Editing IPCC:AR6/WGI/Chapter-2 (section)

==== 2.3.1.4 Atmospheric Circulation ====

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This section focuses on large-scale changes in a subset of components of the atmospheric circulation (Cross-Chapter Box 2.2). [[IPCC:Wg1:Chapter:Chapter-8|Chapter 8]] assesses large-scale as well as regional aspects of circulation components and their impact on the hydrological cycle, while ([[IPCC:Wg1:Chapter:Chapter-11|Chapter 11]] assesses the association of circulation changes and variability with extreme events.

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<span id="the-hadley-and-walker-circulations"></span>
===== 2.3.1.4.1 The Hadley and Walker circulations =====

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The AR5 reported ''low confidence'' in trends in the strength of the Hadley circulation (HC) and the Walker circulation (WC) due to uncertainties in available reanalysis datasets and the large interannual-to-decadal variability of associated circulation patterns. However, AR5 indicated a ''likely'' widening of the tropical belt since the 1970s, albeit with large uncertainty in the magnitude of this change. There was ''high confidence'' that the post-1990s strengthening of the Pacific WC reversed its weakening observed from the mid-19th century to the 1990s.

Paleo reconstructions of rainfall and trade winds extending over the last 100 kyr show an intensification of the NH HC concurrently with a weakening of the SH HC and a southward shift of the inter tropical convergence zone (ITCZ) during Heinrich stadials ([[#Deplazes--2013|Deplazes et al., 2013]]; [[#McGee--2018|McGee et al., 2018]]; [[#Stríkis--2018|Stríkis et al., 2018]]; [[#Wendt--2019|Wendt et al., 2019]]). An intensification of the HC associated with conditions similar to La Niña (northward migrations of both the ITCZ and the SH westerlies) was found in reconstructions for the MH ([[#McGee--2014|McGee et al., 2014]]; [[#Mollier-Vogel--2019|Mollier-Vogel et al., 2019]]). Changes in insolation from the mid to late Holocene favoured a southward migration in the position of the ITCZ and the descending branch of the HC in the NH, approaching its current width and position ([[#Wirth--2013|Wirth et al., 2013]]; [[#Thatcher--2020|Thatcher et al., 2020]]). Tree ring chronologies from the NH mid-latitudes over the last 800 years show that the northern edge of the HC tended to migrate southward during positive phases of ENSO and PDV, with northward shifts during negative phases ([[#Alfaro-Sánchez--2018|Alfaro-Sánchez et al., 2018]]). Between 1400 and 1850 CE the HC over both hemispheres and the ITCZ were displaced southward, consistent with occurrence of drought conditions in several NH regions ([[#Wirth--2013|Wirth et al., 2013]]; [[#Burn--2014|Burn and Palmer, 2014]]; [[#Lechleitner--2017|Lechleitner et al., 2017]]; [[#Alfaro-Sánchez--2018|Alfaro-Sánchez et al., 2018]]; [[#Flores-Aqueveque--2020|Flores-Aqueveque et al., 2020]]). Moreover, several proxy records showed not only inter-hemispheric shifts in the ITCZ but a contraction of the tropical belt during 1400–1850 CE, which followed an expansion during 950–1250 CE ([[#Denniston--2016|Denniston et al., 2016]]; [[#Griffiths--2016|Griffiths et al., 2016]]).

From centennial-scale reanalyses, [[#Liu--2012|Liu et al. (2012)]] and [[#D’Agostino--2017|D’Agostino and Lionello (2017)]] found divergent results on HC extent over the last 150 years, although with unanimity upon an intensification of the SH HC. A substantial discrepancy between HC characteristics in centennial-scale reanalyses and in ERA-Interim ([[#D’Agostino--2017|D’Agostino and Lionello, 2017]]) since 1979 yields significant questions regarding their ability to capture changes in HC behaviour. Taken together with the existence of apparent non-climatic artefacts in the datasets ([[#Nguyen--2015|Nguyen et al., 2015]]), this implies ''low confidence'' in changes in the extent and intensity of HC derived from centennial-scale reanalyses. However, using multiple observational datasets and centennial-scale reanalyses, [[#Bronnimann--2015|Bronnimann et al. (2015)]] identified a southward shift in the NH HC edge from 1945 to 1980 of about 0.25° latitude per decade, consistent with observed changes in global land monsoon precipitation ([[#2.3.1.4.2|Section 2.3.1.4.2]]).

