Editing IPCC:AR6/WGI/Chapter-4 (section)

=== 4.4.4 Response to Short-lived Climate Forcers and Volcanic Eruptions ===

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Mitigation of SLCFs affects future climate projections and could alter the time of emergence of anthropogenic climate change signals. The AR5 assessed that emission reductions aimed at decreasing local air pollution could have a near-term warming impact on climate ( ''high confidence'' ) ( [[#Kirtman--2013|Kirtman et al., 2013]] ). Because of their shorter lifetimes, reductions in emissions of SLCF species mainly influence near-term GSAT trends ( [[#Chalmers--2012|Chalmers et al., 2012]] ; [[#Shindell--2017|Shindell et al., 2017]] ; [[#Shindell--2019|Shindell and Smith, 2019]] ), but on decadal time scales the near-term response to even very large reductions in SLCFs may be difficult to detect in the presence of large internal climate variability ( [[#Samset--2020|Samset et al., 2020]] ). The changes in SLCF emissions during the COVID-19 pandemic has resulted in a small net radiative forcing without a discernible impact on GSAT (Cross-Chapter Box 6.1). SLCF mitigation also leads to a higher GSAT in the mid- to long-term ( [[#Smith--2013|Smith and Mizrahi, 2013]] ; [[#Stohl--2015|Stohl et al., 2015]] ; [[#Hienola--2018|Hienola et al., 2018]] ) and can influence peak warming during the 21st century ( [[#Rogelj--2014|Rogelj et al., 2014]] ; [[#Hienola--2018|Hienola et al., 2018]] ). This section focuses on the total effect of SLCF changes on GSAT projections in the SSP scenarios. A more detailed breakdown of the separate climate effects of SLCF species and precursor species can be found in Sections 6.7.2 and 6.7.3.

A model experiment based on the SSP3-7.0 scenario with aerosols, their precursors, and non-methane tropospheric ozone precursors set to SSP1-1.9 abundances (SSP3-7.0-lowSLCF-highCH4; [[#Collins--2017|Collins et al., 2017]] ) shows a projected multi-model mean GSAT anomaly that is higher by 0.22°C at mid-century (2045-2054) compared to SSP3-7.0 (Figure 4.18; [[#Allen--2020|Allen et al., 2020]] ), but this difference is smaller than the inter-model spread of the SSP3-7.0 projections based on the CMIP6 models. Note the SSP3-7.0-lowSLCF-highCH4 experiment does not perturb methane from SSP3-7.0 concentrations. A modified SSP3-7.0-lowSLCF-lowCH4 scenario that also includes methane mitigation shows a lower GSAT by mid-century compared to SSP3-7.0 ( [[#Allen--2021|Allen et al., 2021]] ).

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[[File:b69103b33e1f4d30dd6d8b7549058f46 IPCC_AR6_WGI_Figure_4_18.png]]

'''Figure 4.18 |''' '''Influence of SLCFs on projected GSAT change.''' Change is shown relative to the 1995–2014 average (left axis) and relative to the 1850–1900 average (right axis). The comparison is for CMIP6 models for the AerChemMIP ( [[#Collins--2017|Collins et al., 2017]] ) SSP3-7.0-lowSLCF-highCH4 experiment (red dashed; note in the original experiment protocol this is called SSP3-7.0-lowNTCF), where concentrations of short-lived species are reduced compared to reference SSP3-7.0 scenario (red solid). Black shows the historical simulation until 2014 for the same 9 models as the projections. The curves show averages over the r1 simulations contributed to the CMIP6 exercise, the shadings around the historical and SSP3-7.0 curves shows 5–95% ranges and the numbers near the top show the number of model simulations.

