Quaternary Climate Variation in West Africa
Summary and Keywords
West Africa is among the most populated regions of the world, and it is predicted to continue to have one of the fastest growing populations in the first half of the 21st century. More than 35% of its GDP comes from agricultural production, and a large fraction of the population faces chronic hunger and malnutrition. Its dependence on rainfed agriculture is compounded by extreme variations in rainfall, including both droughts and floods, which appear to have become more frequent. As a result, it is considered a region highly vulnerable to future climate changes. At the same time, CMIP5 model projections for the next century show a large spread in precipitation estimates for West Africa, making it impossible to predict even the direction of future precipitation changes for this region. To improve predictions of future changes in the climate of West Africa, a better understanding of past changes, and their causes, is needed. Long climate and vegetation reconstructions, extending back to 5−8 Ma, demonstrate that changes in the climate of West Africa are paced by variations in the Earth’s orbit, and point to a direct influence of changes in low-latitude seasonal insolation on monsoon strength. However, the controls on West African precipitation reflect the influence of a complex set of forcing mechanisms, which can differ regionally in their importance, especially when insolation forcing is weak. During glacial intervals, when insolation changes are muted, millennial-scale dry events occur across North Africa in response to reorganizations of the Atlantic circulation associated with high-latitude climate changes. On centennial timescales, a similar response is evident, with cold conditions during the Little Ice Age associated with a weaker monsoon, and warm conditions during the Medieval Climate Anomaly associated with wetter conditions. Land surface properties play an important role in enhancing changes in the monsoon through positive feedback. In some cases, such as the mid-Holocene, the feedback led to abrupt changes in the monsoon, but the response is complex and spatially heterogeneous. Despite advances made in recent years, our understanding of West African monsoon variability remains limited by the dearth of continuous, high- resolution, and quantitative proxy reconstructions, particularly from terrestrial sites.
Modern Climate and the West African Monsoon System
The climate of West Africa is characterized by dramatic south to north decreases in mean annual rainfall, from a high of approximately 3,000 mm/yr along the southwestern coast of the Gulf of Guinea, to 600−1,000 mm/yr across the semi-arid Sahel, to 0 in most years in some parts of the Sahara (Nicholson, 2013; Figure 1). Decreases in rainfall amount are accompanied by reductions in the length of the rainy season as well, with dramatic effects on vegetation. For example, near the coast, many locations lack a clear dry season, whereas at the northern edge of the Sahel, the rainy season may be as short as 1 to 2 months. Most of the humid, coastal zone is occupied by tropical rainforest, which transitions first to woodland and then to grasslands in the Sahel and desert north of 18°N. Mean annual temperatures typically exceed 18°C, with temperatures near the equator reaching averages as high as 26°C. However, the highest temperatures are reached in the Sahara, where they can exceed 40°C. Despite these strong north-to-south climate gradients, the climate (and vegetation) is remarkably homogeneous longitudinally (Nicholson, 2000).
The seasonal climate of West Africa is dominated by the West African monsoon system (Nicholson, 2017). It follows the classic characteristics of a monsoon system, with a seasonal reversal of winds (Trenberth, Stepaniak, & Caron, 2000; Webster et al., 1998) generated by the thermal contrast between the Atlantic cold tongue and the Saharan heat low (Figure 1). In boreal summer, the low intensifies and draws moist, southwesterly winds inland. In winter, the low weakens and the winds reverse, and dry northeasterlies are drawn down to the Guinea coast. The traditional view of the West African monsoon is that rainfall is directly associated with the seasonal passage of the intertropical convergence zone (ITCZ; Griffiths, 1972). However, a more nuanced view of the West African monsoon has emerged (e.g., Gu & Adler, 2004; Nicholson, 2009; Nicholson & Grist, 2003; Thorncroft, Nguyen, Zhang, & Peyrille, 2011). In part, this is due to the observation that West African rainfall has little to do with the ITCZ itself, which typically lies about 1,000 km north of the rainfall maximum (Nicholson, 2009). Instead, rainfall is mostly associated with the upper-level tropical easterly jet (TEJ), the mid-level African easterly jet (AEJ), the monsoon surface westerlies, and associated regions of convection and shear (Nicholson, 2017; Nicholson & Grist, 2003).
On interannual timescales, a precipitation dipole can develop, with anomalies of opposite sign between continental and coastal West Africa (Nicholson, 2000); this phenomenon appears to be related to changes in the seasonal position of the AEJ, with a more northerly (southerly) mean position leading to wetter (drier) conditions inland and drier (wetter) conditions near the coast (Nicholson, 2013; Nicholson & Grist, 2001). In contrast, precipitation anomalies of the same sign across West Africa tend to be associated with changes in the intensity of the tropical rain belt produced by changes in the strength of the AEJ or the TEJ (Nicholson, 2013; Nicholson & Grist, 2001).
The most dramatic change in the instrumental record of West Africa is the well-documented 1970s drying, during which rainfall decreased abruptly by approximately 20% to 40% over just a few years (Dai et al., 2004; Nicholson, 2005; Figure 2). The instrumental record highlights one of the most climatically significant features of the West African monsoon system: the multidecadal persistence of precipitation anomalies. In every year between 1950 and 1967, Sahel rainfall exceeded the long-term mean. After declining abruptly in 1968, rainfall remained below the mean from 1969 to 1997 (Nicholson, 2013). A number of studies have attributed the 1970s drought and its persistence to changes in Atlantic sea surface temperature (SST) gradients (Hastenrath, 1990; Knight, Folland, & Scaife, 2006; Lamb, 1978a, 1978b; Mohino, Janicot, & Bader, 2011; Zhang & Delworth, 2006), possibly associated with a low-frequency mode of Atlantic variability termed the Atlantic Multidecadal Oscillation (AMO; Schlesinger & Ramankutty, 1994). Although some researchers have questioned the importance of Atlantic SSTs in the 1970s Sahel drought (Biasutti, Held, Sobel, & Giannini, 2008; Giannini, Biasutti, & Verstraete, 2008), climate models and observations clearly show strong correlations between an AMO-like Atlantic SST pattern and West African rainfall (Mohino et al., 2011; Rodríguez‐Fonseca et al., 2011; Zhang & Delworth, 2006; Figure 2). Furthermore, an annually resolved lake record of West African rainfall shows a continuous, multidecadal mode of variability spanning the last 2.5 millennia that is coherent and in phase with paleo-AMO reconstructions (Shanahan et al., 2009). These data support the role of tropical Atlantic SSTs in driving multidecadal variability in the West African monsoon, but they do not preclude the possibility that recent changes are significantly influenced by other factors as well (e.g., aerosols, greenhouse gas).
