider spreads than the assessed best estimates and ''very likely'' ranges within this Report. These higher mean ECS and TCR values can be traced to a positive net cloud feedback that is larger in CMIP6 by about 20%. The broader ECS and TCR ranges from CMIP6 also lead the models to project a range of future warming that is wider than the assessed future warming range, which is based on multiple lines of evidence (Cross-Section Box TS.1). However, some of the high-sensitivity CMIP6 models (Section TS.1.2.2) are less consistent with observed recent changes in global warming and with paleoclimate proxy records than models with ECS within the ''very likely'' range. Similarly, some of the low-sensitivity models are less consistent with the paleoclimate data. The CMIP6 models with the highest ECS and TCRs values provide insights into low-likelihood, high-impact futures, which cannot be excluded based on currently available evidence (Cross-Section Box TS.1). Links to chapters 4.3.1, 4.3.4, 7.4.2, 7.5.6

Uncertainties regarding the true value of ECS and TCR are the dominant source of uncertainty in global temperature projections over the 21st century under moderate to high GHG concentrations scenarios. For scenarios that reach net zero CO <sub>2</sub> emissions (Section TS.3.3), the uncertainty in the ERF values of aerosol and other SLCFs contribute substantial uncertainty in projected temperature. Global ocean heat uptake is a smaller source of uncertainty in centennial warming. Links to chapters 7.5.7

The transient climate response to cumulative CO <sub>2</sub> emissions (TCRE) is the ratio between globally averaged surface temperature increase and cumulative CO <sub>2</sub> emissions (see Glossary). This Report reaffirms with ''high confidence'' the finding of AR5 that there is a near-linear relationship between cumulative CO <sub>2</sub> emissions and the increase in global average temperature caused by CO <sub>2</sub> over the course of this century for global warming levels up to at least 2°C relative to 1850–1900. The TCRE falls ''likely'' in the 1.0°C–2.3°C per 1000 PgC range, with a best estimate of 1.65°C per 1000 PgC. This is equivalent to a 0.27°C–0.63°C range with a best estimate of 0.45°C when expressed in units per 1000 GtCO <sub>2</sub> . This range is about 15% narrower than the 0.8°–2.5°C per 1000 PgC assessment of AR5 because of a better integration of evidence across chapters, in particular the assessment of TCR. Beyond this century, there is ''low confidence'' that the TCRE alone remains an accurate predictor of temperature changes in scenarios of very low or net negative CO <sub>2</sub> emissions because of uncertain Earth system feedbacks that can result in further changes in temperature or a path dependency of warming as a function of cumulative CO <sub>2</sub> emissions. Links to chapters 4.6.2, 5.4, 5.5.1

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==== TS.3.2.2 Earth System Feedbacks ====

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The combined effect of all climate feedback processes is to amplify the climate response to forcing ( ''virtually certain'' ). While major advances in the understanding of cloud processes have increased the level of confidence and decreased the uncertainty range for the cloud feedback by about 50% compared to AR5, clouds remain the largest contribution to overall uncertainty in climate feedbacks ( ''high confidence'' ). Uncertainties in the ECS and other climate sensitivity metrics, such as the TCR and TCRE, are the dominant source of uncertainty in global temperature projections over the 21st century under moderate to high GHG emissions scenarios. CMIP6 models have higher mean values and wider spreads in ECS and TCR than the assessed best estimates and ''very likely'' ranges within this Report, leading the models to project a range of future warming that is wider than the assessed future warming range (Section TS.2.2). Links to chapters 7.1, 7.4.2, 7.5

Earth system feedbacks can be categorized into three broad groups: physical feedbacks, biogeophysical and biogeochemical feedbacks, and feedbacks associated with ice sheets. In previous assessments, the ECS has been associated with a distinct set of physical feedbacks (Planck response, water vapour, lapse rate, surface albedo, and cloud feedbacks). In this assessment, a more general definition of ECS is adopted whereby all biogeophysical and biogeochemical feedbacks that do not affect the atmospheric concentration of CO <sub>2</sub> are included. These include changes in natural CH <sub>4</sub> emissions, natural aerosol emissions, N <sub>2</sub> O, ozone, and vegetation, which all act on time scales of years to decades and are therefore relevant for temperature change over the 21st century. Because the total biogeophysical and non-CO <sub>2</sub> biogeochemical feedback is assessed to have a central value that is near zero ( ''low confidence'' ), including it does not affect the assessed ECS but does contribute to the net feedback uncertainty. The biogeochemical feedbacks that affect the atmospheric concentration of CO <sub>2</sub> are not included because ECS is defined as the response to a sustained doubling of CO <sub>2</sub> . Moreover, the long-term feedbacks associated with ice sheets are not included in the ECS owing to their long time scales of adjustment. Links to chapters 5.4, 6.4, 7.4, 7.5, Box 7.1

The net effect of changes in clouds in response to global warming is to amplify human-induced warming, that is, the net cloud feedback is positive ( ''high confidence'' ). Compared to AR5, major advances in the understanding of cloud processes have increased the level of confidence and decreased the uncertainty range in the cloud feedback by about 50% (Figure TS.17a). An assessment of the low-altitude cloud feedback over the subtropical ocean, which was previously the major source of uncertainty in the net cloud feedback, is improved owing to a combined use of climate model simulations, satellite observations, and explicit simulations of clouds, altogether leading to strong evidence that this type of cloud amplifies global warming. The net cloud feedback is assessed to be +0.42 [–0.10 to 0.94] W m <sup>–2</sup> °C <sup>–</sup> <sup>1</sup> . A net negative cloud feedback is ''very unlikely'' . The CMIP5 and CMIP6 ranges of cloud feedback are similar to this assessed range, with CMIP6 having a slightly more positive median cloud feedback ( ''high confidence'' ). The surface albedo feedback and combined water vapour-lapse rate feedback are positive (Figure TS.17a), with ''high confidence'' in the estimated value of each based on multiple lines of evidence, including observations, models and theory (Box TS.6). Links to chapters 7.4.2, Figure 7.14, Table 7.10

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'''Figure TS.17 |''' '''An overview of physical and biogeochemical feedbacks in the climate system.''' ''The intent of this figure is to summarize assessed estimates of physical, biogeophysical and biogeochemical feedbacks on global temperature based on Chapters 5, 6 and 7.'' '''(a)''' Synthesis of physical, biogeophysical and non-carbon dioxide (CO <sub>2</sub> ) biogeochemical feedbacks that are included in the definition of equilibrium climate sensitivity (ECS) assessed in this Technical Summary. These feedbacks have been assessed using multiple lines of evidence including observations, models and theory. The net feedback is the sum of the Planck response, water vapour and lapse rate, surface albedo, cloud, and biogeophysical and non-CO <sub>2</sub> biogeochemical feedbacks. Bars denote the mean feedback values, and uncertainties represent ''very likely'' ranges; '''(b)''' Estimated values of individual biogeophysical and non-CO <sub>2</sub> biogeochemical feedbacks. The atmospheric methane (CH <sub>4</sub> ) lifetime and other non-CO <sub>2</sub> biogeochemical feedbacks have been calculated using global Earth system model simulations from AerChemMIP, while the CH <sub>4</sub> and nitrous oxide (N <sub>2</sub> O) source responses to climate have been assessed for the year 2100 using a range of modelling approaches using simplified radiative forcing equations. The estimates represent the mean and 5–95% range. The level of confidence in these estimates is ''low'' owing to the large model spread. '''(c)''' Carbon-cycle feedbacks as simulated by models participating in the C4MIP of the Coupled Model Intercomparison Project Phase 6 (CMIP6). An independent estimate of the additional positive carbon-cycle climate feedbacks from permafrost thaw, which is not considered in most C4MIP models, is added. The estimates represent the mean and 5–95% range. Note that these feedbacks act through modifying the atmospheric concentration of CO <sub>2</sub> and thus are not included in the definition of ECS, which assumes a doubling of CO <sub>2</sub> <sub>, 4</sub> but are included in the definition and assessed range of the transient climate response to cumulative CO <sub>2</sub> emissions (TCRE). Links to chapters 5.4.7, 5.4.8, Box 5.1, Figure 5.29, 6.4.5, Table 6.9, 7.4.2, Table 7.10

