Climate of the Sahel and West Africa
Summary and Keywords
This article provides an in-depth look at all aspects of the climate of the Sahel, including the pervasive dust in the Sahelian atmosphere. Emphasis is on two aspects: West African monsoon and the region’s rainfall regime. This includes an overview of the prevailing atmospheric circulation at the surface and aloft and the relationship between this and the rainfall regime. Aspects of the rainfall regime that are considered include its unique characteristics, its changes over time, the storm systems that produce rainfall, and factors governing its variability on interannual and decadal time scales. Variability is examined on three time scales: millennial (as seen is the paleo records of the last 20,000 years), multi-decadal (as seen over the last few centuries as seen from proxy data and, more recently, in observations), and interannual to decadal (quantified by observations from the late 19th century and onward). A unique feature of Sahel climate is that is rainfall regime is perhaps the most sensitive in the world and this sensitivity is apparent on all of these time scales.
West Africa1 has one of the most extreme climatic gradients in the world. Within some 1,000 km, the environment transitions from the hyper-arid Sahara Desert to forests with over 1,200 mm of rainfall per year. The Sahel region of West Africa is a semi-arid expanse of grassland, shrubs, and small, thorny trees lying just south of the Sahara. The term “Sahel” is often applied to the general region extending some 5,000 km across the east-west extent of Africa and from the desert to the humid savanna woodland at roughly 10 oN. “Sahel” more properly applies to a smaller region between the latitudes of roughly 14 oN and 18 oN (Figure 1). It includes much of the countries of Mauritania, Senegal, Mali, Niger, Chad, the Sudan, and the northern fringes of Burkina Faso and Nigeria. The Sahel’s highly diverse inhabitants have a long history, its fabled cities such as Timbuktu and Djenne having prospered as centers of trade, education, and political empires many centuries ago. Today the region is home to major cities such as Dakar (Senegal), Niamey (Niger), Bamako (Mali), and Khartoum (Sudan), but most of the inhabitants live in rural areas and practice agriculture. Hence, the vagaries of climate are of great importance to the region. While drought is the predominant concern, extreme rainfall events, which bring rapid flooding, affect both agriculture and urban life. Evidence suggests, in fact, that extreme rainfall has increased in the Sahel, with unprecedented floods ravaging much of the region between 2009 and 2013 (Panthou et al., 2014).
This article focuses on the climate of the Sahel, but all of West Africa from the central Sahara southward to the Guinea Coast along the Atlantic will be considered to some extent. The entire expanse is climatically linked by the dominance of the West African monsoon. The article commences with a geographical overview followed by a description of the mean climate. An important aspect of the climate is the dense dust layer that prevails in much of the region. The discussion proceeds with a description of the West African monsoon, including upper-atmospheric circulation and its relationship to the rainfall regime. Then the rainfall regime is characterized in detail, including the rain-bearing systems that influence the region. Finally, the characteristics of rainfall variability on modern to paleo time scales will be described and potential causes of this variability summarized.
A remarkable feature of West Africa is a vast area of continuous desert that extends over 1,000 km north to south and some 5,000 km east to west. It is called the Sahara in Africa, but the Arabian Desert further east. Along its northern boundary, the region gives way to extratropical, semi-arid Mediterranean lands. Along its southern border lies the semi-arid Sahel. The desert surface is composed primarily of regs, which are stone pavements covering about 68% of the Sahara (LeHouerou, 1986). Another 22% is occupied by ergs, huge sand seas made up of a hierarchy of dunes. Rocky mountain massifs and hamadas, large stone plains, make up the remaining 10%.
In the Sahara most of the “soils” are pedocals, which consist mainly of parent material. Most common are saline solonchaks and calcisols. The former are found in depressions that are seasonally or permanently waterlogged and are found in ancient lake beds. True soils are generally confined to the semidesert, semi-arid, and subhumid regions further south. The most common are weakly developed arenosols, or sand-rich soils. These are the most widespread soils, covering some 50% of the northern Sahel and roughly a quarter of the Sahel overall (FAO, 1991). They are aeolian remnants of more arid periods of the past. The second most common are well-developed lixisols—clay-rich soils that often contain large amounts of iron, aluminum, and titanium oxides. In general, soil fertility is low and the soils are often degraded.
The iron-rich soils are often mixed with bare stretches of hard, rocky, and impervious sandstone or laterite (hardpan) crusts (Figure 2). These usually have a loamy texture and are quite heterogeneous and prone to intense run-off. Water collects on these soils in spotty, temporary ponds or on flood plains. In some areas there are fluvial or lacustrine soils produced from sediments of former lakes and streams. These are often sandy on the surface with clay layers below, so drainage can be poor. The fossil soils tend to be impoverished, but the recent fluvial and lacustrine soils are generally quite fertile.
Overall, West Africa is relatively devoid of high terrain (Figure 3). However, a number of mountain massifs cover the region, mostly in the central and eastern sectors. In the central Sahara and northern Sahel lie the Hoggar of Algeria, Aïr of northern Niger, and Tibesti and Ennedi of northern Chad. Further south lie the Guinea highlands, the Jos plateau of northern Nigeria, and Darfur in the Sudan. Furthest south are the Adamawa Mountains of Cameroon. Peaks in Tibesti, the Hoggar, Darfur, and the Adamawa Mountains reach over 3,000 m elevations. These areas are all climatically important because they markedly enhance rainfall and might also serve as triggers for many of the region’s storms.
Within the Sahel flow six major rivers: the Senegal, Gambia, Niger, Bani, Ubangi, and Chari. The only major lake is Lake Chad, which is fed by the Ubangi-Chari system. Its size was once 25,000 km2, but it is now nearly desiccated as a result of dry conditions since the 1960s and irrigation (Coe & Foley, 2001). The Sahel landscape is dotted with numerous dry lake beds, which are remnants of vast lakes of the mid-Holocene period. The Senegal and Niger originate in the Guinea Highlands. The Niger meanders northward into the central Sahel, where it creates an inland delta (the Niger Bend region) that often overflows into low-lying depressions in the wet years. The river then turns southward, reaching the Atlantic in southern Nigeria.