The first studies of the influence of the AMO on Sahel rainfall suggested that warmer North Atlantic SSTs during the positive phase of the AMO resulted in increased interhemispheric and meridional SSTs and pressure gradients, a northward shift of the ITCZ, and increased rainfall over the Sahel (Hastenrath, 1990; Lamb, 1978a, 1978b). More recent work has highlighted the role of the position of the AEJ and changes in moisture flux associated with low-level westerlies (Martin & Thorncroft, 2014; Nicholson & Grist, 2001; Wang & Gillies, 2011), rather than the ITCZ as the cause of rainfall anomalies in association with the AMO. Other studies emphasized the importance of the AMO for the strength of the Saharan heat low and its impact on moisture flux to the monsoon during the spring and summer through both northeasterly moisture transport from the Mediterranean and low-level southerly winds from the Gulf of Guinea (Fontaine et al., 2010; Martin & Thorncroft, 2014; Peyrillé, Lafore, & Redelsperger, 2007)
Some researchers have also pointed out the potential influence of the tropical Pacific on interannual variability in the West African monsoon (Janicot, Moron, & Fontaine, 1996; Rowell, 2001; Semazzi, Mehta, & Sud, 1988). However, understanding the causes of these remote teleconnections is complicated by the spatial and temporal variability in the West African monsoon response (Ward, 1998). Although there are statistically significant relationships between the El Niño and Southern Oscillation (ENSO) and Sahel rainfall, not all ENSO events produce a signal in West African rainfall, because the connection depends on the background climate state and the “flavor” of ENSO (Joly, Voldoire, Douville, Terray, & Royer, 2007). Particularly important is the onset of the ENSO event, because the teleconnections appear to be most important during the initiation of El Niños or the decay of persistent La Niña events (Joly et al., 2007; Nicholson, 2013). Although the dynamical causes of the ENSO teleconnection are still poorly understood, some modeling experiments suggest that the teleconnection does not involve the Indian Ocean directly (Joly et al., 2007) but is instead related to changes in the Walker circulation and a reduction in the strength of the TEJ (Nicholson & Grist, 2001; Rowell, 2001).
Land-surface feedback plays an important role in the monsoon as well. Originally the feedback was proposed to occur between desertification of North Africa and West African drought (Charney, 1975), but a number of modeling studies have since highlighted the role of vegetation and land-surface properties as positive feedback in monsoon variability (Xue & Shukla, 1993; Zeng, Neelin, Lau, & Tucker, 1999). Land-surface properties influence monsoon rainfall by affecting the fluxes of sensible and latent heat, which in turn influence the properties of the atmospheric boundary layer and the development of convection (Adler, Kalthoff, & Gantner, 2011; Gantner & Kalthoff, 2010; Mohr, David Baker, Tao, & Famiglietti, 2003). These act as positive feedback on rainfall, with dry (wet) conditions leading to reduced (increased) rainfall. The impact of land-surface feedback (especially soil moisture) appears to be particularly strong in semi-arid North Africa (Taylor, de Jeu, Guichard, Harris, & Dorigo, 2012).
The Neogene-Quaternary Aridification of West Africa
Reconstructions from West Africa indicate that conditions were much wetter during the Miocene, with widespread expansion of grasslands and forests (Bakker & Mercer, 1986; Dutton & Barron, 1997; Hoetzel, Dupont, Schefuß, Rommerskirchen, & Wefer, 2013; Jacobs, 2004; Micheels, Bruch, Uhl, Utescher, & Mosbrugger, 2007; Pickford, 2000; Swezey, 2009; Uno, Polissar, Jackson, & deMenocal, 2016) . The transition from the late Neogene (< 7 Ma) to the Quaternary (< 2.8 Ma) was characterized by a shift to progressively drier conditions, resulting in the eventual development of the modern-day Sahara Desert (Bonnefille, 2010; Moussa et al., 2016; Novello et al., 2015; Uno et al., 2016). Terrestrial sedimentological and microbiological evidence suggests that this transition began sometime around 7.5 Ma, when forests in west-central Africa gave way first to heterogeneous forest-grassland systems and then to more homogeneous mixed grasslands after 4.5 Ma (Moussa et al., 2016; Novello et al., 2015, 2017). A late Neogene expansion of grasslands across West Africa is also suggested by increased grass cuticle remains and pollen in sediment cores from the Niger Delta starting at around 7 Ma (Bonnefille, 2010; Morley, 2000; Morley & Richards, 1993). Oxygen isotope data on fossil otoliths show a progressive trend toward heavier values between 7.5 and 3 Ma, consistent with regional aridification (Otero et al., 2011). Recent climate model simulations point to paleogeography as a critical control on the stronger monsoon during the Miocene (Bradshaw et al., 2012; Micheels, Bruch, Uhl, Utescher, & Mosbrugger, 2007). The late Miocene shift to more arid conditions is more controversial, but it appears also to be related to paleogeography—resulting from circulation changes associated with the shrinkage of the Tethys Sea (Zhang et al., 2014).
A number of studies have suggested that a major change in aridity began at ca. 3 Ma, culminating in the modern West African climate system. At Lake Chad, a significant aridity event is evident in diatom and phytolith data at 3.6−2.7 Ma (Moussa et al., 2016; Novello et al., 2015, 2017). Vegetation and paleodischarge reconstructions also indicate that major increases in regional aridity occurred at ca. 3.5−3.2 Ma and 2.6 Ma (Bonnefille, 2010; Dupont & Leroy, 1995; Leroy & Dupont, 1994). Using a series of marine dust records collected from the North African margins, it was suggested that this transition occurred by a series of “jumps” in aridity starting at 2.8 ± 0.2 Ma (deMenocal, 1995, 2004). The timing of this shift corresponds closely in timing to global climate reorganizations associated with high-latitude cooling and the formation of large Northern Hemisphere ice sheets at ca. 3.5−2.5 Ma (Mudelsee & Raymo, 2005), implicating high-latitude climate changes as an important control on the drying of West Africa at the beginning of the Quaternary.
