Quaternary Climate Variation in Eastern Africa
Summary and Keywords
The people of East Africa are particularly vulnerable to the whims of their regional climate. A rapidly growing population depends heavily on rain-fed agriculture, and when the rains deviate from normal, creating severe drought or flooding, the toll can be devastating in terms of starvation, disease, and political instability. Humanity depends upon climate models to ascertain how the climate will change in the coming decades, in response to anthropogenic forcing, to better comprehend what lies in store for East African society, and how they might best cope with the circumstances. These climate models are tested for their accuracy by comparing their output of past climate conditions against what we know of how the climate has evolved. East African climate has undergone dramatic change, as indicated by lake shorelines exposed several tens of meters above present lake levels, by seismic reflection profiles in lake basins displaying submerged and buried nearshore sedimentary sequences, and by the fossil and chemical records preserved in lake sediments, which indicate dramatic past change in lake water chemistry and biota, both within the lakes and in their catchments, in response to shifting patterns of rainfall and temperature. This history, on timescales from decades to millennia, and the mechanisms that account for the observed past climate variation, are summarized in this article. The focus of this article is on paleoclimate data and not on climate models, which are discussed thoroughly in an accompanying article in this volume. Very briefly, regional climate variability over the past few centuries has been attributed to shifting patterns of sea surface temperature in the Indian Ocean. The Last Glacial Maximum (LGM) was an arid period throughout most of East Africa, with the exception of the coastal terrain), and the region did not experience much wetter conditions until around 15,000 years ago (15 ka). A brief return to drier times occurred during the Younger Dryas (YD) (12.9–11.7 ka), and then a wet African Humid Period until about 5 ka, after which the region, at least north of Lake Malawi at ~10º S latitude, became relatively dry again. The penultimate ice age was much drier than the LGM, and such megadroughts occurred several times over the previous 1.3 million years. While the African continent north of the equator experienced, on average, progressively drier conditions over the past few million years, unusually wet periods occurred around 2.7–2.5, 1.9–1.7, and 1.1–0.7 million years ago. By contrast, the Lake Malawi basin at ~10º—14º S latitude has undergone a trend of progressively wetter conditions superimposed on a glacial–dry, interglacial–wet cycle since the Mid-Pleistocene Transition at ~900 ka.
East Africans are particularly dependent on a stable climate, for their economy and personal wellbeing are overwhelmingly dominated by agricultural productivity. Unfortunately the interannual variability in climate is extreme, at times leading to severe drought or flooding, starvation, disease, and political unrest. An important question arises regarding future climate in the region; will it become wetter as some models predict (Vecchi & Soden, 2007), or drier as has been observed over the past few decades and attributed to anthropogenic warming (Funk et al., 2008; Tierney, Ummenhofer, & deMenocal, 2015)? Given the implications for the welfare of future generations in East Africa, climate models that attempt to answer this fundamental question need to be improved to the point where they are consistent in their output and compare favorably with observed trends in rainfall and temperature. There is of course no single, uniform climate trend that will apply to all of East Africa, for the varied topography of the landscape, with its large lakes, rift valleys, and highlands, including snow-capped peaks, results in significant spatial heterogeneity in mean rainfall totals (Mutai & Ward, 2000). Thus, the challenge for the climate modeling community is to accurately predict both the temporal and spatial trends of future climate change in the region.
The climate models can be tested through comparison of their hindcasts against observed climate history, both on a decadal scale from instrumental observations of rainfall and temperature, and on centennial—millennial timescales of past trends derived from various proxy records, recovered primarily from lake and ocean sediment cores, ice cores, and cave deposits. The purpose of this article is to provide an overview of the data that underpin our present understanding of climate change on these longer timescales in East Africa that span the Quaternary, or roughly the last two million years. Climate models are discussed elsewhere in this volume (see Theory and Modelling of the African Humid Period).
Present-Day Climate in East Africa
For the purpose of this article, “East Africa” is considered to extend from the western margin of the East African Rift Valley to the Indian Ocean, and from 5º N to 15º S latitude (Fig. 1).
Given the relatively limited range of temperature change in this tropical setting throughout the year, the most important aspect of East African climate is rainfall, for it has the greatest impact on the distribution of vegetation and habitability on the landscape. The annual range of daily average temperature through the seasons is generally just 3–6 ºC, although the diurnal range can be substantially greater, around 12 ºC (Nicholson, 2011). Daily high temperatures are typically in the range of 27–30 ºC during the summer months throughout much of East Africa, and nightly lows average about 10–12 ºC in the winter months, and of course much colder in the higher elevations.