Natural sources and sinks of non-CO <sub>2</sub> greenhouse gases such as methane (CH <sub>4</sub> ) and nitrous oxide (N <sub>2</sub> O) respond both directly and indirectly to atmospheric CO <sub>2</sub> concentration and climate change, and thereby give rise to additional biogeochemical feedbacks in the climate system. Many of these feedbacks are only partially understood and are not yet fully included in ESMs. There is ''medium confidence'' that the net response of natural ocean and land CH <sub>4</sub> and N <sub>2</sub> O sources to future warming will be increased emissions, but the magnitude and timing of the responses of each individual process is known with ''low confidence'' . Links to chapters 5.4.7

Non-CO <sub>2</sub> biogeochemical feedbacks induced from changes in emissions, abundances or lifetimes of SLCFs mediated by natural processes or atmospheric chemistry are assessed to decrease ECS (Figure TS.17b). These non-CO <sub>2</sub> biogeochemical feedbacks are estimated from ESMs, which since AR5 have advanced to include a consistent representation of biogeochemical cycles and atmospheric chemistry. However, process-level understanding of many biogeochemical feedbacks involving SLCFs, particularly natural emissions, is still emerging, resulting in ''low confidence'' in the magnitude and sign of the feedbacks. The central estimate of the total biogeophysical and non-CO <sub>2</sub> biogeochemical feedback is assessed to be −0.01 [–0.27 to +0.25] W m <sup>–2</sup> °C <sup>–1</sup> (Figure TS.17a). Links to chapters 5.4.7, 5.4.8, 6.2.2, 6.4.5, 7.4, Table 7.10

The combined effect of all known radiative feedbacks (physical, biogeophysical, and non-CO <sub>2</sub> biogeochemical) is to amplify the base climate response (in the absence of feedbacks), also known as the Planck temperature response <sup>[[#footnote-001|20]]</sup> ( ''virtually certain'' ) '''.''' Combining these feedbacks with the Planck response, the net climate feedback parameter is assessed to be –1.16 [–1.81 to –0.51] W m <sup>–2</sup> °C <sup>–1</sup> , which is slightly less negative than that inferred from the overall ECS assessment. The combined water vapour and lapse rate feedback makes the largest single contribution to global warming, whereas the cloud feedback remains the largest contribution to overall uncertainty. Due to the state-dependence of feedbacks, as evidenced from paleoclimate observations and from models, the net feedback parameter will increase (become less negative) as global temperature increases. Furthermore, on long time scales the ice-sheet feedback parameter is ''very likely'' positive, promoting additional warming on millennial time scales as ice sheets come into equilibrium with the forcing. ( ''high confidence'' ) Links to chapters 7.4.2, 7.4.3, Figure 7.14, Table 7.10

The carbon cycle provides for additional feedbacks on climate owing to the sensitivity of land–atmosphere and ocean–atmosphere carbon fluxes and storage to changes in climate and in atmospheric CO <sub>2</sub> (Figure TS.17c). Because of the time scales associated with land and ocean carbon uptake, these feedbacks are known to be scenario dependent. Feedback estimates deviate from linearity in scenarios of stabilizing or reducing concentrations. With ''high confidence'' , increased atmospheric CO <sub>2</sub> will lead to increased land and ocean carbon uptake, acting as a negative feedback on climate change. It is ''likely'' that a warmer climate will lead to reduced land and ocean carbon uptake, acting as a positive feedback (Box TS.5). Links to chapters 4.3.2, 5.4.1–5

Thawing terrestrial permafrost will lead to carbon release ( ''high confidence'' ), but there is ''low confidence'' in the timing, magnitude and the relative roles of CO <sub>2</sub> versus CH <sub>4</sub> as feedback processes. An ensemble of models projects CO <sub>2</sub> release from permafrost to be 3–41 PgC per 1ºC of global warming by 2100, leading to warming strong enough that it must be included in estimates of the remaining carbon budget but weaker than the warming from fossil fuel burning. However, the incomplete representation of important processes, such as abrupt thaw, combined with weak observational constraints, only allow ''low confidence'' in both the magnitude of these estimates and in how linearly proportional this feedback is to the amount of global warming. There is emerging evidence that permafrost thaw and thermokarst give rise to increased CH <sub>4</sub> and N <sub>2</sub> O emissions, which leads to the combined radiative forcing from permafrost thaw being larger than from CO <sub>2</sub> emissions only. However, the quantitative understanding of these additional feedbacks is low, particularly for N <sub>2</sub> O. These feedbacks, as well as potential additional carbon losses due to climate-induced fire feedback are not routinely included in Earth system models. Links to chapters Box 5.1, 5.4.3, 5.4.7, 5.4.8

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=== TS.3.3 Temperature Stabilization, Net Zero Emissions and Mitigation ===

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==== TS.3.3.1 Remaining Carbon Budgets and Temperature Stabilization ====

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'''The near-linear relationship between cumulative CO <sub>2</sub> emissions and maximum global surface temperature increase caused by CO <sub>2</sub> implies that stabilizing human-induced global temperature increase at any level requires net anthropogenic CO <sub>2</sub> emissions to become zero. This near-linear relationship further implies that mitigation requirements for limiting warming to specific levels can be quantified in terms of a carbon budget ( ''high confidence'' ). Remaining carbon budget estimates have been updated since AR5 with methodological improvements, resulting in larger estimates that are consistent with SR1.5. Several factors, including estimates of historical warming, future emissions from thawing permafrost, variations in projected non-CO <sub>2</sub> warming, and the global surface temperature change after cessation of CO <sub>2</sub> emissions, affect the exact value of carbon budgets ( ''high confidence'' ). Links to chapters 1.3.5, Box 1.2, 4.7.1, 5.5'''
Limiting further climate change would require substantial and sustained reductions of GHG emissions. Without net zero CO <sub>2</sub> emissions, and a decrease in the net non-CO <sub>2</sub> forcing (or sufficient net negative CO <sub>2</sub> emissions to offset any further warming from net non-CO <sub>2</sub> forcing), the climate system will continue to warm. There is ''high confidence'' that mitigation requirements for limiting warming to specific levels over this century can be estimated using a carbon budget that relates cumulative CO <sub>2</sub> emissions to global mean temperature increase (Figure TS.18, Table TS.3). For the period 1850–2019, a total of 2390 ± 240 GtCO <sub>2</sub> of anthropogenic CO <sub>2</sub> has been emitted. Remaining carbon budgets (starting from 1 January 2020) for limiting warming to 1.5°C, 1.7°C and 2.0°C are estimated at 500 GtCO <sub>2</sub> , 850 GtCO <sub>2</sub> and 1350 GtCO <sub>2</sub> , respectively, based on the 50th percentile of TCRE. For the 67th percentile, the respective values are 400 GtCO <sub>2</sub> , 700 GtCO <sub>2</sub> and 1150 GtCO <sub>2</sub> . The remaining carbon budget estimates for different temperature limits assume that non-CO <sub>2</sub> emissions are mitigated consistent with the median reductions found in scenarios in the literature as assessed in SR1.5, but they may vary by an estimated ±220 GtCO <sub>2</sub> depending on how deeply future non-CO <sub>2</sub> emissions are assumed to be reduced (Table TS.3). Links to chapters 5.5.2, 5.6, Box 5.2, 7.6

