Editing IPCC:AR6/WGI/Chapter-8 (section)

==== 8.2.2.2 Large-scale Responses in Atmospheric Circulation Patterns ====

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Responses of the large-scale atmospheric circulation to a warming climate are not as well understood as thermodynamic drivers ( [[#Shepherd--2014|Shepherd, 2014]] ). The AR5 identified robust features including a weakening and broadening of tropical circulation with poleward movement of tropical dry zones and mid-latitude jets ( [[#Collins--2013|Collins et al., 2013]] ). These can dominate regional water cycle changes, affecting the availability of freshwater and the occurrence of climate extremes. Atmospheric circulation changes generally dominate the spatial pattern of rapid precipitation adjustments ( [[#8.2.1|Section 8.2.1]] ) to different forcing agents in the tropics ( [[#Bony--2013|Bony et al., 2013]] ; [[#He--2015|He and Soden, 2015]] ; [[#Richardson--2016|T.B. Richardson et al., 2016]] , 2018a; [[#Tian--2017|Tian et al., 2017]] ; X. [[#Li--2018|]] [[#Li--2018|]] [[#Li--2018|Li et al., 2018]] ). Radiative forcing with heterogeneous spatial patterns such as ozone and aerosols (including cloud interactions; Section 6.4.1 and Box 8.1) drive substantial responses in regional atmospheric circulation through uneven heating and cooling effects(L. [[#Liu--2018|]] [[#Liu--2018|]] [[#Liu--2018|]] [[#Liu--2018|Liu et al., 2018]] ; [[#Dagan--2019b|Dagan et al., 2019b]] ; [[#Wilcox--2019|Wilcox et al., 2019]] ). Changes in atmospheric circulation are also driven by slower, evolving patterns of warming and associated changes in temperature and moisture gradients ( [[#Bony--2013|Bony et al., 2013]] ; [[#Samset--2016|Samset et al., 2016]] , 2018a; [[#Ceppi--2018|Ceppi et al., 2018]] ; [[#Ma--2018|Ma et al., 2018]] ). There is strong evidence that large regional water cycle changes arise from the atmospheric circulation response to radiative forcings and associated SST pattern evolution but ''low agreement'' in the sign and magnitude ( [[#Chadwick--2016b|Chadwick et al., 2016b]] ). The role of prolonged weather regimes in determining wet and dry extremes is also better understood since AR5 ( [[#Kingston--2015|Kingston and McMecking, 2015]] ; [[#Schubert--2016|Schubert et al., 2016]] ; D. [[#Richardson--2018|Richardson et al., 2018]] ; [[#Barlow--2019|Barlow et al., 2019]] ). Advances in knowledge of expected large-scale dynamical responses of the water cycle are further assessed in this section (see also Figure 8.21).

Long-term weakening of the tropical atmospheric overturning circulation is expected as climate warms in response to elevated CO <sub>2</sub> ( [[#Collins--2013|Collins et al., 2013]] ). A weaker circulation is required to reconcile global mean low-level water vapour increases (around 7% °C <sup>–1</sup> ) with the smaller global precipitation responses of about 1–3% °C <sup>–1</sup> ( [[#8.2.1|Section 8.2.1]] ). The slowdown can occur in both the Hadley and Walker circulations, but occurs preferentially in the Walker circulation in most climate models (Vecchi and Soden, 2007) but this response has been questioned on the basis of model bias in east Pacific SST ( [[#Seager--2019a|Seager et al., 2019a]] ). Weakening is expected to drive P–E decreases over the western Pacific and increases over the eastern Pacific. However, the driving mechanisms for Walker circulation weakening differ to those involved in determining ENSO variability, so it is too simplistic to interpret changes as an El Niño pattern of regional hydrological cycle extremes ( [[#Sohn--2019|Sohn et al., 2019]] ). Internal variability is also capable of temporarily strengthening the Walker circulation ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1.4.1|Section 2.3.1.4.1]] ; [[#L’Heureux--2013|L’Heureux et al., 2013]] ; [[#Chung--2019|Chung et al., 2019]] ) while regional responses depend on the pattern of warming (Sandeep et al., 2014).

