Editing IPCC:AR6/WGI/Chapter-10 (section)

==== 10.1.3.1 Forcings Controlling Regional Climate ====

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There are important differences in the processes affected by greenhouse gases (GHGs) over land and ocean. Notably, this leads to preferential warming of the land regions, which are themselves skewed towards the Northern Hemisphere (NH).

Variations in solar forcing ( [[IPCC:Wg1:Chapter:Chapter-2#2.2.1|Section 2.2.1]] ) could influence regional climate through its modulation of circulation patterns, although this research field is still hampered by large observational and modelling uncertainties. The 11-year solar cycle has been suggested to affect the leading atmospheric circulation modes of the North Atlantic region in model-based studies ( [[#Gray--2013|Gray et al., 2013]] ; [[#Thiéblemont--2015|Thiéblemont et al., 2015]] ; [[#Sjolte--2018|Sjolte et al., 2018]] ). In particular the solar cycle has been suggested as an important source of near-term predictability of the North Atlantic Oscillation (NAO; [[#Kushnir--2019|Kushnir et al., 2019]] ), while other studies have not found evidence for links between the solar cycle and NAO in observational records ( [[#Ortega--2015|Ortega et al., 2015]] ; [[#Sjolte--2018|Sjolte et al., 2018]] ; [[#Chiodo--2019|Chiodo et al., 2019]] ). On centennial time scales, solar fluctuations were found to be correlated with the Eastern Atlantic Pattern ( [[#Sjolte--2018|Sjolte et al., 2018]] ). Possible influences on winter circulation and temperature over Eurasia ( [[#Chen--2015|Chen et al., 2015]] ) and North America ( [[#Liu--2014|Liu et al., 2014]] ; [[#Li--2018|Li and Xiao, 2018]] ) have also been identified.

An updated assessment of past changes in stratospheric ozone can be found in [[IPCC:Wg1:Chapter:Chapter-2#2.2.5.2|Section 2.2.5.2]] . The AR6 assesses that both GHG and stratospheric ozone depletion have contributed to the expansion of the zonal mean Hadley cell in the Southern Hemisphere (SH) for the period 1981–2000 with ''medium confidence'' [[IPCC:Wg1:Chapter:Chapter-3#3.3.3|Section 3.3.3]] ; [[#Garfinkel--2015|Garfinkel et al., 2015]] ; [[#Waugh--2015|Waugh et al., 2015]] ; [[#Grise--2019|Grise et al., 2019]] ). There is ''medium confidence'' that stratospheric ozone depletion contributed to the strengthening trend of the summer Southern Annular Mode (SAM) for the period 1970–1990, but this influence has been weaker since 2000 ( [[IPCC:Wg1:Chapter:Chapter-3#3.7.2|Section 3.7.2]] ). The poleward shift of the SH westerlies has also been explained by stratospheric ozone depletion ( [[#Solman--2016|Solman and Orlanski, 2016]] ). [[#10.4|Section 10.4]] assesses its role in the multi-decadal increase of rainfall in south-eastern South America and [[#10.6.2|Section 10.6.2]] does so for the occurrence of the Cape Town drought.