Building on CMIP6 results for the effects of reducing SLCF emissions from a baseline of SSP3-7.0, the overall contribution of SLCFs to GSAT changes in the marker SSPs are now quantified using a simple climate model emulator. For consistency with Section 6.7.2 and Figure 6.22, the basket of SLCF compounds considered includes aerosols, ozone, methane, black carbon on snow and hydrofluorocarbons (HFCs) with lifetimes of less than 50 years. In the five marker SSPs considered, the net effect of SLCFs contributes to a higher GSAT in the near, mid- and long term (Table 4.6 and Section 6.7.2). In the SSP1-1.9 and SSP1-2.6 scenarios, SLCFs contribute to a higher GSAT by a central estimate of around 0.3°C compared to 1995–2014 across the three-time horizons. In the long-term, the 0.3C warming due to SLCFs in SSP1-2.6 can be compared to the assessed ''very likely'' GSAT change for this period of 0.5°C–1.5°C ( [[#4.3.4|Section 4.3.4]] and Table 4.5). The SSP2-4.5, SSP3-7.0 and SSP5-8.5 scenarios all show a larger SLCF effect on GSAT in the long term relative to the near term. In SSP3-7.0, the long-term warming due to SLCFs by 0.7°C can be compared with the assessed ''very likely'' GSAT anomaly for this period of 2.0°C –3.7°C ( [[#4.3.4|Section 4.3.4]] ). In summary, it is ''very likely'' that changes in SLCFs contribute to an overall warmer GSAT over the near, mid- and long term in the five SSP scenarios considered (Table 4.6, Section 6.7.2 and Figure 6.22).

In addition to effects on GSAT, SLCFs affect other aspects of the global climate system (Section 6.7.2). The additional warming at high northern latitudes associated with projected reductions in aerosol emissions over the 21st century leads to a more rapid reduction in Arctic sea ice extent in the RCP scenarios ( [[#Gagné--2015|Gagné et al., 2015]] ). Furthermore, mitigation of non-methane SLCFs in the SSP3-7.0-lowSLCF-highCH4 scenario causes an increase in global mean precipitation, with larger regional changes in southern and eastern Asia ( [[#Allen--2020|Allen et al., 2020]] ).

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'''Table''' '''4.6 |''' '''The net effect of SLCFs on GSAT change.''' Changes in 20-year averaged GSAT relative to 1995–2014 for 2021–2040, 2041–2060, and 2081–2100 for the five marker SSP scenarios. Values give the median and, in parentheses, the 5–95% range calculated from a 2237-member ensemble of the two-layer emulator that is driven with the ERF projections, including uncertainties, described in [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] Supplementary Material 7.SM.1.4. The ensemble is constrained to assessed ranges of ECS, TCR, ocean heat content change, GSAT response, and carbon cycle metrics (Section 7.3.5; [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] Supplementary Material 7.SM.2.2). The GSAT contribution of individual forcer responses use the difference between parallel runs of the constrained two-layer model with all anthropogenic forcing and all anthropogenic forcing with the component of interest (e.g., methane) removed ( [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] Supplementary Material 7.SM.2.3). Values are given to one decimal place.

{| class="wikitable"
|-
| '''Time Period'''

| '''SSP1-1.9 (°C)'''

| '''SSP1-2.6 (°C)'''

| '''SSP2-4.5 (°C)'''

| '''SSP3-7.0 (°C)'''

| '''SSP5-8.5 (°C)'''

|-
| Near Term (2021–2040)

| 0.2 (0.1, 0.3)

| 0.2 (0.1, 0.3)

| 0.2 (0.1, 0.3)

| 0.2 (0.1, 0.3)

| 0.3 (0.2, 0.4)

|-
| Mid-Term (2041–2060)

| 0.2 (0.0, 0.4)

| 0.2 (0.0, 0.4)

| 0.3 (0.2, 0.4)

| 0.3 (0.2, 0.4)

| 0.5 (0.3, 0.7)

|-
| Long Term (2081–2100)

| 0.1 (-0.1, 0.4)

| 0.2 (0.0, 0.4)

| 0.3 (0.1, 0.6)

| 0.5 (0.4, 0.8)

| 0.7 (0.4, 1.0)

|}

The main uncertainties in climate effects of SLCFs in the future come from: (i) the uncertainty in anthropogenic aerosol ERF (Section 7.3.3); (ii) uncertainty in the relative emissions of different SLCFs that have warming and cooling effects in the current climate (Section 6.2); and (iii) physical uncertainty including the efficacy of the climate response to SLCFs compared to long-lived GHGs ( [[#Marvel--2016|Marvel et al., 2016]] ; [[#Richardson--2019|Richardson et al., 2019]] ). One example of physical uncertainty is that the shortwave radiative forcing from methane was neglected in previous calculations ( [[#Etminan--2016|Etminan et al., 2016]] ; [[#Collins--2018|Collins et al., 2018]] ), which affects understanding of present day and future methane ERF ( [[#Modak--2018|Modak et al., 2018]] ). Another example of physical uncertainty is projected changes in lightning-NO <sub>x</sub> production, which contribute to future ozone radiative forcing ( [[#Banerjee--2014|Banerjee et al., 2014]] , 2018; [[#Finney--2018|Finney et al., 2018]] ).