Model simulations show a stronger Pacific Walker circulation during the LGM in response to a cooler climate (consistent with an expected weakening in a warmer climate), but a weaker Indian Ocean east – west circulation in response to the exposure of the Sunda and Sahul shelves due to lowered sea level (DiNezio et al., 2011). The latter effect is detectable in proxies for hydroclimate, as well as salinity and sea surface temperature (DiNezio and Tierney, 2013; [[#DiNezio--2018|DiNezio et al., 2018]] ). More relevant to future warming is the mid-Pliocene period (3 million years ago), the last time the Earth experienced CO <sub>2</sub> levels comparable to present (see Cross-Chapter Box 2.4). Sea surface temperature (SST) reconstructions show a weakening of the Pacific zonal gradient and a pattern of warmth consistent with a weaker Walker cycle response (Corvec and Fletcher, 2017; [[#Tierney--2019|Tierney et al., 2019]] ; [[#McClymont--2020|McClymont et al., 2020]] ). Although the Pliocene SST pattern and wet subtropics contrast with present conditions ( [[#Burls--2017|Burls and Fedorov, 2017]] ), the paleoclimate record strengthens evidence that a warmer climate is associated with a weaker Walker circulation (Cross-Chapter Box 2.4; [[IPCC:Wg1:Chapter:Chapter-3#3.3.3|Section 3.3.3]] ).

Since AR5, weakening of the tropical circulation has been explained as a rapid response to increasing CO <sub>2</sub> concentrations and slower response to warming and evolving SST patterns ( [[#He--2017|He and Soden, 2017]] ; [[#Xia--2017|Xia and Huang, 2017]] ; [[#Shaw--2018|Shaw and Tan, 2018]] ; [[#Chemke--2020|Chemke and Polvani, 2020]] ). Large-scale tropical circulation weakens by 3 – 4% in a rapid response to a quadrupling of CO <sub>2</sub> concentrations ( [[#Plesca--2018|Plesca et al., 2018]] ), which suppresses tropospheric radiative cooling, particularly in subtropical ocean subsidence regions ( [[#Bony--2013|Bony et al., 2013]] ; [[#Merlis--2015|Merlis, 2015]] ; [[#Richardson--2016|Richardson et al., 2016]] ). The resulting increased atmospheric stability explains the rapid weakening of the Walker circulation ( [[#Wills--2017|Wills et al., 2017]] ) and Northern Hemisphere Hadley Cell ( [[#Chemke--2020|Chemke and Polvani, 2020]] ). Subsequent surface warming contributes up to a 12% slowing of circulation for a uniform 4°C SST increase, driven by thermodynamic decreases in temperature lapse rate ( [[#Plesca--2018|Plesca et al., 2018]] ).

The regional Inter-tropical Convergence Zone (ITCZ) position, width and strength determine the location and seasonality of the tropical rain belt. Since AR5, multiple studies have linked cross-equatorial energy transport to the mean ITCZ position ( [[#Donohoe--2013|Donohoe et al., 2013]] ; [[#Frierson--2013|Frierson et al., 2013]] ; [[#Bischoff--2014|Bischoff and Schneider, 2014]] ; [[#Boos--2016|Boos and Korty, 2016]] ; [[#Loeb--2016|Loeb et al., 2016]] ; [[#Adam--2018|Adam et al., 2018]] ; [[#Biasutti--2019|Biasutti and Voigt, 2019]] ). Multi-model studies agree that aerosol cooling in the NH led to a southward shift in the ITCZ and tropical precipitation after the 1950s up to the 1980s that is linked with the 1980s Sahel drought (Box 8.1; [[#8.3.2.4|Section 8.3.2.4]] and 10.4.2.1). In particular, aerosol-cloud interaction was identified as a potentially important driver of this shift ( [[#Chung--2017|Chung and Soden, 2017]] ) but this is uncertain since observations suggest that models may overestimate (Malavelle et al. , 2017; Toll et al. , 2017) or underestimate (Rosenfeld et al. , 2019) the aerosol cloud-mediated cooling effects. In addition, greenhouse gas forcing has been invoked in explaining much of the increase in Sahel precipitation since the 1980s through enhanced meridional temperature gradient, with only a secondary role for aerosol ( [[#Dong--2015|Dong and Sutton, 2015]] ).