Since AR5 several studies based upon a range of metrics and different reanalyses products have suggested that the annual mean HC extent has shifted poleward at an approximate rate of 0.1°–0.5° latitude per decade over the last about 40 years ([[#Allen--2017|Allen and Kovilakam, 2017]]; [[#Davis--2017|Davis and Birner, 2017]]; [[#Grise--2018|Grise et al., 2018]]; [[#Staten--2018|Staten et al., 2018]], 2020; [[#Studholme--2018|Studholme and Gulev, 2018]]; [[#Grise--2020|Grise and Davis, 2020]]). The observed widening of the annual mean HC, revealed by a variety of metrics, is primarily due to poleward shift of the Northern Hemisphere HC. There have been stronger upward trends in the NH extent of HC after 1992 (Figure 2.17a). The estimated magnitude of the recent changes based on modern-era reanalyses is not as large as that in AR5, due to apparent biases in older-generation reanalyses ([[#Grise--2019|Grise et al., 2019]]). Moreover, large interannual variability leads to uncertainties in estimates of long-term changes ([[#Nguyen--2013|Nguyen et al., 2013]]; [[#Garfinkel--2015b|Garfinkel et al., 2015b]]; [[#Seviour--2018|Seviour et al., 2018]]; [[#Staten--2018|Staten et al., 2018]]), particularly for the NH given its zonal asymmetries ([[#Staten--2020|Staten et al., 2020]]; [[#Wang--2020|Wang et al., 2020]]). These large-scale features of the HC based on reanalyses agree with estimates revealed from the Integrated Global Radiosonde Archive (IGRA) during 1979–2012 ([[#Lucas--2015|Lucas and Nguyen, 2015]]; [[#Mathew--2016|Mathew et al., 2016]]). Recent trends based on reanalyses indicate a larger seasonal widening in the HC for summer and autumn in each hemisphere, although the magnitude of changes in HC extent is strongly dependent on dataset and metrics used ([[#Grise--2018|Grise et al., 2018]]; Y. [[#Hu--2018|]] [[#Hu--2018|Hu et al., 2018]]; [[#Staten--2018|Staten et al., 2018]]). The shifts in the HC position were accompanied by a narrowing ITCZ over the Atlantic and Pacific basins, with no significant change in its location and increases in the precipitation intensity ([[#Byrne--2018|Byrne et al., 2018]]).

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[[File:edccb51f7c800aa7409e4500c5eec32e IPCC_AR6_WGI_Figure_2_17.png]]

'''Figure''' '''2.17 |''' '''Time series of the annual mean Northern Hemisphere (NH, top curves) and Southern Hemisphere (SH, bottom curves) Hadley cell extent (a) and Hadley cell intensity (b) since 1979.''' Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Trends in the HC intensity since 1979 differ between reanalyses, although there is a tendency toward HC intensification (Figure 2.17b; [[#Nguyen--2013|Nguyen et al., 2013]]; [[#Chen--2014|Chen et al., 2014]]; [[#D’Agostino--2017|D’Agostino and Lionello, 2017]]; R. [[#Huang--2019|Huang et al., 2019]]), which is more marked in the NH than the SH ([[#Studholme--2018|Studholme and Gulev, 2018]]). However, the ability of reanalyses to represent the HC strength has been questioned due to inaccurate representation of latent heating distribution, which is directly related to tropical convection and influences the HC dynamics ([[#Chemke--2019|Chemke and Polvani, 2019]]; [[#Mathew--2019|Mathew and Kumar, 2019]]).

Paleo evidence during the LGM indicates a weaker WC over the Indian Ocean ([[#DiNezio--2018|DiNezio et al., 2018]]; [[#Windler--2019|Windler et al., 2019]]) with a stronger Pacific WC ([[#DiNezio--2013|DiNezio and Tierney, 2013]]). During the Holocene, a transition from a strong WC located more westward during the Early-to-Mid Holocene towards a weak and eastward shifted WC during the late Holocene was inferred from proxy records from the Pacific Warm Pool and South East Asia ([[#Barr--2019|Barr et al., 2019]]; [[#Dang--2020|Dang et al., 2020]]; [[#Griffiths--2020|Griffiths et al., 2020]]), in concurrence with changes in ENSO activity ([[#2.4.2|Section 2.4.2]]). Reconstructions for the CE showed weakened WC during 1000–1250 and since 1850, with an intensified circulation during 1500–1850 CE ([[#Xu--2016|Xu et al., 2016]]; [[#Deng--2017|Deng et al., 2017]]).