Throughout West Africa, climate and vegetation are largely a function of latitude (Figure 4), so that vegetation zones stretch from east to west. As rainfall increases progressively southward from the Sahara there is a gradual transition between semidesert grassland, savanna grassland with low trees and shrubs, savanna woodland, woodland, and forest. Along this gradient, the vegetation becomes increasingly taller and the proportion of woody species (trees, shrubs, bushes) and the amount of ground cover increase.
In the Sahel itself (Figure 5), most grasses are both perennial, generally not taller than 80 cm, and annual. In some areas there are occasional small trees or shrubs. These are often thorny, like the typical acacia. The surface is to a large extent bare soil and the vegetation tends to form a mosaic pattern, clustered in regions of favorable soil or run-off. A widespread example of this is the so-called brousse tigrée (tiger bush, Figure 6), a notable Sahelian feature (Nicholson, 1995). The name comes from the fact that, viewed from the air, the vegetation clusters resemble stripes on a tiger. The clusters of dense vegetation can be 200 m in length and 20 m wide. The size of the cluster is generally about half the size of the surrounding bare soil.
By far the most significant climatic element of West Africa is rainfall. The extremes range from near zero in the hyper-arid eastern Sahara to over 10,000 mm per year (on average) at Debundscha, at the base of Mount Cameroon. In 1919 Debundscha received nearly 15,000 mm of rainfall. Figure 7 shows mean annual rainfall over all of northern Africa. Most of the Sahara receives less than 25 mm per year, with large expanses receiving less than 10 mm. Mean annual rainfall steadily increases southward toward the equator, but nearly everywhere it is less than 2,000 mm. The exceptions are over the Guinea Highlands and the Adamawa Mountains of Cameroon. A small “dry” zone exists along the Guinea Coast, with annual rainfall of less than 1,000 mm.
A striking characteristic of the West African rainfall regime is the roughly east-west (i.e., zonal) orientation of the rainfall isohyets, as illustrated by both annual and August rainfall (Figure 7). This is a consequence primarily of the rain-bearing disturbances, most of which commence in the east and traverse nearly the entire region, steered by the prevailing easterly winds. (This is described in more detail in the section on “Rain-Bearing Disturbances.”) This zonal structure of the rainfall regime is, in turn, largely a consequence of the orientation and latitude of the West African land mass, with dynamical processes largely constrained by the “fixed” boundary conditions imposed by the presence of the Sahara to the north and the ocean to the south.
The pattern of rainfall produces a similar zonal orientation of the vegetation zones. Consequently, numerous authors have produced a climate classification for the region based on the vegetation zonation (Figure 4). The zones commence with the Sahelo-Sahara furthest north, progressively shifting to the Sahelian, Soudanian, Soudano-Guinean and Guinean zones (Wezel, Bohlinger, & Böcker, 2000). The Sahel-Sahara is primarily a semidesert grassland, while a savanna grassland is characteristic of the Sahel. The Soudanian zone corresponds roughly to dry savanna woodland, the Soudano-Guinean zone to moist woodland and the transition to forest. The Guinean zone is primarily forest.
No precise limits can be assigned to these zones either geographically or on the basis of rainfall because varying definitions exist (Wezel et al., 2000) and rainfall in the region has changed markedly over time. A rough idea of the corresponding rainfall regime is given in Table 1. The southern desert boundary is generally given as 50 mm/year, with no preferred seasonality. Semi-arid climates begin at about 20–22 °N. Within the semi-arid zones, rainfall systematically shifts from 50 mm/year with a rainy season lasting 1–2 months, to 1,000 mm/year with a rainy season lasting 3–5 months. In the north, about 80–90% of the rainfall occurs during the wettest quarter of the year; in the south, this figure falls to about 50% (Figure 8). Most of the Soudano-Guinean zone, with 5 to 8 months of rainfall per year, and Guinean zone, with 9 to 12 months, fall within the humid climate classification.
Table 1. Approximate Climatic Characteristics of Four Vegetation Zones in West Africa: Mean Annual Rainfall (mm), Coefficient of Variation (CV, %), Length of the Rainy Season (Months).
Annual Rainfall (mm)
Source: Adapted from Nicholson (1981).
Throughout the semi-arid regions of West African rainfall is highly variable from year to year. The coefficient of variation (the standard deviation of yearly totals deviation by the annual mean) ranges from 20% in the south to 75% to over 100% in the north (Figure 9). Within the Sahara itself it ranges from about 75% to over 150%. In the humid regions to the south it is on the order of 15–20%. Thus, the drier the region, the greater the interannual variability of rainfall.
Figure 10 presents further detail about rainfall seasonality in West Africa. Rainfall tends to be unimodal (i.e., one maximum in the seasonal cycle) over the semi-arid regions, but bimodal in the central Sahara and in the humid Soudano-Guinean and Guinean zones. In unimodal regions maximum rainfall occurs on average in August. The exception is a September peak in the narrow belt in the southernmost sector of the Soudano-Guinean zone. In the Guinean zone rainfall tends to peak in September or October and June or July, with a minimum generally occurring in August. These patterns of seasonality relate to the north-south movement of the West African monsoon (see the section on the “West African Monsoon”). The bimodal in the central Sahara has a different cause. The Sahara marks the transition between the tropical summer rainfall regime and the extratropical winter rainfall regime. For that reason, most areas have a small peak on average in either the boreal summer (more southern areas) or the boreal winter (more northern areas). The central Sahara also receives some rainfall in the transition months from systems developing from a combination of tropical and midlatitude effects. Thus, most of the central Sahara has, on average, a second peak in one of the transition season months.
