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

==== 10.4.2.1 The Sahel and West African Monsoon Drought and Recovery ====

<div id="h3-41-siblings" class="h3-siblings"></div>

The Sahel, fed by the West African monsoon, has experienced severe decadal rainfall variations (Figure 10.11a). Abundant rainfall in the 1950s–1960s was followed by a large negative trend (Figure 10.11b) until at least the 1980s, over which annual rainfall fell by 20–30% ( [[#Hulme--2001|Hulme, 2001]] ). The subsequent partial recovery ( [[#Wang--2021|]] [[#Wang--2021|B. Wang et al., 2021]] ) is more uncertain: rain-gauge studies suggest a return to long-term positive anomalies in the western Sahel in the early 2000s ( [[#Panthou--2018|Panthou et al., 2018]] ), while CHIRPS merged satellite/gauge data show a wetter western Sahel since 1981 ( [[#Bichet--2018a|Bichet and Diedhiou, 2018a]] , b). The recovery has been more significant over the central rather than the western Sahel ( [[#Lebel--2009|Lebel and Ali, 2009]] ; [[#Maidment--2015|Maidment et al., 2015]] ; Sanogo et al., 2015) and a multiple-gauge record supports a greater recovery to the eastern side ( [[#Nicholson--2018|Nicholson et al., 2018]] ). In this attribution example, drivers of the long-term drought and subsequent partial recovery are discussed, including anthropogenic GHG and aerosol emissions, and sea surface temperature (SST) variations that, in part, relate to internal variability. The reader is also referred to assessment in [[IPCC:Wg1:Chapter:Chapter-8#8.3.2.4|Section 8.3.2.4]] . We define the Sahel within 10°N–20°N across to 30°E, consistent with the eastern boundary used in Chapter 8, and the rainy season as spanning June to September.

<div id="_idContainer038" class="Basic-Text-Frame"></div>

[[File:1db32c7ae496c02335d5444aa0a60a8c IPCC_AR6_WGI_Figure_10_11.png]]

'''Figure 10.11''' '''|''' '''Attribution of historic precipitation change in the Sahelian West African monsoon during June to September. (a)''' Time series of CRU TS precipitation anomalies (mm day <sup>–1</sup> , baseline 1955–1984) in the Sahel box (10°N–20°N, 20°W–30°E) indicated in panel '''(b)''' applying the same low-pass filter as that used in Figure 10.10. The two periods used for difference diagnostics are shown in grey columns. (b) Precipitation change (mm day <sup>–1</sup> ) in CRU TS data for 1980–1990 minus 1950–1960 periods. '''(c)''' Precipitation difference (mm day <sup>–1</sup> ) between 1.5× and 0.2× historical aerosol emissions scaling factors averaged over 1955–1984 and five ensemble members of HadGEM3 experiments after [[#Shonk--2020|Shonk et al. (2020)]] . '''(d)''' Sahel precipitation anomaly time series (mm day <sup>–1</sup> , baseline 1955–1984) in Coupled Model Intercomparison Project Phase 6 (CMIP6) for 49 historical simulations with all forcings (red), and thirteen for each of greenhouse gas-only forcing (light blue) and aerosol-only forcing (grey), with a thirteen-point weighted running mean applied (a variant on the binomial filter with weights [1-6-19-42-71-96-106-96-71-42-19-6-1]). The CMIP6 subsample of all forcings matching the individual forcing simulations is also shown (pink). '''(e)''' Precipitation linear trend (% per decade) for (left) decline (1955–1984) and (right) recovery periods (1985–2014) for ensemble means and individual CMIP6 historical experiments (including single-forcing) as in panel (d) plus 34 CMIP5 models (dark blue). Box-and-whisker plots show the trend distribution of the three coupled and the d4PDF atmosphere-only single-model initial-condition large ensembles (SMILEs) used throughout (Chapter 10 and follow the methodology used in Figure 10.6. The two black crosses represent observational estimates from GPCC and CRU TS. Trends are estimated using ordinary least-squares regression. Further details on data sources and processing are available in the chapter data table (Table 10.SM.11).

