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Tectonic Dynamics in the African Rift Valley and Climate Change

Summary and Keywords

The East African Rift System (EARS) transecting the high-elevation East African plateau is one of the most outstanding rift systems on earth. Rifting was caused by a huge uprising mantle plume under East Africa. Two distinct rift branches are distinguished: an older, volcanically very active Eastern Branch and a younger, much less volcanic Western Branch. The Eastern Branch is generally characterized by high elevation, whereas the Western Branch comprises a number of deep rift lakes (e.g., Lake Tanganyika, Lake Malaŵi). These differences reflect different plate strengths, the latter of which are largely governed by differences in how the mantle plume interacted with the East African lithosphere. Much of the topography forming the East African plateau has been caused by the uprising mantle plume. The onset of topographic uplift in the EARS is poorly dated but preceded graben development, the latter of which commenced at ~24 Ma in the Ethiopian Rift, at ~12 Ma in Kenya, and at ~10 Ma in the Western Branch. Increased uplift of the East African plateau since ~15–10 Ma might be connected to climate change in East Africa and human evolution. East Africa experienced cooling starting at 15.5–12.5 Ma that heralded profound faunal changes at 8–5 Ma, when the hominin lineage split from the chimpanzee lineage. The Pliocene is characterized by warm and wet climate between 5.3 and 3.3 Ma transitioning into a period of cooler and more arid conditions after ~3 Ma. The climate in the EARS is controlled by westerly monsoonal flow over equatorial West Africa and easterly monsoonal flow over the Indian Ocean. The uplifting East African plateau intercepted those winds and contributed to the increased aridification of East Africa.

Keywords: East African Rift System, tectonics, mantle plume, uplift, climate change, hominins

Tectonics and Climate in the East African Rift System

The more than 3,000-km-long East African Rift System (EARS) (Fig. 1) cuts across the entire eastern part of the African continent. It is a biodiversity hotspot and, according to Leakey (1973), the “Cradle of Mankind”. The latter proposition highlights plausible relations between tectonics, climate, biodiversity, and hominin evolution.

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Figure 1. East African Rift System (EARS) Superimposed on the Ethiopian and Kenyan Plateaus, Collectively Known as East African Plateau (image GeolMapApp). Red bold line outlines elevation >1,000 m; blue lines delimit Archean Tanzania Craton (dashed blue line shows craton at 1,500 km depth (Adams et al., 2012); black lines indicate rift graben. EARS comprises series of individual graben that link up to form Western and Eastern Branch; rift floors of Eastern Branch with vast volcanic rock accumulations have high elevation, only Turkana graben of Eastern Branch has subdued elevation. Western Branch with small, scattered volcanic provinces and elongated, deep rift lakes. Localities mentioned in text and lines for cross sections in Figure 7 are shown.

The EARS is a diffuse rift system splitting the African Plate into the Somalian Plate in the east and the Nubian Plate in the west (Calais, DeMets, & Nocquet, 2003; Chorowicz, 2005; Saria, Calais, Stamps, Delvaux, & Hartnady, 2014). It is commonly subdivided into an Eastern and a Western Branch (Fig. 1). The Eastern Branch extends from the Afar depression in northern Ethiopia to the Manyara Rift in northern Tanzania. The Eastern Branch dissects the high-elevation Ethiopian and Kenyan plateaus (herein collectively referred to as East African plateau), which are separated by the Turkana depression. The rift basins in the Eastern Branch have high elevations; for instance, Lake Awasa in the Awash valley in the Ethiopian Rift has a surface elevation of 1,708 m asl and Lake Naivasha in the central Kenya Rift has one of 1,884 m asl. Many of the small lakes (100–200 km2 surface area) of the Eastern Branch formed during relatively recent times (mostly since the middle Pleistocene; Salzburger, van Bocxlaer, & Cohen, 2014) and then played an essential role in the sensitivity of the Eastern Branch to climate change (Trauth et al., 2010). The Eastern Branch is magmatically very active, with large shield volcanoes like Kilimanjaro, Mt. Kenya, and Mt. Elgon on the rift flanks.

The Western Branch extends from Lake Albert in Uganda to the Urema Graben in Mozambique (Fig. 1) and is magmatically much less active. Basin floors have, on average, lower elevations, and a number of long and deep rift lakes formed in the basins, indicating great absolute basin subsidence. The lake floor of Lake Tanganyika is ~700 m below sea level, and together with the floor of Lake Baikal and the Dead Sea represent the deepest points on the continents. Some of the large lakes of the Western Branch (e.g., Lake Tanganyika) formed in the middle/late Miocene (Cohen, Soreghan, & Scholz, 1993).

Davies, Vincent, and Beresford (1985) showed that the present-day topography of the East African plateau has a strong modifying effect on regional precipitation patterns by focusing and intensifying the summer monsoonal flow along the East African coast (Fig. 2). Strong surface winds just offshore East Africa are known as the East African (or Somalian) low-level jet. This jet results from the interaction of the winds emanating from a stable high in the southern Indian Ocean and the high elevation of the East African plateau creating a barrier that intensifies the northeastward deflection and overall wind speed. The East African Jet is known for its remarkable steadiness in direction and strength. Another major climate boundary at the western limit of the EARS between the Congo Basin and the Western Branch is the Congo Air Boundary (Fig. 3a), which is a convergence zone that delineates the limit between low-level easterly winds off the Indian Ocean and the westerly monsoonal flow over equatorial West Africa (see the article “Quaternary Climate Variation in Eastern Africa”). The East African Jet and the Congo Air Boundary highlight the complexity of tropical African climate, which is responding to a mixture of coupled ocean-atmosphere processes associated with both the Indian and Atlantic oceans (Camberlin, Janicot, & Poccard, 2001; see also the article “Climate of Eastern Africa”).

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Figure 2. Composite Map Showing Mean Surface Winds for Northern Hemisphere Summer (data from 1968 to 1996, NOAA-CIRES Climate Diagnostics Center, Boulder, Colorado) emanating from stable high near Mascarene islands (bold red contour line delimits area of high sea level pressure >1,020 hPa called Mascarene High). Dashed bold lines with arrows indicate direction of surface winds over Indian Ocean—wind originating from persistent Mascarene High flows northwest toward East Africa and shifts northeastward due to Coriolis force; highest surface winds over Arabian See exceeding 15 m/s, known as East African (or Somalian) low-level jet, occur just offshore Horn of Africa and cannot be accounted for by barometric pressure gradient between India and Indian Ocean. East African low-level jet forms by interaction of wind flowing towards East Africa and topography of East African plateau creating barrier that intensifies northeastward deflection and overall wind speed. Strong East African Jet remarkable for steadiness of direction and strength; wind picks up large amounts of moisture delivering wet maritime air to Indian subcontinent.