West Africa extends from the equatorial tropics to the lower midlatitudes so that temperature in the region is very much a function of latitude. In most areas, January is the coldest month. Mean maximum temperatures in January increase equatorward, ranging from roughly 21 °C in the central Sahara to 33 °C in latitudes nearest the equator (Figure 11). Mean January minimum temperatures similarly increase equatorward, from about 6 °C in the central Sahara to 18 °C or 20 °C in the equatorial latitudes. In July mean maximum temperatures decrease equatorward, being around 40 °C in the central Sahara but on the order of 28 °C to 30 °C in southernmost areas of West Africa. Minimum temperatures in July, however, show little latitudinal variation. Over most of West Africa they are on the order of 21–24 °C.
In contrast to temperate latitudes, the annual cycle of temperature does not necessarily have minima and maxima in the extreme seasons (Figure 12). Factors other than the seasonal march of solar radiation can play important roles. At Accra, Ghana, for example, August is the coldest month, in part because cold water in the Gulf of Guinea reduces temperature. In the semi-arid regions cloud cover during the boreal summer rainy season similarly reduces temperature, so that the highest temperatures occur in April or May, just prior to the rainy season.
The diurnal and annual range of temperature (Figure 13) are also strongly influenced by latitude, but aridity and proximity to the ocean also play a role. Over West Africa the mean annual range is lowest in the near-equatorial latitudes (roughly 3 °C) and increases northward. In the northernmost Sahel it is on the order of 15 °C. Throughout West Africa the diurnal range exceeds the annual. In January the diurnal range is on the order of 15–18 °C. Exceptions are areas along the Atlantic coast, where it is as low as 9–12 °C, and the highlands of the Sudan, where it exceeds 22 °C. In July the diurnal range varies from 6–8 °C in the lowest latitudes to 15 °C in the northern Sahel.
Relative Humidity and Cloud Cover
The patterns of relative humidity and cloud cover are highly variable from one station to the next. However, several generalizations can be made. For one, these tend to roughly follow the rainfall seasonality. Also, they tend to be different north and south of about 10 °N, south of which humid climates prevail. Griffiths (1972) provides specifics for numerous stations).
In the humid climates relative humidity tends to be above 80% and often near 100% in the early morning throughout the year. Afternoon relative humidity is more typically on the order of 60–70%. In the semi-arid regions, including the Sahel, morning relative humidity tends to be on the order of 40% in the dry season but over 90% in the wet season. Afternoon relative humidity tends to be as low as 20% or less during the dry season and on the order of 60–70% during the wet season.
Cloud cover in the humid zone tends to range from 50% in the driest months to 80–100% during the wet season. In the semi-arid zones, it is generally as low as 25–30% on average during the dry months and 70–80% during the wet season, with cloud cover generally decreasing northward.
Surface Wind and Pressure
Figure 14 shows surface wind and sea-level pressure during January, April, July, and October. In January, high pressure prevails over the western Sahara, and throughout most of West Africa wind is from a northeasterly direction, with a stronger northern component in the east than in the west. This wind is referred to as the Harmattan and on average it extends in January equatorward to roughly 7 °N. To the south is a very weak southwesterly flow.
The picture is very different in July when a thermal low exists over the northwestern Sahara, between the Hoggar and the Atlas Mountains. Termed the Saharan Heat Low or the West African Heat Low, it maintains this position from late June through September. Within the course of the system, however, it undergoes changes in intensity and location, occasionally migrating further east. A second region of low pressure exists in the eastern Sahara and over the Arabian Peninsula. The northeasterly wind extends only to about 20 °N, with strong southwesterly flow to the south of this latitude. This is sometimes called the southwest monsoon.
During the transition season months of April and October, high pressure prevails over the Sahara and the northeast Harmattan is the dominant flow, as it is in January. However, it only reaches approximately 10–12 °N in April and October. The main contrast between prevailing winds in April and October is the development of the southwest monsoon flow. It is strong in April but extremely weak in October.
A consequence of the dominance of these two wind systems, the Harmattan and the monsoon, is that prevailing surface wind direction is remarkably constant from month to month over most of West Africa. Near the coast, it is generally southwesterly, or in some regions westerly, throughout the year. North of about 10 °N, these same directions prevail during from roughly April to September or October, with the northeasterly surface winds prevailing in other months.
Figure 15 presents vertical cross-sections of the mean zonal wind over West Africa in these same four months. The dominant feature in all months is the prevalence of easterly winds in the low latitudes and westerlies in the higher latitudes. The belt of westerlies expands with altitude and also exhibits a seasonal displacement towards the summer hemisphere. Several features apparent in the zonal winds play a major role in West Africa climate. One is a midlevel easterly jet stream (the African Easterly Jet, or AEJ) with a core at roughly 650 hPa. It is evident all year, but similarly migrates with the seasons, with the latitude of its core ranging from 2 °N in January to 14 °N in July. The monthly mean speed of the AEJ can attain 16 ms-1, but mean speeds on the order of 10 ms-1 to 12 ms-1 are more typical. A second midlevel jet core is evident in the Southern Hemisphere in October. Nicholson and Grist (2003) suggested distinguishing these as the AEJ-N and the AEJ-S. A second feature is an upper-troposphere easterly jet stream, the Tropical Easterly Jet, or TEJ. It is well developed in July, but a weak jet core is apparent in January and April as well. The third feature, low-level westerly flow near the equator, is apparent throughout the year.
The equatorial westerlies generally represent the southwest monsoon flow. However, in some years a distinct jet core develops at roughly 850 hPa. This is termed the African Westerly Jet, or AWJ, and it is not a simple extension of the monsoon flow. There is a marked directional discontinuity between the monsoon flow and the AWJ, with the southerly component disappearing at the AWJ level. The origin of this jet appears to be inertial instability related to the cross-equatorial pressure gradient (Nicholson & Webster, 2007). Some semblance of the AWJ is seen in the July cross-section in Figure 15. In such cases, westerlies can prevail up to the mid-troposphere, in which case they display the AEJ poleward.