However, a revised statistical analysis of African dust records has been used to suggest more complex and spatially heterogeneous transitions to drier conditions at this time (Trauth, Larrasoaña, & Mudelsee, 2009), which may depend more on conditions at the low latitudes (Caley et al., 2011; Lourens et al., 1996). Breakpoint analysis (Trauth et al., 2009) suggests that the most significant change in West African dust production actually began at ca. 1.62 ± 0.82 Ma (ODP659, 18.08°N, 21.03°W; Tiedemann, Sarnthein, & Shackleton, 1994), over 1 Ma after the onset of Northern Hemisphere glaciation and substantially later than previous estimates (deMenocal, 1991, 1995, 2004). Similar analyses on marine dust records from the Mediterranean (ODP 967, 34.07°N, 32.72°E; Larrasoaña, Roberts, Rohling, Winklhofer, & Wehausen, 2003) and the Arabian Sea (ODP 721/722, 16.68°N, 59.86°E; deMenocal, 1991, 1995, 2004) suggests the main increase in dust occurred at 1.44 ± 0.2 Ma (ODP 967) and 1.86 ± 0.44 Ma (ODP721/722), respectively (Trauth et al., 2009). One proposed explanation for the lag in the development of aridity behind ice volume may be that it reflects a threshold response in the monsoon to gradually increasing ice volume (Trauth et al., 2009).
Orbital Variability During the Plio-Pleistocene
The long-term trend toward more arid conditions was characterized by higher-frequency switches between wet and dry climate states, driven by orbital variations in solar insolation. Discontinuous terrestrial records from the Lake Chad Basin provide evidence for recurrent desertlike conditions alternating with forests and grasslands since at least 6 Ma (southern basin; Moussa et al., 2016) to 8 Ma BP (northern basin; Lebatard et al., 2010; Schuster et al., 2009). Records of sapropel deposition in the eastern Mediterranean also show that periodic intervals of increased monsoon rainfall occurred back as far as 8 Ma (Figure 3; Larrasoaña, Roberts, & Rohling, 2013). The timing of sapropel deposition suggests that these wet intervals occurred during minimum precession, when summer insolation is at a maximum in the Northern Hemisphere (Larrasoaña et al., 2013; Lourens et al., 1996; Rossignol-Strick, 1983). This observation is consistent with the prediction that during times of higher Northern Hemisphere summer insolation, enhanced warming of the continents results in increased ocean–land thermal and pressure gradients and an intensified and more northerly monsoon rain belt (Braconnot, Marzin, Grégoire, Mosquet, & Marti, 2008; de Noblet, Braconnot, Joussaume, & Masson, 1996; Kutzbach, 1981; Montoya, von Storch, & Crowley, 2000; Prell & Kutzbach, 1987). Subsequent work noted that not all precessional minima contain sapropels, but instead form clusters at the 100- and 400-kyr frequencies, providing support for the impact of eccentricity modulation of precession (Hilgen, Lourens, Berger, & Loutre, 1993; Lourens et al., 1996; Lourens, Wehausen, & Brumsack, 2001; Van Vugt, Langereis, & Hilgen, 2001). Other studies have argued for an additional, weaker 40-kyr obliquity (tilt) signal in sapropel thickness (Lourens et al., 1996).
Continuous records spanning the Plio-Pleistocene that can be used to test the influence of orbital variations on the monsoon are rare. The longest continuous record is a reconstruction of dust from ODP 659 on the Cape Verde Plateau which spans the last 5 Myr (Figure 3; Tiedemann et al., 1994). Spectral analysis of the dust record from ODP 659, as well as records from both the Mediterranean (ODP 967) and the Arabian Sea (ODP 721/722), show evidence for continuous precession (19−23 kyr) and obliquity (41 kyr) variations throughout the records, suggesting that the orbitally forced changes in insolation drove changes in monsoon strength across North Africa throughout at least the last 5 Ma, consistent with the sapropel record (deMenocal, 2004; Larrasoaña et al., 2003; Tiedemann et al., 1994). At ODP 659, power in the 100-kyr (eccentricity) band is also evident throughout, but only appears in the Mediterranean and Arabian Sea dust records after ca. 1 Ma (Trauth et al., 2009). The evidence for all three modes of orbital variations extending back to 5 Ma in ODP 659 is significant in that it suggests that their influence on the West African monsoon does not entirely depend on high-latitude forcing, and points instead to low-latitude changes as the primary causes of orbital variability in this system. However, changes in the strength of the orbital variations do indicate that other factors influence the magnitude of the response to orbital variations. For example, variations at both the 41- and 100-kyr frequencies intensify starting at ca. 3.0−2.8 Ma and 3.2–3.0 Ma, broadly consistent in timing with the formation of Northern Hemisphere ice sheets at 2.8 Ma (deMenocal, 2004; Trauth et al., 2009). A second increase in the 100-kyr frequency occurs at 1.6 Ma. At the same time, the amplitude of the precession signal decreases in two steps at 1.6 Ma and starting at around the mid-Pleistocene transition (0.8 Ma; Trauth et al., 2009).
The evidence for precession in West African monsoon records is consistent with model simulations showing a direct low-latitude influence of precession on the monsoon (Braconnot et al., 2008; de Noblet et al., 1996; Kutzbach, 1981; Montoya et al., 2000; Prell & Kutzbach, 1987). When precession is at a minimum (maximum), boreal summer insolation maximized (minimized), driving a stronger (weaker) summer monsoon. This is, in part, a function of the relative differences in the heat capacity of the North African continent and the oceans, which allows North Africa to heat up more rapidly than the oceans, producing meridional pressure and temperature gradients that drive the summer monsoon winds. Increased boreal summer insolation intensifies this effect, driving a stronger monsoon. First demonstrated using general circulation models in the early 1980s (Kutzbach, 1981), it was confirmed in numerous climate models and was termed the Orbital Monsoon Hypothesis (Ruddiman, 2001). The paleoclimate evidence for 100- and 400-kyr frequencies in both sapropels and marine dust records provides strong support for eccentricity modulation of the precessional forcing of monsoon strength.