Rainfall is highly seasonal in East Africa as the tropical rain belt migrates north and south with the sun. Much of the region experiences two rainy seasons each year, the so-called “long rains” in March–April–May (MAM) and the “short rains” in October–December (OND). However, the latitudinal extremes of East Africa experience just one rainy season per year: Lake Malawi at 10–14º S receives most of its rainfall in DJF and Lake Turkana at 4º N is wettest in JJA. The seasonality is traditionally explained by the north–south migration of the Inter-Tropical Convergence Zone (ITCZ), but in fact the situation is much more complex. The migration and seasonal changes of atmospheric jet streams, together with orographic impact on the winds play a major role (Nicholson, 2011). Also important is what is termed the Congo Air Boundary (CAB), a convergence zone that delineates the boundary between low-level easterly winds off the Indian Ocean and the westerly monsoon flow over western equatorial Africa (Fig. 2).
The long rains make up roughly 70% of the annual precipitation over much of the region exposed to two rainy seasons, and they display far less inter-annual variance than the short rains (Camberlin & Philippon, 2002). The inter-annual variability of the short rains is linked to the El Niño—Southern Oscillation (ENSO) and the distribution of sea surface temperature (SST) anomalies in the Indian Ocean. During most ENSO years, the eastern (western) Indian Ocean experiences anomalously cool (warm) SSTs, and an anomalous Walker circulation sets up between Africa and the Maritime Continent. The pattern of anomalously warm (cool) SST in the western (eastern) Indian Ocean is referred to as the positive phase of the Indian Ocean Dipole (IOD) (or Indian Ocean Zonal Mode [IOZM]). A positive IOD leads to weaker westerlies and a reduction in subsidence over East Africa, which in turn enhances convection and rainfall events during OND (Camberlin, Janicot, & Poccard, 2001; Hastenrath, Nicklis, & Greischer, 1993; Mutai & Ward, 2000). The correlation between OND precipitation and ENSO or the IOD is not uniformly strong across all of East Africa. Indeje, Semazzi, and Ogallo (2000) identify eight homogeneous rainfall regions between 4º N and 10º S latitude and from 30º E longitude to the coast which, in general, all experience enhanced OND precipitation in ENSO years, but the correlation between sea surface temperature in the central equatorial Pacific Ocean (Niño-3 SST) and OND rainfall varies from a high of 0.28–0.3 over much of Tanzania and to the north of Lake Victoria, to a low of just 0.11 in a small subregion of lakes and mountains to the west of Lake Victoria (Mutai & Ward, 2000). Thus the residual after ENSO is removed still accounts for 70% or more of the inter-annual OND rainfall variance and is attributed to coherent, large-scale climate processes extending from East Africa westward to Amazonia, with potential impact by Atlantic SST (Mutai & Ward, 2000). Camberlin and Philippon (2002) claim that inter-annual variability in the long rains of MAM does not exhibit a strong relationship with any large-scale climate anomalies. However, total precipitation in the long rains has been diminishing progressively in recent decades in lock step with steady and intense warming of the central Indian Ocean (2–3 times faster than the central Pacific [Williams & Funk, 2011]) and is tentatively attributed to anthropogenic warming. This progressive warming has resulted in a westward expansion of the Indo-Pacific warm pool, enhanced zonal westerlies over the equatorial western Indian Ocean, more intense convection and precipitation over the Indian Ocean, and advection of dry easterly air aloft towards Africa (Funk et al., 2008; Lyon & DeWitt, 2012; Williams & Funk, 2011).
Signals of Past Climate Change in East Africa
Paleoclimate records of the continents are derived from a variety of sources, including lake and bog sediments, glacial ice cores, tree rings, cave stalagmites, and ocean sediment cores from proximal marine environments. Caves are rare in East Africa, although there are some in eastern Tanzania that have the potential for providing excellent records with unparalleled chronological control. Paleoclimate records derived from tree ring analysis are rare in the tropics compared to temperate latitudes, because annual growth rings are not common in most tropical tree species. However, some species are showing promise, and a few dendrochronology records are being developed for East Africa, such as a 200-year record of past rainfall in Zimbabwe derived from Pterocarpus angolensis, a hardwood tree that grows to a height of about 15 m and is widespread throughout much of tropical Africa between the equator and 20º S latitude (Therrell, Stahle, Ries, & Shugart, 2006). And more recently cores recovered from baobab trees (Adansonia digitata L.) in northeastern South Africa have been radiocarbon-dated and analyzed for δ13C to produce a 1,000-year record of past rainfall in the region that correlated reasonably well with an ENSO index based on equatorial Pacific SSTs (Woodborne et al., 2015). Ice core δ18O and aerosol records from the glaciers of Mt. Kilimanjaro span the past 11.5 ky that Thompson et al. (2002) correlate to regional lake levels. There are a few coral records from the western Indian Ocean (Zinke, Pfeiffer, Timm, Dullo, & Brummer, 2009), including a 194-year coral record from the Kenyan coast near Malindi that provides a record of past SST (δ18O) and freshwater input from a nearby river (fluorescence of river-derived fulvic acids) (Cole, Dunbar, McClanahan, & Muthiga, 2000). There are also new records coming out from marine sediment cores recovered off the mouths of the Rufiji and Zambezi off East Africa (Bouimetarhan et al., 2015; Schefuss, Kuhlmann, Mollenhauer, Prange, & Patzold, 2011). However, these are the exceptions; most paleoclimate records from East Africa are derived from sediment cores recovered from lakes and bogs, and these provide much of the basis for the following review.