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'''Figure TS.18 |''' '''Illustration of (a) relationship between cumulative emissions of carbon dioxide (CO''' 2 ''') and global mean surface air temperature increase and (b) the assessment of the remaining carbon budget from its constituting components based on multiple lines of evidence.''' ''The intent of this figure is to show (i) the proportionality between cumulative CO'' 2 ''emissions and global surface air temperature in observations and models as well as the assessed range of the transient climate response to cumulative CO'' 2 ''emissions (TCRE), and (ii) how information is combined to derive remaining carbon budgets consistent with limiting warming to a specific level.'' Carbon budgets consistent with various levels of additional warming are provided in Table 5.8 and should not be read from the illustrations in either panel. In panel (a) thin black line shows historical CO <sub>2</sub> emissions together with the assessed global surface temperature increase from 1850–1900 as assessed in [[IPCC:Wg1:Chapter:Chapter-2|Chapter 2]] (Box 2.3). The orange-brown range with its central line shows the estimated human-induced share of historical warming. The vertical orange-brown line shows the assessed range of historical human-induced warming for the 2010–2019 period relative to 1850–1900 (Chapter 3). The grey cone shows the assessed ''likely'' range for the TCRE ( [[IPCC:Wg1:Chapter:Chapter-5#5.5.1.4|Section 5.5.1.4]] ), starting from 2015. Thin coloured lines show Coupled Model Intercomparison Project Phase 6 (CMIP6) simulations for the five scenarios of the WGI core set (SSP1-1.9, light blue; SSP1-2.6, blue; SSP2-4.5, yellow; SSP3-7.0, red; SSP5-8.5, maroon), starting from 2015 and until 2100. Diagnosed carbon emissions are complemented with estimated land-use change emissions for each respective scenario. Coloured areas show the [[IPCC:Wg1:Chapter:Chapter-4|Chapter 4]] assessed ''very likely'' range of global surface temperature projections and thick coloured central lines show the median estimate, for each respective scenario. These temperature projections are expressed relative to cumulative CO <sub>2</sub> emissions that are available for emissions-driven CMIP6 ScenarioMIP experiments for each respective scenario. For panel (b), the remaining allowable warming is estimated by combining the global warming limit of interest with the assessed historical human-induced warming ( [[IPCC:Wg1:Chapter:Chapter-5#5.5.2.2.2|Section 5.5.2.2.2]] ), the assessed future potential non-CO <sub>2</sub> warming contribution ( [[IPCC:Wg1:Chapter:Chapter-5#5.5.2.2.3|Section 5.5.2.2.3]] ) and the zero emissions commitment (ZEC; [[IPCC:Wg1:Chapter:Chapter-5#5.5.2.2.4|Section 5.5.2.2.4]] ). The remaining allowable warming (vertical blue bar) is subsequently combined with the assessed TCRE (Sections 5.5.1.4 and 5.5.2.2.1) and contribution of unrepresented Earth system feedbacks ( [[IPCC:Wg1:Chapter:Chapter-5#5.5.2.2.5|Section 5.5.2.2.5]] ) to provide an assessed estimate of the remaining carbon budget (horizontal blue bar, Table 5.8). Note that contributions in panel (b) are illustrative and are not to scale. For example, the central ZEC estimate was assessed to be zero. Links to chapters Box 2.3, 5.2.1, 5.2.2, Figure 5.31

There is ''high confidence'' that several factors, including estimates of historical warming, future emissions from thawing permafrost, and variations in projected non-CO <sub>2</sub> warming, affect the value of carbon budgets but do not change the conclusion that global CO <sub>2</sub> emissions would need to decline to net zero to halt global warming. Estimates may vary by ±220 GtCO <sub>2</sub> depending on the level of non-CO <sub>2</sub> emissions at the time global anthropogenic CO <sub>2</sub> emissions reach net zero levels. This variation is referred to as non-CO <sub>2</sub> scenario uncertainty and will be further assessed in the AR6 Working Group III Contribution. Geophysical uncertainties surrounding the climate response to these non-CO <sub>2</sub> emissions result in an additional uncertainty of at least ±220 GtCO <sub>2</sub> , and uncertainties in the level of historical warming result in a ±550 GtCO <sub>2</sub> uncertainty. Links to chapters 5.4, 5.5.2

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'''Table TS.3 |''' '''Estimates of remaining carbon budgets and their uncertainties.''' Assessed estimates are provided for additional human-induced warming, expressed as global surface temperature, since the recent past (2010–2019), ''likely'' amounted to 0.8° to 1.3°C with a best estimate of 1.07°C relative to 1850–1900. Historical CO <sub>2</sub> emissions between 1850 and 2014 have been estimated at about 2180 ± 240 GtCO <sub>2</sub> (1-sigma range), while since 1 January 2015, an additional 210 GtCO <sub>2</sub> has been emitted until the end of 2019. GtCO <sub>2</sub> values to the nearest 50. Links to chapters Table 3.1, 5.5.1, 5.5.2, Box 5.2, Table 5.1, Table 5.7, Table 5.8

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<sup>a</sup> Human-induced global surface temperature increase between 1850–1900 and 2010–2019 is assessed at 0.8–1.3°C ( ''likely'' range; Cross-Section Box TS.1) with a best estimate of 1.07°C. Combined with a central estimate of TCRE (1.65°C per 1000 PgC) this uncertainty in isolation results in a potential variation of remaining carbon budgets of ±550 GtCO <sub>2</sub> , which, however, is not independent of the assessed uncertainty of TCRE and thus not fully additional.

<sup>b</sup> TCRE: transient climate response to cumulative emissions of carbon dioxide, assessed to fall ''likely'' between 1.0–2.3°C per 1000 PgC with a normal distribution, from which the percentiles are taken. Additional Earth system feedbacks are included in the remaining carbon budget estimates as discussed in [[IPCC:Wg1:Chapter:Chapter-5#5.5.2.2.5|Section 5.5.2.2.5]] .

<sup>c</sup> Estimates assume that non-CO <sub>2</sub> emissions are mitigated consistent with the median reductions found in scenarios in the literature as assessed in SR1.5. Non-CO <sub>2</sub> scenario variations indicate how much remaining carbon budget estimates vary due to different scenario assumptions related to the future evolution of non-CO <sub>2</sub> emissions in mitigation scenarios from SR1.5 that reach net zero CO <sub>2</sub> emissions. This variation is additional to the uncertainty in TCRE. The Working Group III Contribution to AR6 will reassess the potential for non-CO <sub>2</sub> mitigation based on literature since SR1.5.

<sup>d</sup> Geophysical uncertainties reported in these columns and TCRE uncertainty are not statistically independent, as uncertainty in TCRE depends on uncertainty in the assessment of historical temperature, non-CO <sub>2</sub> versus CO <sub>2</sub> forcing, and uncertainty in emissions estimates. These estimates cannot be formally combined, and these uncertainty variations are not directly additional to the spread of remaining carbon budgets due to TCRE uncertainty reported in columns three to seven.

<sup>e</sup> Recent emissions uncertainty reflects the ±10% uncertainty in the historical CO <sub>2</sub> emissions estimate since 1 January 2015.