East Africa with its high elevation at the fringe of the West African and Indian monsoons forms a major dry climate anomaly region in the otherwise wet equatorial belt (Nicholson, 2017) (Fig. 3a, b). The orographically controlled Congo Air Boundary and the East African plateau are shielding East African climate from the Atlantic Ocean. It follows that the major question for East African paleoclimate is when high enough topography for intercepting the West African monsoon and creating the East African Jet was established. A proxy-base vegetation reconstruction by Lohmann, Butzin, and Bickert (2015) suggests that the tropical African rain and seasonal forest was much wider and stretched out further to the east in the Tortonian (early late Miocene, ~11–7 Ma ago) (Fig. 3c). This finding correlates nicely with predicted elevation changes in East Africa (Moucha & Forte, 2011). Another important question on geological timescales concerns changes in the sea surface temperature gradient in the Indian Ocean (Cane & Molnar, 2001) as it alters the climate in East Africa through modification of the intensity of the Indian monsoon.

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Figure 3. (a) Climatology of Tropical Africa for August Including Precipitation Rates (Mitchell & Jones, 2005), 925 hPa Winds (Kalnay et al., 1996), and Approximate Locations of Intertropical Convergence Zone (ITCZ) and Congo Air Boundary (CAB) (redrawn from Tierney et al., 2011). Also shown are highest and slowest shear-wave seismic velocity for the range 100–175 km beneath Africa (source) and red contour line showing 1,000-m elevation increase relative to predicted topography at 30 Ma (from Moucha & Forte, 2011). Seismic tomography reflects thermal and compositional variations, high seismic velocities show roots of Archean Congo craton, while low velocities in east underlie EARS. Note that low velocities in east coincide with 1,000-m contour line and high velocities delineate Congo Basin controlling CAB. Close juxtaposition of zonal circulation cells over eastern Atlantic and western equatorial Africa and Indian Ocean cell along CAB makes East Africa very sensitive to changes in zonal circulation and shrinking of rainforest in Congo Basin influences rainfall over East Africa (Nicholson, 2017). Lakes Albert, Tanganyika, Malaŵi, and Victoria shown for reference. (b, c) Distribution of present and Tortonian (11–7 Ma) vegetation relative to plume-driven elevation change. (b) Present-Day Vegetation of Tropical Africa (Lohmann et al., 2015; New et al., 1999); high elevation in EARS marked by xerophytic (dry) woods and savannah. (c) Proxy-based reconstructed of Tortonian vegetation (Lohmann et al., 2015). Note broad tropical rainforest belt west and south of 500-m elevation increase contour; tropical rain and tropical seasonal forest shrunk mainly in N-S but also in E-W from Tortonian to present. Warm grass corresponds to subdesertic Mediterranean-like open vegetation.

Most of East Africa experiences two rainy seasons during the course of the year. These occur in March to May (long rains) and October to November (short rains). The summer rainfall regime, especially over Ethiopia and South Sudan, is loosely linked to the West African monsoon causing pronounced precipitation at the northwestern margin of the East African plateau (Fig. 3a). The Indian monsoon creates low-level northeasterly flow from November to March and predominantly southerly flow from May to October. However, the peak monsoon months (July to September and December to February) correspond to the dry seasons in East Africa. The great lakes of East Africa are also factors in the regional distribution of rainfall (Trauth et al., 2010). Lake Victoria, for instance, creates a mesoscale circulation system resulting in strong rainfall over the lake. While the mean annual rainfall over its catchment is 1,354 mm, it is ~1,800 mm over the lake itself (Yin, Nicholson, & Ba, 2000). The lake also influences convective activity over the Kenyan highlands (Okeyo, 1987).

The greatest recent changes in East African climate appear to have occurred during the long rains, which have been declining since the early 1980s (Nicholson, 2017). Funk et al. (2013) and Lyon (2014) argued that the tropical Indian Ocean is responsible for this trend. The most direct mechanism by which the sea surface temperature changes modulate East African rainfall appears to be a weakening of the East African Jet and moisture transport into East Africa (Nicholson, 2017).

The Eastern Branch of the EARS is relatively dry, and its large savannas (Fig. 3b) host an extraordinary wildlife. In contrast, the Western Branch has important mountain forests and a rich lake biodiversity. Both the lake systems and the “sky islands” (i.e., isolated mountaintops) of the Western Branch have produced an unusual degree of endemism among the incredible overall biodiversity. Examples include the Rwenzori Mountains (Eggermont, Van Damme, & Russell, 2009; Wronski, Apio, Semwanga, & Hausdorf, 2016) or rift lakes such as Lake Tanganyika and Lake Malaŵi (Salzburger et al., 2014).

The major differences in geography, elevation, and hydrography of the EARS are largely related to its tectonic development. East Africa1 sits above a major mantle upwelling, the African Superswell (Fig. 4), and is breaking apart along the EARS. The African plume is rising up vertically from the core-mantle boundary (at a depth of approximately 2,900 km) and in the upper mantle below the lithosphere starts flowing subhorizontally to the northeast beneath East Africa (Bagley & Nyblade, 2013) (Fig. 5). Nyblade, Owens, Gurrola, Ritsema, and Langston (2000) showed a depression of the 410-km discontinuity in the upper mantle beneath the eastern side of the Tanzania Craton and suggested that a more than 200-km-thick plume head underlies northeast Zambia (Fig. 5). Geophysical and geochemical data demonstrate that mantle upwelling and magmatism developed above an anomalously hot and thus buoyant asthenosphere (Hart, WoldeGabriel, Walter, & Merrtzmann, 1989). This might be the reason why Africa has the highest mean elevation of all continents and why especially the Eastern Branch with its vast volcanic rock accumulations has such elevated rift floors.

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Figure 4. Schematic Cross Section of Dynamic Earth Through Rotation Axis Depicting Two Major Mantle Upwellings Forming Two Antipodal Domes Underneath East-Central Pacific and East Africa (after Courtilliot et al., 2003); Louisville hotspot in South Pacific, Afar in northern EARS, Hawaii in northern Pacific, and Reunion in southern Indian Ocean provide geographic reference. Deep upwellings originate at lowermost mantle boundary (D′′ layer); current convection in lower mantle driven by sinking and disintegrating of cold, denser material at subduction zones adding heterogeneity to lower mantle; downwellings circumventing two large areas centered on roughly antipodal equatorial regions situated under Africa and east-central Pacific where hot, less dense, and seismically slower material (two superplumes) rises. Note that boundary between upper and lower mantle is 660 km discontinuity (boundary defined by phase change in olivine and chemical changes that increase density and seismic velocity by 6–11%; Christensen, 1995); 410 km discontinuity in upper mantle marked by phase change in olivine causing velocity increase of 5–6%; zone between 410 and 660 km discontinuities called mantle transition zone (Ringwood, 1994).