Decadal–Centennial Climate Variability Over the Past Two Millennia
There are a dozen or so high-resolution, reasonably well-dated lake sediment records from East Africa that reveal a complex history, both spatially and temporally, of rainfall over the past 2,000 years. Among these are records of: Mg/Ca preserved in endogenic calcite from Lake Edward (Uganda/Congo) (Russell & Johnson, 2007); lake-level and salinity fluctuations in Lake Naivasha (Kenya) inferred from sediment stratigraphy and the species compositions of fossil diatom and midge assemblages (Verschuren, Laird, & Cumming, 2000); leaf wax δD in Lake Challa on the eastern slope of Mt. Kilimanjaro, Tanzania (Tierney, Russell, Sinninghe-Damste, Huang, & Verschuren, 2011) and Sacred Lake on the northeastern slope of Mt. Kenya (Konecky et al., 2014); the mass accumulation rate of terrigenous sediment in Lake Malawi (Johnson & McCave, 2008); diatom assemblages, water content, and organic carbon content in Lake Tanganyika (Stager, Cocquyt, Bonnefille, Weyhenmeyer, & Bowerman, 2009); carbonate abundance in cores from Lake Turkana in northern Kenya (Halfman, Johnson, & Finney, 1994); % shallow water diatoms in a core from Lake Victoria (Stager, Ryves, Cumming, Meeker, & Beer, 2005); sediment magnetic susceptibility in a core from Lake Masoko in the Rungwe Mountains of southern Tanzania (Garcin et al., 2007); and a combination of lithological parameters (magnetic susceptibility, % organic carbon, % inorganic carbon, authigenic mineralogy) in cores from two crater lakes in western Uganda (Russell, Verschuren, & Eggermont, 2007).
The first high-resolution record to be published for East Africa was from the Crescent Island crater in Lake Naivasha, Kenya (Verschuren et al., 2000). It displays lake level fluctuations of several tens of meters over the past 11 centuries, and a rather surprising result, that this part of Kenya was relatively wet during the Little Ice Age. Verschuren et al. observed that variations in lake level were associated with sunspot activity, with high lake levels having occurred during the Maunder, Sporer, and Wolf sunspot minima. While this observation is intriguing, no mechanism was proposed by Verschuren et al. to explain the link between rainfall and solar intensity, and their actual correspondence cannot be proven, given the chronological uncertainty associated with radiocarbon dates on this timescale.
Subsequent high-resolution records from East Africa have also revealed decades or centuries of relatively wet or dry conditions that, while not highly consistent, were similar to the Lake Naivasha record in their dominant timescale of variability. Drought was widespread throughout much of East Africa at ~0–200 ce, 550–900 ce, and 1000–1200 ce (Konecky et al., 2014; Russell et al., 2007). However, several papers claimed a zonal pattern of anti-phased relationship between lake level histories in the western and eastern parts of East Africa during the Little Ice Age (1500–1850 ce), when lakes in the Albertine Rift were low and lakes to the east (Victoria, Naivasha, Challa) were high (Russell et al., 2007; Tierney et al., 2011). Tierney et al. (2013) presented a Monte Carlo empirical orthogonal function (MCEOF) analysis of seven lake records from East Africa, coupled with a GPCC climate model EOF of ten-year smoothed instrumental rainfall data in the region and found a good correlation between the spatial EOF of the decadal rainfall pattern and the 1,000-year MCEOF1 of the lake level histories (Fig. 3).