Methodological improvements and new evidence result in updated remaining carbon budget estimates. The assessment in AR6 applies the same methodological improvements as in SR1.5, which uses a recent observed baseline for historic temperature change and cumulative emissions. Changes compared to SR1.5 are therefore small: the assessment of new evidence results in updated median remaining carbon budget estimates for limiting warming to 1.5°C and 2°C being the same and about 60 GtCO <sub>2</sub> smaller, respectively, after accounting for emissions since SR1.5. Meanwhile, remaining carbon budgets for limiting warming to 1.5°C would be about 300–350 GtCO <sub>2</sub> larger if evidence and methods available at the time of AR5 would be used. If a specific remaining carbon budget is exceeded, this results in a lower probability of keeping warming below a specified temperature level and higher irreversible global warming over decades to centuries, or alternatively a need for net negative CO <sub>2</sub> emissions or further reductions in non-CO <sub>2</sub> greenhouse gases after net zero CO <sub>2</sub> is achieved to return warming to lower levels in the long term. Links to chapters 5.5.2, 5.6, Box 5.2

Based on idealized model simulations that explore the climate response once CO <sub>2</sub> emissions have been brought to zero, the magnitude of the zero CO <sub>2</sub> emissions commitment (ZEC, see Glossary) is assessed to be ''likely'' smaller than 0.3°C for time scales of about half a century and cumulative CO <sub>2</sub> emissions broadly consistent with global warming of 2°C. However, there is ''low confidence'' about its sign on time scales of about half a century. For lower cumulative CO <sub>2</sub> emissions, the range would be smaller yet with equal uncertainty about the sign. If the ZEC is positive on decadal time scales, additional warming leads to a reduction in the estimates of remaining carbon budgets, and vice versa if it is negative. Links to chapters 4.7.1, 5.5.2

Permafrost thaw is included in estimates together with other feedbacks that are often not captured by models. Limitations in modelling studies combined with weak observational constraints only allow ''low confidence'' in the magnitude of these estimates (Section TS.3.2.2). Despite the large uncertainties surrounding the quantification of the effect of additional Earth system feedback processes, such as emissions from wetlands and permafrost thaw, these feedbacks represent identified additional risk factors that scale with additional warming and mostly increase the challenge of limiting warming to specific temperature levels. These uncertainties do not change the basic conclusion that global CO <sub>2</sub> emissions would need to decline to net zero to halt global warming. Links to chapters 5.4.8, 5.5.2, Box 5.1

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<span id="ts.3.3.2-carbon-dioxide-removal"></span>
==== TS.3.3.2 Carbon Dioxide Removal ====

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'''Deliberate carbon dioxide removal (CDR) from the atmosphere has the potential to compensate for residual CO <sub>2</sub> emissions to reach net zero CO <sub>2</sub> emissions or to generate net negative CO <sub>2</sub> emissions. In the same way that part of current anthropogenic net CO <sub>2</sub> emissions are taken up by land and ocean carbon stores, net CO <sub>2</sub> removal will be partially counteracted by CO <sub>2</sub> release from these stores ( ''very high confidence'' ). Asymmetry in the carbon cycle response to simultaneous CO <sub>2</sub> emissions and removals implies that a larger amount of CO <sub>2</sub> would need to be removed to compensate for an emission of a given magnitude to attain the same change in atmospheric CO <sub>2</sub> ( ''medium confidence'' ). CDR methods have wide-ranging side-effects that can either weaken or strengthen the carbon sequestration and cooling potential of these methods and affect the achievement of sustainable development goals ( ''high confidence'' ). Links to chapters 4.6.3, 5.6'''
Carbon dioxide removal (CDR) refers to anthropogenic activities that deliberately remove CO <sub>2</sub> from the atmosphere and durably store it in geological, terrestrial or ocean reservoirs, or in products. Carbon dioxide is removed from the atmosphere by enhancing biological or geochemical carbon sinks or by direct capture of CO <sub>2</sub> from air. Emissions pathways that limit global warming to 1.5°C or 2°C typically assume the use of CDR approaches in combination with GHG emissions reductions. CDR approaches could be used to compensate for residual emissions from sectors that are difficult or costly to decarbonize. CDR could also be implemented at a large scale to generate global net negative CO <sub>2</sub> emissions (i.e., anthropogenic CO <sub>2</sub> removals exceeding anthropogenic emissions), which could compensate for earlier emissions as a way to meet long-term climate stabilization goals after a temperature overshoot. This Report assesses the effects of CDR on the carbon cycle and climate. Co-benefits and trade-offs for biodiversity, water and food production are briefly discussed for completeness, but a comprehensive assessment of the ecological and socio-economic dimensions of CDR options is left to the WGII and WGIII reports. Links to chapters 4.6.3, 5.6

CDR methods have the potential to sequester CO <sub>2</sub> from the atmosphere ( ''high confidence'' ). In the same way part of current anthropogenic net CO <sub>2</sub> emissions are taken up by land and ocean carbon stores, net CO <sub>2</sub> removal will be partially counteracted by CO <sub>2</sub> release from these stores, such that the amount of CO <sub>2</sub> sequestered by CDR will not result in an equivalent drop in atmospheric CO <sub>2</sub> ( ''very'' ''high confidence'' ). The fraction of CO <sub>2</sub> removed from the atmosphere that is not replaced by CO <sub>2</sub> released from carbon stores – a measure of CDR effectiveness – decreases slightly with increasing amounts of removal ( ''medium confidence'' ) and decreases strongly if CDR is applied at lower atmospheric CO <sub>2</sub> concentrations ( ''medium confidence'' ). The reduction in global surface temperature is approximately linearly related to cumulative CO <sub>2</sub> removal ( ''high confidence'' ). Because of this near-linear relationship, the amount of cooling per unit CO <sub>2</sub> removed is approximately independent of the rate and amount of removal ( ''medium confidence'' ). Links to chapters 4.6.3, 5.6.2.1, Figure 5.32, Figure 5.34

Due to non-linearities in the climate system, the century-scale climate–carbon cycle response to a CO <sub>2</sub> removal from the atmosphere is not always equal and opposite to its response to a simultaneous CO <sub>2</sub> emission ( ''medium'' confidence). For CO <sub>2</sub> emissions of 100 PgC released from a state in equilibrium with pre-industrial atmospheric CO <sub>2</sub> levels, CMIP6 models simulate that 27± 6% (mean ± 1 standard deviation) of emissions remain in the atmosphere 80–100 years after the emissions, whereas for removals of 100 PgC only 23 ± 6% of removals remain out of the atmosphere. This asymmetry implies that an extra amount of CDR is required to compensate for a positive emission of a given magnitude to attain the same change in atmospheric CO <sub>2</sub> . Due to ''low agreement'' between models, there is ''low confidence'' in the sign of the asymmetry of the temperature response to CO <sub>2</sub> emissions and removals. Links to chapters 4.6.3, 5.6.2.1, Figure 5.35

Simulations with ESMs indicate that under scenarios where CO <sub>2</sub> emissions gradually decline, reach net zero and become net negative during the 21st century (e.g., SSP1-2.6), land and ocean carbon sinks begin to weaken in response to declining atmospheric CO <sub>2</sub> concentrations, and the land sink eventually turns into a source (Figure TS.19). This sink-to-source transition occurs decades to a few centuries after CO <sub>2</sub> emissions become net negative. The ocean remains a sink of CO <sub>2</sub> for centuries after emissions become net negative. Under scenarios with large net negative CO <sub>2</sub> emissions (e.g., SSP5-3.4-OS) and rapidly declining CO <sub>2</sub> concentrations, the land source is larger than for SSP1-2.6 and the ocean also switches to a source. While the general response is robust across models, there is ''low confidence'' in the timing of the sink-to-source transition and the magnitude of the CO <sub>2</sub> source in scenarios with net negative CO <sub>2</sub> emissions. Carbon dioxide removal could reverse some aspects climate change if CO <sub>2</sub> emissions become net negative, but some changes would continue in their current direction for decades to millennia. For instance, sea level rise due to ocean thermal expansion would not reverse for several centuries to millennia ( ''high confidence'' ) (Box TS.4). Links to chapters 4.6.3, 5.4.10, 5.6.2.1, Figure 5.30, Figure 5.33