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Figure 5. Block Diagram Showing Northeast-Directed Mantle Flow (Solid Arrows on Vertical Side) Around Tanzania Craton (Dashed Arrows on Horizontal Side) Associated with African Superplume (outline of Tanzania Craton as in Fig. 1); colored vertical side shows compressional seismic wave speed variations in mantle (seismic tomography similar to taking CT scan of Earth); red colors indicate warmer and weaker material characterizing African superplume structure rising from lower mantle under northern Zambia and flowing to northeast beneath eastern Africa; blue and green shades mean colder and stiffer rocks; 410- and 660-km discontinuities shown for reference (redrawn from Bagley & Nyblade, 2013). African Plate moves northward, plume first interacted with EARS in Ethiopia and subsequently with regions further south.

The mantle plume underneath East Africa controls the current topography in East Africa, and topography in turn has a profound impact on climate (Fig. 3b). These relations are harder to decipher in the geological past. Nonetheless, many researchers showed that climate change since about 15 Ma had a significant influence on the floral (Fig. 3c) and faunal evolution and also on the appearance of our first ancestors (Bonnefille, 2010; Bromage & Schrenk, 1999; Potts & Faith, 2015; Trauth et al., 2007; Vrba, 1999; White, 1995). The oldest hominin fossils were discovered in 7–6 Ma deposits in Africa (Brunet et al., 2002; Haile-Selassie et al., 2001; Levin, 2015; Simpson et al., 2015), an age that broadly fits with the chimpanzee-hominin divergence time estimated to be at 8–6 Ma (Langergarber et al., 2012; Patterson, Richter, Gnerre, Lander, & Reich, 2006; Steiper & Young, 2006). However, climatic and environmental factors that have been driving the chimpanzee-hominin split remain poorly known. A challenging problem here is to relate tectonics and climate change to shifts in paleoenvironments and species evolution (Behrensmeyer, 2006; Levin, 2015; Potts, 2007). Because local tectonic and climatic processes control the ecology of any rift valley, faunal patterns may reflect environments within restricted areas rather than regional or global climate.

An additional problem is that most available long-term climate records are based on marine core data from basins proximal to Africa (deMenocal, 2004; Feakins & deMenocal, 2010; Leroy & Dupont, 1994; Shackleton, 1995; Tiedemann, Sarnthein, & Shackleton, 1994), while most of the late Miocene and Pliocene hominins and other fauna occur in continental deposits of the EARS. In response to this issue, researchers increasingly attempt to obtain climate proxy records directly from outcropping deposits (Cerling, Harris, & Leakey, 2005; Cerling, 2014; Kingston et al., 2002). More recently, in order to bypass the problematic characteristics of outcrops (such as weathering, diagenesis, and stratigraphic discontinuity), cores are being drilled into lakes and/or ancient lake sediments to allow capturing of relatively pristine and continuous climate records that can be correlated to fossil records (Cohen et al., 2016; Lyons et al., 2015).

Tectonic Development of the East African Rift System

Continental Rifting Processes

Continental rift zones are sites of lithospheric extension that occurs in response to far-field plate forces (Fig. 6). The extension of the crust is achieved through normal faulting that thins the brittle crust. Where the brittle crust breaks apart continental rifts develop. In mature continental rifts the lower crust is hot and weak and thins via viscous flow (Buck, 1986; Ring, 2014). Dense lithospheric mantle rocks rise upward to replace the thinning crust, enhancing subsidence in the fault-bounded basins. The lithospheric mantle also extends and thins, and the thus created space is filled by uprising hotter asthenosphere, thereby transferring heat to the lithosphere beneath the extending region, reducing rock density and subsequently causing regional, time-dependent uplift over tens of millions of years (Şengör & Burke, 1978). Mantle upwelling may be enhanced by plumes and/or small-scale mantle convection induced by steep thickness gradients at the transition between thinned and non-thinned lithosphere (Ebinger, van Wijk, & Keir, 2013).

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Figure 6. Simplified, Schematic Sketch Showing Changes in Lithospheric Structure During Rifting; note that magma-rich rift (Eastern Branch) and magma-poor rift (Western Branch) are shown in same sketch. Pre-rift lithosphere at either end of cross section modified by faulting and magmatism in rift; extension thins lithosphere causing subsidence; asthenospheric inflow beneath rift replaces denser lithospheric mantle causing long-term uplift, note elevated transition zone between brittle upper crust and ductile lower crust in rift zone caused by thermal perturbation, especially in magma-rich part of rift (after Ebinger & Scholz, 2012).

Continental rift zones are typically made up by a series of asymmetric graben. From their very inception, rift zones show regular along-axis structural segmentation into graben bounded on one or both sides by large-offset border faults (Ebinger & Scholz, 2012). The border faults are flanked by localized uplifts that may rise 3 km above the surrounding regional elevations (Fig. 1), in some extreme cases like the Rwenzori Mountains in the Albertine Rift more than 4 km (Ring, 2008). In general, border fault lengths, rift-flank uplift, and basin dimensions increase with increasing strength of the lithosphere (Weissel & Karner, 1989). This is why young rift segments in strong lithosphere are characterized by long, skinny graben with deep rift lakes bounded by impressive escarpments.

Two principal mechanisms cause high topography in extensional settings: (1) rift-flank uplift and (2) mantle flow triggering a thermal perturbation that causes density changes in the upper mantle. The first mechanism is a flexural response in the footwall of large-scale normal (border) faults that counteracts subsidence of the graben floor in the hanging wall of the normal fault. Kusznir and Ziegler (1992) showed that flexural topography creates localized uplift of the relatively narrow rift flanks. The second mechanism results in far broader uplift that occurs above the underlying mantle thermal perturbation.

Geodynamics of the East African Rift System

The EARS comprises several discrete and diachronous rift sectors (Fig. 1). Field data have shown that the Eastern and Western branches follow weak belts of Proterozoic mountain building and skirt around the strong Archean Tanzania Craton with its thick lithosphere (Ring, 1994). Geophysical imaging shows that the deep roots of the Tanzania Craton play a major role in guiding the upwelling hot mantle emanating from the top of the African Superswell in the upper few hundred kilometers of the mantle (Ebinger & Sleep, 1998) (Fig. 5). The Tanzanian Craton (Fig. 1) is underlain by a broad anomaly characterized by low seismic velocity extending across the 410 km discontinuity down to the transition zone between the upper and lower mantle at 660 km. This anomaly is interpreted to record high temperatures and the presence of melt, consistent with the spreading of a mantle plume head beneath the craton (Adams, Nyblade, & Weeraratne, 2012; Nyblade et al., 2000). Below the transition zone, this plume head probably connects with the African Superswell, a large-scale low seismic velocity anomaly extending upward from the core-mantle boundary (Fig. 4).

The earliest basaltic volcanism in the EARS occurred between 45 and 39 Ma in southwest Ethiopia and northernmost Kenya (Morley et al., 1992) when kimberlites were emplaced in the crust surrounding the rift. The widespread distribution of the kimberlites suggests heating and mantle metasomatism along the asthenosphere–lithosphere boundary (~120 km depth) long (~20 Myr) before any regional extension and surface expression of rifting (Ebinger et al., 2013). Rifting has progressed to incipient seafloor spreading in the Afar depression (Corti, 2009). Further south, the Eastern Branch is superposed on the broad Ethiopian and Kenyan plateaus. The Turkana depression between the two plateaus marks a failed Mesozoic rift system, allowing the possibility that the plateaus are part of one large zone of uplift extending from southern Africa to the Red Sea (Nyblade & Robinson, 1994).