The two easternmost lakes (Naivasha and Challa) fell into a rainfall regime that extended along the East African coast into the Horn of Africa, with high lake stands when the sea surface temperatures were in a positive mode of the Indian Ocean Dipole (IOD) (warm west, cool east, which often occurs in El Niño years), while lakes to the west (Tanganyika, Edward, Masoko, and Malawi) were at low stands (Fig. 3).
However, as more records have been generated, the anti-phased relationship of east-versus-west may not be holding up. The δD record from Sacred Lake, Mt. Kenya, in the eastern province (Konecky et al., 2014), for example, appears to be very much in line with the Mg/Ca record from Lake Edward in the western province (Russell & Johnson, 2005) (Fig. 4).
In addition, a careful look at the records from Lake Edward and Lake Naivasha (Verschuren et al., 2000), which served as the initial impetus for contrasting records from east and west (Russell et al., 2007), are not that far out of phase (Fig. 4). They may actually be recording coinciding patterns of pluvials and drought. However, it is impossible to test this possibility, given the inadequate resolution of the chronology in relation to the decadal pattern of change. If in fact most of the records from East Africa are indeed coincidental, the history of rainfall in the region has been responding quite uniformly to Indian Ocean SSTs, much as depicted by Goddard and Graham (1999) and Nicholson (2014) for historical rainfall. This shows two major rainfall regimes in East Africa, with tropical East (Southern) Africa experiencing wet (dry) conditions during positive IODs (Fig. 5).
With the discovery that a new paleo-temperature proxy for marine sediments, TEX86, (Schouten, Hopmans, Schefuss, & Sinninghe Damsté, 2002) applies to some lake sediments as well (Powers et al., 2004), high-resolution temperature records spanning the last few centuries have been developed for Lake Malawi and Lake Tanganyika. A 700-year temperature record from Lake Malawi displays a range of 24–27 ºC, with cool temperatures during the Little Ice Age and substantial warming during the past century that agrees well with instrumental observations and is attributed, at least in part, to anthropogenic causes (Powers et al., 2011). A 1,500-year temperature record from Lake Tanganyika exhibits a range of 22.5–26 ºC that closely parallels the Lake Malawi record as well as Northern Hemisphere temperature reconstructions for the same period, and shows unprecedented warming in the last few decades due to anthropogenic forcing (Tierney et al., 2010) (Fig. 6).
Climate Change in East Africa Since the Last Glacial Maximum (LGM)
Well-dated paleoclimate records extending back to the LGM (~19–26 ka) are now available from a number of lakes in East Africa and from off the Rufiji (Kenya) and Zambezi (Mozambique–South Africa border) Rivers, that are based on a variety of signals in the sediments, including assemblages of diatoms, ostracods and pollen; inorganic and organic geochemical composition; lithological characteristics; and exposed and submerged shoreline features.
Lake Turkana is a closed-basin lake in northern Kenya that has fluctuated some 200 m in elevation since the LGM, providing several lines of evidence for a climate history that is typical of most of East Africa. Shoreline terraces exposed as high as 100 m above present lake level (at which elevation the lake overflows into the White Nile drainage to the west) have been dated, and document wetter conditions than present in the basin during the AHP, between 12.5 and 5 ka (Fig. 7) (Butzer, Isaac, & Washbourn-Kamau, 1972; Garcin, Melnick, Strecker, Olago, & Tiercelin, 2012). Seismic reflection profiles from the southern basin of the lake, coupled with dated sediment cores, display submerged shoreline features that indicate considerably drier conditions than today during the LGM, and even more arid times, involving complete desiccation of the lake, at 18–19 ka and 17.5–17 ka. This would place lake level about 100 m below present (Fig. 8) (Morrissey & Scholz, 2014).
The timing of the desiccation events roughly coincides with Heinrich Stadial Event 1 (H1), a time of widespread aridity throughout much of Africa (Stager, Ryves, Chase, & Pausata, 2011), although with some exceptions, especially in southern Africa (Thomas, Burrough, & Parker, 2012). Evidence for fluctuating climate in sediment cores from Lake Turkana have been derived from fossil diatom assemblages (Halfman, Jacobson, Cannella, Haberyan, & Finney, 1992), abundance of total organic carbon and total inorganic carbon, and physical parameters such as grain size and sediment bulk density (Halfman et al., 1994; Morrissey & Scholz, 2014).