Another factor that could substantially alter projections in the near-term would be the occurrence of a large explosive volcanic eruption, or even a decadal to multi-decadal sequence of small-to-moderate volcanic eruptions as witnessed over the early 21st century ( [[#cross-chapter-box-4.1|Cross-Chapter Box 4.1]] ; [[#Santer--2014|Santer et al., 2014]] ). An eruption similar to the last large tropical eruption, Mount Pinatubo in the Philippines in June 1991, is expected to cause substantial Northern Hemisphere (NH) cooling, peaking between 0.09°C and 0.38°C and lasting for three to five years, as indicated by climate model simulations over the past millennium (e.g., [[#Jungclaus--2010|Jungclaus et al., 2010]] ). Phase 3 of Paleoclimate Modelling Intercomparison Project (PMIP3) simulated a significant NH cooling in response to individual volcanic events (peaks between 0.1°C and 0.5°C, depending on model, during the first year after the eruption) that lasts for three to five years. On a regional scale, the double volcanic events that occurred in 536 and 540 CE resulted in a cooling of 2°C ( [[#Buntgen--2016|Büntgen et al., 2016]] ; [[#Toohey--2016|Toohey et al., 2016]] ).

Since AR5, there has been growing progress in understanding the climate impacts of volcanic eruptions. Volcanic forcing is regarded as the dominant driver of forced variability in preindustrial surface air temperature ( [[#Schurer--2013|Schurer et al., 2013]] , 2014). Large eruptions in the tropics and high latitudes were primary drivers of interannual-to-decadal temperature variability in the Northern Hemisphere during the past 2,500 years, with cooling persisting for up to ten years after some of the largest eruptive episodes ( [[#Sigl--2015|Sigl et al., 2015]] ). Repeated clusters of volcanic eruptions can induce a net negative radiative forcing that results in a centennial- and global-scale cooling trend via a decline in mixed-layer oceanic heat content ( [[#McGregor--2015|McGregor et al., 2015]] ). The response to multi-decadal changes in volcanic forcing (representing clusters of eruptions) shows similar cooling in both simulations and reconstructions of NH temperature. Volcanic eruptions generally result in decreased global precipitation for up to a few years following the eruption ( [[#Iles--2014|Iles and Hegerl, 2014]] , 2015; [[#Man--2014|Man et al., 2014]] ), with climatologically wet regions drying and climatologically dry regions wetting ( ''medium confidence'' ), which is opposite to the response under global warming ( [[#Held--2006|Held and Soden, 2006]] ; [[#Iles--2013|Iles et al., 2013]] ; [[#Zuo--2019a|Zuo et al., 2019a]] , b). El Niño-like warming appears after large volcanic eruptions, as seen in both observations ( [[#Adams--2003|Adams et al., 2003]] ; [[#McGregor--2010|McGregor et al., 2010]] ; [[#Khodri--2017|Khodri et al., 2017]] ) and climate model simulations ( [[#Ohba--2013|Ohba et al., 2013]] ; [[#Pausata--2015|Pausata et al., 2015]] ; [[#Colose--2016|Colose et al., 2016]] ; [[#Stevenson--2016|Stevenson et al., 2016]] ; [[#Khodri--2017|Khodri et al., 2017]] ; [[#Predybaylo--2017|Predybaylo et al., 2017]] ; [[#Zuo--2018|Zuo et al., 2018]] ). The large tropical eruptions are coincident with positive Indian Ocean dipole events ( [[#Maher--2015|Maher et al., 2015]] ).

