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

== 2.4 Changes in Modes of Variability ==

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Modes of climate variability that are important for large-scale climate on interannual and longer timescales (Cross-Chapter Box 2.2) are assessed herein and defined and summarized in Annex IV. Though the modes of variability discussed here are assessed and classified on the basis of physical variables, it is important to recognize that the distribution and function of various components of the terrestrial and marine biospheres are modified in response to them.

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=== 2.4.1 Annular Modes ===

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==== 2.4.1.1 Northern Annular Mode (NAM)/North Atlantic Oscillation (NAO) ====

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The AR5 reported that the shift towards a positive NAO, a mode of variability in the North Atlantic that is closely related to the hemispheric-scale NAM, from the 1950s to the 1990s was largely reversed by more recent changes ( ''high confidence'' ). Moreover, periods of persistent negative or positive NAO states observed during the latter part of the 20th century were not unusual, based on NAO reconstructions spanning the last half millennium ( ''high confidence'' ).

New multi-millennial-timescale NAO reconstructions are derived from marine and lake sediments, speleothems, tree rings and ice cores ( [[#Ortega--2015|Ortega et al., 2015]] ; [[#Faust--2016|Faust et al., 2016]] ; [[#Sjolte--2018|Sjolte et al., 2018]] ). NAO variability over the past 8 kyr suggests the presence of a significant 1.5 kyr periodicity ( [[#Darby--2012|Darby et al., 2012]] ). Positive NAO conditions dominated during the MH, while the prevailing NAO sign during the early and late Holocene was negative according to most reconstructions ( [[#Olsen--2012|Olsen et al., 2012]] ; [[#Ait%20Brahim--2019|Ait Brahim et al., 2019]] ; [[#Røthe--2019|Røthe et al., 2019]] ; [[#Hernández--2020|Hernández et al., 2020]] ). For the CE, reconstructions developed since AR5 indicate no dominant NAO phase during 1000–1300 CE ( [[#Jones--2014|Jones et al., 2014]] ; [[#Baker--2015|Baker et al., 2015]] ; [[#Ortega--2015|Ortega et al., 2015]] ; [[#Lasher--2019|Lasher and Axford, 2019]] ; [[#Hernández--2020|Hernández et al., 2020]] ), with either negative ( [[#Baker--2015|Baker et al., 2015]] ; [[#Faust--2016|Faust et al., 2016]] ; [[#Mellado-Cano--2019|Mellado-Cano et al., 2019]] ) or a more variable phase of the NAO ( [[#Jones--2014|Jones et al., 2014]] ; [[#Ortega--2015|Ortega et al., 2015]] ; [[#Sjolte--2018|Sjolte et al., 2018]] ; [[#Cook--2019|Cook et al., 2019]] ) between 1400 and 1850 CE. Several instrument-based NAO reconstructions extending back to the 17th and 18th centuries highlight the presence of multi-decadal variations in the NAO phases ( [[#Cornes--2013|Cornes et al., 2013]] ; [[#Cropper--2015|Cropper et al., 2015]] ), although these studies have limitations considering the seasonality of the centres of action and the locations of the stations used. Recent reconstructions of the large scale sea level pressure field yield more robust NAO analysis, showing a persistently negative NAO phase from the 1820s to the 1870s, with positive values dominating during the beginning of the 20th century followed by a declining trend over 1920–1970, with a recovery thereafter to a period of consistently high values between 1970 and the early 1990s ( [[#Delaygue--2019|Delaygue et al., 2019]] ; [[#Mellado-Cano--2019|Mellado-Cano et al., 2019]] ). Based on the evaluation of several NAO reconstructions for recent centuries, [[#Hernández--2020|Hernández et al. (2020)]] highlighted that the strong positive NAO phases of the 1990s and early 21st century were not unusual.

The predominantly positive phase during the 1990s was followed by partial reversal and a tendency towards stronger variability in boreal winter NAM and NAO since the late 1990s ( [[#Pinto--2012|Pinto and Raible, 2012]] ; [[#Hanna--2018|Hanna et al., 2018]] ). This is particularly evident in December NAO ( [[#Hanna--2015|Hanna et al., 2015]] ) and NAM ( [[#Overland--2015|Overland and Wang, 2015]] ) indices, and is not unusual on multi-decadal time scales ( [[#Woollings--2018a|Woollings et al., 2018a]] ). Since the 1990s, a statistically significant summer NAO decline was reported, which is, to a lesser extent, also evident in the winter NAO, linked to an enhanced blocking activity over Greenland ( [[#Hanna--2015|Hanna et al., 2015]] , 2016; [[#Wachowicz--2021|Wachowicz et al., 2021]] ). However, this was moderated by persistent positive NAO values since 2015 (Annex IV.2.1). Based on observations and reanalysis datasets, multi-decadal variations were found for the NAM patterns: the Atlantic centre remained unchanged throughout 1920–2010, whereas the Pacific centre was stronger during 1920–1959 and 1986–2010 and weaker during 1960–1985 ( [[#Gong--2018|Gong et al., 2018]] ). Multidecadal changes were also observed in the position of the centres of action of the NAO ( [[#Moore--2013|Moore et al., 2013]] ; [[#Zuo--2016|Zuo et al., 2016]] ).

In summary, positive trends for the NAM/NAO winter indices were observed between the 1960s and the early 1990s, but these indices have become less positive or even negative thereafter ( ''high confidence'' ). The NAO variability in the instrumental record was ''very likely'' not unusual in the millennial and multi-centennial context.