In contrast to precession and eccentricity, the evidence for a strong obliquity signal in proxy monsoon records from North Africa is at odds with direct insolation forcing; obliquity has its greatest influence at high latitudes and is relatively small in the tropics. For example, calculated changes in tropical insolation due to changes in obliquity are an order of magnitude smaller than those resulting from precession (Laskar et al., 2004). One explanation for this apparent discrepancy is remote forcing of the monsoon from the high latitudes, where the influence of obliquity is much larger. However, the evidence for a significant obliquity signal in monsoon records prior to the formation of Northern Hemisphere ice sheets precludes changes in ice volume as a driver for the teleconnections (deMenocal, 1995, 2004; Larrasoaña et al., 2013; Trauth et al., 2009). Instead, experiments with a climate model of intermediate complexity (EC-Bilt) suggest that remote forcing of the North African monsoon can occur as a result of high-latitude insolation-driven changes in convergence over southern Asia, which generates stronger southwesterly winds and greater moisture advection into North Africa during summer (Tuenter, Weber, Hilgen, & Lourens, 2003). An alternative hypothesis is that the obliquity signal is a direct response to local or interhemispheric meridional insolation gradients. Rossignol-Strick (1983) proposed a monsoon index (M) to explain the obliquity signal in Mediterranean sapropels that was based on the summer insolation difference between the equator and 21°N, reasoning that this was crucial for generating meridional pressure gradients associated with the monsoon. More recently, single forcing experiments conducted using the fully coupled EC-Earth model were used to suggest that the obliquity signal in North African monsoon records is the result of the summer intertropical insolation gradient (Bosmans, Drijfhout, Tuenter, Hilgen, & Lourens, 2015; Bosmans, Hilgen, Tuenter, & Lourens, 2015). These simulations showed that obliquity-driven changes in this gradient strengthen the Hadley circulation, enhance cross-equatorial moisture transport during the summer and result in a stronger and northward shifted summer monsoon system, independent of changes at high latitudes. Support for this hypothesis comes from the small time lags (e.g., 1.5–2.6 kyr) between changes in obliquity and monsoon reconstructions over the last 160 kyr (Caley et al., 2011; Lourens et al., 1996). However, it is also notable that, as with eccentricity, a strong increase in the amplitude of the 41-kyr variability in ODP 659 occurs at ca. 3 Ma, broadly coincident with the formation of Northern Hemisphere ice sheets, implying a contribution from the high latitudes. This may suggest that the ice sheets play a role in enhancing this signal through their influence on the ITCZ (Chiang & Friedman, 2012).
The Last Glacial Cycle
The influence of orbital variations on the strength of the monsoon is particularly evident in proxy climate records spanning the last glacial cycle (Figure 4). In both marine and terrestrial records, glacial intervals (e.g., Marine Isotope Stages 2 and 6) are associated with more arid conditions, an equatorward expansion of arid and semi-arid landscapes, and the replacement of forests with grasslands. In contrast, interglacials are characterized by the northward expansion of vegetation zones, including the replacement of grasslands by forests and the establishment of new grasslands into deserts, the development of extensive lakes, wetlands, and fluvial systems, and even the migration and expansion of humans and mammals across the Sahara (Braconnot, Joussaume, De Noblet, & Ramstein, 2000; Prentice & Jolly, 2000). These changes reflect shifts in both the intensity of monsoon rainfall and shifts in its northward extent, which is thought to have reached at least 25°N during interglacials (Dupont, 2011; Larrasoaña et al., 2013; Lézine, Hély, Grenier, Braconnot, & Krinner, 2011). Hydroclimatic variations in continuous marine records from offshore West Africa (Tjallingii et al., 2008; Weldeab, Lea, Schneider, & Andersen, 2007) are also consistent with records of global methane production from trapped gases in polar ice cores, which are interpreted as reflecting changes in the Northern Hemisphere tropical hydrological cycle (Schmidt, Shindell, & Harder, 2004), indicating that the changes recorded in West African sites reflect variations in the climate of the global tropics (Figure 4).
These records highlight the role of both precession and eccentricity in monsoon variability through the last glacial cycle. Most marine hydroclimate records from West Africa show a strong 21-kyr periodicity, with peaks in wetness occurring in phase with maxima in precession and Northern Hemisphere summer insolation. However, the magnitude of these wet intervals declines with decreasing eccentricity, with the smallest-amplitude wet−dry cycles when eccentricity is at a minimum and the largest-amplitude changes when eccentricity is at a maximum (Grant et al., 2016; Larrasoaña et al., 2013; Revel et al., 2010; Tjallingii et al., 2008; Weldeab et al., 2007). Thus, the wettest conditions occur during the interglacials (i.e., Marine Isotope Stages 1, 5), though intervening intervals of wetter conditions occur in association with maxima in precession at ca. 105, 80, and 55 kyr BP.
Wet conditions associated with local summer insolation maxima are also indicated by sapropel formation in the eastern Mediterranean, with S1 forming during the Holocene African Humid Period (AHP; deMenocal et al., 2000), and S3−S5 occurring between 80 and 130 kyr BP in conjunction with maxima in precession (Grant et al., 2016). However, during the intervening maxima in precession ca. 30 and 55 kyr BP, when changes in insolation were comparatively weaker, sapropels are absent in records from the eastern Mediterranean, indicating that changes in river outflow to the Mediterranean were much lower and that the monsoon did not supply substantial moisture to northernmost Africa (Rohling, Marino, & Grant, 2015). Compilations of dated lake, tufa, and fluvial deposits also show variations that are consistent with this model of an eccentricity-modulated control on precession, though they lack the dating control to unambiguously attribute wet phases to individual insolation peaks. However, careful studies of these deposits suggest that they can be grouped into stratigraphic units separated by deflational surfaces (Armitage et al., 2007; Drake et al., 2008; Kleindienst et al., 2008; Smith et al., 2004; Smith, Hawkins, Asmerom, Polyak, & Giegengack, 2007; Szabo, Haynes, & Maxwell, 1995), consistent with a series of relatively short wet intervals superimposed on hyperarid conditions.
The close correspondence between eccentricity-modulated precession and marine proxy records suggests that exceptionally high summer insolation during the peak of the last interglacial (ca. 125 kyr BP) should have produced a monsoon that was substantially stronger than that of the Holocene (Figure 4). Though terrestrial hydrologic data for the Eemian is much scarcer than for the Holocene AHP, quantitative estimates of monsoon runoff in the eastern Mediterranean (Osborne et al., 2008; Rohling et al., 2004) and North African precipitation (Bergner, Trauth, & Bookhagen, 2003; Hoelzmann, Kruse, & Rottinger, 2000; Kieniewicz & Smith, 2009; Kowalski et al., 1989) are consistent with this hypothesis. For example, mean annual precipitation is estimated to have been around 100 mm/yr in the northeastern Sahara (23−29°N) and 450 mm/yr in the Sahel (18°N) during the Holocene (Kuper & Kröpelin, 2006). In contrast, Eemian precipitation is estimated to have been as high as 410−670 mm/yr in the northeast Sahara and 500−900 mm/yr in the Sahel (Hoelzmann et al., 2000).