Several other low-level jet streams affect the region. The best known is the Bodélé Jet in northern Chad. It arises as the easterlies are channeled through the gap between the Tibesti and Ennedi massifs (Todd, Washington, Raghavan, Lizcano, & Knippertz, 2008; Washington & Todd, 2005). The mean core speed of the Bodélé Jet is 8 ms-1, with maximum intensity in January (Figure 15). It transports massive amounts of dust toward the Atlantic Ocean. A second low-level westerly jet, which resides over the equatorial Atlantic, also appears to influence West African climate. Termed the West African Westerly Jet, or WAWJ, it brings moisture into the latitudes south of the Sahel, but its speed exhibits a correlation with Sahel rainfall (Grodsky et al., 2003). A boundary layer feature termed the Nocturnal Low-Level Jet (NNLJ) also influences rainfall over the region by way of moisture transport (Abdou, Parker, Brooks, Kalthoff, & Lebel, 2010; Lothon, Said, Lohou, & Campistron, 2008; Sultan, Janicot, & Drobinski, 2007). Its direction is northeasterly when embedded within the Harmattan but southwesterly when embedded within the monsoon flow.
Aerosols and Dust Storms
An outstanding characteristic of Sahel climate is a deep layer of mineral dust produced by wind erosion of the dry and barren land surface. It persists much of the year, generally from late spring to early fall. The frequent squalls occurring in the region mobilize the dust and entrain it into the atmosphere. Easterly waves help to generate dust (Knippertz & Todd, 2010; Jones, Mahowald, & Luo, 2003). Dust is transported across the Atlantic and into Europe (Figure 16). When a dust storm occurs, visibility is reduced to near zero. Much of the dust is deposited on the surface during these storms or during rain events. Figure 17 shows before and after photos of a swimming pool in Niamey during a July storm. Within 30 minutes of the storm’s onset, the pool’s water changed from crystal clear to an opaque deep orange-brown. In addition to such transient storms, a dust haze often persists near the surface for weeks at a time, especially during the boreal winter. This dramatically reduces visibility and poses a health hazard.
The source of the mineral dust is dry beds of numerous Holocene lakes that are spread across the Sahel. The biggest source is from the Bodélé Depression, the site of the expanded Lake Chad some 5,000 years ago. West Africa overall generates roughly 50% of the world’s total atmospheric mineral dust content (Prospero, Ginoux, Torres, Nicholson, & Gill, 2002). The amount of dust is strongly anticorrelated with rainfall during the previous year (Prospero & Lamb, 2003). The dust was fairly localized and transient during the 1950s, but increased steadily as a result of the drought in the early 1970s and early 1980s (N’tchayi Mbourou, Bertrand, & Nicholson, 1997).
The concentration of dust is higher near the top of the boundary layer than near the surface, where convection helps to cleanse the atmosphere. The residual dust lies in a more stable air layer in the upper boundary layer, forming a deep layer of hot and dry air that is characterized by a red haze (Figure 18) (Carlson & Prospero, 1972; Prospero et al., 2002). This is termed the Saharan Air Layer, or SAL (Figure 19). This layer is vast, typically some 2,000–3,000 km in east-west extent, and it can extend to 5,000 km (Dunion & Velden, 2004). The SAL resides between roughly 900 mb/1,800 m and 500 mb/5,500m near the Atlantic coast (Carlson & Prospero, 1972), but it extends considerably higher over the continent (Parker, Thorncroft, Burton, & Diongue-Niang, 2005).
The SAL is bound by subsidence inversions at its top and bottom (Figure 19). These help to maintain the integrity of the layer and its thermal structure and its relatively uniform conditions of high heat content and low moisture content. These inversions help to stabilize the atmosphere, suppressing convection. The dust itself further impacts Sahel climate. It affects the atmospheric radiation budget (Zipser et al., 2009), creates an “elevated heat pump” (Lau, Kim, Sud, & Walker, 2009), and serves as ice and cloud condensation nuclei (deMott et al., 2003). The SAL impacts the AEJ and African Easterly Wave development (Thorncroft et al., 2003), systems that strongly affect the production of rainfall. However, the net effect on rainfall, both on mean climate and its variations on annual to multidecadal scales, is still unclear (Nicholson, 2015).
The West African Monsoon
Classic Picture of West African Climate
In the classic picture of West African climate, rainfall and its seasonality are associated with the annual displacement of a surface feature, the Intertropical Convergence Zone (ITCZ). This zone represents the location where the northeast Harmattan and southwesterly monsoon converge; it is marked by a solid line in each frame of Figure 14. This zone “follows the sun,” migrating northward into West Africa in the boreal summer and southward toward the equator in the austral summer. It reaches its extreme latitudinal positions in August and January. It is generally assumed that rainfall is associated with this zone, arising from local instability but facilitated by the low-level convergence of the ITCZ. The resultant picture of West African weather zones is shown in Figure 20, with maximum rainfall occurring south of the ITCZ where the moist layer is sufficiently deep to support intense cloud development.
This scenario stems back to the time when rainfall in the tropics was thought to result mainly from local thunderstorms. Moreover, it appears to explain the annual march of rainfall over Africa (Figure 21), with the region of intense rains migrating from the Northern Hemisphere to the Southern Hemisphere between August and January, and reversing its course in January.
The Current Monsoon Paradigm
Research during the last decade has greatly modified the picture just described and the current paradigm is that of a monsoon system. The circulation over West Africa clearly exhibits the most basic characteristics of a monsoon: a pronounced seasonal wind shift that is produced by thermodynamic contrasts between the land (i.e., the Sahara) and the ocean (i.e., the equatorial Atlantic). The southwesterly flow extends from the Atlantic cold tongue (cool water close to the equator in the Gulf of Guinea) to the Saharan heat low. This flow originates in the Southern Hemisphere as a consequence of the southeasterly flow associated with the South Atlantic High, taking on a westerly component as it crosses the equator.