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'''Figure TS.19 | Carbon sink response in a scenario with net carbon dioxide (CO''' 2 ''') removal from the atmosphere.''' ''The intent of this figure is to show how atmospheric CO'' 2 ''evolves under negative emissions and its dependence on the negative emissions technologies. It also shows the evolution of the ocean and land sinks.'' Shown are CO 2 flux components from concentration-driven Earth system model (ESM) simulations during different emissions stages of SSP1–2.6 and its long-term extension. (a) Large net positive CO <sub>2</sub> emissions, (b) small net positive CO <sub>2</sub> emissions, (c–d) net negative CO <sub>2</sub> emissions, and (e) net zero CO <sub>2</sub> emissions. Positive flux components act to raise the atmospheric CO <sub>2</sub> concentration, whereas negative components act to lower the CO <sub>2</sub> concentration. Net CO <sub>2</sub> emissions and land and ocean CO <sub>2</sub> fluxes represent the multi-model mean and standard deviation (error bar) of four ESMs (CanESM5, UKESM1, CESM2-WACCM, IPSL-CM6a-LR) and one Earth system model of intermediate complexity (Uvic ESCM). Net CO <sub>2</sub> emissions are calculated from concentration-driven ESM simulations as the residual from the rate of increase in atmospheric CO <sub>2</sub> and land and ocean CO <sub>2</sub> fluxes. Fluxes are accumulated over each 50-year period and converted to concentration units (parts per million, or ppm). Links to chapters 5.6.2.1, Figure 5.33

Carbon dioxide removal methods have a range of side effects that can either weaken or strengthen the carbon sequestration and cooling potential of these methods and affect the achievement of sustainable development goals ( ''high confidence'' ). Biophysical and biogeochemical side-effects of CDR methods are associated with changes in surface albedo, the water cycle, emissions of CH <sub>4</sub> and N <sub>2</sub> O, ocean acidification and marine ecosystem productivity ( ''high confidence'' ). These side-effects and associated Earth system feedbacks can decrease carbon uptake and/or change local and regional climate and in turn limit the CO <sub>2</sub> sequestration and cooling potential of specific CDR methods ( ''medium confidence'' ). Deployment of CDR, particularly on land, can also affect water quality and quantity, food production and biodiversity ( ''high confidence'' ). These effects are often highly dependent on local context, management regime, prior land use, and scale ( ''high confidence'' ). The largest co-benefits are obtained with methods that seek to restore natural ecosystems or improve soil carbon sequestration ( ''medium confidence'' ). The climate and biogeochemical effects of terminating CDR are expected to be small for most CDR methods ( ''medium confidence'' ). Links to chapters 4.6.3, 5.6.2.2, Figure 5.36, 8.4.3, 8.6.3

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==== TS.3.3.3 Relating Different Forcing Agents ====

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'''When including other GHGs, the choice of emissions metric affects the quantification of net zero GHG emissions and their resulting temperature outcome ( ''high confidence'' ). Reaching and sustaining net zero GHG emissions typically leads to a peak and decline in temperatures when quantified with the global warming potential over a 100-year period (GWP-100). Carbon-cycle responses are more robustly accounted for in emissions metrics compared to AR5 ( ''high confidence'' ). New emissions metric approaches can be used to generate equivalent cumulative emissions of CO <sub>2</sub> for short-lived greenhouse gases based on their rate of emissions. Links to chapters 7.6.2''' '''Over 10- to 20-year time scales, the temperature response to a single year’s worth of current emissions of short-lived climate forcers (SLCFs) is at least as large as that of CO <sub>2</sub> , but because the effects of SLCFs decay rapidly over the first few decades after emission, the net long-term temperature response to a single year’s worth of emissions is predominantly determined by cumulative CO <sub>2</sub> emissions.''' '''Emissions reductions in 2020 associated with COVID-19 containment led to small and positive global ERF; however, global and regional climate responses to the forcing are undetectable above internal variability due to the temporary nature of emissions reductions. Links to chapters 6.6, Cross-Chapter Box 6.1'''
The relative climate effects of different forcing agents are typically quantified using emissions metrics that compare the effects of an idealised pulse of 1 kg of some climate forcing agent against a reference climate forcing agent, almost always CO <sub>2</sub> . The two most prominent pulse emissions metrics are the global warming potential (GWP) and global temperature change potential (GTP) (see Glossary). The climate responses to CO <sub>2</sub> emissions by convention include the effects of warming on the carbon cycle, so for consistency these also need to be determined for non-CO <sub>2</sub> emissions. The methodology for doing this has been placed on a more robust scientific footing compared to AR5 ( ''high confidence'' ). Methane from fossil fuel sources has slightly higher emissions metric values than those from biogenic sources since it leads to additional fossil CO <sub>2</sub> in the atmosphere ( ''high confidence'' ). Updates to the chemical adjustments for CH <sub>4</sub> and N <sub>2</sub> O emissions (Section TS.3.1) and revisions in their lifetimes result in emissions metrics for GWP and GTP that are slightly lower than in AR5 ( ''medium confidence'' ). Emissions metrics for the entire suite of GHGs assessed in the AR6 have been calculated for various time horizons. Links to chapters 7.6.1, Table 7.15, Table 7.SM.7

New emissions metric approaches, such as GWP* and Combined-GTP (CGTP), relate changes in the emissions rate of short-lived greenhouse gases to equivalent cumulative emissions of CO <sub>2</sub> (CO <sub>2</sub> -e). Global surface temperature response from aggregated emissions of short-lived greenhouse gases over time is determined by multiplying these cumulative CO <sub>2</sub> -e by TCRE (see Section TS.3.2.1). When GHGs are aggregated using standard metrics such as GWP or GTP, cumulative CO <sub>2</sub> -e emissions are not necessarily proportional to future global surface temperature outcomes ( ''high confidence'' ) Links to chapters 7.6.1, Box 7.3

Emissions metrics are needed to aggregate baskets of gases to determine net zero GHG emissions. Generally, achieving net zero CO <sub>2</sub> emissions and declining non-CO <sub>2</sub> radiative forcing would halt human-induced warming. Reaching net zero GHG emissions quantified by GWP-100 typically leads to declining temperatures after net zero GHGs emissions are achieved if the basket includes short-lived gases, such as CH <sub>4</sub> . Net zero GHG emissions defined by CGTP or GWP* imply net zero CO <sub>2</sub> and other long-lived GHG emissions and constant (CGTP) or gradually declining (GWP*) emissions of short-lived gases. The warming evolution resulting from net zero GHG emissions defined in this way corresponds approximately to reaching net zero CO <sub>2</sub> emissions, and would thus not lead to declining temperatures after net zero GHG emissions are achieved but to an approximate temperature stabilization ( ''high confidence'' ). The choice of emissions metric hence affects the quantification of net zero GHG emissions, and therefore the resulting temperature outcome of reaching and sustaining net zero GHG emissions levels ( ''high confidence'' ). Links to chapters 7.6.1.4, 7.6.2, 7.6.3

As pointed out in AR5, ultimately, it is a matter for policymakers to decide which emissions metric is most applicable to their needs. This Report does not recommend the use of any specific emissions metric, as the most appropriate metric depends on the policy goal and context (see Chapter 7, [[IPCC:Wg1:Chapter:Chapter-7#7.6|Section 7.6]] ). A detailed assessment of GHG metrics to support climate change mitigation and associated policy contexts is provided in the WGIII contribution to the AR6.