The weaker insolation forcing during Marine Isotope Stage 3 (MIS 3; ca. 30−80 kyr BP) is consistent with proxy evidence indicating smaller and more spatially variable wet intervals associated with maximum precession at this time (Grant et al., 2016). In most of the marine records from West Africa, MIS 3 is the weakest of the insolation-driven wet stages of the past 150 kyr BP (Grant et al., 2016; Larrasoaña et al., 2013; Revel et al., 2010; Tjallingii et al., 2008; Weldeab et al., 2007). And, as indicated previously, it is the only interval when precession maxima are not associated with Mediterranean sapropels (Rohling et al., 2015). There is also evidence that the precipitation response to insolation changes was spatially heterogeneous at this time. For example, records from northernmost Africa provide an inconsistent view of hydroclimate changes at this time, with some suggesting a slight return to wetter conditions (Revel et al., 2010) and others showing an uninterrupted period of aridity (Grant et al., 2016). The reasons for these inconsistencies are unclear but could reflect either a smaller northward penetration of the monsoon during MIS 3 or changes in the relative influence of the Indian Ocean, compared with insolation-driven changes in the monsoon (Grant et al., 2016). However, these discrepancies are not limited to the northernmost Sahara. A leaf wax isotope reconstruction from ODP site 659 (18.08°N, 21.33°W) is dominated by insolation forcing when variations in insolation are large, but during MIS 3, when insolation forcing is weak, the record appears to be dominated by the influence of SSTs (Kuechler, Schefuß, Beckmann, Dupont, & Wefer, 2013). At a location farther south along the West African margin, a leaf wax δ13C record from GeoB9528-3 (09.166°N, 17.66°) shows vegetation changes over the last ~ 200 kyr that are dominated by changes in the Atlantic overturning circulation rather than insolation (Castañeda et al., 2009). One possible explanation for these differences is that during intervals when weak insolation forcing does not dominate the monsoon system (such as MIS 3), multiple factors play a role in influencing the monsoon, leading to regional differences in the character and magnitude of the hydroclimate response (Singarayer & Burrough, 2015).
There is evidence that these orbitally forced changes in the climate of West Africa had significant impacts on the development, evolution, and migration of early hominins (Larrasoaña, 2012; Larrasoaña et al., 2013). Many of the mid to late Pleistocene paleolacustrine, tufa, and fluvial deposits of the Sahara are associated with Early and Middle Stone Age archaeological deposits (Haynes, Maxwell, El Hawary, Nicoll, & Stokes, 1997; Hill, 2001; Kleindienst et al., 2008; Mandel & Simmons, 2001; McHugh, Breed, Schabers, McCauley, & Szabo, 1988; Smith et al., 2004, 2007) suggesting that these pluvial intervals allowed early humans to expand across the modern-day North African deserts. However, marine records indicate that the wet events were relatively short-lived, with peak humidity in the Sahara lasting as little as ~ 8 ka (Larrasoaña et al., 2013). This would have fragmented habitats and populations, forcing them to retreat to ecological refugia, increasing resource competition, and enhancing reproductive isolation, processes that have been suggested as drivers for the development of modern Homo sapiens (Mellars, 2006).
Late Pleistocene wet intervals, however brief, may have also provided gateways to allow humans to migrate to Eurasia (Castañeda et al., 2009; Larrasoaña et al., 2013; Osborne et al., 2008). For example, wet conditions during the Eemian (130−115 kyr BP) coincide with a major period of hominin dispersal out of Africa at ca. 125 kyr BP (Drake et al., 2008; Grün et al., 2005; Lahr & Foley, 1998; Osborne et al., 2008; Stringer, 2002). This may have been facilitated by the development of large active rivers across North Africa that provided a continuous route to the Mediterranean and Europe (Armitage et al., 2007; Drake et al., 2008; Osborne et al., 2008). The second major migration of hominins out of Africa occurred between ~ 40 and 60 kyr BP (Mellars, 2006). Although there is evidence that this was a time of relatively wetter conditions at some locations, conditions at this time were significantly drier than in the Eemian. One explanation for this is that resource scarcity and competition associated with the monsoon collapse after 80 kyr BP drove the significant technological advances needed to successfully disperse from Africa into Europe (Mellars, 2006), particularly in the Sahara (Larrasoaña et al., 2013).
In addition to orbital-scale climate variations, most records from West Africa also show significant variability on millennial timescales that are correlated with changes in North Atlantic climate (Figure 5). These records display a strong signature of North Atlantic Heinrich stadials (HSs), prominent millennial-scale North Atlantic cooling events that terminate in the deposition of ice rafted debris (IRD) layers (Bond et al., 1999; Broecker, Bond, Klas, Clark, & McManus, 1992; Heinrich, 1988). Numerous studies have shown that the impacts of these events are global, including substantial drying throughout much of the Northern Hemisphere tropics (Clement & Peterson, 2008; Hemming, 2004). They are hypothesized to be associated with reductions in the strength of the Atlantic meridional overturning circulation (AMOC; Kageyama, Paul, Roche, & Van Meerbeeck, 2010), which plays a crucial role in the distribution of oceanic heat and salinity (Houghton, 1996; Trenberth & Solomon, 1994). In tropical and subtropical Africa, the spatial signature of these events is controversial, with some proxy reconstructions suggesting that HSs were associated with an interhemispherically symmetric contraction of the tropical rain belt (Collins et al., 2011; Stager, Ryves, Chase, & Pausata, 2011) rather than an asymmetric response associated with a southward shift in the ITCZ, as suggested by models (Kageyama et al., 2010).