The current image of the monsoon contrasts with the classic picture of West African climate in the diminished importance of the ITCZ and the inclusion of several jet streams and the Saharan Heat Low. (Overviews are presented by Nicholson & Grist, 2003; Lebel, Diedhiou, & Laurent, 2003; Gu, Adler, Huffman, & Curtis, 2004; Parker et al., 2005; Zhang, Woodworth, & Gu, 2006; Nicholson, 2009; Peyrille, Lafore, & Redelsperger, 2007; Thorncroft, Nguyen, Zhang, & Peyrille, 2011.) Several representations of the monsoon were developed independently by these authors, but they all are substantially similar in their most salient features.
Most of the recent research on the West African monsoon stems from several field experiments and research programs. These include the African Monsoon Multidisciplinary Analysis (AMMA) experiment that took place in 2006 (Janicot et al., 2008; Redelsperger et al., 2006); the associated model intercomparisons project (ALMIP) (Boone et al., 2009); the AMMA Catch Experiment, which extended AMMA southward into Benin (Lebel et al., 2009); and the JET2000 Experiment, which focused on the AEJ (Thorncroft et al., 2003).
Thorncroft et al. (2011) identify four key phases of the annual cycle of the monsoon, distinguished by the location of peak rainfall. These are the oceanic, coastal, transitional, and Sahelian phases. In the oceanic phase, between November and mid-April, a broad rain belt lies just north of the equator. A cold tongue develops just south of the equator in the Atlantic. During the subsequent coastal phase, which generally prevails to mid-June, peak rainfall lies over the ocean but in the near-coastal region around 4–5 °N. During the transitional phase, which occurs in early July, a decrease in rainfall is observed and the peak rainfall occurs around 6 ° N. Lebel et al. (2003) term these three phases the “oceanic regime.” The transition phase is brief, followed in mid-July by an abrupt shift of the rainfall peak to around 10 °N (Lebel et al., 2003; Sultan & Janicot, 2000), marking the onset of the Sahelian phase. The onset of this phase is also marked by an intensification and northward shift of the Saharan Heat Low (Ramel, Gallee, & Messager, 2006; Sijikumar, Roucou, & Fontaine, 2006) and a northward shift of the AEJ (Gu et al., 2004). This phase, termed the continental regime by Lebel et al. (2003), persists through September.
Figure 22 presents a schematic of the monsoon structure in the Sahelian phase at the height of the Sahel rainy season. The main circulation features are the TEJ in the upper-troposphere and the midlevel AEJ. The monsoon also includes three areas of convergence and two meridionally (north-south) oriented vertical circulation cells. A deep meridional cell consists of rising motion in the latitudes between axes of the AEJ and the TEJ and subsidence over the central Sahara and near the equator. This cell is essentially a local manifestation of a Hadley-type circulation. A shallow meridional cell is centered over and driven by the Saharan Heat Low and also corresponds to the surface position of the ITCZ. Convergence is associated with the ITCZ and the equatorial flank of the AEJ. Convergence also occurs at low levels when the southwest monsoon flow encounters the Atlantic coast. This appears to feed into the ascent of the deep cell.
The regions of wind convergence are also regions of moisture convergence. This facilitates the development of precipitation within the deep meridional cell and near the coast, but little rainfall is associated with the convergence within the shallow meridional cell. Within the core of the deep cell lies a deep region of moist air in which relative humidity is 60–80% throughout the troposphere.
The vertical motion field associated with this circulation is shown in Figure 23 (Nicholson, 2009). Deep ascent, with a maximum around 450 hPa, occurs in the latitudes bounded by the two jet axes. This corresponds to the core of the rain belt. Above the ITCZ is a second region of ascent, the ascending branch of the shallow meridional cell. Subsidence overrides the shallow region of ascent. A zone of subsidence also separates the ITCZ from the rainfall maximum associated with the ascent of the deep meridional cell. This effectively “decouples” the ITCZ from the region of significant rainfall. On average, the Sahel lies within this region of subsidence. In very wet years the two cells merge, creating a very broad band of rainfall (Nicholson, 2008).
The independence of the ITCZ and rainfall maximum is a critical point, because many authors equate the ITCZ with the rainfall or cloudiness maximum. However, their seasonal migration is quite different and year-to-year shifts are also quite independent (Grist & Nicholson, 2001). Peak rainfall (Figure 21) shifts from 10 °N in August to 15 °S in January, while at the same time the surface convergence zone shifts from 20 °N to 7 °N, never reaching the Southern Hemisphere. To emphasize this point, various authors refrain from using the term “ITCZ” with respect to West Africa and make specific mention of the region of high rainfall. Nicholson (2009) refers to this as the tropical rain belt, Zhang et al. (2006) use the term “rain band,” and Ross and Krishnamurti (2007) use the term “equatorial rain belt.”
An additional component of the West African monsoon is the African Easterly Wave (AEW). This comprises disturbances in the zonal wind field with typical wavelengths on the order of 2,000 to 5,000 km (Leroux, Hall, & Kiladis, 2010). They propagate along two east-west tracks (Zawislak & Zipser, 2010) (Figure 24) and occur in two distinct frequency bands of 3–5 days and 6–9 days (Diedhiou, Janicot, Viltard, de Felice, & Laurent, 1999). Those along the northerly track (∼18–20 °N) are seldom associated with precipitation. Waves on the southern track, at ∼9–11 °N, are usually, but not always, associated with convection and precipitation (Druyan, Fulakeza, & Lonergan, 2006; Pytharoulis & Thorncroft, 1999). The waves organize convection and probably enhance it, contributing to the development of intense rain-bearing systems (Laing, Carbone, Levizzani, & Tuttle, 2008). They also are important triggers for the development of squall lines (Fink & Reiner, 2003).