Magmatism of Rift Sectors

Thermal processes can be linked to the degree and timescales of plume involvement in the rifting processes (Ebinger, 2005). Partial melting in rifts is largely controlled by the temperature and volatile content of the underlying asthenosphere, as well as the degree of decompression, which depends on the geometry and rate of lithospheric thinning (White et al., 1987). Partial melting can therefore be used as a proxy of maturity of a rift sector (Ebinger et al., 2013). We follow Ebinger and Scholz (2012) and describe those differences from the most juvenile, primitive stage in the Western Branch and the southern tip of the Eastern Branch to the most evolved stage in the Ethiopian Rift and Afar. The differences in volcanism, uplift and subsidence of the two EARS branches largely reflect the way the mantle plume was channeled underneath East Africa (Fig. 5). The young rift sectors in the Western Branch and the southern part of the Eastern Branch have not been significantly affected by plume-related heating, the entire lithosphere is strong and has not been thinned to any significant extent. The further north one goes, the longer the rift sectors have been affected by plume-related heating—there is more pronounced igneous activity, the lower crust is weak, and rifting is more advanced.

Western Branch

The individual rift basins of the Western Branch are long (~100–150 km) and narrow (~50–70 km). Volcanism occurs in four isolated centers—from north to south, the Toro-Ankole, Virunga, South Kivu, and Rungwe volcanic fields (Fig. 1). Volcanism commenced at 10 ± 2 Ma largely coeval with rift faulting (Ring & Betzler, 1995). There is no trend in ages of the volcanic rocks (Ebinger, 1989). However, rift faulting opened, for instance, the Malaŵi Rift in a zipper-like fashion from north to south since about 5–4 Ma (Ring & Betzler, 1995). Overall, the morphological and magmatic evolution of the Western Branch indicates a strong plate that has been very modestly thinned. Border faults penetrate the entire lower crust, which is consistent with deep seismicity, pronounced basin floor subsidence and deep lakes, as well as significant rift-flank uplift (Fig. 7a).

Tanzanian Divergence Zone of Eastern Branch

The Eastern Branch shows a striking progression of magmatism (and rift evolution) from south to north. The most juvenile rift sector is the Tanzanian Divergence Zone. In contrast to the narrow (~50 km wide) and well-defined Kenyan rift sector, the Tanzanian Divergence Zone is ~300–400 km wide and consists of three separate graben—from west to east, the Eyasi, the Manyara, and the Pangani rifts (Fig. 7b). There are also normal faults in between these three graben, showing how diffuse extension is accommodated. It appears that the Archaean Tanzania craton and a strong lithospheric domain, the Massai block (Fig. 1), largely restricted fault slip, and this resistance caused splaying of the major rift faults of the southward propagating Eastern Branch (Corti, Iandelli, & Cerca, 2013; Ebinger, Djomani, Mbede, Foster, & Dawson, 1997).

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Figure 7. Simplified Lithospheric Cross Sections Through Western and Eastern Rift Branches Showing Different Styles of Rifting, Lithospheric Attenuation, and Magmatism (localities of cross sections see Fig. 1) (modified from Chorowicz, 2005). (a) Albertine Rift of Western Branch; lithosphere hardly thinned and lower crust not ductily deforming; chemistry of magmatic rocks primitive and volume of magmatism very limited; pronounced rift-flank uplift created >5,000-m-high Rwenzori Mountains in Albertine Rift. (b) Tanzania Divergence Zone of Eastern Branch; note wide area affected by rifting caused by strong lithosphere of Masai Block; lithospheric structure similar to (a) indicating early stages of rifting. (c) Central Kenya Rift; lithosphere notably attenuated and thinned resulting in viscous, weak lower crust in which high-angle upper crustal normal faults flatten out; note off-axis volcanism forming large volcanic edifices like Mt. Kenya on rift flank. (d) Main Ethiopian Rift; most advanced stage of continental rifting resulting in strongly thinned lithosphere; style of rifting becomes more symmetric and rifting strongly assisted by pronounced magmatism.

Individual rift basins are half graben bounded by a faulted rift escarpment on one side and a flexural warp on the other (Foster, Ebinger, Mbede, & Rex, 1997). Each basin is ~100 km long and ~50 km wide (Ebinger et al., 1997). The rift basins contain thin (<3 km) sequences of synrift volcanics and sediments. Most rift basins in northern Tanzania are younger than ~1 Ma (Ring, Schwartz, Bromage, & Sanaane, 2005; Schwartz et al., 2012). Volcanic activity began at ~8–4 Ma in the center of the Tanzanian Divergence Zone. After ~4.5 Ma, volcanic activity was transferred to the outer graben (Foster et al., 1997). In the Manyara Rift, volcanism migrated southward with ages of ~4.9–1.5 Ma for the northern Manyara Rift and ~1.5–0.7 Ma for the southern part (Dawson, 2008). Earthquakes are distributed throughout the entire ~35 km of crust (Foster & Jackson, 1998).

There is minor extension across the rift basins, and age data fit into the general pattern for the Eastern Branch, indicating that magmatic activity predated tectonism. The composition of the lavas indicates small volumes of melt beneath a ~140-km-thick lithosphere, compatible with geophysical data showing a few kilometers of crustal thinning and a narrow zone of thinned mantle lithosphere directly beneath the rift zone (Green, Achauer, & Meyer, 1991) (Fig. 7b). All data are compatible with a strong lithospheric plate.

Kenya Rift of Eastern Branch

Further north, the Kenya Rift is distinctly older and more mature than the Tanzanian Divergence Zone. The basins are still asymmetric half graben, extension is of the order of 10 km, and the lithosphere has thinned to ~90 km (Mechie, Keller, Prodehl, Khan, & Gaciri, 1997) (Fig. 7c). Volcanism and rift faulting commenced by ~23 Ma (Morley et al., 1992) in broad downwarps that subsequently became the sites of half graben. The volcanics in the Kenya Rift are voluminous (>100,000 km3; King, 1978). Between 14 and 11 Ma, a period of intense volcanism filled and overflowed the rift basins. Another period of pronounced volcanism between 5 and 2 Ma produced noticeably more evolved lavas, again overspilling the rift basins (Morley et al., 1992). A stunning feature of young volcanism are the large off-rift shield volcanoes such as Mt. Elgon and Mt. Kenya in the central Kenya Rift, as well as Kilimanjaro at the transition zone between the Kenya Rift and the Tanzanian Divergence Zone (Fig. 1).

Earthquakes are restricted to the upper 15 km of the crust (Ibs von Seht, Blumenstein, Wagner, Hollnack, & Wohlenberg, 2001). Seismic and gravity data show a central rift zone of high-velocity, high-density material interpreted as upper crustal dikes. These dikes are probably connected to more than 2 Ma old volcanoes in the central rift, which are cut by a set of prominent, young faults (Ebinger et al., 2013).