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==== 2.4.1.2 Southern Annular Mode (SAM) ====

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The AR5 concluded that it was ''likely'' that the SAM had become more positive since the 1950s and that this increase was unusual in the context of the prior 400 years ( ''medium confidence'' ). Both AR5 and SROCC reported statistically significant trends in the SAM during the instrumental period for the austral summer and autumn.

Several studies have attempted to reconstruct the evolution of the SAM during the Holocene using proxies of the position and strength of the SH zonal winds, although with no clear consensus regarding the timing and phase of the SAM ( [[#Hernández--2020|Hernández et al., 2020]] ). The early Holocene was dominated by SAM positive phases ( [[#Moreno--2018|Moreno et al., 2018]] ; [[#Reynhout--2019|Reynhout et al., 2019]] ), consistent with increasing westerly wind strength ( [[#Lamy--2010|Lamy et al., 2010]] ), with some reconstructions showing significant centennial and millennial variability but no consistent trend after 5 ka ( [[#Hernández--2020|Hernández et al., 2020]] ). For the CE, enhanced westerly winds occurred over 0–1000 CE, as reflected in increased burning activity in Patagonia ( [[#Turney--2016a|Turney et al., 2016a]] ) and tree ring records from southern New Zealand ( [[#Turney--2016b|Turney et al., 2016b]] ) imply a predominantly positive SAM phase. Pollen records and lake sediments from Tasmania, southern mainland Australia, New Zealand and southern South America, inferred the period of 1000 to 1400 CE to be characterized by anomalously dry conditions south of 40°S, implying a positive SAM ( [[#Moreno--2014|Moreno et al., 2014]] ; [[#Fletcher--2018|Fletcher et al., 2018]] ; [[#Evans--2019|Evans et al., 2019]] ; [[#Matley--2020|Matley et al., 2020]] ). Nevertheless, proxy reconstructions of the SAM based on temperature-sensitive records from tree rings, ice cores, lake sediments and corals spanning the mid-to-polar latitudes show alternating positive and negative phases (Figure 2.35).

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[[File:328f0580e5a1117c67759075c3d6a122 IPCC_AR6_WGI_Figure_2_35.png]]

'''Figure 2.35''' '''|''' '''Southern Annular Mode (SAM) reconstruction over the last millennium. (a)''' SAM reconstructions as seven-year moving averages (thin lines) and 70-year LOESS filter (thick lines). '''(b)''' Observed SAM index during 1900–2019. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Prolonged periods of negative SAM values were identified during the period 1400–1700 CE in several reconstructions (Figure 2.35, [[#Villalba--2012|Villalba et al., 2012]] ; [[#Abram--2014|Abram et al., 2014]] ; [[#Dätwyler--2018|Dätwyler et al., 2018]] ), with a minimum identified during the 15th century ( [[#Hernández--2020|Hernández et al., 2020]] ), although some disagreements exist between proxy records before 1800 CE ( [[#Hessl--2017|Hessl et al., 2017]] ). [[#Abram--2014|Abram et al. (2014)]] concluded that the mean SAM index during recent decades is at its highest levels for at least the last 1 kyr. Similarly, the summer SAM reconstruction by [[#Dätwyler--2018|Dätwyler et al. (2018)]] indicates a strengthening over the last 60 years that is outside the ''very likely'' range of the last millennium natural variability. The largest 30- and 50-year trends in the annual SAM index occurred at the end of the 20th century (after 1969 and 1950 respectively), indicating that the recent increase in the SAM is unprecedented in the context of at least the past three centuries ( [[#Yang--2018|Yang and Xiao, 2018]] ).

Before the mid-1950s, SAM indices derived from station-based datasets, and centennial reanalyses show pronounced interannual and decadal variability but no significant trends, with low correlation between SAM indices due to the diversity across different datasets and sensitivity to the definition used for the index calculation ( [[#Barrucand--2018|Barrucand et al., 2018]] ; [[#Schneider--2018|Schneider and Fogt, 2018]] ; J. [[#Lee--2019|]] [[#Lee--2019|Lee et al., 2019]] ). Various SAM indices exhibit significant positive trends since the 1950s, particularly during austral summer and autumn ( [[#Barrucand--2018|Barrucand et al., 2018]] ; [[#Schneider--2018|Schneider and Fogt, 2018]] ; J. [[#Lee--2019|]] [[#Lee--2019|Lee et al., 2019]] ), unprecedented for austral summer over the last 150 years (J.M. [[#Jones--2016|]] [[#Jones--2016|Jones et al., 2016]] ; [[#Fogt--2020|Fogt and Marshall, 2020]] ). This indicates a strengthening of the surface westerly winds around Antarctica, related to both the position and intensity of the subpolar jet in the SH ( [[#2.3.1.4.3|Section 2.3.1.4.3]] ; [[#Ivy--2017|Ivy et al., 2017]] ; [[#IPCC--2019|IPCC, 2019]] ). The SAM trends have slightly weakened after about 2000 ( [[#Fogt--2020|Fogt and Marshall, 2020]] ).

In summary, historical station-based reconstructions of the SAM show that there has been a robust positive trend in the SAM index, particularly since 1950 and for the austral summer ( ''high confidence'' ). The recent positive trend in the SAM is ''likely'' unprecedented in at least the past millennium, although ''medium confidence'' arises due to the differences between proxy records before 1800 CE.