There are only four temperature records from East Africa derived from TEX86 and branched GDGT analyses that extend from the present back to the LGM: Lake Malawi (Powers et al., 2005), Lake Tanganyika (Tierney et al., 2008), Lake Challa (Sinninghe-Damste, Ossebar, Schouten, & Verschuren, 2012), and Sacred Lake, Mt. Kenya (Loomis, Russell, Ladd, Street-Perrott, & Sinninghe Damsté, 2012). They all show the LGM to have been on the order of 3–5 ºC cooler than today (Fig. 9), all consistent with previous estimates of LGM cooling based on pollen records from East Africa (Bonnefille et al., 1992; van Zinderen Bakker and Coetzee, 1972).
The highest of the three lakes, Sacred Lake, at 2,240 masl, warmed ~5.1 ºC, while Malawi and Tanganyika, both at elevations below 1,000 masl, warmed 3.4º and 2.9ºC, respectively (Fig. 9). The Lake Challa record is different from the other three lake records, in that it displays substantial cooling from a relatively warm LGM until 15 ka, which may have resulted from meltwater input off the glaciers of Mt. Kilimanjaro (Sinninghe-Damste et al., 2012). Given its unique character that may reflect highly localized conditions, the Challa temperature record is not considered further in this review. Lakes Malawi, Tanganyika, and Sacred all began to warm gradually soon after the LGM, but at different rates and with no significant cooling during Heinrich Stadial Event 1 (H1). The rate of warming began to accelerate in all three lakes around 14.5 ka and continued until the YD, when they all cooled ~1.5ºC. A short TEX86 record from Lake Albert that spans the YD shows a dramatic cooling of ~3ºC (Berke et al., 2014). In Lake Victoria, a warming trend that began immediately after the lake began to refill at 15 ka, after complete desiccation in the Late Pleistocene (Johnson et al., 1996), shows a hesitation in the warming rather than a reversal to cooler temperatures during the YD (Berke, Johnson, Werne, Grice, et al., 2012) (Fig 9). Most of the lakes in East Africa displayed their warmest temperatures in the mid Holocene, around 5 ka, and have since cooled by 1–2ºC (Berke, Johnson, Werne, Schouten, et al., 2012) (Fig. 9), until recent decades when anthropogenic warming has exceeded the mid-Holocene high temperatures. The reason for the thermal maximum at 5 ka is not fully understood; Berke, Johnson, Werne, Schouten et al. attributed it to a time of maximum September insolation, which is the month when all of the large lakes in East Africa, both north and south of the equator, undergo maximum seasonal warming.
The history of East African precipitation since the LGM was first deduced from dated lake shoreline and pollen records, and in recent years has been much improved by δD and δ13C records from specific compounds preserved in terrestrial leaf wax extracted from lake sediment cores. Two such records, from Lake Malawi (Castañeda, Werne, & Johnson, 2007) and Lake Tanganyika (Tierney et al., 2008), indicate the LGM to have been much drier than present (Fig. 10).
According to these records, both basins remained relatively dry from the LGM, through H1, until about 16 ka, when they became progressively wetter until the YD, when they reverted to much drier conditions (Fig. 10). (Some authors have referred to the leaf wax δD record from Lake Malawi, which displays relatively depleted values during the LGM, as evidence for wet conditions in the basin at that time. However, Konecky et al.  provide clear evidence for this record being a reflection of atmospheric circulation and source area of precipitation, which in this basin displays enriched δD values at times of relatively high, not low, precipitation.) A leaf wax δD record from Lake Challa exhibits depleted values in the early stage of the LGM (~23–21 ka) suggesting somewhat wetter conditions than the present day (Tierney et al., 2011) (Fig. 10). The Challa δD record indicates substantial drying during the YD, consistent with the Tanganyika and Malawi records, as well as another significantly dry period around 4 ka, just after the AHP (Fig. 10). However, another record of rainfall history since the LGM in the Challa basin is derived from the Branched and Isoprenoid Tetraether (BIT) Index of soil bacterial versus aquatic archaeal membrane lipids extracted from the sediment (Verschuren et al., 2009), and it looks more like the Malawi and Tanganyika records between the LGM and 10 ka than like the Challa δD record, with a relatively dry LGM, at least from 23 to 19 ka, with an onset of wetter conditions around 16 ka, a distinctively dry YD, and the wettest conditions at Challa around 10–11 ka (Fig. 10). DiNezio and Tierney (2013) report that several lakes in easternmost tropical Africa were wetter during the LGM than today, and a few that record no change between the LGM and today. DiNezio and Tierney attribute the higher rainfall to exposure of the Sunda Shelf in the eastern Indian Ocean during the LGM low stand of sea level, a consequent reduction in atmospheric convection and pronounced aridity over the Maritime Continent, and enhanced rainfall over the western equatorial Indian Ocean, which extended onto the African continent.