Considering instrumental records, there is considerable interdecadal variability in the strength of the WC, resulting in time-period dependent magnitude and even sign of trends ([[#Carilli--2015|Carilli et al., 2015]]; [[#Bordbar--2017|Bordbar et al., 2017]]; [[#Hou--2018|Hou et al., 2018]]), with some studies reporting weakening over the 20th century (e.g., [[#Power--2011|Power and Kociuba, 2011]]; [[#Liu--2019|Liu et al., 2019]]), while others reported strengthening (Z. [[#Li--2020|]] [[#Li--2020|Li et al., 2020]]), particularly over the last 30–40 years (e.g., [[#Hu--2013|Hu et al., 2013]]; [[#L’Heureux--2013|L’Heureux et al., 2013]]; [[#Yim--2017|Yim et al., 2017]]). Based on estimation of changes in mid-tropospheric velocity from changes in observed cloud cover, [[#Bellomo--2015|Bellomo and Clement (2015)]] suggest a weakening and eastward shift of the WC over 1920–2010, however the robustness of this signal is questionable due to high uncertainty in the ship-reported cloud data used before 1954. Using centennial-scale 20CR reanalysis [[#Tseng--2019|Tseng et al. (2019)]] showed that the vertical westerly wind shear over the western Pacific does not indicate any long-term change during 1900–1980, but shows a marked increase since the 1980s that is not present in ERA-Interim and JRA-55, again calling into question the ability of centennial-scale reanalyses to capture tropical circulation changes. Recent strengthening together with a westward shift of the WC ([[#Bayr--2014|Bayr et al., 2014]]; [[#Ma--2016|Ma and Zhou, 2016]]) was identified across several reanalysis products and observational datasets, and using different metrics for quantifying WC. Nevertheless, satellite observations of precipitation and analyses of upper tropospheric humidity suggest substantially weaker strengthening of the WC than implied by reanalyses ([[#Chung--2019|Chung et al., 2019]]). This recent strengthening in the WC is associated with enhanced precipitation in the tropical western Pacific, anomalous westerlies in the upper troposphere, strengthened downwelling in the central and eastern tropical Pacific, and anomalous surface easterlies in the western and central tropical Pacific ([[#Dong--2013|Dong and Lu, 2013]]; [[#McGregor--2014|McGregor et al., 2014]]; [[#Choi--2016|Choi et al., 2016]]). Positive trends in sea level pressure over the eastern Pacific and concurrent negative trends over the Indonesian region result in a pattern implying a shift towards a La Niña-like WC regime, with strengthening of the Pacific Trade Winds mainly over 1979–2012 ([[#L’Heureux--2013|L’Heureux et al., 2013]]; [[#England--2014|England et al., 2014]]; [[#Sohn--2016|Sohn et al., 2016]]; [[#Zhao--2019|Zhao and Allen, 2019]]). Seasonal assessment of the WC showed significant changes in the vertical westerly wind shear over the Pacific during the austral summer and autumn implying a strengthening ([[#Clem--2017|Clem et al., 2017]]).

In summary, there has been a ''likely'' widening of the Hadley circulation since the 1980s, mostly due to its extension in the NH, although there is only ''medium confidence'' in the extent of the changes. This has been accompanied by a strengthening of the Hadley circulation, particularly in the NH (''medium confidence''). There is ''low confidence'' in the estimation of long-term trends in the strength of the Walker circulation, which are time period dependent and subject to dataset uncertainties. Trends since 1980 are better characterized and consistent with a ''very likely'' strengthening that resembles a La Niña-like Walker circulation and a westward shift of the Walker circulation, although with ''medium confidence'' in the magnitude of the changes, arising from the differences between satellite observations and reanalysis products.

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<span id="global-monsoon-gm-changes"></span>

===== 2.3.1.4.2 Global monsoon (GM) changes =====

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The AR5 reported a weakening of the global monsoon (GM) circulation as well as a decrease of global land monsoon rainfall over the second half of the 20th century. Nevertheless, there was ''low confidence'' in the observed circulation trends due to uncertainties in reanalysis products and in the definition of the monsoon area. From a paleo perspective, AR5 only assessed regional monsoon changes.

New research based on high-resolution proxies reinforces previous findings on the influence of orbital cycles on GM variability on millennial time scales. The intensity of the monsoon systems is generally out of phase between hemispheres, being associated with the precession cycle (about 21–23 kyr) ([[#An--2015|An et al., 2015]]; P.X. [[#Wang--2017|Wang et al., 2017]]; [[#Seth--2019|Seth et al., 2019]]), with intensified NH monsoon systems during precession minima ([[#Toucanne--2015|Toucanne et al., 2015]]; [[#Wagner--2019|Wagner et al., 2019]]). The eccentricity forcing (about 100 kyr cycle) shows stronger GM during interglacial periods (P.X. [[#Wang--2014|]] [[#Wang--2014|Wang et al., 2014]], 2017; [[#An--2015|An et al., 2015]]; [[#Mohtadi--2016|Mohtadi et al., 2016]]). Changes in obliquity (about 41 kyr cycle) modify the strength of monsoon systems, with increased summer monsoon rainfall when obliquity is maximal (Y. [[#Liu--2015|]] [[#Liu--2015|Liu et al., 2015]] b; [[#Mohtadi--2016|Mohtadi et al., 2016]]). Millennial scale variability in GM during the LDT was also linked to the occurrences of Heinrich stadials, resulting in weakened NH monsoons and intensified SH monsoons ([[#An--2015|An et al., 2015]]; P.X. [[#Wang--2017|Wang et al., 2017]]; [[#Margari--2020|Margari et al., 2020]]).

An intensification of the NH monsoons in the early to mid-Holocene with increased precipitation and regional expansions of rainfall areas identified through a variety of proxy records is shown by [[#Biasutti--2018|Biasutti et al. (2018)]] and P.X. [[#Wang--2017|Wang et al. (2017)]]. The response for the SH monsoons during this period indicates a weakening in both summer and winter precipitation (P.X. [[#Wang--2014|]] [[#Wang--2014|Wang et al., 2014]], 2017; [[#Sachs--2018|Sachs et al., 2018]]). A decline in GM precipitation and a retraction of the northern fringes of monsoon areas was inferred from the mid-Holocene onwards, with some regions experiencing wetter conditions during the mid to late Holocene compared with present and a strengthening of the SH monsoons (P.X. [[#Wang--2014|]] [[#Wang--2014|Wang et al., 2014]], 2017; [[#Sachs--2018|Sachs et al., 2018]]). For the CE, GM reconstructions exhibit inter-hemispheric contrast during the period 950–1250 CE, with intensified NH monsoons and weakened SH monsoons, and the opposite pattern during 1400–1850 CE (P.X. [[#Wang--2014|]] [[#Wang--2014|Wang et al., 2014]]; [[#An--2015|An et al., 2015]]).