The global surface temperature response following a climate change mitigation measure that affects emissions of both short- and long-lived climate forcers depends on their lifetimes, their ERFs, how fast and for how long the emissions are reduced, and the thermal inertia in the climate system. Mitigation, relying on emissions reductions and implemented through new legislation or technology standards, implies that emissions reductions occur year after year. Global temperature response to a year’s worth of current emissions from different sectors informs about the mitigation potential (Figure TS.20). Over 10- to 20-year time scales, the influence of SLCFs is at least as large as that of CO <sub>2</sub> , with sectors producing the largest warming being fossil fuel production and distribution, agriculture, and waste management. Because the effects of the SLCFs decay rapidly over the first few decades after emission, the net long-term temperature effect from a single year’s worth of current emissions is predominantly determined by CO <sub>2</sub> . Fossil fuel combustion for energy, industry and land transportation are the largest contributing sectors on a 100-year time scale ( ''high confidence'' ). Current emissions of CO <sub>2</sub> , N <sub>2</sub> O and SLCFs from East Asia and North America are the largest regional contributors to additional net future warming on both short ( ''medium confidence'' ) and long time scales (10 and 100 years, respectively) ( ''high confidence'' ). Links to chapters 6.6.1, 6.6.2, Figure 6.16

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'''Figure TS.20 |''' '''Global surface temperature change 10 and 100 years after a one-year pulse of present-day emissions.''' ''The intent of this figure is to show the sectoral contribution to present-day climate change by specific climate forcers, including carbon dioxide (CO'' 2 '') as well as short-lived climate forcers (SLCFs).'' The temperature response is broken down by individual species and shown for total anthropogenic emissions (top) , and sectoral emissions on 10-year (left) and 100-year time scales (right) . Sectors are sorted by (high-to-low) net temperature effect on the 10-year time scale. Error bars in the top panel show the 5–95% range in net temperature effect due to uncertainty in radiative forcing only (calculated using a Monte Carlo approach and best estimate uncertainties from the literature). Emissions for 2014 are from the Coupled Model Intercomparison Project Phase 6 (CMIP6) emissions dataset, except for hydrofluorocarbons (HFCs) and aviation H 2 O, which rely on other datasets (see Section 6.6.2 for more details). CO <sub>2</sub> emissions are excluded from open biomass burning and residential biofuel use. Links to chapters 6.6.2, Figure 6.16

COVID-19 restrictions led to detectable reductions in global anthropogenic emissions of nitrogen oxides (NO x ) (about 35% in April 2020) and fossil CO <sub>2</sub> (7%, with estimates ranging from 5.8% to 13.0%), driven largely by reduced emissions from the transportation sector ( ''medium confidence'' ). There is ''high confidence'' that, with the exception of surface ozone, reductions in pollutant precursors contributed to temporarily improved air quality in most regions of the world. However, these reductions were lower than what would be expected from sustained implementation of policies addressing air quality and climate change ( ''medium confidence'' ). Overall, the net global ERF from COVID-19 containment was likely small and positive for 2020 (with a temporary peak value less than 0.2 W m <sup>–2</sup> ), thus temporarily adding to the total anthropogenic climate influence, with positive forcing (warming influence) from aerosol changes dominating over negative forcings (cooling influence) from CO <sub>2</sub> , NO ''x'' and contrail cirrus changes. Consistent with this small net radiative forcing, and against a large component of internal variability, Earth system models show no detectable effect on global or regional surface temperature or precipitation ( ''high confidence'' ). Links to chapters Cross Chapter Box 6.1

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'''Box TS.7 | Climate and Air Quality Responses to Short-lived Climate Forcers in Shared''' '''Socio-economic''' '''Pathways'''

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'''Future changes in emissions of short-lived climate forcers (SLCFs) are expected to cause an additional global mean warming, with a large diversity in the end-of-century response across the WGI core set of Shared Socio-economic Pathways (SSPs), depending upon the level of climate change and air pollution mitigation (Box TS.7, Figure 1). This additional warming is either due to reductions in cooling aerosols for air pollution regulation or due to increases in methane (CH <sub>4</sub> ), ozone and hydrofluorocarbons (HFCs). This additional warming is stable after 2040 in SSPs associated with lower global air pollution as long as CH <sub>4</sub> emissions are also mitigated, but the overall warming induced by SLCF changes is higher in scenarios in which air quality continues to deteriorate (induced by growing fossil fuel use and limited air pollution control) ( ''high confidence'' ).''' '''Sustained CH <sub>4</sub> mitigation reduces global surface ozone, contributing to air quality improvements, and also reduces surface temperature in the longer term, but only sustained CO <sub>2</sub> emissions reductions allow long-term climate stabilization ( ''high confidence'' ). Future changes in air quality (near-surface ozone and particulate matter, or PM) at global and local scales are predominantly driven by changes in ozone and aerosol precursor emissions rather than climate ( ''high confidence'' ). Air quality improvements driven by rapid decarbonization strategies, as in SSP1-1.9 and SSP1-2.6, are not sufficient in the near term to achieve air quality guidelines set by the World Health Organization in some highly polluted regions ( ''high confidence'' ). Additional policies (e.g., access to clean energy, waste management) envisaged to attain United Nations Sustainable Development Goals bring complementary SLCF reduction. Links to chapters 4.4.4, 6.6.3, 6.7.3, Box 6.2'''
The net effect of SLCF emissions changes on temperature will depend on how emissions of warming and cooling SLCFs will evolve in the future. The magnitude of the cooling effect of aerosols remains the largest uncertainty in the effect of SLCFs in future climate projections. Since the SLCFs have undergone large changes over the past two decades, the temperature and air pollution responses are estimated relative to the year 2019 instead of 1995–2014.

'''Temperature Response'''
In the next two decades, it is ''very likely'' that SLCF emissions changes will cause a warming relative to 2019, across the WGI core set of SSPs (see Section TS.1.3.1), in addition to the warming from long-lived GHGs. The net effect of SLCF and HFC changes in global surface temperature across the SSPs is a ''likely'' warming of 0.06°C–0.35°C in 2040 relative to 2019. This near-term global mean warming linked to SLCFs is quite similar in magnitude across the SSPs due to competing effects of warming (CH <sub>4</sub> , ozone) and cooling (aerosols) forcers (Box TS.7, Figure 1). There is greater diversity in the end-of-century response among the scenarios. SLCF changes in scenarios with no climate change mitigation (SSP3-7.0 and SSP5-8.5) will cause a warming in the ''likely'' range of 0.4°C–0.9°C in 2100 relative to 2019 due to increases in CH <sub>4</sub> , tropospheric ozone and HFC levels. For the stringent climate change and pollution mitigation scenarios (SSP1-1.9 and SSP1-2.6), the cooling from reductions in CH <sub>4</sub> , ozone and HFCs partially balances the warming from reduced aerosols, primarily sulphate, and the overall SLCF effect is a ''likely'' increase in global surface temperature of 0.0°C–0.3°C in 2100, relative to 2019. With intermediate climate change and air pollution mitigations, SLCFs in SSP2-4.5 add a ''likely'' warming of 0.2°C–0.5°C to global surface temperature change in 2100, with the largest warming resulting from reductions in aerosols. Links to chapters 4.4.4, 6.7.3