Understanding of how ITCZ width and strength respond to a warming climate has improved since AR5 ( [[#Byrne--2016|Byrne and Schneider, 2016]] ; [[#Harrop--2016|Harrop and Hartmann, 2016]] ; [[#Popp--2017|Popp and Silvers, 2017]] ; [[#Dixit--2018|Dixit et al., 2018]] ; [[#Zhou--2020|Zhou et al., 2020]] ). Studies suggest that convection gets stronger and more focused within the core of the ITCZ ( [[#Lau--2015|Lau and Kim, 2015]] ; [[#Byrne--2018|Byrne et al., 2018]] ). This leads to drying on the equatorward edges of the ITCZ and a moistening tendency in the ITCZ core ( [[#Byrne--2016|Byrne and Schneider, 2016]] ). Feedbacks involving clouds have been identified as an important mechanism leading to tightening and strengthening of the ITCZ ( [[#Popp--2017|Popp and Silvers, 2017]] ; [[#Su--2017|Su et al., 2017]] , 2019, 2020; [[#Talib--2018|Talib et al., 2018]] ). Stronger ascent in the core amplifies the ‘wet get wetter’ response while reduced moisture inflow near the ITCZ edges reduces this response below the 7% °C <sup>–1</sup> Thermodynamic increase in moisture transport. Thus, there is a range of evidence and ''medium agreement'' for strengthening and contraction of the ITCZ with warming that sharpens contrasts between wet and dry regimes. However, understanding of how the regional ITCZ location responds in a warming climate is not robust ( [[#8.4.2.1|Section 8.4.2.1]] ) with ''limited evidence'' of distinct regional responses to GHG forcing including a northward shift over eastern Africa and the Indian Ocean and a southward shift in the eastern Pacific and Atlantic oceans ( [[#Mamalakis--2021|Mamalakis et al., 2021]] ). Paleoclimate evidence highlights the distinct regional ITCZ responses to hemispheric asymmetry in volcanic and orbital forcing ( [[#McGee--2014|McGee et al., 2014]] ; [[#Boos--2016|Boos and Korty, 2016]] ; [[#Colose--2016|Colose et al., 2016]] ; [[#Denniston--2016|Denniston et al., 2016]] ; [[#PAGES%20Hydro2K%20Consortium--2017|PAGES Hydro2K Consortium, 2017]] ; [[#Singarayer--2017|Singarayer et al., 2017]] ; [[#Atwood--2020|Atwood et al., 2020]] ) and rapid (&gt;1° latitude over decades) shifts in the ITCZ and regional monsoons in response to AMOC collapse cannot be ruled out (Sections 8.6.1.1 and 5.1.3).

Monsoons are key components of the tropical overturning circulation that can be understood as a balance between net energy input (e.g., radiative and turbulent fluxes) and the export of moist static energy. This is determined by contrasting surface heat capacity between ocean and land and modified through changes in atmospheric dynamics, tropical tropospheric stability and land surface properties ( [[#Jalihal--2019|Jalihal et al., 2019]] ). Thermodynamic increases in moisture transport are expected to increase monsoon strength and area (Christensen et al., 2013). Since AR5, evidence continues to demonstrate that monsoon circulation is sensitive to spatially varying radiative forcing by anthropogenic aerosols (Hwang et al. , 2013; R.J. Allen et al. , 2015; Z. Li et al. , 2016b) and GHGs (Dong andSutton, 2015). Changes in SST patterns also play a role ( [[#Guo--2016|Guo et al., 2016]] ; W. [[#Zhou--2019|]] [[#Zhou--2019|]] [[#Zhou--2019|Zhou et al., 2019]] ; [[#Cao--2020|Cao et al., 2020]] ) by altering cross-equatorial energy transports and land–ocean temperature contrasts. This evidence continues to support a thermodynamic strengthening of monsoon precipitation that is partly offset by slowing of the tropical circulation but with ''weak evidence'' and ''low agreement'' for regional aspects of circulation changes. Disagreement between paleoclimate and modern observations, physical theory and numerical simulations of global monsoons have been partly reconciled ( [[IPCC:Wg1:Chapter:Chapter-3#3.3.3.2|Section 3.3.3.2]] ) through improved understanding of regional processes ( Harrison et al. , 2015; R. Bhattacharya et al. , 2017; Bhattacharya et al. , 2018; Biasutti et al. , 2018; D’Agostino et al. , 2019; Jalihal et al. , 2019; Seth et al. , 2019 ), although interpreting past changes in the context of future projections requires careful account of differing forcings and feedbacks (D’Agostinoet al., 2019). Assessment of past changes and future projections in regional monsoons are provided in Sections 2.3.1.4.2, 8.3.2.4 and 8.4.2.4.