Early work on AEWs emphasized their development from the AEJ via combined barotropic-baroclinic instability and their role in organizing and promoting convection. This picture has changed over the last decade. Extensive evidence has been presented that there is mutual interaction between the waves and convection and that convection can serve as a trigger for wave disturbances. However, barotropic-baroclinic instability appears to be required for their continued development.
Most of the rainfall over West Africa is associated with large-scale systems that propagate westward across the continent. These are larger than individual thunderstorms but smaller than synoptic scale phenomena. These systems have various names, including line squalls, organized convective systems (OCSs), and cloud clusters (Fink, Vincent, & Emert, 2006). These are collectively termed Mesoscale Convective Systems (MCSs) by some authors, while others (e.g., Nesbitt & Zipser, 2003) reserve this label for “cold” systems, that is, those in which ice is present in the upper levels of cloud.
A typical MCS is shown in Figure 25 and compared with other systems. An MCS is essentially a large, continuous area of deep cloud (at least 2,000 km) in which one or more areas of convective precipitation are embedded. The average size of MCSs is 10,000 km. Most definitions require the presence of ice in the upper cloud layers. MCSs account for up to 90% of the rainfall over the Sahel and 50% of the rainfall in the tropics overall, although they comprise only 2% of all rain-bearing features in the tropics (Nesbitt, Cipelli, & Rutledge, 2006).
MCSs are in a constant state of evolution. These systems produce intense convective rainfall mainly during the afternoon, with most convective events lasting three hours or less over land (Ricciardulli & Sardeshmukh, 2002). As the system evolves in the later hours of the day and into the night, the cloud anvil that tops the system spreads and produces a large area of stratiform cloud (Nesbitt & Zipser, 2003). Rain from the stratiform cloud layers typically occurs at night and for a longer period of time, but the rain rate is roughly one-fourth the rain rate associated with convective clouds. Stratiform precipitation accounts for 73% of the rain area and contributes roughly 40% of the total rainfall for the tropics as a whole (Schumacher & Houze, 2003). Over the Sahel the contribution to total rainfall ranges from less than 20% in the south to more than 80% in the north (Jackson, Nicholson, & Klotter, 2009).
The mesoscale convective complexes include both rapidly moving squalls and slower moving nonsquall systems (Schumacher & Houze, 2003). A study at Niamey, Niger (13.5 °N), in 2006 found that MCSs with squalls accounted for about 90% of the rain there, though the squalls were present in only 17% of rain events (Ferreira, Rickenbach, Guy, & Williams, 2009).
Although most of the rainfall in the Sahel occurs in the boreal summer, unseasonal rains can occur in the transition seasons and even in the heart of the dry season. These most often affect the western Sahel and can bring as much as 25 mm to the Sahel in the middle of the dry season. A case in point is the “heug” rains of Mauritania. These can persist for days on end (Nicholson, 1981). The rainfall associated with them can extend as far south as the Guinea Coast.
Most of this unseasonal rainfall is associated with systems that develop as a result of tropical-extratropical interactions. The two systems most commonly described are the “tropical plume” (Knippertz & Fink, 2008; Knippertz, Fink, Reiner, & Speth, 2003; Knippertz & Martin, 2005) and the Soudano-Saharan depression (Dubief, 1947). Both include a diagonal trough emanating from the midlatitude westerlies at upper levels toward the tropics (Figure 26). Both are also assumed to overlie a surface tropical disturbance, often an AEW, but Schepanski and Knippertz (2011) suggest this is not always the case. Further examination is warranted because such systems may have played a much greater role in the Sahel in past centuries, contributing to some of the historical wetter episodes experienced by the region (Nicholson, 1981).
Climatic Variability and Change
Rainfall Conditions from Historical to Modern Times
One of the most significant aspects of West African, and particularly Sahelian, climate is its extreme variability on time scales ranging from years to millennia. During the peak of the late Pleistocene glacial period, some 18,000 years ago, the Sahara had expanded southward to roughly 10 °N. Tropical woodlands and forests had all but disappeared, existing only in mountain refuges (Nicholson & Flohn, 1980). By the mid-Holocene, some 5,000 years ago, conditions were much wetter than present and most of the Sahara was a savanna environment. Lakes dotted the landscape across its east-west extent. Lake Chad was ten times its present size.
During the past millennium significant changes have also occurred and these have been documented in more detail. Information from lakes, historical chronicles, explorers’ journals, and other documentary sources suffice to provide very general detail prior to the 16th century and higher temporal resolution in the centuries that followed. Episodes of wetter conditions probably prevailed in the Sahel and other semi-arid zones throughout much of the 8th through 13th centuries and again during the 16th through 18th centuries, or possibly earlier (Nicholson, 1979). Important kingdoms such as the Songhai and Mali empires thrived in West Africa, making the region a center of commence and education. Ibn Battuta’s (1929) journals of his 14th century travels suggest that a savanna environment then existed in now-desert areas of northern Mauritania. At the same time, caravans traveled directly across the El Djouf desert, where a lack of wells and oases would now preclude such travel. In much of the 8th through 15th centuries, rain-fed agriculture was practiced in sandy plains of Mauritania where today’s aridity would make this impossible. An overall more humid climate probably persisted well into the 1700s, but commencing in the mid-1600s a series of disastrous droughts occurred. Drought was widespread in the region in the 1680s, ca. 1738–1756, the 1770s, and the 1790s. More localized and less intense droughts also occurred in the 1640s and 1710s (Nicholson, 1980).
The 1790s drought appeared to mark a major transition to more arid conditions throughout West Africa. This culminated in persistent and ravaging drought that was most severe from around 1828 to 1839. This era was actually part of a major, continent-wide period of anomalously dry conditions. Figure 27 shows the rainfall conditions over 90 regions of the continent, using a 7-class index described in Nicholson (2001), Nicholson, Dezfuli, and Klotter (2012), and Nicholson, Klotter, and Dezfuli (2012). By mid-century more normal conditions had returned and in the 1870s and 1880s abnormally good rains prevailed in the Sahel. Touareg harvests in the most northern Sahel were consistently good and the Niger floods were consistently high, often overflowing the river’s banks and flooding the surrounding low-lying depressions. Prosperity was so good that wheat was exported from the Niger Bend region. This period ended fairly abruptly around 1895, with a steady decline leading to drought that was particularly severe in 1913 and 1914.