The role of SST forcing in the rainfall decline is assessed first. Competing mechanisms from equatorial Atlantic SSTs and inter-hemispheric SST gradients regulate decadal variability in the Sahel ( [[#Nicholson--2013|Nicholson, 2013]] ), alternatively explained by tropical warming leading to Sahel drought, while North Atlantic warming promotes increased rainfall ( [[#Rodríguez-Fonseca--2015|Rodríguez-Fonseca et al., 2015]] ). The SST influence has been formalized in an AMV framework ( [[#Giannini--2013|Giannini et al., 2013]] ; [[#Martin--2014|Martin and Thorncroft, 2014]] ; [[#Martin--2014|Martin et al., 2014]] ; [[#Park--2015|Park et al., 2015]] ), suggesting that relative North Atlantic SST warming increases the Northern Hemisphere differential warming, enhancing Sahel rainfall. The AMV influence is supported by CMIP5 initialized decadal hindcasts ( [[#Gaetani--2013|Gaetani and Mohino, 2013]] ; [[#Mohino--2016|Mohino et al., 2016]] ; [[#Sheen--2017|Sheen et al., 2017]] ), which outperform empirical predictions based on persistence. Some caution is needed since the full magnitude of internal variability is not captured in most CMIP5 models, as poor resolution prevents reproduction of AMV teleconnection responses ( [[#Vellinga--2016|Vellinga et al., 2016]] ), and the magnitude of AMV-related SST variation may be underestimated in CMIP5 ( [[IPCC:Wg1:Chapter:Chapter-3#3.7.7|Section 3.7.7]] , which also assesses that the AMV may be partially forced). The influence of PDV has been studied to a lesser extent, with the PDV positive phase having a negative impact on Sahel rainfall in combined observational/CMIP5 analysis ( [[#Villamayor--2015|Villamayor and Mohino, 2015]] ). The closer match between the observed rainfall declining trend and those in an atmosphere-only SMILE, in which SSTs are matched to observations, compared to three coupled SMILEs in which they are not, suggests that the underlying ocean surface might be essential in driving the decline (Figure 10.11e).

In terms of anthropogenic emissions, regional aerosol emissions from Europe, and to a lesser extent from Asia, have been shown in a global model to weaken Sahel precipitation either through a weakened Saharan heat low or via the Walker circulation ( [[#Dong--2014|Dong et al., 2014]] ). Greenhouse gases (GHGs) and anthropogenic aerosol can be considered together to control ITCZ position based on temperature asymmetry at the hemispheric scale. GHGs increase Sahel precipitation, while aerosol reduces it (in coupled slab-ocean model experiments by [[#Ackerley--2011|Ackerley et al. (2011)]] following [[#Biasutti--2006|Biasutti and Giannini (2006)]] ). This effect is stronger when models account for aerosol–cloud interactions ( [[#Allen--2015|Allen et al., 2015]] ). Perturbed physics GCM ensembles suggests that aerosol emissions were the main driver of observed drying over 1950–1980 ( [[#Ackerley--2011|Ackerley et al., 2011]] ), supported by CMIP5 single-forcing experiments ( [[#Polson--2014|Polson et al., 2014]] ). A coherent drying signal in CMIP5 over the extended 1901–2010 period has also been found, although smaller than the observed trend ( [[#Knutson--2018|Knutson and Zeng, 2018]] ). By applying aerosol scaling factors to the historical period in order to sample the uncertainty in CMIP5 aerosol radiative forcing, [[#Shonk--2020|Shonk et al. (2020)]] found differences of 0.5 mm day <sup>–1</sup> for Gulf of Guinea rainfall between strong and weak aerosol experiments as illustrated in Figure 10.11c, although the drying appears further south than observed due to model bias.