In AR5, uncertainty due to future volcanic activity was not considered in the assessment of the CMIP5 21st century climate projections ( [[#Taylor--2012|Taylor et al., 2012]] ; [[#O’Neill--2016|O’Neill et al., 2016]] ). Since AR5, there has been considerable progress in quantifying the impacts of volcanic eruptions on decadal climate prediction and longer-term climate projections ( [[#Meehl--2015|Meehl et al., 2015]] ; [[#Swingedouw--2015|Swingedouw et al., 2015]] , 2017; [[#Timmreck--2016|Timmreck et al., 2016]] ; [[#Bethke--2017|Bethke et al., 2017]] ; [[#Illing--2018|Illing et al., 2018]] ). By exploring 60 possible volcanic futures under RCP4.5, it has been demonstrated that the inclusion of time-varying volcanic forcing may enhance climate variability on annual-to-decadal time scales ( [[#Bethke--2017|Bethke et al., 2017]] ). Consistent with a tropospheric cooling response, the change in ensemble spread in the volcanic cases is skewed towards lower GSAT relative to the non-volcanic cases ( [[#cross-chapter-box-4.1|Cross-Chapter Box 4.1]] , Figure 1). In these simulations with multiple volcanic forcing futures there is: (i) an increase in the frequency of extremely cold individual years; (ii) an increased likelihood of decades with negative GSAT trend (decades with negative GSAT trends become 50% more commonplace); (iii) later anthropogenic signal emergence (the mean time at which the signal of global warming emerges from the noise of natural climate variability is delayed almost everywhere) ( ''high confidence'' ); and (iv) a 10% overall reduction in global land monsoon precipitation and a 20% overall increase in the ensemble spread ( [[#Man--2021|Man et al., 2021]] ).

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'''Cross-Chapter Box 4.1 | The Climate Effects of Volcanic Eruption'''

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'''Contributing Authors:''' Sarah L. Connors (France/United Kingdom), Amanda Maycock (United Kingdom), Peter W. Thorne (Ireland/United Kingdom), Nicolas Bellouin (United Kingdom/France), Ingo Bethke (Norway/Germany), Deliang Chen (Sweden), Annalisa Cherchi (Italy), Alejandro Di Luca (Australia/Canada/Argentina), Piers Forster (United Kingdom), Nathan P. Gillett (Canada), Darrell S. Kaufmann (The United States of America), June-Yi Lee (Republic of Korea), Elizaveta Malinina (Canada/Russian Federation), Seung-Ki Min (Republic of Korea), Johannes Quaas (Germany), Alex C. Ruane (The United States of America), Jean-Baptiste Sallée (France), Sonia I. Seneviratne (Switzerland), Chris Smith (United Kingdom), Matthew Toohey (Canada, Germany/Canada), Andrew Turner (United Kingdom), Cunde Xiao (China), Tianjun Zhou (China)

Before the industrial period, explosive volcanic eruptions were the largest source of forced climate variability globally on interannual to centennial time scales ( [[IPCC:Wg1:Chapter:Chapter-2#2.2|Section 2.2]] ). While usually omitted from scenarios used for future climate projections, as they are unpredictable, volcanic eruptions have the potential to influence future climate on multi-annual to decadal time scales and affect many climatic impact drivers (as defined in Sections 12.1 and 12.3). Since AR5, more comprehensive paleo evidence and observations, as well as improved modelling have advanced understanding of the climate response to past volcanic eruptions. Building on multiple chapter assessments, this box synthesizes how volcanic eruptions affect climate and considers implications of possible future events.

'''How frequent are volcanic eruptions?'''

Proxy records show that large volcanic eruptions with effective radiative forcing (ERF) more negative than –1 W m <sup>–2</sup> occurred on average twice a century throughout the last 2500 years, the most recent being Pinatubo in 1991 ( [[IPCC:Wg1:Chapter:Chapter-2#2.2.2|Section 2.2.2]] ). About eight larger eruptions (ERF stronger than –5 W m <sup>–2</sup> ) also occurred during this period (Figure 2.2), notably Tambora about 1815 and Samalas about 1257. A Samalas-type eruption may occur one to two times per millennium on average ( [[#Newhall--2018|Newhall et al., 2018]] ). Typically, three in every four centuries have experienced at least one eruption stronger than –1 W m <sup>–2</sup> (Pinatubo or larger). The volcanic aerosol burden was 14% lower during the 20th century compared to the average of the preceding 24 centuries ( [[IPCC:Wg1:Chapter:Chapter-2#2.2.2|Section 2.2.2]] ), whereas the 13th century was among the most volcanically active, with four eruptions exceeding that of Pinatubo-1991 ( [[#Sigl--2015|Sigl et al., 2015]] ).