Most records from West Africa suggest drying during HSs, although the magnitude varies as a function of proxy, location, and event (Figure 5). Records spanning the last glacial cycle with sufficient resolution to capture multiple HSs are currently limited to marine sediments. Of these, the clearest signature of drying comes from marine sediment reconstructions off the Atlantic coast of West Africa (Bradtmiller et al., 2016; Collins et al., 2011, 2013; Itambi, Von Dobeneck, & Adegbie, 2010; Itambi, Von Dobeneck, Mulitza, Bickert, & Heslop, 2009; Jullien et al., 2007; Mulitza et al., 2008; Niedermeyer et al., 2009, 2010; Tierney, Pausata, & deMenocal, 2017). A compilation of these dust records spanning the past 60 kyr BP shows a coherent meridional pattern, with a small HS dust signal in the north, off the coast of Mauritania (21°N, GeoB7920-2; Tjallingii et al., 2008), large HS dust events at 15°N (GeoB9508-5; Mulitza et al., 2008), a reduced signature at 12°N (GeoB9526-5; Zarriess et al., 2011) and no dust at 9°N (GeoB9528-3; Castañeda et al., 2009). Collins and coworkers (2013) interpreted these variations as reflecting changes in the position of the Sahara-Sahel boundary, which shifts as far south as 12°N during HS 1, 2, 4, and 5 (and slightly less during HS 3). The longest of these records (GeoB7920-2) shows dry intervals correlative with the largest Greenland stadials back to 120 kyr (Tjallingii et al., 2008). Confirmation that these dusty events reflect intervals of exceptional drying (as opposed to changes in wind strength) comes from a 40-kyr length record of leaf wax isotopes, which shows more positive δDwax and reduced runoff during HS 1−4 (Niedermeyer et al., 2010). Although the area is less well studied, existing data suggests that HSs produced dry conditions in humid tropical West Africa as well (Lézine, Assi‐Kaudjhis, Roche, Vincens, & Achoundong, 2013; Lezine & Cazet, 2005; Shanahan et al., 2015; Weldeab, 2012; Weldeab et al., 2007).
Glacial millennial-scale warmings (i.e., Dansgaard-Oeschger [D-O], interstadials) also occur in conjunction with changes in Atlantic circulation, though there is limited proxy evidence for these events in records from West Africa. Most terrestrial systems lack the resolution, length, or age control to unambiguously record these events. And although numerous marine records positioned along the West African coast provide clear signatures of HSs through MIS 2 and 3, they are of insufficient resolution to clearly record the shorter and higher frequency D-O events. The best evidence for D-O events comes from δ18O variations in Globigerinoides ruber from marine core MD2707 in the eastern Gulf of Guinea (Weldeab, 2012). This record shows reduced discharge during each Greenland stadial over the past 75 kyr BP. Furthermore, single foraminifera δ18O variations indicate that the seasonal contrast in salinity during D-O stadials was significantly reduced, consistent with a reduction in monsoon strength during these events (Weldeab, 2012).
Numerous studies have attempted to better understand the underlying dynamics behind the millennial-scale variability by imposing large freshwater fluxes to the North Atlantic in general circulation models to perturb the strength of the AMOC (Kageyama et al., 2010). These freshwater hosing experiments have been applied using a variety of models, boundary conditions, and mean states (e.g., preindustrial versus LGM) and produce broadly similar responses: weakening or collapse of the AMOC induces cooling over the North Atlantic by reducing northward heat transport in the Atlantic, leading to a southward shift of the ITCZ and drying over North Africa.
The mechanism controlling the linkages between North Atlantic temperature and the West African precipitation likely resides in the atmosphere, rather than the ocean (Chiang & Friedman, 2012). In climate model experiments forced with freshwater inputs to the North Atlantic, high-latitude cooling is advected to the northern extratropics by the westerlies and propogated to the tropics through the wind-evaporation-sea surface temperature (WES) feedback (Mahajan, Saravanan, & Chang, 2009). As this cooling expands over North Africa, cooler temperatures reduce humidity and moist static energy and increase subsidence over the Sahara, with feedbacks acting to enhance cooling over North Africa. A reduction in low-level moisture increases moist stability of the atmosphere and weakens convection, particularly over the Sahel (Liu, Chiang, Chou, & Patricola, 2014). Anomalous cooling over North Africa also reduces the meridional pressure gradient, weakening the humid southwesterlies in summer and strengthening the dry northeasterlies in winter. At the same time, the reduced temperature gradient weakens the TEJ and causes the AEJ to strengthen and shift southward. These processes act both to shift the tropical rain belt equatorward and to reduce rainfall by transporting moisture away from West Africa (Chiang & Friedman, 2012).
The best characterized of the millennial-scale events is HS 1 (Figure 6), which occurred at ca. 18−15 kyr BP in conjunction with a well-documented collapse of the AMOC (McManus, Francois, Gherardi, Keigwin, & Brown-Leger, 2004). Most of the records from West Africa are derived from marine sediments because of the difficulties in obtaining continuous terrestrial records spanning the deglaciation, when conditions were significantly drier than today (Shanahan et al., 2006); the exceptions are the records from Lake Bosumtwi (Shanahan et al., 2015, 2016), Barombi Mbo (Lebamba, Vincens, & Maley, 2012; Maley & Brenac, 1998), and Bambili (Lézine, Assi‐Kaudjhis, et al., 2013), which are located along the humid Gulf of Guinea coast (Figure 6). The majority of existing records show drier conditions during HS 1, though the magnitude of the signal varies, with the strongest drying to the south of ca. 18°N and west of the prime meridian. North and east of this, the signal of HS 1 is much smaller or undetectable in many of the proxy records.
A comparison between the proxy reconstructions and freshwater hosing experiments (Kageyama et al., 2013) shows broad similarities and some notable differences (Figure 6). There is good agreement between the pattern of a declining HS 1 signal in the proxy records and the northward decrease in the model precipitation anomaly along the eastern Atlantic coast of West Africa. And the strong HS 1 drying evident in the Lake Bosumtwi record is qualitatively consistent with the large model simulated precipitation deficits in the western portion of the Gulf of Guinea coast. Similarly, the evidence for increased precipitation and steep gradients in rainfall anomalies in the far eastern Gulf of Guinea are broadly consistent with the small or absent HS 1 signal in proxy records, though one might expect to still see negative anomalies, particularly at sites like Barombi Mbo (Lebamba et al., 2012; Maley & Brenac, 1998). Differences may reflect the limitations of the proxy records. For example, at GeoB9528-3 (Castañeda et al., 2009), the proxy HS 1 signal is small but the model suggests some of the largest precipitation anomalies along the adjacent coast. Resolving these spatial patterns will require a greater density of records, particularly within the Gulf of Guinea and on the continent.