Both natural and anthropogenic aerosols are often emitted at a regional scale, have a short atmospheric lifetime (from a few hours to several days; Section 6.1), are dispersed regionally and affect climate at a regional scale through radiative cooling/heating and cloud microphysical effects (Chapter 8; [[#Rotstayn--2015|Rotstayn et al., 2015]] ; [[#Sherwood--2015|Sherwood et al., 2015]] ). The majority of aerosols scatter solar radiation, but with strong regional variations ( [[#Shindell--2009|Shindell and Faluvegi, 2009]] ) that lead to regional radiative effects of up to two orders of magnitude larger than the global average ( [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|B. Li et al., 2016]] ; K. [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|Li et al., 2016]] ; [[#Mallet--2016|Mallet et al., 2016]] ). Black carbon, instead, is known to absorb solar radiation, leading to regional atmospheric warming patterns due to its inhomogeneous spatial distribution ( [[#Gustafsson--2016|Gustafsson and Ramanathan, 2016]] ). Patterns of forcing generally follow those of aerosol burden. However, temperature and precipitation responses are both local and remote (Z. [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|Li et al., 2016]] ; [[#Kasoar--2018|Kasoar et al., 2018]] ; L. [[#Liu--2018|]] [[#Liu--2018|Liu et al., 2018]] ; [[#Samset--2018|Samset et al., 2018]] ; [[#Thornhill--2018|Thornhill et al., 2018]] ; [[#Westervelt--2018|Westervelt et al., 2018]] ). For instance, changes in aerosol concentrations in the NH have been reported to modulate monsoon precipitation in West Africa and the Sahel ( [[#Undorf--2018|Undorf et al., 2018]] ; [[#10.4.2.1|Section 10.4.2.1]] ) and in Asia (H. [[#Zhang--2018|]] [[#Zhang--2018|Zhang et al., 2018]] ; [[#10.6.3|Section 10.6.3]] ).

Natural aerosols include mineral dust, volcanic aerosol and sea salt. The feedback processes between climate and mineral dust as well as sea salt are assessed in Section 6.4, while the volcanic aerosol is dealt with in Cross-Chapter Box 4.1. Mineral dust created by wind erosion of arid and semi-arid surfaces dominates the aerosol load over some areas. The major sources of contemporary dust are located in the arid topographic basins of northern Africa, Middle East, Central and south-west Asia, the Indian subcontinent, and East Asia ( [[#Prospero--2002|Prospero et al., 2002]] ; [[#Ginoux--2012|Ginoux et al., 2012]] ) and emissions are controlled by changes in surface winds, precipitation, and vegetation ( [[#Ridley--2014|Ridley et al., 2014]] ; W. [[#Wang--2015|]] [[#Wang--2015|Wang et al., 2015]] ; [[#DeFlorio--2016|DeFlorio et al., 2016]] ; [[#Evan--2016|Evan et al., 2016]] ; [[#Pu--2018|Pu and Ginoux, 2018]] ). Dust both scatters and absorbs radiation and serves as a nuclei of warm and cold clouds (Chapter 6). The surface direct radiative effect is likely negative over land and ocean, especially when the assumed solar absorption by dust is large ( [[#Miller--2014|Miller et al., 2014]] ; [[#Strong--2015|Strong et al., 2015]] ). Surface temperature and precipitation adjust to the direct radiative effect with both sign and magnitude depending on the dust absorptive properties. Dust often cools the surface, but in regions such as the Sahara surface air temperature increases as the shortwave absorption by dust is increased, leading to increases of surface temperature over the major reflective dust sources ( [[#Miller--2014|Miller et al., 2014]] ; [[#Solmon--2015|Solmon et al., 2015]] ; [[#Strong--2015|Strong et al., 2015]] ; [[#Jin--2016|Jin et al., 2016]] ; [[#Sharma--2017|Sharma and Miller, 2017]] ).

Volcanic eruptions load the atmosphere with large amounts of sulphur, which is transformed through chemical reactions and micro-physics processes into sulphate aerosols (Cross-Chapter Box 4.1; [[#Stoffel--2015|Stoffel et al., 2015]] ; [[#LeGrande--2016|LeGrande et al., 2016]] ). If the plume reaches the stratosphere, sulphate aerosols can remain there for months or years (about two to three for large eruptions) and can then be transported to other areas by the Brewer-Dobson circulation. If the eruption occurs in the tropics, its plume is dispersed across the Earth in a few years, while if the eruption occurs in the high latitudes, aerosols mainly remain in the same hemisphere ( [[#Pausata--2015|Pausata et al., 2015]] ). The global temperature response observed after the last five major eruptions of the last two centuries is estimated to be around –0.2°C ( [[#Swingedouw--2017|Swingedouw et al., 2017]] ), in association with a general decrease of precipitation ( [[#Iles--2017|Iles and Hegerl, 2017]] ). Nevertheless, the statistical significance of the regional response remains difficult to evaluate over the historical era ( [[#Bittner--2016|Bittner et al., 2016]] ; [[#Swingedouw--2017|Swingedouw et al., 2017]] ) due to the small sampling of large volcanic eruptions over this period and the fact that the signal is superimposed upon relatively large internal variability ( [[#Gao--2018|Gao and Gao, 2018]] ; [[#Dogar--2019|Dogar and Sato, 2019]] ). Evidence from paleoclimate observations is therefore crucial to obtain a sufficient signal-to-noise ratio ( [[#Sigl--2015|Sigl et al., 2015]] ). Reconstructed modes of climate variability based on proxy records allow evaluation of the influence on those modes ( [[#Zanchettin--2013|Zanchettin et al., 2013]] ; [[#Ortega--2015|Ortega et al., 2015]] ; [[#Sjolte--2018|Sjolte et al., 2018]] ; [[#Michel--2020|Michel et al., 2020]] ).