The image in Figure 27 was reconstructed from hydrological and documentary information. However, toward the end of this period quantitative rainfall measurements were adequate to confirm most of these trends as of the mid-nineteenth century. Long-term time series for the Sahel and the Guinea Coast are shown in Figure 28 (from Nicholson et al., 2018).
Two characteristics of the variability are readily seen in these time series. The first is the major shift to dry conditions around 1968, with relatively dry conditions persisting since that time. Rainfall in the Guinea Coast has recovered to some extent in recent years, but the Sahel rainfall has exceeded the long-term mean in only a handful of years. The second characteristic is the similarity of the two time series on decadal time scales, with high rainfall in the 1950s and driest conditions in the 1980s. On annual time scales, however, rainfall anomalies are frequently of opposite sign in the two regions, as described in the section on “Land Surface Effects.”
Figure 29 shows the spatial pattern of rainfall anomalies in these two decades. In the 1950s rainfall was above average in most of West Africa, with rainfall exceeding the long-term (nearly century-long) average by over 20% in much of the Sahel. Similar conditions had not existed since the 1880s. However, along the Guinea Coast the pattern in the 1950s was quite varied and, as seen in Figure 28, wet years are interspersed with dry years. In the 1980s, the rainfall deficit averaged more than 20% throughout the Sahel and on the order of 10—20% over much of the Guinea Coast region.
Figures 28 and 29 illustrate two almost unique characteristics of rainfall variability in the Sahel: the vast spatial coherence and the tendency for year-to-year persistence of strong anomalies. Rainfall anomalies tend to be of the same sign throughout the region north of roughly 10 °N, leading in some years to extremely widespread drought. Rainfall in the Sahel is also remarkably persistent from year to year, with rainfall being above the long-term mean every year from 1950 to 1967, then below the mean in every year but one from 1978 to 1986. A major regime shift appears to have occurred in 1968 (Losada et al., 2012). Rainfall along the Guinea Coast shares these characteristics to some extent, but the anomalies are neither as spatially coherent nor as persistent from year to year as those in the Sahel.
A Conceptual Model of Rainfall Variability Over the Sahel
Spatial coherence of rainfall variability over West Africa is exceedingly strong and much of the interannual variability can be described by two spatial modes (Figure 30) (Nicholson, 2008, 2009). These are evident on both modern and historical time scales (Nicholson, 2014). One is a pattern of opposition between the Sahel and the Guinea Coast, a pattern often referred to as a rainfall dipole. The second comprises anomalous conditions of the same sign through the region.
Nicholson and Grist (2001) developed a conceptual model of these modes that relates the dipole to a latitudinal shift of the tropical rain belt over West Africa and relates the mode with anomalies of uniform sign to an intensification or weakening of the tropical rain belt. In both cases, wet conditions in the Sahel are related to an intensification of the Tropical Easterly Jet. In the case of the dipole, Nicholson and Webster (2007) demonstrated that the wet Sahel/dry Guinea coast pattern is linked to inertial instability. This in turn is linked to the cross-equatorial pressure gradient, with an intensification of both the Saharan Heat Low and the South Atlantic High. The impact of the inertial instability is the development of a strong African Westerly Jet, displacing the AEJ northward, along with the rain belt, into the Sahel. The dipole appears to have become much less prominent in recent decades (Losada et al., 2012).
Large-Scale Forcing of Rainfall Variability in the Sahel
While the changes in the circulation regime over West Africa provide the direct forcing of rainfall variability, the question arises as to what produces these circulation changes. Most work has centered on sea-surface temperatures. Some of the earliest work was that of Lamb (1978a, 1978b), who linked Sahel drought to an SST dipole in the equatorial and subtropical Atlantic. Folland, Palmer, and Parker (1986) demonstrated a broader-scale link to the drought, interhemispheric SST gradient.
It has been well established that rainfall along the Guinea Coast is strongly forced by SSTs to the south, in the equatorial Atlantic (e.g., Wagner & DaSilva, 1994). However, there has been vigorous debate about the role of SSTs in Sahel rainfall variability, with links to tropical sectors of the Pacific, Atlantic, and Indian Ocean; to the North Atlantic; to the Mediterranean; and to the Antarctic being demonstrated by various studies (Nicholson, 2013). On the synoptic time scale, cold air surges from the Mediterranean and the Madden-Julian Oscillation also appear to play a role (Janicot & Sultan, 2001; Lavender & Matthews, 2009; Vizy & Cook, 2014). One of the more controversial conclusions concerns the role of ENSO, with several observational studies (e.g., Nicholson & Kim, 1997; Rasmusson & Arkin, 1993; Ropelewski & Halpert, 1987, 1989) concluding there is no link, and others (e.g., Joly, Voldoire, Douville, Terray, & Royer, 2007; Rowell, 2001; Semazzi, Mehta, & Sud, 1988) claiming a strong relationship. Ward (1998) has suggested that ENSO’s importance in the Sahel is on interannual (year-to-year) time scales.
The plethora of demonstrated relationships with SSTs and the discrepancy between various authors can probably be accounted for by several factors. As for el Niño, it appears to have the opposite effect during the height of the rainy season (when it decreases rainfall) and in the early and late rains (when it increases rainfall) (Nicholson, Somé, & Kone, 2000). Also, the timing of the ENSO event is critical to the development of a Sahel teleconnection (Joly et al., 2007). Further, the relationship to SSTs is highly nonstationary, so that conclusions concerning SSTs as a forcing factor have some dependence on the period of analysis (Losada et al., 2012; Mohino et al., 2011; Paeth & Friederich, 2004). Finally, the influence varies across the east-west extent of the Sahel (Bader & Latif, 2003).