The history of rainfall in these basins differ significantly through the Holocene (Fig. 10). The Challa and Tanganyika δD records look quite similar, with wet conditions during the AHP, terminating abruptly around 5 ka. The Victoria record shows a gradual trend towards drier conditions from 11 ka to 3 ka, with no abrupt shift at the end of the AHP. The timing of the termination of the AHP at Lake Edward appears slightly earlier than at Tanganyika and Challa, but was quite abrupt as well. The Holocene history of moisture in the Malawi basin is completely different from the lakes to the north. While wetter than the YD, the early Holocene was relatively dry in this Southern Hemisphere lake and highly variable compared to the late Holocene, with the wettest conditions having existed at about 3 ka and 5 ka. Rainfall in the Lake Malawi basin apparently responded to precessional forcing of Southern Hemisphere summer insolation, which was at a minimum around 9 ka (Castañeda et al., 2007). It is interesting to note that rainfall apparently was elevated during H1 and the YD in the lower Zambezi River catchment to the south of Lake Malawi, which Schefuss et al. (2011) attributed to a southward shift in the ITCZ during these stadials (rather than a compression of the latitudinal range of annual transit of the ITCZ as proposed by Trauth et al., 2007) causing summer rains to have been abbreviated in the Malawi basin, while elevated in the lower Zambezi.
Thus, the overall picture for East Africa is that the LGM was cooler and drier than present for most of the region, with the exception of the African coastline, extending inland to portions of the eastern arm of the Rift Valley where the LGM was relatively wet (DiNezio & Tierney, 2013; Singarayer & Burrough, 2015). While East Africa began to warm from 20 ka on, wetter conditions did not set in until around 15–16 ka, at the end of H1. The Younger Dryas (YD) was relatively cool and dry throughout East Africa. Moist conditions, coinciding with the African Humid Period (AHP), lasted from the end of the YD until ~4–5 ka. The AHP coincides with a precessional maximum in Northern Hemisphere summer insolation. It enhanced both the East African Monsoon and the West African Monsoon by elevated summer heating of the African interior, thereby increasing the thermal gradient between land and sea and promoting more intense convective rainfall over Africa, not only north of the equator, but as far south as Lake Rukwa at 8º S latitude (Claussen & Gayler, 1997; deMenocal et al., 2000; Kutzbach & Street-Perrott, 1985; Liu, Harrison, Kutzbach, & Otto-Blisner, 2004).
The African Humid Period is not observed in the early Holocene of Lake Malawi (10–15º S) and farther south in Africa, where precessional forcing of the African monsoon would be driven by austral summer insolation, which is anti-phased with Northern Hemisphere summer insolation (Gasse, Chalie, Vincens, Williams, & Williamson, 2008). Climate models indicate that the main drivers of this climate history since the LGM are orbital variation in insolation, the changing concentration of CO2 in the atmosphere, and the expansion/contraction of the continental ice sheets (Otto-Bliesner et al., 2014).
East African Climate Prior to the LGM
Very few East African paleoclimate records are available that extend much beyond the LGM. There are a number of outcrops of lacustrine sequences exposed in Ethiopia, Kenya, and northern Tanzania that indicate large lakes were in existence around 2.7–2.5 million years ago (Ma), 1.9–1.7 Ma and from 1.1–0.9 Ma (Trauth, Maslin, Deino, & Strecker, 2005) (Fig. 11).
These were periods of extreme climate variability that Trauth et al. (2007) correlate to maxima in the 400 ka eccentricity cycle of NH insolation. Another period of demonstrably wetter conditions than today is revealed in the Lake Naivasha basin, where a 60-m-thick unit of diatomites and associated lake deposits are interbedded with tephras dating between 60 and 176 ka (Trauth, Deino, Bergner, & Strecker, 2003). Lake high stands are apparent around the Marine Isotope Series (MIS) 6–5 transition, peaking around 136±3 ka, with subsequent high stands occurring at 10–11 ka intervals, suggesting half-precessional forcing when insolation maxima coincide with the spring and fall seasons of high rainfall (Trauth et al., 2003).