Assuming implementation and efficient enforcement of both the Kigali Amendment to the Montreal Protocol on Substances that Deplete the Ozone Layer and current national plans result in limiting emissions (as in SSP1-2.6), the effects of HFCs on global surface temperature, relative to 2019, would remain below +0. 02°C from 2050 onwards versus about +0.04°C–0.08°C in 2050 and +0.1°C–0.3°C in 2100 considering only national HFC regulations decided prior to the Kigali Amendment (as in SSP5-8.5) ( ''medium confidence'' ). Links to chapters 6.6.3, 6.7.3

'''Air Quality Responses'''
Air pollution projections range from strong reductions in global surface ozone and PM (e.g., SSP1-2.6, with stringent mitigation of both air pollution and climate change) to no improvement and even degradation (e.g., SSP3-7.0 without climate change mitigation and with only weak air pollution control) ( ''high confidence'' ). Under the SSP3-7.0 scenario, PM levels are projected to increase until 2050 over large parts of Asia, and surface ozone pollution is projected to worsen over all continental areas through 2100 ( ''high confidence'' ). In SSP5-8.5, a scenario without climate change mitigation but with stringent air pollution control, PM levels decline through 2100, but high CH <sub>4</sub> levels hamper the decline in global surface ozone at least until 2080 ( ''high confidence'' ). Links to chapters 6.7.1

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'''Box TS.7, Figure 1 |'''

'''Effects of short-lived climate forcers (SLCFs) on global surface temperature and air pollution across the WGI core set of Shared Socio-economic Pathways (SSPs).''' ''The intent of this figure is to show the climate and air quality (surface ozone and particulate matter smaller than 2.5 microns in diameter, or PM'' ''2.5'' '') response to SLCFs in the SSP scenarios for the near and long-term.'' Effects of net aerosols, tropospheric ozone, hydrofluorocarbons (HFCs; with lifetimes less than 50 years), and methane (CH <sub>4</sub> ) are compared with those of total anthropogenic forcing for 2040 and 2100 relative to year 2019. The global surface temperature changes are based on historical and future evolution of effective radiative forcing (ERF) as assessed in [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] of this Report. The temperature responses to the ERFs are calculated with a common impulse response function (RT) for the climate response, consistent with the metric calculations in [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] (Box 7.1). The RT has an equilibrium climate sensitivity of 3.0°C for a doubling of atmospheric CO <sub>2</sub> concentration (feedback parameter of –1.31 W m <sup>–2</sup> °C <sup>–1</sup> ). The scenario total (grey bar) includes all anthropogenic forcings (long- and short-lived climate forcers, and land-use changes). Uncertainties are 5–95% ranges. The global changes in air pollutant concentrations (ozone and PM 2.5 ) are based on multimodel Coupled Model Intercomparison Project Phase 6 (CMIP6) simulations and represent changes in five-year mean surface continental concentrations for 2040 and 2098 relative to 2019. Uncertainty bars represent inter-model ±1 standard deviation. Links to chapters 6.7.2, 6.7.3, Figure 6.24

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'''Box TS.8 | Earth System Response to Solar Radiation Modification'''

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'''Since AR5, further modelling work has been conducted on aerosol-based solar radiation modification (SRM) options such as stratospheric aerosol injection, marine cloud brightening, and cirrus cloud thinning <sup>[[#footnote-000|21]]</sup> and their climate and biogeochemical effects. These investigations have consistently shown that SRM could offset some of the effects of increasing greenhouse gases on global and regional climate, including the carbon and water cycles ( ''high confidence'' ). However, there would be substantial residual or overcompensating climate change at the regional scales and seasonal time scales ( ''high confidence'' ), and large uncertainties associated with aerosol–cloud–radiation interactions persist. The cooling caused by SRM would increase the global land and ocean CO <sub>2</sub> sinks ( ''medium confidence'' ), but this would not stop CO <sub>2</sub> from increasing in the atmosphere or affect the resulting ocean acidification under continued anthropogenic emissions ( ''high confidence'' ). It is ''likely'' that abrupt water cycle changes will occur if SRM techniques are implemented rapidly. A sudden and sustained termination of SRM in a high CO <sub>2</sub> emissions scenario would cause rapid climate change ( ''high confidence'' ). However, a gradual phase-out of SRM combined with emissions reduction and carbon dioxide removal (CDR) would avoid these termination effects ( ''medium confidence'' ). Links to chapters 4.6.3, 5.6.3. 6.4.6, 8.6.3 .'''
Solar radiation modification (SRM) refers to deliberate, large-scale climate intervention options that are studied as potential supplements to deep mitigation, for example, in scenarios that overshoot climate stabilization goals. SRM options aim to offset some of the warming effects of GHG emissions by modification of Earth’s shortwave radiation budget. Following SR1.5, the SRM assessed in this Report also includes some options, such as cirrus cloud thinning, that alter the longwave radiation budget.

SRM contrasts with climate change mitigation activities, such as emissions reductions and CDR, as it introduces a ‘mask’ to the climate change problem by altering Earth’s radiation budget, rather than attempting to address the root cause of the problem, which is the increase in GHGs in the atmosphere. By masking only the climate effects of GHG emissions, SRM does not address other issues related to atmospheric CO <sub>2</sub> increase, such as ocean acidification. This Report assesses physical understanding of the Earth system response to proposed SRM, and the assessment is based primarily on idealized climate model simulations. There are other important considerations, such as risk to human and natural systems, perceptions, ethics, cost, governance, and trans-boundary issues and their relationship to the United Nations Sustainable Development Goals – issues that the WGII (Chapter 16) and WGIII (Chapter 14) Reports address. Links to chapters 4.6.3

SRM options include those that increase surface albedo, brighten marine clouds by increasing the amount of cloud condensation nuclei, or reduce the optical depth of cirrus clouds by seeding them with ice nucleating particles. However, the most commonly studied approaches attempt to mimic the cooling effects of major volcanic eruptions by injecting reflective aerosols (e.g., sulphate aerosols) or their precursors (e.g., sulphur dioxide) into the stratosphere. Links to chapters 4.6.3, 5.6.3, 6.4.6

SRM could offset some effects of greenhouse gas-induced warming on global and regional climate, but there would be substantial residual and overcompensating climate change at the regional scale and seasonal time scales ( ''high confidence'' ). Since AR5, more modelling work has been conducted with more sophisticated treatment of aerosol-based SRM approaches, but the uncertainties in cloud–aerosol–radiation interactions are still large ( ''high confidence'' ). Modelling studies suggest that it is possible to stabilize multiple large-scale temperature indicators simultaneously by tailoring the deployment strategy of SRM options ( ''medium confidence'' ) but with large residual or overcompensating regional and seasonal climate changes. Links to chapters 4.6.3

SRM approaches targeting shortwave radiation are ''likely'' to reduce global mean precipitation, relative to future CO <sub>2</sub> emissions scenarios, if all global mean warming is offset. In contrast, cirrus cloud thinning, targeting longwave radiation, is expected to cause an increase in global mean precipitation ( ''medium confidence'' ). If shortwave approaches are used to offset global mean warming, the magnitude of reduction in regional precipitation minus evapotranspiration (P–E) (Box TS.5), which is more relevant to freshwater availability, is smaller than precipitation decrease because of simultaneous reductions in both precipitation and evapotranspiration ( ''medium confidence'' ). Links to chapters 4.6.3, 8.2.1, 8.6.3 .