A strong factor in decadal-scale fluctuations appears to be the Atlantic Multidecadal Oscillation (AMO) (Ting, Kushnir, Seager, & Li, 2011; Zhang & Delworth, 2006). The Sahel appears to be particularly sensitive to the AMO. This may be a result of the strong relationship between the cross-equatorial pressure gradient and Sahel rainfall, as described in section on a “Conceptual Model of Rainfall Variability over the Sahel.” The AMO and this pressure gradient over West Africa are closely linked (Nicholson & Webster, 2007). This is a likely factor in the decadal scale persistence of rainfall anomalies.
Clearly rainfall variability in the Sahel responds significantly to SST forcing. The resultant rainfall reflects collectively the impact of the world’s ocean, with the dominant factor or factors varying over time. The relevant point is the role these play in modifying the local circulation regime most closely associated with interannual variability, the intensity of the TEJ, and the latitudinal location of the AEJ. More detail about the relationship to SSTs can be found in Biasutti, Held, Sobel, and Giannini (2008) and Rodriguez-Fonseca et al. (2015).
Land Surface Effects
When drought appeared in the Sahel in the late 1960s and early 1970s, Charney (1975) put forth a long-debated hypothesis that the drought was anthropogenic. He speculated that albedo changes linked to overgrazing modified the radiation balance in such a way that increased subsidence was required to produce a radiative equilibrium. This would negatively affect the production of rainfall. He later (Charney, Quirk, Chow, & Kornfield, 1977) suggested that soil moisture could exert a similar feedback on rainfall. Most meteorologists working in Africa refused to accept the idea that the drought was caused by human beings. In fact, much observational evidence clearly refuted the idea. However, it was noted that Sahel rainfall is characterized by an almost unique persistence of wet or dry conditions over a decade or longer. Nicholson (2000), Entekhabi (1995), and others suggested that this persistence might be related to land-surface atmosphere feedback.
Eventually the concept of this feedback as a factor in interannual variability, particularly in the persistence or intensification of drought, became accepted by the broader community (e.g., Giannini, Biasutti, & Verstraete, 2008; Rodriguez-Fonseca et al., 2015). Several studies, some based on models and some on observations, identified the Sahel as a “hot spot” in land-atmosphere interaction. Moreover, there is definitive observational evidence that land surface effects, including soil moisture and vegetation distribution and mineral dust, can affect the convective systems that bring rainfall to West Africa (see review in Nicholson, 2015).
Summary and Conclusions
The climate of the West Africa has many unusual features. These include the pronounced latitudinal orientation of climate zones and precipitation, the extraordinary degree of interannual variability, the high degree of interannual persistence in the Sahelian sector, and the extreme spatial coherence of rainfall anomalies. The region’s rainfall regime appears to have abruptly changed around 1968 and the factors leading to this change are not yet fully understood.
Thermal conditions in the region are generally a function of latitude. However, the annual cycle of precipitation and cloudiness and proximity to the coast can modulate this relationship. The diurnal variation of temperature exceeds the annual throughout sub-Saharan West Africa.
The common denominator through the region is the West African monsoon. North of roughly 10 °N, the seasonal cycle of rainfall is unimodal with the rainfall maximum occurring throughout the region in August. This peak is linked to the northward displacement of the monsoon during the boreal summer. Further south in the Soudano-Guinean and Guinean zones, rainfall is bimodal. Rainfall maxima occur in the boreal spring and the boreal autumn, coinciding with the passage of the monsoon through this region. Most of the rain is associated with a small number very intense storms.
Because of the dominance of the monsoon and the east-west transit of the major rain-bearing systems, the spatial coherence of annual rainfall anomalies is exceedingly strong. Year-to-year fluctuations tend to affect the region as a whole. The spatial coherence is strong enough that two spatial modes of variability prevail. In one, annual rainfall fluctuations are opposite in sign in the Sahelian zones and the Guinean zones. In the second mode, rainfall anomalies are mostly of the same sign through West Africa.
These patterns reflect the underlying regional forcing of interannual variability, that is, the forcing by major circulation features over West Africa. The most important factors are the intensity of the Tropical Easterly Jet, the latitude of the African Easterly Jet, and the intensity of the equatorial westerly flow. The remote factors regulating the regional circulation are less well understood. There is consensus that sea-surface temperatures play the most important role on interannual and longer-term decadal rainfall fluctuations, but there is much less agreement concerning the ocean regions of greatest influence. These appear to change over time and differ for interannual versus decadal time scales. All the world’s oceans as well as the Mediterranean Sea appear to have an impact on West Africa’s rainfall regime. The land surface, including the dust that it sends into the atmosphere, also appears to have an influence on the rainfall regime, possibly intensifying and prolonging drought and modulating individual weather systems.
A notable change affecting the entire region occurred around 1968. Since that time drier conditions have prevailed throughout West Africa. There has not been a complete recovery of the rainfall regime subsequent to the major droughts of the 1970s and 1980s. The reasons for this regime change are not yet fully understood.
Our understanding of West African meteorology has greatly changed in recent decades. The framework for understanding regional dynamics has shifted from changes in the Intertropical Convergence Zone to an emphasis on the components of the West African monsoon. This includes the upper-level circulation as well as the Saharan Heat Low. The importance of several rainfall-bearing systems has been well established, including Mesoscale Convective Systems and diagonal troughs. Future work is needed to understand the remote modulation of the monsoon and its characteristics on seasonal, interannual, and decadal time scales and how these changes translate into changes in the rain-bearing systems and their links to variability.
The author would like to thank Doug Klotter for his development of figures for the manuscript. This work was supported in part by a grant from the National Science Foundation, No. 1445605.
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(1.) Geographically speaking, the countries of Cameroon, Chad and the Sudan are not part of West Africa. However, they are characterized by the same climatic and meteorological features as West Africa, so that the term will be loosely applied to include these regions as well.