Though there is abundant evidence for wetter conditions in North Africa during the Holocene consistent with the Orbital Monsoon Hypothesis (Figure 7), there also remain a number of unanswered questions about the nature and drivers of Holocene variations in the West African monsoon. First, existing paleolake, pollen, and archaeological evidence points to a dramatic northward shift of the summer monsoon extent during the Holocene AHP (15−5 kyr BP), with substantially increased rainfall amount and a northward expansion of the monsoon influence to at least 30°N (Lézine et al., 2011; Shanahan et al., 2015; Tierney et al., 2017). And although models are able to simulate an intensification of the monsoon in response to increases in boreal summer insolation, they are unable to produce the northward extent of monsoon rainfall during the AHP indicated by proxy reconstructions (Harrison et al., 2014; Krinner et al., 2012). A number of studies have investigated the role of positive feedback involving SST, land surface, vegetation, and most recently dust in amplifying the monsoon response to local insolation changes (Braconnot, Joussaume, Marti, & De Noblet, 1999; Claussen & Gayler, 1997; Ganopolski, Kubatzki, Claussen, Brovkin, & Petoukhov, 1998; Kutzbach, Bonan, Foley, & Harrison, 1996; Kutzbach & Liu, 1997). These experiments demonstrate that positive feedback can produce a substantially larger and northward-shifted monsoon system under mid-Holocene insolation, highlighting the importance of feedback mechanisms in estimating the monsoon response to external forcing. However, to date, even models incorporating soil and vegetation feedback have had difficulty reproducing the full northward monsoon extent during the AHP (Krinner et al., 2012; Levis, Bonan, & Bonfils, 2004).
A second unresolved issue involves the transient nature of the monsoon response to insolation forcing over the Holocene. Orbital precession changes gradually, and models predict precipitation should change gradually in response to this forcing (Kutzbach & Liu, 1997; Liu et al., 2007; Liu, Wang, Gallimore, Notaro, & Prentice, 2006). However, proxy records of dust generation from coastal northwestern Africa suggest that transitions in and out of the AHP were abrupt, changing in as little as a few centuries and suggesting a nonlinear monsoon response to insolation forcing (deMenocal et al., 2000). Early modeling efforts pointed to vegetation feedback as a potential driver of nonlinear monsoon response to insolation forcing (Claussen et al., 1999) and subsequent experiments showed that positive vegetation-climate feedback could lead to dynamically unstable systems that are susceptible to rapid transitions between “green” and “brown” states (Renssen, Brovkin, Fichefet, & Goosse, 2006). Other model experiments have suggested that the abrupt changes in dust may reflect a response to vegetation thresholds, rather than a nonlinear monsoon response (Liu et al., 2006, 2007). More recently, a growing number of studies have suggested that the monsoon response to Holocene insolation changes was either gradual or spatially heterogeneous (Kroepelin et al., 2008; Shanahan et al., 2015; Tierney et al., 2017), strongly dependent upon local hydroclimate (Vincens, Buchet, & Servant, 2010), and with locally abrupt changes occurring later toward the south along the Gulf of Guinea coast (Shanahan et al., 2015; Figure 7). The latter study posits that sites at the northern limit of the peak Holocene monsoon change abruptly as the monsoon system contracts equatorward with waning insolation. Near the coast, the peak northward extension of the monsoon is unimportant, and sites only record the gradual weakening of monsoon strength in direct response to insolation forcing. Between the two, the monsoon response is complex, and reflects a combination of gradual changes in monsoon strength and the southward contraction of the monsoon front, resulting in spatially heterogeneous changes in monsoon rainfall (Shanahan et al., 2015). A more comprehensive evaluation of these hypotheses is currently hindered by a lack both of continuous terrestrial proxy data spanning the terminal AHP and of high-resolution transient simulations with coupled GCMs.
The Last Two Millennia
The last two millennia of climate variability on the African continent have recently been reviewed as part of the PAGES 2K synthesis effort (Nash et al., 2016). This effort highlighted the dearth of high-resolution, well-dated paleoclimate records spanning this time interval, especially for West Africa. The lack of records reflects the challenge in finding sediment records with well-preserved core tops and from areas where climatic and anthropogenic influences can be clearly separated. Within the semi-arid Sahel, the situation is even more complicated because many lacustrine and wetland systems dried out during the drought of the 1950s to 1980s. Still, there remain some potential locations, particularly across humid equatorial West Africa, where more such reconstructions could be produced. Furthermore, new methods of proxy climate reconstruction that are potentially less susceptible to the confounding influences of anthropogenic land-use change, such as hydrogen isotope analysis of sedimentary leaf waxes (Sachse et al., 2012), may help with the interpretation of these records.
The best-dated and highest-resolution record of the last two millennia in West Africa comes from the annually laminated sediments of Lake Bosumtwi (6.5°N) in the northern coastal zone of Ghana (Shanahan et al., 2009; Figure 8). Oxygen isotope variations in authigenic lake carbonate (δ18Olw) indicate that conditions became gradually drier over much of the first millennium, reaching a minimum centered at ca. 950 CE, followed by a return to wetter conditions at the end of the Medieval Climate Anomaly (MCA). Dry conditions returned during the Little Ice Age (LIA) between 1450 and 1750 CE, with peak aridity occurring at ca. 1600 CE. From 1750 CE onward, conditions became more humid and remained that way until present. The authors note that the recent “Sahel drought” of the mid to late 20th century is superimposed on the interval of overall wetter conditions that occurred over the last century (Shanahan et al., 2009). The magnitudes of naturally occurring droughts, such as those that occurred during the LIA and the MCA, greatly exceed any drought in the historical record.
The first millennium drying trend at Lake Bosumtwi is supported by records from across West Africa, all of which show a general trend toward drier conditions over this time period (Nash et al., 2016; Figure 8). For example, both the pollen records from Lake Tizong (7.3°N; Lézine, Holl, et al., 2013) and Lake Bambili (5.9°N; Lézine, Assi‐Kaudjhis, et al., 2013) show a reduction in arboreal pollen at ca. 500 CE, although at the more southerly site, this is followed by a second, much larger transition at ca. 900 CE. A transition to drier conditions at around the same time is also evident in the ostracod Sr/Ca record from Kajemarum Oasis (13.3°N; Street-Perrott et al., 2000) and in a record of dust accumulation in a marine record from offshore Senegal (GeoB9501, 16.8°N, not shown; Mulitza et al., 2010). In the alkaliphilic diatom record from Lake Ossa (3.8°N; Nguetsop, Servant-Vildary, Servant, & Roux, 2010), the transition to drier conditions occurs slightly later (ca. 660 CE) because of a brief return to wetter conditions between 530 and 870 CE.