Anthropogenic aerosols play a key role in climate change (Chapter 6). Although the global mean optical depth caused by anthropogenic aerosols did not change from 1975 to 2005 (Chapter 6), the regional pattern changed dramatically between Europe and eastern Asia ( [[#Fiedler--2017|Fiedler et al., 2017]] , 2019; [[#Stevens--2017|Stevens et al., 2017]] ). Large regional differences in present-day aerosol forcing exist with consequences for regional temperature, hydrological cycle and modes of variability (Chapter 8, [[#10.6|Section 10.6]] ). Examples of regions with a notable role for anthropogenic aerosol forcing are the Indian monsoon region ( [[#10.6.3|Section 10.6.3]] ) and the Mediterranean basin [[#10.6.4|Section 10.6.4]] ). Anthropogenic aerosols are also very relevant in many urban areas (Box 10.3; [[#Gao--2016|Gao et al., 2016]] ; [[#Kajino--2017|Kajino et al., 2017]] ).

The SRCCL assessed that nearly three-quarters of the land surface is under some form of land use, particularly in agriculture and forest management ( [[#Jia--2019|Jia et al., 2019]] ). The effects of land management on climate are much less studied than land cover effects although net cropland has changed little over the past 50 years, while land management has continuously changed ( [[#Jia--2019|Jia et al., 2019]] ). [[IPCC:Wg1:Chapter:Chapter-7#7.3.4.1|Section 7.3.4.1]] assesses the global influence of both land use and irrigation on the effective radiative forcings. Land cover changes and land management can influence climate locally, such as the urban heat island and non-locally as in the case of increased rainfall downwind of a city ( [[#Jia--2019|Jia et al., 2019]] ; Box 10.3) or the monsoon circulation affected by irrigation ( [[#10.6.3|Section 10.6.3]] ). The influence of land cover changes and land management on regional climate extremes is assessed in [[IPCC:Wg1:Chapter:Chapter-11#11.1.6|Section 11.1.6]] .

It is ''very likely'' that the global land surface air temperature response to urbanization is negligible ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1.1.3|Section 2.3.1.1.3]] ). However, there is evidence that urbanization may regionally amplify the air temperature response to climate change in different climatic zones ( [[#Mahmood--2014|Mahmood et al., 2014]] ), either under present ( [[#Doan--2016|Doan et al., 2016]] ; [[#Kaplan--2017|Kaplan et al., 2017]] ; X. [[#Li--2018|]] [[#Li--2018|]] [[#Li--2018|Li et al., 2018]] ) or future climate conditions ( [[#Argüeso--2014|Argüeso et al., 2014]] ; [[#Kim--2016|Kim et al., 2016]] ; [[#Kusaka--2016|Kusaka et al., 2016]] ; [[#Grossman-Clarke--2017|Grossman-Clarke et al., 2017]] ; [[#Krayenhoff--2018|Krayenhoff et al., 2018]] ). For instance, in northern Belgium, [[#Berckmans--2019|Berckmans et al. (2019)]] found that including urbanization scenarios for the near future (up to 2035) have a comparable influence on minimum temperature (increasing it by 0.6°C) to that of the GHG-induced warming under RCP8.5.

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