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=== 2.4.2 El Niño–Southern Oscillation (ENSO) ===

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The AR5 reported with ''medium confidence'' that ENSO-like variability existed, at least sporadically, during the warm background state of the Pliocene. It was also found ( ''high confidence'' ) that ENSO has remained highly variable during the last 7 kyr with no discernible orbital modulation. The AR5 concluded that large variability on interannual to decadal timescales, and differences between datasets, precluded robust conclusions on any changes in ENSO during the instrumental period. The SROCC reported epochs of strong ENSO variability throughout the Holocene, with no indications of a systematic trend in ENSO amplitude, but with some indication that the ENSO amplitude over 1979–2009 was greater than at any point in the period from 1590–1880 CE. It was also reported that the frequency and intensity of El Niño events in the period from 1951–2000 was high relative to 1901–1950.

[[#Manucharyan--2014|Manucharyan and Fedorov (2014)]] found that ENSO-like variability has been present, at least sporadically, during epochs of millions of years (including the MPWP; Cross-Chapter Box 2.4), with proxy records indicating that this was the case even when cross-Pacific SST gradients were much weaker than present. There is substantial disagreement between proxy records for ENSO activity during the early Holocene ( [[#Zhang--2014|Zhang et al., 2014]] ; [[#White--2018|White et al., 2018]] ), and for ENSO activity and mean state at the LGM ( [[#Leduc--2009|Leduc et al., 2009]] ; [[#Koutavas--2012|Koutavas and Joanides, 2012]] ; [[#Sadekov--2013|Sadekov et al., 2013]] ; [[#Ford--2015|Ford et al., 2015]] , 2018; [[#Zhu--2017|Zhu et al., 2017]] ; [[#Tierney--2020|Tierney et al., 2020]] ). A number of studies ( [[#Cobb--2013|Cobb et al., 2013]] ; H. [[#McGregor--2013|]] [[#McGregor--2013|McGregor et al., 2013]] ; [[#Carré--2014|Carré et al., 2014]] ; [[#Emile-Geay--2016|Emile-Geay et al., 2016]] ; [[#Thompson--2017|Thompson et al., 2017]] ; [[#Tian--2017|Tian et al., 2017]] ; [[#White--2018|White et al., 2018]] ; [[#Grothe--2019|Grothe et al., 2019]] ) have found that ENSO was substantially weaker than at present at various times in the mid-Holocene within the period from 6 to 3 ka, with stronger decreases in variability revealed by remote proxies than by those close to the core region of ENSO activity. However, [[#Karamperidou--2015|Karamperidou et al. (2015)]] find that weakening in ENSO-related variability in eastern Pacific proxies does not necessarily correspond to weakening in central Pacific proxies. [[#Barrett--2018|Barrett et al. (2018)]] concluded that multi-proxy reconstructions are more efficient at identifying eastern Pacific than central Pacific events. This suggests that a weakening of proxy-based signals may indicate an along-equatorial shift in ENSO activity rather than a weakening of ENSO during some periods. Following the period of weak ENSO variability in the mid-Holocene, a number of studies find an increase in ENSO activity which, depending upon the study, commences between 4.4 and 3 ka ( [[#Koutavas--2012|Koutavas and Joanides, 2012]] ; [[#Cobb--2013|Cobb et al., 2013]] ; [[#Zhang--2014|Zhang et al., 2014]] ; S. [[#Chen--2016|]] [[#Chen--2016|Chen et al., 2016]] ; [[#Emile-Geay--2016|Emile-Geay et al., 2016]] ; [[#Thompson--2017|Thompson et al., 2017]] ; [[#Du--2021|Du et al., 2021]] ).

Numerous studies (J. [[#Li--2013|]] [[#Li--2013|Li et al., 2013]] ; S. [[#McGregor--2013|]] [[#McGregor--2013|McGregor et al., 2013]] ; [[#Rustic--2015|Rustic et al., 2015]] ; [[#Hope--2017|Hope et al., 2017]] ; Y. [[#Liu--2017|]] [[#Liu--2017|]] [[#Liu--2017|Liu et al., 2017]] ) find substantial variability in ENSO activity on multi-decadal to centennial timescales over the last 500 to 1 kyr (Figure 2.36). Different proxies show a wide spread in the specific timing and magnitude of events in the pre-instrumental period (e.g., [[#Dätwyler--2019|Dätwyler et al., 2019]] ). Most investigators find that ENSO activity in recent decades was higher than the most recent centuries prior to the instrumental period. [[#Grothe--2019|Grothe et al. (2019)]] also found that ENSO variance of the last 50 years was 25% higher than the average of the last millennium, and was substantially higher than the average of the mid- to late-Holocene. S. McGregor et al., (2010, 2013) looked for common variance changes in pre-existing ENSO proxies, finding stronger ENSO variance for the 30-year period 1979–2009 compared to any 30-year period within the timespan 1590–1880 CE. This finding also holds when adding more recently developed ENSO proxies (Figure 2.36). [[#Koutavas--2012|Koutavas and Joanides (2012)]] , [[#Ledru--2013|Ledru et al. (2013)]] and [[#Thompson--2017|Thompson et al. (2017)]] identify various periods within the range 1000 BCE to 1300 CE when ENSO activity was greater than in the following centuries, and more closely comparable to the mid-20th century onwards behaviour.