'''Past climate responses to volcanic activity'''

Major eruptions drive a range of climate system responses for several years depending upon whether the eruption occurs in the tropics (stratospheric aerosol dispersion into both hemispheres) or the extratropics (dispersion into the hemisphere of eruption) owing to the Brewer-Dobson circulation. The climatic response also depends on the effective injection height, sulphur mass injected, and time of year of the eruption ( [[#Marshall--2019|Marshall et al., 2019]] , 2020). These factors determine the total mass, lifetime and optical properties of volcanic aerosol in the stratosphere and influence the stratospheric aerosol optical depth (SAOD). The ERF from volcanic stratospheric aerosol is assessed to be –20 ± 5 W m <sup>–2</sup> per unit sAOD (Section 7.3.4.6).

Due to the direct radiative effect of volcanic stratospheric aerosols, large volcanic eruptions lead to an overall decrease of GSAT, which can extend to multi-decadal or century time scales in the case of clustered volcanism ( [[IPCC:Wg1:Chapter:Chapter-3#3.3.1.1|Section 3.3.1.1]] ; [[#Schurer--2013|Schurer et al., 2013]] ; [[#McGregor--2015|McGregor et al., 2015]] ; [[#Sigl--2015|Sigl et al., 2015]] ; [[#Kobashi--2017|Kobashi et al., 2017]] ; [[#Zambri--2017|Zambri et al., 2017]] ; [[#Brönnimann--2019|Brönnimann et al., 2019]] ; [[#Neukom--2019|Neukom et al., 2019]] ). Large eruptions also increase the frequency of extremely cold individual years and the likelihood of cooling trends occurring in individual decades ( [[IPCC:Wg1:Chapter:Chapter-3#cross-chapter-box-3.1|Cross-Chapter Box 3.1]] and [[#4.4.4|Section 4.4.4]] ; [[#Paik--2018|Paik and Min, 2018]] ). Re-dating of ice core chronologies now confirms that the coldest decades of the past approximately 2000 years are the outcome of volcanic eruptions ( [[#Sigl--2015|Sigl et al., 2015]] ; [[#Buntgen--2016|Büntgen et al., 2016]] ; [[#Toohey--2016|Toohey et al., 2016]] ; [[#Neukom--2019|Neukom et al., 2019]] ). CMIP5 and CMIP6 models reproduce the decreased GSAT that follows periods of intense volcanism. New reconciliations between simulations and proxy-based reconstructions of past eruptions have been achieved through better Earth System Model representation of volcanic plume chemical compositions ( [[#Legrande--2016|Legrande et al., 2016]] ; [[#Marshall--2020|Marshall et al., 2020]] ; F. [[#Zhu--2020|]] [[#Zhu--2020|Zhu et al., 2020]] ). Yet, remaining disagreements reflect differences in the volcanic forcing datasets used in the simulations ( ''medium confidence'' ) ( [[IPCC:Wg1:Chapter:Chapter-3#3.3.1.1|Section 3.3.1.1]] and Figure 3.2c).