Scholz et al. (2003) identified a promising coring site from a high-resolution seismic reflection survey of the Kavala Island Ridge in central Lake Tanganyika, where they succeeded in getting a core that extends back perhaps 100 ka, and possibly to MIS 6 (~135 ka). There was no way to date the sediment beyond a radiocarbon age of 45 ka at 350 cm depth in core. However, the 11 14C dates above 350 cm depth in core fell close to a straight line, which the authors extrapolated to estimate an age of 79 ka at the top of a major disconformity at 940 cm depth in core. A dense, gray mud, very low in organic carbon content, underlay the unconformity and was interpreted to reflect subaerial exposure, indicating a lake level more than 350 m below present lake level. Scholz et al. (2003) speculated that this occurred during MIS 6 and concluded that the penultimate glacial was more extreme than the LGM in the Tanganyika basin. They further deduced that Tanganyika was cooler and drier than present from 58 ka through the LGM, and that the lake experienced brief peaks in primary production that may coincide with Heinrich Stadials 1–6.
Two piston cores recovered from slightly farther south in Lake Tanganyika were analyzed for TEX86 and leaf wax δD to develop the history of rainfall and temperature in the basin back to 60 ka (Tierney et al., 2008). The records display as much variability between 60 ka and the LGM as from the LGM to the present, with temperatures, for example, as warm at 60 ka as during the Holocene thermal maximum at 5 ka (Fig. 12).
The δD record displays a number of brief excursions to drier conditions that Tierney et al. (2008) attribute to Heinrich Stadial events H1, H4, H5, and H6, as did Scholz et al. (2003), although no such shifts to aridity were observed during H2 or H3. If these correlations are true, East African climate responded to millennial-scale climatic variability in the Northern Hemisphere during MIS 3. One mechanism by which this may happen is through cooling of Indian Ocean surface water during NH stadials, thereby lowering the latent heat flux from the Indian Ocean and moisture transport onto the East African landscape (Tierney et al., 2008).
The longest, high-resolution, nearly continuous record of past climate change in East Africa comes from the Lake Malawi Drilling Project. Two sites were drilled in 2005; one in the central basin in 593 m water depth recovered a nearly complete sediment record to 384 m below lake floor, and the other in the north basin from 395 m water depth, where three cores were recovered from a ~40 m—thick sequence of hemipelagic silty clay overlying a transgressive, nearshore sand (Scholz et al., 2005). The first papers to be published from the drilling project focused on the top 88 m of the sequence, representing approximately the past 150,000 years. This sequence contains 3 intervals of calcareous, ostracod—bearing sediment that indicate major low stands in the lake, the most severe being 550 m below the present lake level. The ages (and drop in lake level) of these low stands were estimated to have been 74–78 ka (−350 m), 95–115 ka (−550 m), and 128–135 ka (−550 m) (Scholz et al., 2007). These “megadroughts” were far more severe than the relatively dry conditions in the Malawi basin during the LGM and are attributed to the higher amplitude of precessional forcing at ~100 ka compared to 20 ka. The hydroclimate became considerably more humid and less variable in the Malawi basin after ~70 ka, as it also did in the Tanganyika basin and at Lake Bosumtwi, Ghana (West Africa) (Scholz et al., 2007). The timing of megadroughts in Lake Malawi subsequently had to be adjusted to considerably older ages (perhaps MIS 6) with the discovery of Toba ash (75 ka) at ~28 m below lake floor (Lane, Chorn, & Johnson, 2013), which had a tentative age of ~60 ka in the initial age model of Scholz et al. (2007).
The severity of the megadroughts in the Malawi basin had a major impact on the landscape. Pollen burial rates in the lake sediment were extremely low under these conditions, and the abundance of charred particles in the sediment was also greatly diminished, indicating insufficient vegetation on the landscape to cause major fires (Cohen et al., 2007). The saline, alkaline lake would have had a maximum depth of 125 m, compared to ~700 m today. Such conditions would have changed the locations and relative proportions of rocky, sandy, and muddy nearshore environments, and would have had severe impact on the biodiversity of cichlid fish and other biota in the lake (Cohen et al., 2007). Lake Malawi currently holds more species of fish (~1,000) than any other lake on Earth, and over 90% of them are endemic to the lake. Their history of speciation was undoubtedly impacted by the dramatic shifts in lake level and water chemistry that have been documented by the Malawi Drilling Project.
Brown, Johnson, Scholz, Cohen, and King (2007) analyzed one of the north basin cores in very high resolution with scanning X-ray fluorescence (XRF). They focused on the Zr/Ti ratio as an indicator of wind-swept volcanogenic sediment delivered to the lake by northerly winds in austral summer from the Rungwe volcanic field, when the ITCZ lies to the south of Lake Malawi. The record displays sharp peaks in this ratio during the YD, H1, and several other stadials between 10 and 50 ka (Fig. 13).
This is strong evidence for southward displacement of the ITCZ during stadial conditions in the Northern Hemisphere and underlines the significant impact that Northern Hemisphere climate dynamics exert on tropical East Africa, even at centennial–millennial timescales.