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[[File:bb143b9338be6a32e60692bdf649080d IPCC_AR6_WGI_Figure_2_36.png]]

'''Figure 2.3''' '''6 |''' '''Reconstructed and historical variance ratio of El Niño–Southern Oscillation (ENSO). (a)''' 30-year running variance of the reconstructed annual mean Niño 3.4 or related indicators from various published reconstructions. '''(b)''' Variance of June–November Southern Oscillation Index (SOI) and April–March mean Niño 3.4 (1981–2010 base period) along with the mean reconstruction from (a). Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Since AR5, updates to datasets used widely in prior ENSO assessments resulted in substantial and important revisions to observed tropical Pacific SST data ( [[#2.3.1.1|Section 2.3.1.1]] ). In particular, ERSSTv4, and then ERSSTv5, addressed known SST biases in ERSSTv3 in the equatorial Pacific which affected the derived mean state and amplitude of indices based on that dataset ( [[#Huang--2015|Huang et al., 2015]] ). During the instrumental period, there is no robust indication of any significant century-scale trend in the east-west SST gradient across the equatorial Pacific Ocean, with periods when gradients have been stronger and weaker than the long-term average on decadal timescales, associated with a predominance of La Niña or El Niño events respectively. The frequency of El Niño and La Niña events is also subject to considerable decadal variability (e.g., [[#Hu--2013|Hu et al., 2013]] ) but with no indication of a long-term signal in the frequency of events. The ENSO amplitude since 1950 has increased relative to the 1910–1950 period, as confirmed by independent proxy records (e.g., [[#Gergis--2009|Gergis and Fowler, 2009]] ), the Southern Oscillation Index (SOI) ( [[#Braganza--2009|Braganza et al., 2009]] ) and SSTs (e.g., [[#Ohba--2013|Ohba, 2013]] ; [[#Yu--2013|Yu and Kim, 2013]] ), although there is a spread between different proxy and instrumental sources as to the magnitude of that increase (Figure 2.36). The El Niño events of 1982–1983, 1997–1998 and 2015–2016 had the strongest anomalies in the Niño 3.4 SST index since 1950. Their predominance was less evident from indices based on de-trended data such as the Oceanic Niño Index (ONI) (which still ranked them as the three strongest events since 1950, but only by a small margin), and in the SOI. [[#Huang--2019a|]] [[#Huang--2019a|B. Huang et al. (2019a)]] also note that analyses based upon buoy and Argo data, which are only available since the 1990s, are more capable of resolving strong events than analyses which do not include such data.

Prior to the 1950s, SST observations in the tropical Pacific were much sparser and hence uncertainties in Niño indices are much larger ( [[#Huang--2020|]] [[#Huang--2020|B. Huang et al., 2020]] ). SOI data and some newer SST-based studies show high ENSO amplitude, comparable to the post-1950 period, in the period from the mid-late 19th century to about 1910, but proxy indicators generally indicate that the late 19th and early 20th century were less active than the late 20th century (Figure 2.36). Yu and Kim’s (2013) implementation of the ONI found a number of events with the ONI above 1.5°C between 1888 and 1905, then no such events until 1972, whilst the SOI indicates comparable or stronger events to the three strongest post-1950 events in 1896 and 1905. [[#Giese--2011|Giese and Ray (2011)]] also found a number of such events between 1890 and 1920 in the SODA ocean reanalysis, corroborated further by [[#Huang--2020|]] [[#Huang--2020|B. Huang et al. (2020)]] and [[#Vaccaro--2021|Vaccaro et al. (2021)]] , who found that the strength of the 1877–1878 event was comparable with that of the 1982–1983, 1997–1998 and 2015–2016 events. There have also been a number of strong La Niña events (e.g., 1973–1974, 1975–1976 and 2010–2011), with few clear analogues in the 1920–1970 period; the proxy-based analysis of [[#McGregor--2010|McGregor et al. (2010)]] indicates that the mid-1970s La Niña period was also extreme in a multi-centennial context. There is no indication that the frequency of high-amplitude events since the 1970s reflects a long-term trend which can be separated from multi-decadal variability, given apparent presence of several high-amplitude events in the late 19th and early 20th centuries, and the relatively large uncertainty in pre-1950 SST data in the tropical Pacific region.

There is a distinction (Annex IV.2.3.1) between El Niño events centred in the eastern Pacific (‘Eastern Pacific’ (EP) or ‘classical’ events) and those centred in the Central Pacific (‘Central Pacific’ (CP) or ‘Modoki’ events), which have different typical teleconnections (e.g., [[#Ashok--2007|Ashok et al., 2007]] ; [[#Ratnam--2014|Ratnam et al., 2014]] ; [[#Capotondi--2015|Capotondi et al., 2015]] ; [[#Timmermann--2018|Timmermann et al., 2018]] ). A number of studies, using a range of indicators, have found an increase in recent decades of the fraction of CP El Niño events, particularly after 2000 ( [[#Yu--2013|Yu and Kim, 2013]] ; [[#Lübbecke--2014|Lübbecke and McPhaden, 2014]] ; [[#Pascolini-Campbell--2015|Pascolini-Campbell et al., 2015]] ; [[#Jiang--2018|Jiang and Zhu, 2018]] ). [[#Johnson--2013|Johnson (2013)]] found that the frequency of CP El Niño events had increased (although not significantly) over the 1950–2011 period, being accompanied by a significant increase in the frequency of La Niña events with a warm (as opposed to cool) western Pacific warm pool. A coral-based reconstruction starting in 1600 CE ( [[#Freund--2019|Freund et al., 2019]] ) found that the ratio of CP to EP events in the last 30 years was substantially higher than at any other time over the last 400 years. Variations in the proportion of CP and EP events have also been found in earlier periods, with [[#Carré--2014|Carré et al. (2014)]] finding a period of high CP activity around 7 ka.