All data indicate a higher degree of lithospheric extension in the central Kenya Rift as compared to the Western Branch and the Tanzanian Divergence Zone. Therefore, the lithosphere–asthenosphere boundary is shallower, the volume of decompression melting has increased, and the lithosphere has weakened.

Ethiopian Rift of Eastern Branch

The review of the various rift segments shows that with increasing maturity of the rift sectors the lithosphere thins and the amount of decompressional melt increases (Fig. 6), creating weaker lithosphere. In the EARS these features should be best expressed in the Ethiopian Rift, which links into the Afar Depression with the spreading centers of the Red Sea and Gulf of Aden.

Mohr (1983) estimated that volcanic rocks amount to about 300,000 km3 and produced the vast Ethiopian Plateau. Most of the volcanics erupted between 32 and 21 Ma with a short-lived period of extensive and aerially widespread flood basalts at 31–30 Ma (Corti, 2009). Immediately after the peak of flood basalt emplacement, a number of large shield volcanoes developed from 30 to 10 Ma on the surface of the volcanic plateau (Kieffer et al., 2004). Initial basin formation and minor extension occurred between 24 and 11 Ma (Corti, 2009). The main phase of extension in the Main Ethiopian Rift started at ~11 Ma and is related to the onset of seafloor spreading in the Gulf of Aden and the initiation of extension in Kenya (Corti, 2009). The region of active extension narrowed with time to a ~20 km wide zone near the center of the rift since ~2 Ma. This period of rifting is correlated to seafloor spreading in the Red Sea (Calais et al., 2003).

The crust thins from ~40 km underneath the Ethiopian plateau to 35–24 km underneath the Ethiopian Rift. Crustal thickness is higher in the south (35–33 km) than in the north near Afar (26–24 km). The lithospheric mantle underlying the Ethiopian Rift is either very thin or entirely missing, especially in the northern part of the rift and in Afar (Dugda et al., 2005).

Collectively, these findings indicate that the crust beneath the rifted regions in Ethiopia has been extensively modified by magmatic processes and by addition of mafic rocks to the middle and lower crust (Fig. 7d). The drastic thinning of the lithospheric mantle is probably a combined effect of lithospheric extension and plume-related thermal erosion (i.e., heat causes partial melting of the lithospheric mantle and thereby transforms it into asthenosphere; this transformation is associated with a density decrease of about 100 kg/m3). Beneath Afar, the mantle structure is much akin to that of mid-ocean ridge systems, and magmatism and faulting are focused to the central rift axis (Corti, 2009).


Rift faults started to form at ~24–23 Ma, with major extension since ~14 to 11 Ma in Ethiopia and Kenya and at ~12–10 Ma in the Western Branch (Corti, 2009; Ebinger, 1989; Ebinger et al., 2000). These ages show the importance of the 14–10 Ma time window for rift extension. At ~0.4 Ma a major tectonic reorganization occurred, and extension became more oblique, resulting in pronounced strike-slip faulting and local uplift (Ring, Betzler, & Delvaux, 1992; Strecker & Bosworth, 1991). Rift-flank uplift at ~0.4 Ma was important and responsible for the formation of a long-wavelength depression hosting Lake Victoria in between the two EARS branches (Ebinger, 1989).

The high degrees of magmatism in the Kenyan and Ethiopian rifts are related to the underlying mantle plume. The timing of magmatism shows that the African Plate slowly moved northward above the stationary mantle plume (Calais et al., 2003). Magmatism and rift maturation progressed southward, first affecting Ethiopia and Kenya, the Western Branch, northern Tanzania, and finally southern Tanzania (Fig. 5).

The EARS is not only an archetypal rift system, it also seems to be a typical example of an active rift (Şengör & Burke, 1978). In active rifts rising mantle plumes cause regional updoming and thermal erosion of the lithospheric mantle, and these processes are the major drivers initiating plate divergence. There is a distinct lag time between the rising plume and the ensuing surface processes (e.g., elevation change, faulting, volcanism). Plume-related processes also drive topography primarily through density changes, and a fascinating corollary of this is that the plume-driven topography may have been an important cause for climate change in East Africa, thereby influencing the emergence and evolution of mankind.

Climate Change in East Africa

The climate variation over the past ~1.3 Myr has been reviewed by Johnson (see the article “Quaternary Climate Variation in Eastern Africa”), showing, for instance, dramatic lake-level fluctuations in Lake Malaŵi between 1.3 and 0.8 Ma. Here we focus on the less well-known time before 1.3 Ma since about 15 Ma. In a generalized, technical attempt, we separate this period into Miocene cooling, the Pliocene warm (wet) period and then late Pliocene/Pleistocene cooling. We also discuss discrepancies between global climate events and local, mesoscale rift climate variability. Most of the arguments are based on floral and faunal evidence.

Miocene Cooling

Globally, the Neogene was a period of long-term cooling and increased climate variability. Based on δ‎18O records in the Atlantic Ocean, Miller, Mountain, Wright, and Browning (2011) proposed an important cooling period at 15.5–12.5 Ma, foreshadowing the first glaciations in the Northern Hemisphere by 10–6 Ma (Jansen, Sjoholm, Bleil, & Erichsen, 1990), and a second major cooling stage at 2.95–2.52 Ma. The middle Miocene cooling preceded a first expansion of dry-adapted C4 (some tropical grasses, sedges) grasslands in tropical Africa at the demise of C3 plants (mostly trees) (Barboni, 2014; Cerling, 2014; Cerling et al., 2011) and is therefore considered to be a period characterized by the gradual replacement of subtropical forested environments by more seasonal, open country “savanna-mosaic” habitats allowing large space occupied by grasses in East Africa (Bonnefille, 2010) (Fig. 3c). Stable isotope data demonstrate that the proportion of C4 biomass in East Africa gradually increased since about 9.4 Myr (Cerling, 1992; Cerling et al., 1997). In contrast, tropical West Africa west of the Congo Air Boundary had a vegetation cover with high tree density throughout the middle/late Miocene.

The best exposed middle Miocene sediment sections in the EARS are in the Baringo Basin of the Kenya Rift (Fig. 1). At about 15 Ma, δ‎13C values of tooth enamel of rhinocerotids and elephantids suggest a minor dietary component of C4 grasses (Morgan, Kingston, & Marinoa, 1994). A 13-Ma-old pollen assemblage contains highland forest tree taxa (Bonnefille, 1995), and δ‎13C data of palaeosols indicate a strong C3 signal (Cerling, 1992). These data are in line with Jacobs and Kabuye (1987), who documented that environments in the Baringo Basin were mostly tropical moist or wet forest in the tropical lower montane or premontane forest category during the 15.5–12.5 Ma period. The grass macrofossils and pollen from the Baringo Basin are contemporaneous to the greatest grass pollen abundance in the Gulf of Aden at 12–11 Ma ago (Bonnefille, 2010). Rodent and primate assemblages from the middle Miocene indicate faunal stability in the Baringo Basin in a relatively wet environment characterized by C3 plants (Hill, Leakey, Kingston, & Ward, 2002; Hill et al., 1991; Winkler, 2002).