Direct observations highlight that the GM land precipitation, particularly over the NH, experienced a slight increase from 1900 through the early 1950s, followed by an overall decrease from the 1950s to the 1980s, and then an increase to present ([[#Kitoh--2013|Kitoh et al., 2013]]; [[#Wang--2018|]] [[#Wang--2018|B. Wang et al., 2018]], 2021; X. [[#Huang--2019b|]] [[#Huang--2019|Huang et al., 2019]] b). This highlights the existence of multi-decadal variations in the NH monsoon circulation patterns and precipitation intensity ([[#Wang--2013|Wang et al., 2013]]; P.X. [[#Wang--2014|]] [[#Wang--2014|Wang et al., 2014]], 2017; [[#Monerie--2019|Monerie et al., 2019]]). An overall increase in monsoon precipitation during extended boreal summer (JJAS) over the NH since 1979 is revealed by GPCP ([[#Deng--2018|Deng et al., 2018]]; [[#Han--2019|Han et al., 2019]]) and CMAP for 1980–2010 ([[#Jiang--2016|Jiang et al., 2016]]). SH summer monsoon behaviour is dominated by strong interannual variability and large regional differences ([[#Kitoh--2013|Kitoh et al., 2013]]; [[#Lin--2014|Lin et al., 2014]]; [[#Jiang--2016|Jiang et al., 2016]]; [[#Kamae--2017|Kamae et al., 2017]]; [[#Deng--2018|Deng et al., 2018]]; [[#Han--2019|Han et al., 2019]]), with no significant trends reported by GPCP and CMAP ([[#Deng--2018|Deng et al., 2018]]). Uncertainty predominantly arises from the observed increase in tropical precipitation seasonality ([[#Feng--2013|Feng et al., 2013]]) and the estimation of GM precipitation over the ocean areas, leading to a large apparent spread across datasets ([[#Kitoh--2013|Kitoh et al., 2013]]; [[#Kamae--2017|Kamae et al., 2017]]).

In summary, observed trends during the last century indicate that the GM precipitation decline reported in AR5 has reversed since the 1980s, with a ''likely'' increase mainly due to a significant positive trend in the NH summer monsoon precipitation (''medium confidence''). However, GM precipitation has exhibited large multi-decadal variability over the last century, creating ''low confidence'' in the existence of centennial-length trends in the instrumental record. Proxy reconstructions show a ''likely'' NH monsoons weakening since the mid-Holocene, with opposite behaviour for the SH monsoons.

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<span id="extratropical-jets-storm-tracks-and-blocking"></span>

===== 2.3.1.4.3 Extratropical jets, storm tracks, and blocking =====

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The AR5 reported a ''likely'' poleward shift of storm tracks and jet streams since the 1970s from different datasets, variables and approaches. These trends were consistent with the HC widening and the poleward shifting of the circulation features since the 1970s. There was ''low confidence'' in any large-scale change in blocking.

Proxy records consistent with modelling results imply a southward shift of the storm tracks over the North Atlantic during the LGM ([[#Raible--2021|Raible et al., 2021]]). A variety of proxies are available for the changes in the position of the extratropical jets/westerlies during the Holocene. Recent syntheses of moisture-sensitive proxy records indicate drier-than-present conditions over mid-latitudes of western North America ([[#Hermann--2018|Hermann et al., 2018]]; [[#Liefert--2020|Liefert and Shuman, 2020]]) during the MH, which together with a weakened Aleutian Low ([[#Bailey--2018|Bailey et al., 2018]]) implies that the winter North Pacific jetstream was shifted northward. A synthesis of lines of evidence from the SH indicates that the westerly winds were stronger over 14–5 ka, followed by regional asymmetry after 5 ka ([[#Fletcher--2012|Fletcher and Moreno, 2012]]). There is no consensus on the shifts of the SH westerlies with some studies implying poleward migrations ([[#Lamy--2010|Lamy et al., 2010]]; [[#Voigt--2015|Voigt et al., 2015]]; [[#Turney--2017|Turney et al., 2017]]; [[#Anderson--2018|Anderson et al., 2018]]) and others suggesting an equatorward shift ([[#Kaplan--2016|Kaplan et al., 2016]]) in the MH.