There is no robust indication of any changes in ENSO teleconnections over multi-centennial timescales ( [[#Hernández--2020|Hernández et al., 2020]] ) despite multi-decadal variability. [[#Shi--2018|Shi and Wang (2018)]] found that teleconnections with the broader Asian summer monsoon, including the Indian and the East Asian monsoon, were generally stable since the 17th century during the developing phase of the monsoon, and showed substantial decadal variability, but no clear trend, during the decaying phase. They also found that the weakening of teleconnections between the Indian summer monsoon and ENSO in recent decades had numerous precedents over the last few centuries. [[#Räsänen--2016|Räsänen et al. (2016)]] also found substantial decadal variability, but little trend, in the strength of the relationship between ENSO and monsoon precipitation in South East Asia between 1650 and 2000. [[#Dätwyler--2019|Dätwyler et al. (2019)]] identified a number of multi-decadal periods with apparently changed teleconnections at times over the last 400 years.

In the instrumental period, teleconnections associated with ENSO are well known to vary on decadal to multi-decadal timescales (e.g., [[#He--2013|He et al., 2013]] ; [[#Lee--2015|Lee and Ha, 2015]] ; [[#Ashcroft--2016|Ashcroft et al., 2016]] ; [[#Jin--2016|Jin et al., 2016]] ; [[#Wang--2019|Wang et al., 2019]] ). [[#Yun--2018|Yun and Timmermann (2018)]] found that decadal variations in teleconnections between ENSO and the Indian monsoon did not extend beyond what would be expected from a stochastic process. Many observed decadal changes in teleconnections in the instrumental period are consistent with a shift to more central Pacific El Niño events ( [[#Evtushevsky--2018|Evtushevsky et al., 2018]] ; [[#Yeh--2018|Yeh et al., 2018]] ; [[#Yu--2018|Yu and Sun, 2018]] ; [[#Zhao--2019|Zhao and Wang, 2019]] ). Effects of the PDV ( [[#Kwon--2013|Kwon et al., 2013]] ; S. [[#Wang--2014|]] [[#Wang--2014|Wang et al., 2014]] ; [[#Dong--2018|Dong et al., 2018]] ) and the AMV ( [[#Kayano--2019|Kayano et al., 2019]] ) can also modulate ENSO teleconnections, and affect the frequency of CP versus EP events ( [[#Ashok--2007|Ashok et al., 2007]] ). [[#Chiodi--2015|Chiodi and Harrison (2015)]] found that teleconnections over the most recent decades are broadly consistent with those over the last 100 years. Variability in teleconnections can also occur on timescales longer than characteristic PDV timescales (e.g., [[#Gallant--2013|Gallant et al., 2013]] ).

In summary, there is ''medium confidence'' that both ENSO amplitude and the frequency of high-magnitude events since 1950 are higher than over the period from 1850 and possibly as far back as 1400, but ''low confidence'' that they are outside the range of variability over periods prior to 1400, or higher than the average of the Holocene as a whole. Overall, there is no indication of a recent sustained shift in ENSO or associated features such as the Walker Circulation, or in teleconnections associated with these, being beyond the range of variability on decadal to millennial timescales. A high proportion of El Niño events in the last 20–30 years have been based in the central, rather than eastern Pacific, but there is ''low confidence'' that this represents a long-term change.

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=== 2.4.3 Indian Ocean Basin and Dipole Modes ===

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The AR5 did not provide an assessment of the Indian Ocean Dipole (IOD) records based on paleo reconstructions. For the instrumental era, AR5 reported that there were no trends in the IOD behaviour. However, the strength of the Indian Ocean Basin-wide (IOB) mode, quantified by the basin-scale averaged SST index, increased in all assessed periods except 1979–2012, but the AR5 neither quantified trends nor provided a confidence statement.

For the LGM, enhanced equatorial Indian Ocean productivity in marine sediment records was associated with strengthened westerly jets, in line with a shallower central-western Indian Ocean thermocline and stronger negative IOD events ( [[#Punyu--2014|Punyu et al., 2014]] ). The LGM Indian Ocean basin was substantially modified by the exposure of the tropical shelves ( [[#DiNezio--2018|DiNezio et al., 2018]] ) and this has been associated with an Indian Ocean ‘El Niño‘ ( [[#Thirumalai--2019|Thirumalai et al., 2019]] ). [[#Wurtzel--2018|Wurtzel et al. (2018)]] contend that during the LDT, including the Younger Dryas event, Indian Ocean precipitation did not mirror the zonal asymmetry observed in Indian Ocean SSTs in the Holocene based upon a speleothem record from Sumatra, but instead reflected shifts in moisture transport pathways and sources. Using Seychelles corals, representing the western pole of the IOD, spanning from the MH to present, [[#Zinke--2014|Zinke et al. (2014)]] identified changes in seasonality, with the lowest seasonal SST range in the MH and then again around 2 ka, while the largest seasonal range occurred around 4.6 ka and then again during the near-present (1990–2003). Reconstructions from fossil corals for the eastern Indian Ocean point to stronger negative IOD SST anomalies due to the enhanced upwelling and cooling driven by a stronger monsoon with enhanced anomalous easterly winds in the eastern Indian Ocean during the MH ( [[#Abram--2020b|Abram et al., 2020b]] ). [[#Niedermeyer--2014|Niedermeyer et al. (2014)]] from the analysis of stable isotopes in terrestrial plant waxes, suggest that the period 6.5 ka to 4.5 ka was characterized by an anomalously positive IOD mean state. During various parts of the Holocene, periods of a mean positive IOD-like state were associated with increased IOD variability, including events that exceed the magnitude of the strongest events during the instrumental period ( [[#Abram--2020a|Abram et al., 2020a]] ).