Bonnefille (2010) summarized data from fossils leaves, twigs, and fruits at three sites in the Baringo Basin for estimating paleorainfall. At about 12.6 Ma, the assemblages indicate the presence of lowland to submontane forests with a predicted rainfall of about 1,000–1,200 mm/a during the middle Miocene (modern precipitation in the same area averages at 900 mm/a at 1,300 m of elevation). At 10–9 Ma, the data suggest more open vegetation with a pronounced dry season and a predicted rainfall of approximately 500–700 mm/a. At the youngest site (7.2–6.8 Ma old), a woodland/dry forest would have existed with an estimated rainfall of 900–1,000 mm/a. Peaks of grass pollen occurred at 10.5 Ma and 7 Ma Bonnefille (2010). Overall, all data support a change toward more arid C3-dominated habitats in the Kenya Rift since about 10 Ma.

The next key interval in climate evolution is the 8–5 Ma period in the late Miocene when a dramatic change in African biota took place. Shrinkage of the equatorial forests coincided with the expansion of C4 plants (Cerling et al., 1997, 2005). The highest abundance of C4 biomass was reached between 8.5 and 6.5 Ma (Kingston et al., 1994), whereas δ‎13C data from tooth enamel show that the shift from C3 to C4 dominated diet in equids and elephantids took place at around 8 to 7 Ma (Cerling et al., 1997). After 7 Ma, the proportion of tree pollen declined significantly, reaching minimum values at about 6.5 to 6 Ma (Bonnefille, 2010). By 6.5 to 6 Ma there was the minimum tree cover density over the entire African tropics.

Climate change at 8 to 5 Ma resulted in the emergence of the faunal elements that would dominate the late Cenozoic to the present. Sediments dated between ~7.5 and 6.2 Ma in the Baringo Basin show the earliest change toward modern faunas. Associated with the fauna are fossil wood fragments that represent six taxa (Kingston et al., 2002). Living representatives of the fossil taxa show fairly limited environmental tolerance, suggesting that the fossil forest grew in a lowland or upland forest having floral affinities to West and Central Africa. Maximum tree heights were >50 m, a height typical for wet or moist forest communities (Menaut, LePage, & Abadie, 1995) (drier African forests and montane forests have tree heights <35–40 m; Richards, 1996). In the Turkana Basin of the Eastern Branch, two very well-developed luvisols document a period of increased aridity between ~6.7 and 5 Ma and depositional stasis at ~6.5–5.2 Ma (Wynn, 2004). Vegetation throughout the interval appears to have been a mosaic of floodplain savannas dissected by gallery woodland.

In contrast to the trend of increased dryness in the Kenya Rift, an extensive rift lake, Paleolake Obweruka, formed by 8 Ma in the Western Branch (Pickford, Senut, & Adoto, 1993) and occupied an area from the northern end of present Lake Albert to the southern end of Lake Edward. Paleolake Obweruka was comparable in size and depth to present Lake Tanganyika. The principal drainage into the lake was from the Kenyan Highlands to the east, while its outlet was to the Atlantic Ocean. Paleolake Obweruka persisted for about 5.5 Ma. The regional surroundings of Paleolake Obweruka consisted of humid biotopes, including tropical forests and forest-savannah mosaics typical of the Congolian district (Pickford et al., 1993). Strong enrichment of iron in thick lateritic soils, vertical and lateral transport of dissolved iron by groundwater, and its precipitation at the groundwater-lakewater chemocline is related to a wet tropical climate when Paleolake Obweruka formed (Roller, Hornung, Hinderer, & Ssemmanda, 2010). However, a drier Acacia woodland is indicated at about 6 Ma (Bonnefille, 2010). Overall, there is an evolution from semi-arid climate before ~8 Ma to a wet climate with times of high seasonality after Paleolake Obweruka had formed. The distinctly wetter late Miocene climate in the Albertine Rift as compared to the Baringo Basin in the Eastern Branch may indicate (1) an end-member case of an overall diversified and heterogeneous climate in East Africa, (2) that the Western Branch was wetter than the Eastern Branch, or (3) the importance of local rift mesoclimate.

Bonnefille (2010) argued that paleobotanical data suggest that at least since the middle Miocene a diversified and heterogeneous landscape with different types of plant ecosystems and vegetation types existed and environments ranged from evergreen and deciduous forests to wet and dry woodlands. This is in part also documented by the change from wet to arid conditions in the Kenya Rift in the middle to late Miocene, while in the Albertine Rift of the Western Branch a change from semi-arid conditions to wet and variable climate resulted in the formation of Paleolake Obweruka by ~8 Ma.

Overall, faunal assemblages during the 8–5 Ma period are interpreted as transitional, reflecting replacement of middle Miocene communities by more modern and advanced populations representative of contemporary African communities (Kingston et al., 2002). In East Africa this climate shift also resulted in the emergence of hominis. Arguably the major evolutionary breakthrough of human evolution was the origin of bipedal locomotion between 8 and 6 Ma. By ~5 Ma faunas were distinctly different to the middle Miocene ones in East Africa (Hill, 1999). These profound changes are primarily linked to the gradual replacement of subtropical environments by more seasonal open country “savanna-mosaic” habitats considered characteristic of the latest Miocene driven by the gradual change to cooler and dryer climate (Cerling, 2014; Cerling et al., 1997). Globally, the final phase of this climate trend in East Africa coincided with the Messinian salinity crisis at 5.9–5.5 Ma (Duggen et al., 2003). However, it should be kept in mind that Miocene fossil remains are rare in the EARS, and the tempo and mode of change during this important transitional period is therefore poorly known (Kingston et al., 2002).

Pliocene Warm Period

The next broad climate epoch in East Africa was the Pliocene warm (wet) period between 5.3 and 3.3 Ma (Tiedemann et al., 1994). It largely coincides with the development of strong E-W zonal (Walker) circulation. This tropical climate reorganization commenced between 4.5 and 4 Ma and was characterized by altered surface water gradients and ocean circulation, possibly connected to restriction of the Panamanian and Indonesian seaways (Cane & Molnar, 2001; Molnar, 2008; Pfister, Stocker, Rempfer, & Ritz, 2014). Global surface temperature was approximately 3°C warmer than at present, and atmospheric CO2 concentrations were about 30% higher (Ravelo, Andreasen, Lyle, Lyle, & Wara, 2004). The Pliocene warm period had a pronounced impact on lacustrine conditions in the EARS. Faunal and climatic proxies for the Pliocene warm period show wet conditions coinciding with extensive highstands of Lake Tanganyika and Paleolake Obweruka and the formation of Lake Malaŵi in the Western Branch, as well as Paleolake Lonyumun in the Turkana Basin and Paleolake Turasha at about 4.5–4.2 Ma in the Kenya Rift of the Eastern Branch (Fig. 8a).