During 950–1400 CE, hydroclimate indicators suggest a northward shift of Pacific storm tracks over North America ([[#McCabe-Glynn--2013|McCabe-Glynn et al., 2013]]; [[#Steinman--2014|Steinman et al., 2014]]) which was comparable in magnitude to that over 1979–2015 (J. [[#Wang--2017a|]] [[#Wang--2017|Wang et al., 2017]] a). Storm tracks over the North Atlantic-European sector shifted northward as indicated by multi-proxy indicators over the North Atlantic ([[#Wirth--2013|Wirth et al., 2013]]; [[#Orme--2017|Orme et al., 2017]]) and Mediterranean ([[#Roberts--2012|Roberts et al., 2012]]). Reconstructed westerly winds in the SH suggest a poleward shift ([[#Lamy--2010|Lamy et al., 2010]]; [[#Schimpf--2011|Schimpf et al., 2011]]; [[#Goodwin--2014|Goodwin et al., 2014]]; [[#Koffman--2014|Koffman et al., 2014]]; [[#Moreno--2018|Moreno et al., 2018]]), with latitudinal change comparable to that during recent decades ([[#Swart--2012|Swart and Fyfe, 2012]]; [[#Manney--2018|Manney and Hegglin, 2018]]).

Multiple reanalyses show that since 1979 the subtropical jet wind speeds have generally increased in winter and decreased in summer in both hemispheres, but the trends are regionally dependent ([[#Pena-Ortiz--2013|Pena-Ortiz et al., 2013]]; [[#Manney--2018|Manney and Hegglin, 2018]]; S.H. [[#Lee--2019|]] [[#Lee--2019|Lee et al., 2019]]). Over NH mid-latitudes, the summer zonal wind speeds have weakened in the mid-troposphere ([[#Francis--2012|Francis and Vavrus, 2012]]; [[#Coumou--2014|Coumou et al., 2014]], 2015; [[#Haimberger--2017|Haimberger and Mayer, 2017]]). Meanwhile there are indications of enhanced jetstream meandering in boreal autumn at the hemispheric scale ([[#Francis--2015|Francis and Vavrus, 2015]]; [[#Di%20Capua--2016|Di Capua and Coumou, 2016]]), whereas the regional arrangement of meandering depends on the background atmospheric state ([[#Cohen--2020|Cohen et al., 2020]]). These meandering trends, however, are sensitive to the metrics used ([[#Screen--2013|Screen and Simmonds, 2013]]; [[#Hassanzadeh--2014|Hassanzadeh et al., 2014]]; [[#Cattiaux--2016|Cattiaux et al., 2016]]; [[#Vavrus--2018|Vavrus, 2018]]). Hypothesized links to Arctic warming are assessed in Cross-Chapter Box 10.1.

Multiple reanalyses and radiosonde observations show an increasing number of extratropical cyclones over the NH since the 1950s ([[#Chang--2016|Chang and Yau, 2016]]; X.L. [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|Wang et al., 2016]]). The positive trends are generally consistent among reanalyses since 1979, though with considerable spread ([[#Tilinina--2013|Tilinina et al., 2013]]; X.L. [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|Wang et al., 2016]]). In recent decades the number of deep extratropical cyclones has increased over the SH (Section 8.3.2.8.1 and Figure 8.12; [[#Reboita--2015|Reboita et al., 2015]]; X.L. [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|Wang et al., 2016]]), while the number of deep cyclones has decreased in the NH in both winter and summer ([[#Neu--2013|Neu et al., 2013]]; [[#Coumou--2015|Coumou et al., 2015]]; [[#Chang--2016|Chang et al., 2016]]; J. [[#Wang--2017a|]] [[#Wang--2017|Wang et al., 2017]] a ; [[#Gertler--2019|Gertler and O’Gorman, 2019]]). The regional changes for different intensity extratropical cyclones are assessed in Section 8.3.2.8.1. The assessment of trends is complicated by strong interannual to decadal variability, sensitivity to dataset choice and resolution ([[#Tilinina--2013|Tilinina et al., 2013]]; [[#Lucas--2014|Lucas et al., 2014]]; X.L. [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|Wang et al., 2016]]; [[#Pepler--2018|Pepler et al., 2018]]; [[#Rohrer--2018|Rohrer et al., 2018]]) and cyclone identification/tracking methods ([[#Neu--2013|Neu et al., 2013]]; [[#Grieger--2018|Grieger et al., 2018]]). Thus there is overall ''low'' ''confidence'' for recent changes in global extratropical storm tracks.