From the coral δ <sup>18</sup> O record from the Seychelles over 1854–1994, [[#Du--2014|Du et al. (2014)]] showed a 3–7 year dominant period associated with the IOB in response to ENSO forcing. They identified multi-decadal variability in the IOB, with more active IOB phases during 1870–1890, 1930–1955, and 1975–1992, while decadal variability in the IOB dominated during 1940–1975 ( [[#Du--2014|Du et al., 2014]] ). ''Evidence'' for changes in IOB characteristics during earlier periods (e.g., MH, LGM) is ''limited'' .

The role of decadal to multi-decadal variability has recently emerged as an important aspect of the IOD with many indications of the effects of Pacific Ocean processes on IOD variability through atmospheric and oceanic mechanisms ( [[#Dong--2016|Dong et al., 2016]] ; [[#Krishnamurthy--2016|Krishnamurthy and Krishnamurthy, 2016]] ; [[#Zhou--2017|Zhou et al., 2017]] ; [[#Jin--2018|Jin et al., 2018]] ). Positive events in the 1960s and 1990s were linked to a relatively shallow eastern Indian Ocean thermocline, and a primarily negative IOD state in the 1970s and 1980s was related to a deeper thermocline ( [[#Ummenhofer--2016|Ummenhofer et al., 2016]] ). Positive IOD events may have increased in frequency during the second half of the 20th century ( [[#Abram--2020a|Abram et al., 2020a]] , b). Earlier observations of apparent changes in the frequency and/or magnitude of the IOD events are considered unreliable, particularly prior to the 1960s ( [[#Hernández--2020|Hernández et al., 2020]] ). Although the seasonal evolution and the type of ENSO ( [[#2.4.2|Section 2.4.2]] ) may influence the character of the IOD ( [[#Guo--2015|Guo et al., 2015]] ; [[#Zhang--2015|Zhang et al., 2015]] ; [[#Fan--2016|Fan et al., 2016]] ), the occurrence of some IOD events may be independent of ENSO ( [[#Sun--2015|Sun et al., 2015]] ).

To summarize, there is ''low confidence'' in any multi-decadal IOD variability trend in the instrumental period due to data uncertainties especially before the 1960s. In addition to data uncertainties, understanding of the IOB variability during the instrumental period is also limited by large-scale warming of the Indian Ocean. Neither the IOD nor the IOB have exhibited behaviour outside the range implied by proxy records ( ''low confidence'' ).

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=== 2.4.4 Atlantic Meridional (AMM) and Zonal Modes (AZM) ===

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The AR5 reported no changes in the Atlantic Meridional Mode (AMM) during the 20th century or shorter periods thereof. For the Atlantic Zonal Mode (AZM), also referred as the Atlantic Niño, the AR5 reported increases during the 1950–2012 period but neither assessed trends nor provided a confidence statement. The AR5 did not assess paleo evidence for the AZM and AMM.

Paleo-reconstructions of these two modes remain rather limited. Nonetheless, the interhemispheric cross-equatorial SST gradients linked to changes in ITCZ locations characteristic of the AMM has been found during the LGM, Heinrich Stadial 1, and the MH, with the largest shift occurring during HS1 ( [[#McGee--2014|McGee et al., 2014]] ). Similarly, the dipole-like SST pattern in the South Atlantic subtropics, which is related to the AZM ( [[#Venegas--1996|Venegas et al., 1996]] ; [[#Morioka--2011|Morioka et al., 2011]] ; [[#Nnamchi--2016|Nnamchi et al., 2016]] , 2017; [[#Lübbecke--2018|Lübbecke et al., 2018]] ; [[#Rouault--2018|Rouault et al., 2018]] ; [[#Foltz--2019|Foltz et al., 2019]] ), has been reconstructed using SST proxies from marine sediment cores during the past 12 kyr ( [[#Wainer--2014|Wainer et al., 2014]] ). The reconstructed index captures two significant cold events that occurred during the 12.9–11.6 ka and 8.6–8.0 ka periods in the South Atlantic ( [[#Wainer--2014|Wainer et al., 2014]] ).