Lyons et al. (2015) presented the first results of the full 384 m sediment sequence, representing the past 1.3 million years, from the Lake Malawi Drilling Project. The authors derived a lake level history based on a principal component analysis of several sediment properties (lightness of color, % total organic carbon (TOC), δ13C of TOC, carbonate abundance, and saturated bulk density) that are strongly influenced by water depth. The resultant PC1 was quantified in terms of water depth by stratigraphic correlations of seismic reflectors of known sediment composition (PC1 value) at the drill site to coeval submerged and buried shoreline features, such as progradational deltas, at known depth below the present lake surface. The analysis indicates that lake level dropped 24 times in the past 1.3 million years, to at least 200 m below present level, and to 400 m below present level 15 times. Lake level was extremely variable and usually lower than today from 1.3 Ma to 800 ka (Fig. 14).
After the Mid-Pleistocene Transition at ~800 ka, the lake experienced more prolonged high stands and low stands. Age control in the drill core is limited between 75 ka (Toba ash) and 593 ka (the younger of two Ar–Ar dates in the drill core), so the exact timing of high stands and low stands was not determined in this interval.
Johnson et al. (2016) subsequently presented a history of temperature and rainfall for the 1.3 million year Malawi drill core based on measurements of TEX86, leaf wax δ13C, and calcite content (Fig. 15).
The temperature record displays a strong 100 ka cycle since the Mid Pleistocene ition (MPT) that allowed an age model to be generated for the 75–593 ka age interval by aligning the warm (cold) intervals with interglacial (glacial) periods. The leaf wax record displays relatively dry (wet) glacials (interglacials) and a progression overall toward wetter conditions in the Malawi basin since the MPT. This is in stark contrast to the trend towards progressively drier conditions in North Africa over the past few miTransllion years (Cerling, 1992; deMenocal, 2004), and raises interesting questions regarding early human migration and evolution. The very high-resolution record of calcite content in the core derived from scanning XRF, an indicator of past aridity in the basin, shows only limited evidence for precessional variability in the Malawi basin. Johnson et al. attribute this to two possible causes. First, forcing on rainfall between austral summer insolation and the long-term, east–west pattern of equatorial Indian Ocean SSTs appear to be incompatible, which at least for the past 130,000 years have been out of phase. Second, the Malawi basin lies close to the dipole boundary of African rainfall response to the IOD, and it is conceivable that the basin at times experiences the rainfall regime of southern Africa, which is anomalously dry (wet) during positive (negative) IOD, and at times experiences the rainfall regime of equatorial East Africa, which responds in the opposite manner to changing IOD.
As the Malawi drill core becomes better dated and as more of the East African lakes get drilled for long paleoclimate records, the timing of wet and dry conditions throughout East Africa, and their causes, will become better known. More temperature records based on TEX86 and perhaps other proxies will be generated. It is clear that East African climate has undergone dramatic change throughout the Quaternary Period, in terms of both frequency and intensity. Orbital forcing of East African climate at precessional, obliquity, and eccentricity periods is apparent and likely cause relatively cool, dry conditions throughout most of East Africa when NH summer insolation is at a minimum. Centennial-to-millennial-scale change in precipitation and temperature is observed and is tentatively linked to a southward shift in the ITCZ/tropical rain belt during NH stadials, and reduced moisture transport off the Indian Ocean onto the African continent. The decadal, East African rainfall response to shifting SST patterns over the Indian Ocean is still to be fully resolved; on the one hand, it appears to be bipolar, with: (1) high (low) OND rainfall along the coast and Horn of Africa when the Indian Ocean Dipole is positive (negative) (i.e., relatively warm (cool) in the western (eastern) equatorial Indian Ocean and (2) low (high) OND rainfall in the East African interior (Fig. 3). On the other hand, as more data become available in the region, this east–west anti-phased response to the IOD appears to have exceptions. At times, all of East Africa appears to experience higher OND rainfall during positive IOD. In either case (east–west anti-phased vs. in-phase response), over 70% of the variance in regional OND rainfall is still not accounted for.
As we face the most dramatic change in global climate that humankind has ever experienced, a better comprehension of the dynamics of East African climate is emphatically needed to enable more accurate predictions. This will entail the recovery of more high-resolution, precisely dated, paleoclimate records, coupled with more refined climate models that can accurately reproduce the history of climate derived from these records. These models, tested against the past, will provide our best estimate of what is in store for the people of East Africa and will underpin the most effective policies for adapting to the new order of rainfall and temperature in the region.
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