A consistent poleward shift of the tropospheric extratropical jets since 1979 is reported by multiple reanalyses (Figure 2.18; [[#Davis--2012|Davis and Rosenlof, 2012]]; [[#Davis--2013|Davis and Birner, 2013]]; [[#Pena-Ortiz--2013|Pena-Ortiz et al., 2013]]; [[#Manney--2018|Manney and Hegglin, 2018]]), and radiosonde winds ([[#Allen--2012|Allen et al., 2012]]). This is generally consistent with the previously reported shifts retrieved from satellite temperature observations ([[#Fu--2011|Fu and Lin, 2011]]; [[#Davis--2012|Davis and Rosenlof, 2012]]). After the 1960s the magnitude of meridional shifts in extratropical jets over both the North Atlantic and North Pacific in August is enhanced compared to multi-century variability ([[#Trouet--2018|Trouet et al., 2018]]). Despite some regional differences ([[#Woollings--2014|Woollings et al., 2014]]; [[#Norris--2016|Norris et al., 2016]]; J. [[#Wang--2017a|]] [[#Wang--2017|Wang et al., 2017]] a ; [[#Xue--2017|Xue and Zhang, 2017]]; [[#Ma--2018|Ma and Zhang, 2018]]; [[#Melamed-Turkish--2018|Melamed-Turkish et al., 2018]]), overall poleward deflection of storm tracks in boreal winter over both the North Atlantic and the North Pacific was identified during 1979–2010 ([[#Tilinina--2013|Tilinina et al., 2013]]). Over the SH extra-tropics there is a similarly robust poleward shift in the polar jet since 1979 ([[#Pena-Ortiz--2013|Pena-Ortiz et al., 2013]]; [[#Manney--2018|Manney and Hegglin, 2018]]; [[#WMO--2018|WMO, 2018]]), although after 2000 the December–January–February (DJF) tendency to poleward shift of the SH jet stream position ceased ([[#Banerjee--2020|Banerjee et al., 2020]]). The general poleward movement in midlatitude jet streams ([[#Lucas--2014|Lucas et al., 2014]]) is consistent with the expansion of the tropical circulation ([[#2.3.1.4.1|Section 2.3.1.4.1]]). The changes of extratropical jets and westerlies are also related to the annular modes of variability ([[#2.4|Section 2.4]] and Annex IV).

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[[File:fd4b5ce5996815b4bafa896df1e3d7fb IPCC_AR6_WGI_Figure_2_18.png]]

'''Figure 2.1''' '''8 |''' '''Trends in ERA5 zonal-mean zonal wind speed.''' Shown are '''(a)''' DJF (December–January–February); '''(b)''' MAM (March–April–May); '''(c)''' JJA (June–July–August); and '''(d)''' SON (September–October–November). Climatological zonal winds during the data period are shown in solid contour lines for westerly winds and in dashed lines for easterly. Trends are calculated using OLS regression with significance assessed following AR(1) adjustment after [[#Santer--2008|Santer et al. (2008)]] (‘×’ marks denote non-significant trends). Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Robust trends in blocking have only been found in certain regions and specific seasons during recent decades. Increases in blocking frequency have occurred over low-latitude regions in the North Atlantic in boreal winter ([[#Davini--2012|Davini et al., 2012]]), the South Atlantic in austral summer ([[#Dennison--2016|Dennison et al., 2016]]) and the southern Indian Ocean in austral spring ([[#Schemm--2018|Schemm, 2018]]). Over the subpolar North Atlantic sustained periods of positive Greenland blocking were identified during 1870–1900 and from the late 1990s to 2015 ([[#Hanna--2015|Hanna et al., 2015]]). Further analysis of association of Greenland blocking with the NAM is provided in [[#2.4.1.1|Section 2.4.1.1]]. Meanwhile, a reduced blocking frequency has been found over winter in Siberia ([[#Davini--2012|Davini et al., 2012]]) and the south-western Pacific in austral spring ([[#Schemm--2018|Schemm, 2018]]). Over eastern European Russia and western Siberia (40°E–100°E) a tendency towards longer blocking events was reported by [[#Luo--2016|Luo et al. (2016)]] for 2000–2013 and by [[#Tyrlis--2020|Tyrlis et al. (2020)]] for 1979–2017. Inter-annual variance in the number of blocking events over the SH ([[#Oliveira--2014|Oliveira et al., 2014]]) and North Atlantic ([[#Kim--2015|Kim and Ha, 2015]]) has enhanced. Blocking events and their trends are sensitive to choice of datasets, calculation periods and methods ([[#Cheung--2013|Cheung et al., 2013]]; [[#Barnes--2014|Barnes et al., 2014]]; [[#Pepler--2018|Pepler et al., 2018]]; [[#Rohrer--2018|Rohrer et al., 2018]]; [[#Woollings--2018b|Woollings et al., 2018b]]; [[#Kononova--2020|Kononova and Lupo, 2020]]). As a result, hemispheric and global trends in blocking frequency have overall ''low'' ''confidence.''

In summary, the total number of extratropical cyclones has ''likely'' increased since the 1980s in the NH (''low confidence''), but with fewer deep cyclones particularly in summer. The number of strong extratropical cyclones has ''likely'' increased in the SH (''medium confidence''). The extratropical jets and cyclone tracks have ''likely'' been shifting poleward in both hemispheres since the 1980s with marked seasonality in trends (''medium confidence''). There is ''low confidence'' in shifting of extratropical jets in the NH during the mid-Holocene and over 950–1400 CE to latitudes that ''likely'' were similar to those since 1979. There is ''low confidence'' in observed global-scale changes in the occurrence of blocking events.

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<span id="surface-wind-and-sea-level-pressure"></span>

===== 2.3.1.4.4 Surface wind and sea level pressure =====

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The AR5 concluded that surface winds over land had generally weakened. The ''confidence'' for both land and ocean surface wind trends was ''low'' owing to uncertainties in datasets and measures used. Sea level pressure (SLP) was assessed to have ''likely'' decreased from 1979–2012 over the tropical Atlantic and increased over large regions of the Pacific and South Atlantic, but trends were sensitive to the period analysed.