During the observational period the AZM and AMM (Figure 2.37) are related to the AMV largely controlling the interhemispheric gradient of the SST at decadal to multi-decadal timescales ( [[#Tokinaga--2011|Tokinaga and Xie, 2011]] ; [[#Polo--2013|Polo et al., 2013]] ; [[#Svendsen--2014a|Svendsen et al., 2014a]] ; [[#Li--2015|Li et al., 2015]] ; [[#Lübbecke--2018|Lübbecke et al., 2018]] ). The AZM interannual variability is enhanced ( [[#Lübbecke--2018|Lübbecke et al., 2018]] ; [[#Foltz--2019|Foltz et al., 2019]] ), and is more strongly related to ENSO, during the negative phase of the AMV ( [[#Martín-Rey--2014|Martín-Rey et al., 2014]] , 2018; [[#Polo--2015|Polo et al., 2015]] ; [[#Nnamchi--2020|Nnamchi et al., 2020]] ). The AZM displayed a persistent weakened variability over the 1960–2009 period associated with declined cold tongue upwelling ( [[#Tokinaga--2011|Tokinaga and Xie, 2011]] ), which became pronounced since 2000 ( [[#Prigent--2020a|Prigent et al., 2020a]] , b). Despite these multi-decadal fluctuations, there is ''limited evidence'' for any sustained change in the AMM ( [[#Chang--2011|Chang et al., 2011]] ; [[#Martín-Rey--2018|Martín-Rey et al., 2018]] ) and AZM ( [[#Martín-Rey--2018|Martín-Rey et al., 2018]] ; [[#Nnamchi--2020|Nnamchi et al., 2020]] ) during the instrumental period. The AZM and AMM interact on interannual timescales ( [[#Servain--1999|Servain et al., 1999]] ; [[#Foltz--2010|Foltz and McPhaden, 2010]] ; [[#Pottapinjara--2019|Pottapinjara et al., 2019]] ) leading in 2009 to extremes of both modes in which the negative phase of the AMM ( [[#Foltz--2012|Foltz et al., 2012]] ; [[#Burmeister--2016|Burmeister et al., 2016]] ) preceded an equatorial cold tongue cold event that was unprecedented in the prior 30 years ( [[#Foltz--2010|Foltz and McPhaden, 2010]] ; [[#Burmeister--2016|Burmeister et al., 2016]] ).

In summary, ''confidence'' is ''low'' in any sustained changes to the AZM and AMM variability in instrumental observations. There is ''very'' ''low confidence'' in changes of the paleo AZM and AMM due to extremely limited availability of paleo reconstructions.

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<span id="pacific-decadal-variability-pdv"></span>
=== 2.4.5 Pacific Decadal Variability (PDV) ===

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Pacific Decadal Variability (PDV) refers to the ocean-atmosphere climate variability over the Pacific Ocean at decadal-to-interdecadal time scales and is usually described by the Pacific Decadal Oscillation (PDO) or the Inter-decadal Pacific Oscillation (IPO) indices. The AR5 and SROCC reported a large shift of the PDO in the late 1970s, with a predominantly positive phase until the end of the 1990s, being mainly negative afterwards. There was no significant change assessed in the PDO during the instrumental period as a whole, and no confidence level was assigned. Changes in the pre-instrumental era PDO were not assessed in AR5.

The existence of the PDV in the centuries prior to the instrumental period is evidenced by a variety of proxy records based on tree rings ( [[#Biondi--2001|Biondi et al., 2001]] ; [[#D’Arrigo--2015|D’Arrigo and Ummenhofer, 2015]] ), corals ( [[#Felis--2010|Felis et al., 2010]] ; [[#Deng--2013|Deng et al., 2013]] ; [[#Linsley--2015|Linsley et al., 2015]] ) and sediments ( [[#Lapointe--2017|Lapointe et al., 2017]] ; [[#O’Mara--2019|O’Mara et al., 2019]] ). There is little coherence between the various paleo-proxy indices prior to the instrumental record, and neither these nor the instrumental records provide indications of a clearly defined spectral peak ( [[#Chen--2015|Chen and Wallace, 2015]] ; M. [[#Newman--2016|]] [[#Newman--2016|Newman et al., 2016]] ; [[#Henley--2017|Henley, 2017]] ; L. [[#Zhang--2018|]] [[#Zhang--2018|]] [[#Zhang--2018|]] [[#Zhang--2018|Zhang et al., 2018]] ; [[#Buckley--2019|Buckley et al., 2019]] ). For instance, spectral analysis from millennia length PDV reconstructions shows spectral peaks at multi-decadal, centennial and bi-centennial time scales ( [[#Beaufort--2017|Beaufort and Grelaud, 2017]] ), while only multi-decadal oscillations can be detected in the shorter (less than 400 years into the past) paleoclimate reconstructions. A variety of proxies suggest a shift in the PDV from the early-mid Holocene, which was characterized by a persistently negative phase of the PDO (i.e., weak Aleutian Low), to the late Holocene, and more variable and more positive PDO (i.e., strong Aleutian Low) conditions. This shift at around 4.5 ka is also evident in the PDO periodicities, changing from bidecadal and pentadecadal variability in the early Holocene to only pentadecadal periodicities in the late Holocene ( [[#Hernández--2020|Hernández et al., 2020]] ). Several proxy records indicate that the strengthening in the Aleutian Low inferred since the late 17th century is unprecedented over the last millennium (Z. [[#Liu--2017|]] [[#Liu--2017|]] [[#Liu--2017|Liu et al., 2017]] ; [[#Osterberg--2017|Osterberg et al., 2017]] ; [[#Winski--2017|Winski et al., 2017]] ), in line with an increase in PDV low-frequency variability ( [[#Williams--2017|Williams et al., 2017]] ; [[#Hernández--2020|Hernández et al., 2020]] ).