Tectonic Dynamics in the African Rift Valley and Climate ChangeClick to view larger

Figure 8. Paleogeographic Maps of EARS Lakes Showing Hydrographic Configurations of Major Lakes and Rivers at 4.5–4 Ma and 2–1.8 Ma (after Salzburger et al., 2014); solid blue lines indicate major perennial rivers; dashed blue lines show intermittent rivers; Lamu marine embayment in Turkana depression controlled by Cretaceous rift.

A shift to lacustrine conditions at ~4.5 Ma in the Kenya Rift is marked the formation of Paleolake Turasha in the central Kenya Rift and Paleolake Lonyumun that occupied much of the Turkana Basin further north (Feibel, 1993, 2011). In the Ethiopian Rift, stable isotope analysis of pedogenetic carbonates provided evidence of 30–70% C4 plants (Woldegabriel, Ambrose, Barboni, Stewart, & White, 2009). In contrast, abundant fossil wood suggests a C3 woodland setting (White et al., 2009). In the Eyasi Rift at the southwestern end of the Eastern Branch, grass pollen data suggest a woodland or riverine forest along an ancient river drainage with dry open grassland (Bonnefille & Riollet, 1987). Tree pollen frequencies increased significantly from 6 to 3.6 Ma, and forests expanded in the early Pliocene (Bonnefille, 2010).

The northern Malaŵi Rift and the Albertine Rift of the Western Branch contain Pliocene sediment sections yielding climate proxies. In the early Pliocene of northern Malaŵi Rift, a pronounced transgression (formation?) of Lake Malaŵi occurred by 4.5–4 Ma and lasted until ~3.7 Ma (Betzler & Ring, 1995). The fauna in the northern Malaŵi Rift during this period is dominated by a high percentage of terrestrial species (~90%; Sandrock et al., 2007), and the assemblages suggest a tropical, semi-arid bushland or tropical grassland environment. Toward the end of the Pliocene, Lake Malaŵi retreated, a fluviatile system was re-established in the northern Malaŵi Rift (Betzler & Ring, 1995), and the faunas indicate a strong trend toward arid grassland that persisted into the Pleistocene (Sandrock, Kullmer, Schrenk, Juwayeyi, & Bromage, 2007). Lüdecke et al. (2016) reported a long-term Plio-Pleistocene δ‎13C record from pedogenic carbonate and suidae teeth. Consistent δ‎13C values of −9‰ indicate a C3-dominated closed environment with regional patches of C4 grasslands. The overall fraction of woody cover of 60–70% reflects more forest canopy in the Malaŵi Rift than in the Eastern Branch in the Pliocene (Lüdecke et al., 2016).

Fossil wood from Paleolake Obweruka shows that both wet forests and woodland were present from 5 to 4 Ma. The wettest forest is shown at a single locality dated 4.2–4 Ma (Bonnefille, 2010). Overall, in the Western Branch the past vegetation was spatially diverse and showed significant changes between drier and wetter types during the 6–4 Ma time interval. Stable isotope proxy data from tooth enamel of hippopotami suggest substantial but variable amounts of C3 biomass implying wet and variable climate before ~3.5 Ma in the Albertine Rift (Brachert et al., 2010).

Overall it appears that wet conditions kicked in by ~4.5 Ma. In the Eastern Branch there seems to be a mix of C3 and C4 environments, whereas the Western Branch appears to show a stronger trend toward C3 biomass. A first abrupt decline in the spread of forests in East Africa is observed at 3.9 Ma and became pronounced after 3.5 Ma (Bonnefille, 2010).

A second main evolutionary step in human evolution at the end of the Pliocene warm period was the inception of cultural evolution at 3.3 Ma at the end of the warm period. Harmand et al. (2015) reported stone tools from the Turkana graben dated at 3.31–3.21 Ma. This is broadly coeval with animal bones from Ethiopia, which have stone-inflicted cut marks, dated at ~3.39 Ma (McPherron et al., 2010). The youngest Pliocene hominin fossil at 2.8–2.75 Ma is from the northern Ethiopian Rift and represents the so-far-earliest appearance of the genus Homo (Villmoare et al., 2015). The data appear to show that the end of warm and wet conditions in the Pliocene at ~3.3 Ma caused a major evolutionary step in human evolution and ultimately led to the appearance of Homo at ~2.8 Ma.

Late Pliocene/Pleistocene Cooling

The most significant global climate event at the end of the Pliocene was the onset of significant Northern Hemisphere glaciation and a shift in the dominant period of climate oscillation, both of which commenced by ~2.8 Ma, as indicated by general δ‎18O trends (Bartoli, Hönisch, & Zeebe, 2011; Martinez-Boti et al., 2015). The enrichment in δ‎18O values of marine benthic foraminifera beginning in the mid-Pliocene indicates decreased temperatures, preferential evaporation of the lighter isotope 16O, and its retention in spreading continental ice sheets. Global cooling and an associated aridification trend in East Africa is marked by several punctuations at 3.5–3.35 Ma, 2.5–2 Ma, and 1.8–1.6 Ma (Wynn, 2004). Wet phases prior to 2.8 Ma occurred every 400 ka; afterwards wet phases appear every 800 ka, with periods of precessional-forced extreme climate variability at 2.7–2.5 Ma, 1.9–1.7 Ma, and 1.1–0.9 Ma (Trauth et al., 2007; Wynn, 2004). The changes at ~2.8 Ma reflect a shift from orbital precession (23 ka) to obliquity (41 ka) as the overarching determinant of variability in solar heating (insolation), nearly concurrent with the shift in the δ‎18O record (Tiedemann et al., 1994).

A second step in the development of the Walker circulation occurred between 2 and 1.5 Ma, established strong circulation, a steeper sea surface E-W temperature gradient across the Pacific, and overall initiation of the modern tropical climate system (Ravelo et al., 2004). Note that this second step in the evolution of the Walker circulation is also not temporally linked with the onset of Northern Hemisphere glaciation and commenced about 1 Ma later.

By ~1 Ma, glacial-interglacial cycles intensified—the mid-Pleistocene revolution (Berger & Jansen, 1994)—and low-amplitude 41 ka obliquity-forced climate cycles of the earlier Pleistocene were progressively replaced by high-amplitude 100 ka cycles. Overall, while the onset of significant Northern Hemisphere glaciation occurred as subtropical conditions in East Africa began to cool, changes in the Pliocene tropical climate system were to some degree independent (Ravelo et al., 2004).

Trauth et al. (2010) showed that the development of amplifier lakes in the Eastern Branch at about 3–2.8 Ma significantly contributed to the exceptional sensitivity of the Eastern Branch to climate change compared to elsewhere on the African continent. Amplifier lakes are small, high-elevation, high-precipitation, high-evaporation lakes in rift graben with steep walls and flat graben floors forcing distinct climate gradients between high rainfall on the rift flanks and extreme heat/high evaporation in the rift basins. Amplifier lakes respond rapidly to moderate climate shifts (Trauth et al., 2010) and thus emphasize the importance of local rift mesoclimate.