If SRM is used to cool the planet, it would cause a reduction in plant and soil respiration and slow the reduction of ocean carbon uptake due to warming ( ''medium confidence'' ). The result would be an enhancement of the global land and ocean CO <sub>2</sub> sinks ( ''medium confidence'' ) and a slight reduction in atmospheric CO <sub>2</sub> concentration relative to unmitigated climate change. However, SRM would not stop CO <sub>2</sub> from increasing in the atmosphere or affect the resulting ocean acidification under continued anthropogenic emissions ( ''high confidence'' ). Links to chapters 5.6.3

The effect of stratospheric aerosol injection on global temperature and precipitation is projected by models to be detectable after one to two decades, which is similar to the time scale for the emergence of the benefits of emissions reductions. A sudden and sustained termination of SRM in a high GHG emissions scenario would cause rapid climate change and a reversal of the SRM effects on the carbon sinks ( ''high confidence'' ). It is also ''likely'' that a termination of strong SRM would drive abrupt changes in the water cycle globally and regionally, especially in the tropical regions by shifting the Inter-tropical Convergence Zone and Hadley cells. At the regional scale, non-linear responses cannot be excluded, due to changes in evapotranspiration. However, a gradual phase-out of SRM combined with emissions reductions and CDR would avoid larger rates of changes ( ''medium confidence'' ). Links to chapters 4.6.3, 5.6.3, 8.6.3 .

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'''Box TS.9 | Irreversibility, Tipping Points and Abrupt Changes'''

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'''The present rates of response of many aspects of the climate system are proportionate to the rate of recent temperature change, but some aspects may respond disproportionately. Some climate system components are slow to respond, such as the deep ocean overturning circulation and the ice sheets (Box TS.4). It is ''virtually certain'' that irreversible, committed change is already underway for the slow-to-respond processes as they come into adjustment for past and present emissions.''' '''The paleoclimate record indicates that tipping elements exist in the climate system where processes undergo sudden shifts toward a different sensitivity to forcing, such as during a major deglaciation, where 1°C degree of temperature change might correspond to a large or small ice-sheet mass loss during different stages (Box TS.2). For global climate indicators, evidence for abrupt change is limited, but deep ocean warming, acidification and sea level rise are committed to ongoing change for millennia after global surface temperatures initially stabilize and are irreversible on human time scales ( ''very high confidence'' ). At the regional scale, abrupt responses, tipping points and even reversals in the direction of change cannot be excluded ( ''high confidence'' ). Some regional abrupt changes and tipping points could have severe local impacts, such as unprecedented weather, extreme temperatures and increased frequency of droughts and forest fires.''' '''Models that exhibit such tipping points are characterized by abrupt changes once the threshold is crossed, and even a return to pre-threshold surface temperatures or to atmospheric carbon dioxide concentrations does not guarantee that the tipping elements return to their pre-threshold state. Monitoring and early warning systems are being put into place to observe tipping elements in the climate system. Links to chapters 1.3, 1.4.4, 1.5, 4.3.2, Table 4.10, 5.3.4, 5.4.9, 7.5.3, 9.2.2, 9.2.4, 9.4.1, 9.4.2, 9.6.3, Cross-chapter Box 12.1'''
Understanding of multi-decadal reversibility (i.e., the system returns to the previous climate state within multiple decades after the radiative forcing is removed) has improved since AR5 for many atmospheric, land surface and sea ice climate metrics following sea surface temperature recovery. Some processes suspected of having tipping points, such as the Atlantic Meridional Overturning Circulation (AMOC), have been found to often undergo recovery after temperature stabilization with a time delay ( ''low confidence'' ). However, substantial irreversibility is further substantiated for some cryosphere changes, ocean warming, sea level rise, and ocean acidification. Links to chapters 4.7.2, 5.3.3, 5.4.9, 9.2.2, 9.2.4, 9.4.1, 9.4.2, 9.6.3

Some climate system components are slow to respond, such as the deep ocean overturning circulation and the ice sheets. It is ''likely'' that under stabilization of global warming at 1.5°C, 2.0°C or 3.0°C relative to 1850–1900, the AMOC will continue to weaken for several decades by about 15%, 20% and 30% of its strength and then recover to pre-decline values over several centuries ( ''medium confidence'' ). At sustained warming levels between 2°C and 3°C, there is ''limited evidence'' that the Greenland and West Antarctic ice sheets will be lost almost completely and irreversibly over multiple millennia; both the probability of their complete loss and the rate of mass loss increases with higher surface temperatures ( ''high confidence'' ). At sustained warming levels between 3°C and 5°C, near-complete loss of the Greenland Ice Sheet and complete loss of the West Antarctic Ice Sheet is projected to occur irreversibly over multiple millennia ( ''medium confidence'' ); with substantial parts or all of Wilkes Subglacial Basin in East Antarctica lost over multiple millennia ( ''low confidence'' ). Early-warning signals of accelerated sea level rise from Antarctica could possibly be observed within the next few decades. For other hazards (e.g., ice-sheet behaviour, glacier mass loss and global mean sea level change, coastal floods, coastal erosion, air pollution, and ocean acidification) the time and/or scenario dimensions remain critical, and a simple and robust relationship with global warming level cannot be established ( ''high confidence'' ). Links to chapters 4.3.2, 4.7.2, 5.4.3, 5.4.5, 5.4.8, 8.6, 9.2, 9.4, Box 9.3, Cross-Chapter Box 12.1

For global climate indicators, evidence for abrupt change is limited. For global warming up to 2°C above 1850–1900 levels, paleoclimate records do not indicate abrupt changes in the carbon cycle ( ''low confidence'' ). Despite the wide range of model responses, uncertainty in atmospheric CO <sub>2</sub> by 2100 is dominated by future anthropogenic emissions rather than uncertainties related to carbon–climate feedbacks ( ''high confidence'' ). There is no evidence of abrupt change in climate projections of global temperature for the next century: there is a near-linear relationship between cumulative CO <sub>2</sub> emissions and maximum global mean surface air temperature increase caused by CO <sub>2</sub> over the course of this century for global warming levels up to at least 2°C relative to 1850–1900. The increase in global ocean heat content (Section TS.2.4) will likely continue until at least 2300 even for low emissions scenarios, and global mean sea level will continue to rise for centuries to millennia following cessation of emissions (Box TS.4) due to continuing deep ocean heat uptake and mass loss of the Greenland and Antarctic ice sheets ( ''high confidence'' ). Links to chapters 2.2.3; Cross-Chapter Box 2.1; 5.1.1; 5.4; Cross-Chapter Box 5.1; Figures 5.3, 5.4, 5.25, and 5.26; 9.2.2; 9.2.4

The response of biogeochemical cycles to anthropogenic perturbations can be abrupt at regional scales and irreversible on decadal to century time scales ( ''high confidence'' ). The probability of crossing uncertain regional thresholds increases with climate change ( ''high'' '''''Box TS.9'''''

''confidence'' ). It is ''very unlikely'' that gas clathrates (mostly methane) in deeper terrestrial permafrost and subsea clathrates will lead to a detectable departure from the emissions trajectory during this century. Possible abrupt changes and tipping points in biogeochemical cycles lead to additional uncertainty in 21st century atmospheric GHG concentrations, but future anthropogenic emissions remain the dominant uncertainty ( ''high confidence'' ). There is potential for abrupt water cycle changes in some high emissions scenarios, but there is no overall consistency regarding the magnitude and timing of such changes. Positive land surface feedbacks, including vegetation, dust, and snow, can contribute to abrupt changes in aridity, but there is only ''low confidence'' that such changes will occur during the 21st century. Continued Amazon deforestation, combined with a warming climate, raises the probability that this ecosystem will cross a tipping point into a dry state during the 21st century ( ''low confidence'' ). (Section TS.3.2.2) Links to chapters 5.4.3, 5.4.5, 5.4.8, 5.4.9, 8.6.2, 8.6.3, Cross-Chapter Box 12.1

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