The PDO and IPO indices are significantly correlated during the instrumental period, showing regime shifts in the 1920s, 1940s, 1970s and around 1999. Positive PDV phases were observed from the 1920s to the mid-1940s and from the late 1970s to the late 1990s, while negative phases occurred from mid-1940s until the late 1970s, and since 1999 (Figure 2.38; [[#Han--2014|Han et al., 2014]] ; [[#Chen--2015|Chen and Wallace, 2015]] ; M. [[#Newman--2016|]] [[#Newman--2016|Newman et al., 2016]] ). The associated spatial patterns are quite similar, but the PDO pattern exhibits stronger SST anomalies in the extra-tropical North Pacific than the IPO ( [[#Chen--2015|Chen and Wallace, 2015]] ). The strength and structure of the SST patterns also differ among the periods (M. [[#Newman--2016|]] [[#Newman--2016|Newman et al., 2016]] ).

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[[File:76139e2272e794efb7b34b688784344c IPCC_AR6_WGI_Figure_2_38.png]]

'''Figure 2.38''' '''|''' '''Indices of multi-decadal climate variability from 1854–2019 based upon several sea surface temperature data products.''' Shown are the indices of the AMV and PDV based on area averages for the regions indicated in Annex IV. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Instrumental observations are sparse prior to 1950, and thus the fidelity of any PDV index derived for the second part of the 19th century and early decades of the 20th century is relatively low (Figure 2.38; [[#Deng--2013|Deng et al., 2013]] ; [[#Wen--2014|Wen et al., 2014]] ). This results in ''low agreement'' in the classification of the PDO/IPO phase among several indices, even during recent years with the availability of high-quality data. Nevertheless, the teleconnection patterns are robust regardless of the index used to characterize the PDO ( [[#McAfee--2017|McAfee, 2017]] ). Analysis of time series of PDO and IPO highlights the (multi-) decadal nature of this mode of variability with no significant trends, but highlights a recent switch from a positive to a negative phase since 1999/2000 across all indicators ( [[#England--2014|England et al., 2014]] ; [[#Henley--2017|Henley, 2017]] ).

In summary, the PDV in the instrumental record is dominated by (multi-)decadal-scale shifts between positive and negative phases over the last 150 years with no overall trend ( ''high confidence'' ). There is ''low confidence'' in paleo-PDV reconstructions due to discrepancies among the various available time series in terms of phasing and timing. However, there is ''high confidence'' in the occurrence of a shift from predominantly negative to positive PDO conditions from the middle to the late Holocene.

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<span id="atlantic-multi-decadal-variability"></span>
=== 2.4.6 Atlantic Multi-decadal Variability ===

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The AR5 reported no robust changes in Atlantic Multi-decadal Variability (AMV) reconstructions based on paleo records due to low consistency between different AMV reconstructions prior to 1900. AR5 concluded that there have been no significant trends in the AMV index during the instrumental period and there was difficulty in interpreting the AMV signal because of the long-term underlying SST warming trend.

The AR5 conclusions about large uncertainties in AMV paleo reconstructions ( [[#Hernández--2020|Hernández et al., 2020]] ) have been reinforced by recent studies of tree rings (J. [[#Wang--2017b|]] [[#Wang--2017|Wang et al., 2017]] b ), Greenland ice ( [[#Chylek--2012|Chylek et al., 2012]] ), and corals ( [[#Kilbourne--2014|Kilbourne et al., 2014]] ; [[#Svendsen--2014b|Svendsen et al., 2014b]] ; J. [[#Wang--2017b|]] [[#Wang--2017|Wang et al., 2017]] b ). The AMV exhibited a generally positive state over the first millennium of the CE ( [[#Mann--2009|Mann et al., 2009]] ; [[#Singh--2018|Singh et al., 2018]] ). Paleo reconstructions over the last millennium consistently show a negative AMV phase during 1400–1850 CE and a positive phase during 900–1200 CE ( [[#Mann--2009|Mann et al., 2009]] ; J. [[#Wang--2017b|]] [[#Wang--2017|Wang et al., 2017]] b ; [[#Singh--2018|Singh et al., 2018]] ), consistent with warmer surface temperatures from tropical Atlantic records ( [[#Kilbourne--2014|Kilbourne et al., 2014]] ).

Instrumental observations show that AMV is characterized by basin-wide warm and cool periods with an average variation in SST of about 0.4°C, but with larger variations in the North Atlantic subpolar gyre. Despite small differences in indices used to define the AMV (Annex IV), they all show warm periods occurring approximately between 1880–1900, 1940–1960, and from the mid-1990s to present, with cool periods in between (Figure 2.38) but no overall sustained change during the instrumental period ( [[#Booth--2012|Booth et al., 2012]] ; [[#Gulev--2013|Gulev et al., 2013]] ; [[#Bellomo--2018|Bellomo et al., 2018]] ).

The oceanic changes are seen in salinity and temperature variations over the upper 3000 m of the North Atlantic ( [[#Polyakov--2005|Polyakov et al., 2005]] ; [[#Keenlyside--2015|Keenlyside et al., 2015]] ), and in sea level variations in the western North Atlantic along the Gulf Stream passage ( [[#McCarthy--2015|McCarthy et al., 2015]] ). The pattern and strength of the AMV differs among the periods (e.g., [[#Svendsen--2014b|Svendsen et al., 2014b]] ; [[#Reynolds--2018|Reynolds et al., 2018]] ) and there are indications that there may have been a shift since 2005 toward a negative phase of the AMV ( [[#Robson--2016|Robson et al., 2016]] ).

In summary, no sustained change in AMV indices has been observed over the instrumental period ( ''high confidence'' ). However, instrumental records may not be long enough to distinguish any oscillatory behaviour from trends in the AMV. There is ''low confidence'' in the paleo AMV reconstructions due to a paucity of high-resolution records.

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