Most of the West African records also show a shift toward wetter conditions over the last few centuries, suggesting that this, too, reflects a large-scale regional climate response (Figure 8). With the exception of the Lake Tizong pollen record (Lézine, Holl, et al., 2013), which only recovers slightly, the records show above-average wetness over the last two to three centuries, including some of the wettest intervals of the last two millennia (Nash et al., 2016). The lack of a strong recovery in the Tizong pollen record may reflect local or regional differences in the climate or it may reflect anthropogenic influences on vegetation cover.
The climate of the intervening time period (500−1600 CE) is more complex. The record from Lake Bosumtwi suggests that most of the MCA was dry, and the most arid conditions of the Holocene were achieved during this interval (Shanahan et al., 2009; Figure 8). Conditions were more humid from 1200 to 1400 CE and shifted to more arid during the LIA. At Lake Tizong, low arboreal pollen concentrations indicate that overall, conditions were dry throughout this period, but, as in the Lake Bosumtwi record, the lowest arboreal pollen concentrations occur coincident with the MCA and the LIA, with an increase in trees during the intervening time interval (Lézine, Holl, et al., 2013). Similarly, at Kajemarum Oasis, two phases of aridity occur, separated by an interval of wetter conditions (Street-Perrott et al., 2000). The timing of these dry events is broadly similar to that of the Bosumtwi record, though the intervening wet interval occurs more abruptly and starts earlier (970 CE). In contrast, the two southern sites (Bambili, Ossa) show a single, extended period of more arid conditions that starts later (~ 870 CE) and peaks at around the time that sites to the north show wetter conditions (Lézine, Assi‐Kaudjhis, et al., 2013; Nguetsop et al., 2010; Figure 8).
The indications that conditions were dry during the LIA throughout West Africa are consistent with other records from the Northern Hemisphere tropical Atlantic that suggest LIA aridity (Haug, Hughen, Sigman, Peterson, & Röhl, 2001; Hodell et al., 2005; Lane, Horn, Orvis, & Thomason, 2011; Stansell, Steinman, Abbott, Rubinov, & Roman-Lacayo, 2013). At the same time, records from the Southern Hemisphere tropics appear to show wetter conditions, consistent with a southward migration of the ITCZ (Diaz et al., 2011; Ledru et al., 2013; Vuille et al., 2012). Several studies have suggested that the LIA may have been caused by a centennial-scale reduction in the strength of the AMOC. For example, the Florida Current, which feeds from the Caribbean into the Gulf Stream, was reduced by ~ 10% during the LIA (Lund, Lynch-Stieglitz, & Curry, 2006). At the same time, decreased 14C activity of waters off the North Iceland Shelf also suggests a reduction in Gulf Stream transport (Wanamaker et al., 2012). Model experiments demonstrate that reductions in the AMOC can result in a southward displacement of the ITCZ, consistent with paleoprecipitation reconstructions (Vellinga & Wood, 2002). Model experiments suggest that sustained reductions in AMOC during the LIA may have been triggered by some combination of external (e.g., volcanic, solar) forcings (Krebs & Timmermann, 2007; Miller et al., 2012) or by internal atmosphere ocean interactions, and sustained by feedbacks involving the coupled sea ice ocean system (Landrum et al., 2013).
In contrast, the causes of MCA aridity, as well as the differences in the MCA-LIA transition between semi-arid and humid tropical West Africa, are unclear. Previous studies have suggested that the MCA was characterized by either a strengthening of the AMOC (Trouet et al., 2009) and/or a more positive state of the AMO (Mann et al., 2009), both of which would be expected to produce wetter conditions over West Africa. Furthermore, sites from tropical South America show a north−south dipole pattern during the MCA (Ledru et al., 2013; Vuille et al., 2012), consistent with a northward shift of the ITCZ driven by a strengthening of the AMOC (Diaz et al., 2011; Ledru et al., 2013; Vuille et al., 2012) but inconsistent with the more complex signature in tropical Africa. MCA precipitation anomalies also do not seem to be explained by remote forcing from the tropical Pacific. Although climate variations in the tropical Pacific do have a significant influence on interannual variability in the West African monsoon (Janicot et al., 1996; Rowell, 2001; Semazzi et al., 1988), reconstructions suggest a more la Niña-like state during the MCA (Cobb, Charles, Cheng, & Edwards, 2003; Graham et al., 2007; Mann et al., 2009), which is expected to produce wetter, not drier, conditions in West Africa (Janicot, Trzaska, & Poccard, 2001). A more complete understanding of the origins of West African climate variability during the MCA will require the development of more high-resolution records from the region.
Over the past few decades, significant advances have been made to the understanding of the climate of West Africa and its variability on timescales ranging from decades to millennia. These advances improve the understanding of the monsoon response to external forcing but also highlight the limitations of the current proxy data and models in evaluating past changes. Based on existing research, the following observations can be made:
• The modern West African climate system appears to have been established at, or just prior to, the start of the Quaternary, in conjunction with a combination of changes in paleogeography and the establishment of Northern Hemisphere ice sheets.
• Climate variations on timescales of tens to hundreds of thousands of years dominate records of past climate and highlight the influence of orbital variations in insolation on West African climate. The evidence for orbital variations in records spanning as far back as 8 Ma points to direct forcing of West African climate by low-latitude seasonal insolation changes and insolation-driven low-latitude temperature gradients. However, there is also evidence that changes in temperature and ice volume in the high latitudes may have played a role in enhancing and modifying these signals.
• Positive land surface feedbacks play a role in the magnitude and persistence of the climate changes on a variety of timescales. They likely influence how the monsoon transitions between wet and dry states. However, the data suggest that the idea of an abrupt monsoon response to gradual forcing (e.g., at the end of the AHP) may be locally restricted to the northern edge of the monsoon system. Elsewhere, the response was complex and in some cases, locally abrupt.
• There is a strong teleconnection between high-latitude millennial-scale events, such as HSs, and West African rainfall that appears to be reproduced by the current generation of coupled general circulation models. However, the proxy response to individual events is spatially and temporally inconsistent. It is unclear why these differences occur. The limited availability of proxy data, particularly on land and in equatorial West Africa, restricts our ability to evaluate these inconsistencies.
• There is a dearth of high-resolution records from the last millennium in West Africa. Those that exist show qualitatively consistent signals, but differences with records from elsewhere in the tropical Atlantic highlight the need for more high-resolution records, particularly for the MCA.
The author’s research for this paper was supported by a series of small grants from the Jackson School of Geosciences, including a Chevron Centennial Fellowship. The author is grateful to all those who have contributed the data used in this synthesis.
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