Although incomplete, proxy records show large impacts upon contemporary society from eruptions such as 1257 Samalas and 1815 Tambora, the latter resulting in ‘the year without a summer’ with multiple harvest failures across the Northern Hemisphere (e.g., [[#Raible--2016|Raible et al., 2016]] ). Comparing CMIP5 multi-model simulations with observations has improved understanding of the hydrological responses to 20th century eruptions, particularly global land monsoon drying, and associated uncertainties ( [[IPCC:Wg1:Chapter:Chapter-3#3.3.2.3|Section 3.3.2.3]] ). Global mean land precipitation decreases for up to a few years following the eruption, with climatologically wet regions drying and dry regions wetting (Sections 3.3.2.3 and 4.4.4). Changes in monsoon circulations occur with a general weakening of tropical precipitation (Section 8.5.2.3) and a decrease in extreme precipitation over global monsoon regions (Section 11.4.4). Monsoon precipitation in one hemisphere tends to be enhanced by eruptions occurring in the other hemisphere or reduced if they occur in the same hemisphere (Sections 3.3.2.3 and 8.5.2.3). Volcanic eruptions have been linked to the onset of El Niño followed by La Niña although this connection remains contentious ( [[#Adams--2003|Adams et al., 2003]] ; [[#Bradley--2003|Bradley et al., 2003]] ; [[#McGregor--2010|McGregor et al., 2010]] ; [[#Khodri--2017|Khodri et al., 2017]] ; F. [[#Liu--2018|]] [[#Liu--2018|]] [[#Liu--2018|]] [[#Liu--2018|Liu et al., 2018]] ; [[#Sun--2019|Sun et al., 2019]] ; [[#Paik--2020|Paik et al., 2020]] ; [[#Predybaylo--2020|Predybaylo et al., 2020]] ). Volcanic activity could drive short-term (one-to-three-year) positive changes in the annual SAM index through modulations in the extratropical temperature gradient and wave driving of the polar stratosphere ( [[#Yang--2018|Yang and Xiao, 2018]] ). In the cryosphere, Arctic sea ice extent increases for years to decades ( [[#Gagné--2017a|Gagné et al., 2017a]] ), and modelling indicates that sea ice/ocean feedbacks can prolong cooling long after volcanic aerosols are removed ( [[#Miller--2012|Miller et al., 2012]] ). On annual time scales, the ocean buffers the atmospheric response to volcanic eruptions by storing the cooling in the ocean subsurface, then feeding it back to the atmosphere. Large eruptions affect ocean heat content and thermosteric sea level over decadal-to-centennial scales (Section 9.2.2.1).

'''Potential implications on 21st century projections'''

Given the unpredictability of individual eruptions, volcanic forcing is prescribed as a constant background loading in CMIP6 models ( [[#Eyring--2016|Eyring et al., 2016]] ). This means the effects of potential large volcanic eruptions are largely absent from model projections, and few studies have addressed the potential implications on 21st century warming. One study considered future scenarios with hypothetical volcanic eruptions consistent with levels of Common Era volcanic activity ( [[#Bethke--2017|Bethke et al., 2017]] ) under RCP4.5 and found that climate projections could be substantially altered ( [[#cross-chapter-box-4.1|Cross-Chapter Box 4.1]] , Figure 1). Although temporary, close to pre-industrial level temperatures could be experienced globally for a few years after a 1257 Samalas-sized eruption. Several other key climate indicators are also changed substantially, consistent with evidence from past events. [[#Bethke--2017|Bethke et al. (2017)]] suggest that an eruption early in the 21st century could delay the timing of crossing 1.5°C global warming by several years. Clustered eruptions would have substantial impact upon GSAT evolution throughout the century ( [[#cross-chapter-box-4.1|Cross-Chapter Box 4.1]] , Figure 1), and could have far-reaching implications, as observed for past eruptions. For near-term response options, decadal prediction models can update 21st-century projections once a volcanic eruption occurs ( [[#Timmreck--2016|Timmreck et al., 2016]] ).

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'''Cross-Chapter Box 4.1, Fig'''

'''ure 1 |'''

'''Potential impact of volcanic eruption on future global temperature change.''' CMIP5 projections of possible 21st-century futures under RCP4.5 after a 1257 Samalas magnitude volcanic eruption in 2044, from [[#Bethke--2017|Bethke et al. (2017)]] . '''(a)''' Volcanic ERF of the most volcanically active ensemble member, estimated from SAOD. '''(b)''' Annual mean global surface air temperature. Ensemble mean (solid) of future projections including volcanoes (blue) and excluding volcanoes (red) with 5–95% range (shading) and ensemble minima/maxima (dots); evolution of the most volcanically active member (black). Data created using a SMILE approach with NorESM1 in its CMIP5 configuration. See Sections 2.2.2 and 4.4.4 for more details. Further details on data sources and processing are available in the chapter data table (Table 4.SM.1).

'''Summary'''

It is ''likely'' that at least one large eruption will occur during the 21st century. Such an eruption would reduce GSAT for several years, decrease global mean land precipitation, alter monsoon circulation, modify extreme precipitation, and change the profile of many regional climatic impact-drivers. A low-likelihood, high-impact outcome would be several large eruptions that would greatly alter the 21st century climate trajectory compared to SSP-based ESM projections.

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