Terrestrial in situ wind datasets have been updated and the quality-control procedures have been improved, with particular attention to homogeneity and to better retaining true extreme values ([[#Dunn--2012|Dunn et al., 2012]], 2014, 2016). Global mean land wind speed (excluding Australia) from HadISD for 1979–2018 shows a reduction (stilling) of 0.063 m s <sup>–1</sup> per decade ([[#Azorin-Molina--2019|Azorin-Molina et al., 2019]]). Trends are broadly insensitive to the subsets of stations used. Although the meteorological stations are unevenly distributed worldwide and sparse in South America and Africa, the majority exhibit stilling particularly in the NH (Figure 2.19). Regionally, strong decreasing trends are reported in central Asia and North America (–0.106 and –0.084 m s <sup>–1</sup> per decade respectively) during 1979–2018 ([[#McVicar--2012|McVicar et al., 2012]]; [[#Vautard--2012|Vautard et al., 2012]]; J. [[#Wu--2018|]] [[#Wu--2018|Wu et al., 2018]]; [[#Azorin-Molina--2019|Azorin-Molina et al., 2019]]). This stilling tendency has reversed after 2010 and the global mean surface winds have strengthened ([[#Zeng--2019b|Zeng et al., 2019b]]; [[#Azorin-Molina--2020|Azorin-Molina et al., 2020]]), although the robustness of this reversal is unclear given the short period and interannual variability ([[#Kousari--2013|Kousari et al., 2013]]; [[#Kim--2015|Kim and Paik, 2015]]; [[#Azorin-Molina--2019|Azorin-Molina et al., 2019]]).

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[[File:26ea185ca5ced7098b8a78c982776822 IPCC_AR6_WGI_Figure_2_19.png]]

'''Figure 2.19''' '''|''' '''Trends in surface wind speed. (a)''' Station observed winds from the integrated surface database (HadISD v2.0.2.2017f); '''(b)''' Cross-Calibrated Multi-Platform wind product; '''(c)''' ERA5; and '''(d)''' wind speed from the Objectively Analyzed Air-Sea Heat Fluxes dataset, release 3 (OAFLUX, release 3). White areas indicate incomplete or missing data. Trends are calculated using OLS regression with significance assessed following AR(1) adjustment after [[#Santer--2008|Santer et al. (2008)]]; ‘×’ marks denote non-significant trends. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Over the ocean, datasets demonstrate considerable disagreement in surface wind speed trends and spatial features ([[#Kent--2013|Kent et al., 2013]]). Global ocean surface winds from NOCv2.0 demonstrate upward trends of about 0.11 m s <sup>–1</sup> per decade (1979–2015) with somewhat smaller trends from WASwind for 1979–2011 ([[#Azorin-Molina--2017|Azorin-Molina et al., 2017]], 2019). The trends are consistent until 1998, but diverge thereafter. Both ERA5 and JRA-55 reanalyses show consistently increasing global marine wind speeds over 1979–2015, though flattening since 2000, whereas MERRA-2 agrees until 1998, but then exhibits increased variability and an overall decrease in the last two decades ([[#Azorin-Molina--2019|Azorin-Molina et al., 2019]]). This agrees with estimates by [[#Sharmar--2021|Sharmar et al. (2021)]] showing upward ocean wind trends from 1979 to 2000 which are consistent in ERA-Interim, ERA5 and MERRA-2, but disagree with CFSR trends for the same period. Over 2000–2019 all reanalyses show diverging tendencies. An updated multiplatform satellite database (comprising data from altimeters, radiometers, and scatterometers) from 1985–2018 shows small increases in mean wind speed over the global ocean, with the largest increase observed in the Southern Ocean ([[#Young--2019|Young and Ribal, 2019]]), consistent with signals in ERA-Interim, ERA5 and MERRA-2 ([[#Sharmar--2021|Sharmar et al., 2021]]). Overall, most products suggest positive trends over the Southern Ocean, western North Atlantic and the tropical eastern Pacific since the early 1980s.

The modern era reanalyses exhibit SLP increases over the SH subtropics with stronger increases in austral winter over 1979–2018. Over the NH, SLP increased over the mid-latitude Pacific in boreal winter and decreased over the eastern subtropical and mid-latitude North Atlantic in boreal summer. Discrepancies in the low-frequency variations during the first half of the 20th century exist in the centennial-scale reanalysis products ([[#Befort--2016|Befort et al., 2016]]). Overall, modern reanalysis datasets support the AR5 conclusion that there is no clear signal for trends in the strength and position of the permanent and quasi-permanent pressure centres of action since the 1950s. Instead, they highlight multi-decadal variations. Large-scale SLP is strongly associated with the changes in modes of variability ([[#2.4|Section 2.4]] and Annex IV).

In summary, since the 1970s a worldwide weakening of surface wind has ''likely'' occurred over land, particularly marked in the NH, with ''low confidence'' in a recent partial recovery since around 2010. Differences between available wind speed estimates lead to ''low confidence'' in trends over the global ocean as a whole but with most estimates showing strengthening globally over 1980–2000 and over the last four decades in the Southern Ocean, western North Atlantic and the tropical eastern Pacific.