Since AR5, understanding of poleward expansion of the Hadley Cells has improved ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1.4.1|Section 2.3.1.4.1]] ) but its role in subtropical drying is limited to the zonal mean and dominated by ocean regions ( [[#Byrne--2015|Byrne and]] [[#O’Gorman--2015|O’Gorman, 2015]] ; [[#Grise--2016|Grise and Polvani, 2016]] ; [[#He--2017|He and Soden, 2017]] ; [[#Schmidt--2017|Schmidt and Grise, 2017]] ; [[#Siler--2018|Siler et al., 2018]] ; [[#Chemke--2019|Chemke and Polvani, 2019]] ; [[#Grise--2020|Grise and Davis, 2020]] ). Over subtropical land, evolving SST patterns and land–ocean warming contrasts, that are partly explained by rapid responses to CO <sub>2</sub> increases, can dominate aspects of the atmospheric circulation response (Byrne and O’Gorman, 2015; [[#He--2015|He and Soden, 2015]] ; [[#Chadwick--2017|Chadwick et al., 2017]] ; H. [[#Yang--2020|]] [[#Yang--2020|Yang et al., 2020]] ) and resultant regional water cycle changes, particularly for projected drying in semi-arid, winter-rainfall dominated subtropical climates (Deitch et al. , 2017; Brogli et al. , 2019; Seager et al. , 2019b; Zappa et al. , 2020) . Poleward expansion of the tropical belt is expected to drive a corresponding shift in mid-latitude storm tracks, but the controlling mechanisms differ between hemispheres. Southern Hemisphere expansion is driven by GHG forcing and amplified by stratospheric ozone depletion, while weaker Northern Hemisphere expansion in response to GHG forcing is modulated by tropospheric ozone and aerosol forcing, particularly black carbon (Davis et al. , 2016; Grise et al. , 2019; Watt-Meyer et al. , 2019; Zhao et al. , 2020) . However, internal variability is found to dominate observed responses in the NH, precluding attribution to radiative forcing ( [[#D’Agostino--2020a|D’Agostino et al., 2020a]] ). Paleoclimate evidence of poleward expansion and weakening of westerly winds in both hemispheres in the warmer Pliocene is linked to reduced equator-to-pole thermal gradients and ice volume ( [[#Abell--2021|Abell et al., 2021]] ).

The influence of amplified Arctic warming on mid-latitude regional water cycles is not well understood based on simple physical grounds due to the large number of competing physical processes (Cross-Chapter Box 10.1). The thermal gradient between polar and lower latitude regions decreases at lowlevels due to Arctic warming amplification. However, at higher altitudes, the corresponding thermal gradient increases with warming due to cooling of the Arctic stratosphere and this is consistent with a strengthening of the winter jet stream in both hemispheres, yet there is ''low agreement'' on the precise mechanisms (Vallis et al., 2015; [[#Vihma--2016|Vihma et al., 2016]] ). Changes in the strength of the polar stratospheric vortex can also alter the mid-latitude circulation in winter, but responses are not consistent across models ( [[#Oudar--2020a|Oudar et al., 2020a]] ). Nevertheless, thermodynamic strengthening of moisture convergence into weather systems and polar regions is robust ( [[#8.2.2.1|Section 8.2.2.1]] ) and remains valid despite weak understanding of atmospheric circulation change.

In summary, there is ''high confidence'' that altered atmospheric wind patterns in response to radiative forcing and evolving surface temperature patterns will affect the regional water cycle in most regions. Mean tropical circulation is expected to slow with global warming ( ''high confidence'' ) but temporary multi-decadal strengthening is possible due to internal variability ( ''medium confidence'' ). Slowing of the tropical circulation reduces the meridional P–E gradient over the Pacific and can partly offset thermodynamic amplification of P–E patterns and strengthening of monsoons ( ''high confidence'' ) but regional characteristics of tropical rain belt changes are not well understood. There is ''medium confidence'' in processes driving strengthening and tightening of the ITCZ that increase the contrasts between wet and dry tropical weather regimes and seasons. There is ''high confidence'' in understanding of how radiative forcing and global warming drive a poleward expansion of the subtropics and mid-latitude storm tracks but only ''low confidence'' in how poleward expansion influences drying of subtropical and mid-latitude climates. There is ''low confidence'' in understanding how Arctic warming amplification affects mid-latitude regional water cycles but ''high confidence'' that thermodynamic strengthening of precipitation within weather systems and in monsoons and polar regions is robust to large-scale circulation changes.

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