For the partial recovery in West African monsoon and Sahel rainfall since the late 1980s, a detection study using three reanalyses ( [[#Cook--2015|Cook and Vizy, 2015]] ) shows a connection to increasing Saharan temperatures at a rate two to four times greater than the tropical mean, also confirmed by multiple observational and satellite-based data ( [[#Zhou--2016|Zhou and Wang, 2016]] ; [[#Vizy--2017|Vizy and Cook, 2017]] ) and the review of [[#Cook--2019|Cook and Vizy (2019)]] . Reanalyses are also noted to significantly underestimate the Saharan warming ( [[#Zhou--2016|Zhou and Wang, 2016]] ). Saharan warming causes a stronger thermal low and more intense monsoon flow, providing more moisture to the central and eastern Sahel, supported by CMIP5 models ( [[#Lavaysse--2016|Lavaysse et al., 2016]] ), although not all models capture the observed rainfall–heat–low relationship. Sahel rainfall is also incorrectly located in prototype versions of a few CMIP6 models, related to tropospheric temperature biases ( [[#Martin--2017|Martin et al., 2017]] ). Amplified Saharan warming has increased the wind shear, leading to a tripling of extreme storms since 1982, which may partially explain the recovery ( [[#Taylor--2017|Taylor et al., 2017]] ). Instead, observations, multiple models and SST-sensitivity experiments with AGCMs have suggested that stronger Mediterranean Sea evaporation enhances low-level moisture convergence to the Sahel, increasing rainfall ( [[#Park--2016|Park et al., 2016]] ). Meanwhile, an AGCM study suggested that GHGs alone (in the absence of SST warming) could cause Sahel rainfall recovery, with an additional role for anthropogenic aerosol ( [[#Dong--2015|Dong and Sutton, 2015]] ); recent changes in North Atlantic SSTs, although substantial, did not exert a significant impact on the recovery. Large spread in the recovery in a five-member AGCM ensemble suggests that atmospheric internal variability cannot be discounted ( [[#Roehrig--2013|Roehrig et al., 2013]] ).

Consistent timing of the southward ITCZ shift during the decline period in CMIP3 and CMIP5 historical simulations supports the role of external forcing, chiefly anthropogenic aerosol ( [[#Hwang--2013|Hwang et al., 2013]] ). The evolution of the observed decline and recovery is largely followed by the CMIP5 multi-model mean, further supporting the role of external drivers ( [[#Giannini--2019|Giannini and Kaplan, 2019]] ). Updated results from CMIP6 for historical simulations with all and single forcings are represented in Figure 10.11d,e showing smaller trends than those observed. [[#Giannini--2019|Giannini and Kaplan (2019)]] attempted to unify the driving mechanisms for decline and recovery based on singular-value decomposition of observed and modelled SSTs. Since the 1950s, tropical warming arising from GHGs and North Atlantic cooling from aerosol led to regional stabilization, suppressing Sahel rainfall. The subsequent reduction in aerosol emissions then led to North Atlantic warming and recovery of Sahel rainfall. Such mechanisms continue into the near-term future in idealized and modified RCP experiments, with scenarios featuring more aggressive reductions in aerosol emissions, or including aerosol–cloud interactions, favouring a greater northward shift of rainfall ( [[#Allen--2015|Allen, 2015]] ; [[#Westervelt--2017|Westervelt et al., 2017]] , 2018; [[#Scannell--2019|Scannell et al., 2019]] ). There is paleoclimate evidence of changes to Sahel rainfall in the past, in particular with enhancement of the West African monsoon during the mid-Holocene. However, the mechanisms governing such a change have been shown to be largely dynamical in nature ( [[#D’Agostino--2019|D’Agostino et al., 2019]] ), suggesting that the mid-Holocene cannot be used to inform the credibility of changes due to greenhouse warming.

There is ''very high confidence'' ( ''robust evidence'' and ''high agreement'' ) that patterns of 20th-century ocean and land surface temperature variability have caused the Sahel drought and subsequent recovery by adjusting meridional gradients. There is ''high confidence'' ( ''robust evidence'' and ''medium agreement'' ) that the changing temperature gradients that perturb the West African monsoon and Sahel rainfall are themselves driven by anthropogenic emissions: warming by GHG emissions was initially restricted to the tropics but suppressed in the North Atlantic due to nearby emissions of sulphate aerosols, leading to a reduction in rainfall. The North Atlantic subsequently warmed following the reduction of aerosol emissions, leading to rainfall recovery.

<div id="10.4.2.2" class="h3-container"></div>

<span id="the-south-eastern-south-america-summer-wetting"></span>