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===== 2.3.1.4.5 Stratospheric polar vortex and sudden warming events =====

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The AR5 assessed changes in the polar vortices and reported a ''likely'' decrease in the lower-stratospheric geopotential heights over Antarctica in spring and summer at least since 1979.

Multiple definitions for the polar vortex strength and sudden stratospheric warming (SSW) events have been proposed and compared ([[#Butler--2015|Butler et al., 2015]]; [[#Palmeiro--2015|Palmeiro et al., 2015]]; [[#Waugh--2017|Waugh et al., 2017]]; [[#Butler--2018|Butler and Gerber, 2018]]), and new techniques identifying daily vortex patterns and SSWs have been developed (D.M. [[#Mitchell--2013|]] [[#Mitchell--2013|Mitchell et al., 2013]]; [[#Kretschmer--2018|Kretschmer et al., 2018]]). Errors in reanalysis stratospheric winds were assessed and discrepancies in stratospheric atmospheric circulation and temperatures between reanalyses, satellites and radiosondes have been reported (D.M. [[#Mitchell--2013|]] [[#Mitchell--2013|Mitchell et al., 2013]]; [[#Duruisseau--2017|Duruisseau et al., 2017]]).

The northern stratospheric polar vortex has varied intra-seasonally and with altitude during recent decades. Multiple reanalysis and radiosonde datasets show that the midwinter lower stratospheric geopotential height (150 hPa) over the polar region north of 60°N has increased significantly since the early 1980s ([[#Bohlinger--2014|Bohlinger et al., 2014]]; [[#Garfinkel--2017|Garfinkel et al., 2017]]). This signal extends to the middle and upper stratosphere. In January-February zonal winds north of 60°N at 10 hPa have been weakening ([[#Kim--2014|Kim et al., 2014]]; [[#Kretschmer--2018|Kretschmer et al., 2018]]). Daily atmospheric circulation patterns over the northern polar stratosphere exhibit a decreasing frequency of strong vortex events and commensurate increase in more-persistent weak events, which largely explains the observed significant weakening of the vortex during 1979–2015 ([[#Kretschmer--2018|Kretschmer et al., 2018]]). The northern polar vortex has weakened in early winter but strengthened during late winter ([[#Bohlinger--2014|Bohlinger et al., 2014]]; [[#Garfinkel--2015a|Garfinkel et al., 2015a]], 2017; [[#Ivy--2016|Ivy et al., 2016]]; [[#Seviour--2017|Seviour, 2017]]; [[#Kretschmer--2018|Kretschmer et al., 2018]]). In the middle and upper stratosphere, a strengthening trend of the northern polar vortex during DJF has occurred since 1998, contrasting the weakening trend beforehand (D. [[#Hu--2018|]] [[#Hu--2018|Hu et al., 2018]]). The position of the polar vortex also has long-term variations, exhibiting a persistent shift toward Northern Siberia and away from North America in February over the period 1979–2015 ([[#Zhang--2016|Zhang et al., 2016]]; J. [[#Zhang--2018|]] [[#Zhang--2018|]] [[#Zhang--2018|]] [[#Zhang--2018|Zhang et al., 2018]]). Multiple measures show similar location changes ([[#Seviour--2017|Seviour, 2017]]).

Sudden stratospheric warming (SSW), a phenomenon of rapid stratospheric air temperature increases (sometimes by more than 50°C in 1–2 days), is tightly associated with the reversal of upper stratospheric zonal winds, and a resulting collapse or substantial weakening of the stratospheric polar vortex ([[#Butler--2015|Butler et al., 2015]]; [[#Butler--2018|Butler and Gerber, 2018]]) and on average occurs approximately 6 times per decade in the NH winter ([[#Charlton--2007|Charlton et al., 2007]]; [[#Butler--2015|Butler et al., 2015]]). The SSW record from all modern reanalyses is very consistent. There is a higher occurrence of major midwinter SSWs in the 1980s and 2000s with no SSW events during 1990–1997 ([[#Reichler--2012|Reichler et al., 2012]]; [[#Butler--2015|Butler et al., 2015]]). An assessment of multi-decadal variability and change in SSW events is sensitive to both chosen metric and methods ([[#Palmeiro--2015|Palmeiro et al., 2015]]). Due to the lack of assimilation of upper air data, the centennial-scale reanalyses do not capture SSW events, even for the most recent decades ([[#Butler--2015|Butler et al., 2015]], 2017) and hence cannot inform on earlier behaviour. There has been considerably less study of trends in the SH stratosphere polar vortex strength despite the interest in the ozone hole and the potential impact of the SH stratosphere polar vortex strength on it. The occurrence of SSW events in the SH is not as frequent as in the NH, with only 3 documented events in the last 40 years ([[#Shen--2020|Shen et al., 2020]]).

In summary, it is ''likely'' that the northern lower stratospheric polar vortex has weakened since the 1980s in midwinter, and its location has shifted more frequently toward the Eurasian continent. The short record and substantial decadal variability yields ''low confidence'' in any trends in the occurrence of SSW events in the NH winter and such events in the SH are rare.

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