In general, East Africa became more arid and cooler after ~2.8 Ma. Forest fragmentation appears to have been important in speciation and diversification. At 2.7 Ma, a pronounced decline in tree pollen density indicates an abrupt retreat of the tropical rainforest in East Africa (Bonnefille, 2010). Vrba (1988, 1999) preceded a major turnover pulse in EARS mammals at ~2.8 Ma. In recent studies, faunal turnover events were established to have occurred around 2.8 Ma at Afar (DiMaggio et al., 2015) and around 2.0–1.75 Ma on a larger African scale (Bibi & Kiessling, 2015). These turnover events took place during periods of extreme climate variability and important lake highstands (Fig. 8b).

Climate Forced by Tectonics?

Links Between Tectonics and Climate

A major faunal shift in East Africa took place between 8 and 5 Ma and postdated cooling at 15.5–12.5 Ma, widespread volcanism, and major rift faulting in the Kenyan and Ethiopian rifts at ~14–11 Ma and the formation of the Western Branch at ~12–10 Ma. There is a temporal coincidence between middle Miocene cooling and widespread rift faulting. The formation of the Western Branch had a profound impact on East African lacustrine evolution as at least two major long-lived lakes (Tanganyika and Obweruka) formed in the middle to late Miocene. The Pliocene warm (wet) period had a drastic impact on lacustrine conditions, and more large lakes formed by 4.5–4 Ma (e.g., Malaŵi, Lonyumun, and Turasha) (Fig. 8a). However, at this time there is no known tectonic driver for climate change. Nonetheless, the formation of lakes Lonyumun and Malaŵi and the important lake stage of Paleolake Obweruka at 4.5–4 Ma coincided with the first step in the development of strong Walker circulation. If this first step in Walker circulation was indeed related to processes restricting and finally closing the Indonesian and Panamanian seaways, then global tectonic processes can be linked to major shifts in climate and hydrography in East Africa in the early Pliocene.

The formation of lakes Malaŵi, Turasha, and Lomuyun demonstrates the importance of the Pliocene warm (wet) period for the hydrography of the Eastern Branch. After the Pliocene warm period numerous (small) amplifier lakes formed in large parts of the high-elevation Eastern Branch during the post-3–2.8 Ma period of intensified climate variability (Trauth et al., 2010). In general, the Eastern Branch basins remained much smaller and less stable for developing faunas comparable to those of the tectonically controlled deep rift lakes of the Western Branch. This demonstrates the overarching role of tectonics in climate change and biodiversity evolution in East Africa as the different tectonic characteristics are primarily controlled by differences of how the mantle plume interacted with the lithosphere. As shown by Brachert et al. (2010), the large water bodies entrenched between uplifted rift flanks developing in the strong lithosphere of the Western Branch had a profound impact on local rift climate by creating a valley air circulation system controlled by the deep and elongated rift basins filled by large lakes. The dominating role of local rift mesoclimate may have been especially important for the uplifting Rwenzori Mountains at and after 2.5 Ma. The amplifier lakes of the Eastern Branch also stress the significance of local climate effects.

Major tectonic changes at and after ~2.5 Ma involved not only uplift of Rwenzori Mountains but also an important stage of faulting and volcanism along the central rift axis in Ethiopia (Corti, 2009). Faulting and volcanism in Ethiopia has been linked to sea floor spreading in the Red Sea (Calais et al., 2003), the latter of which may have also caused a pulse of extension and faulting in other parts of the EARS, which in turn would have resulted in enhanced uplift in the footwalls of major normal faults (e.g., Rwenzori Mountains) (Ring, 2008).

A final tectonic reorganization at ~0.4 Ma caused local strike-slip faulting and rift-flank uplift in the EARS (Strecker & Bosworth, 1991; Ring et al., 1992). This event ultimately resulted in the formation of Lake Victoria because river outflow to the west was blocked off by rift-flank uplift (Ebinger, 1989). At about this time the Rwenzori Mountains had another major uplift phase, and glaciers started to grow on top of the mountains (Kaser & Ostmaston, 2001). The development of the Lake Victoria cichlid superflock after the last glacial maximum at ~12.5 ka shows hydrographic connections between Lake Victoria and the Western Branch lakes Albert, Edward, and Kivu (Elmer, Reggio, Wirth, Salzburger, & Meyer, 2009). Presently, there is more limited hydrographic connectivity between the Western Branch lakes and Lake Victoria via the Nile River (Schultheiß, van Bocxlaer, Riedel, von Rintelen, & Albrecht, 2014).

Possible Explanations for Climate Change in East Africa

What drove climate change in the middle to late Miocene ultimately controlling the faunal turnover at 8–5 Ma and the emergence of our ancestors at 7–6 Ma? Paleoanthropologists and geologists commonly argue that climate change causing a change from tropical forest to savanna-type habitats (shift from C4 to C3 environments) were the main drivers of the chimpanzee-hominin split. A review of Cenozoic vegetation and climate by Bonnefille (2010) concluded that an expansion of savanna/grassland occurred by about 10 Ma in East Africa as a consequence of the sustained cooling period at 15.5–12.5 Ma. Another pronounced change in vegetation took place at 6.3–6 Ma and was marked by a decrease in tree cover across entire tropical Africa. We discuss the relative roles of the developing topography in East Africa and global tectonic events.

In general, topography in East Africa was caused by both rift-flank uplift and mantle flow–related uplift. Sepulchre et al. (2006) suggested that “massive uplift of eastern African topography” led to pronounced reorganization of atmospheric circulation and caused increased aridification in East Africa. Sepulchre et al. (2006) argued that rift-flank elevation was in the range of 1,500–5,200 m. These heights are true for the current EARS. However, the development of pronounced and localized topography of the rift flanks has not been dated, and paleoelevation is largely unconstrained. Current work strongly suggests that pronounced rift flank uplift occurred much too recently to explain middle to late Miocene climate change (Bauer et al., 2010, 2013, 2016; Foster & Gleadow, 1993; Ring, 2008, 2014).

A more significant factor that may have influenced East African climate in the middle to late Miocene appears to be the topography exerted by mantle flow related to the African Superswell (Figs. 3 and 5). Moucha and Forte (2011) designed a numerical simulation based on flow pattern changes in the mantle that allows reconstructing the amplitude and timing of topographic uplift above the African Superswell through time. Their results suggest that geographically extensive (~5 x 106 km2) topographic uplift in excess of 500 m of elevation started to develop by 15 Ma and became pronounced at 10 Ma. Wichura, Bousquet, Oberhänsli, Strecker, and Trauth (2010) documented elevated topography in the Kenya Rift by 13.5 Ma and Gani, Gani, and Abdelsalam (2007) between 10 and 6 Ma on the Ethiopian Plateau. Major surface uplift of the East African plateau is probably causally connected with extensive volcanism in the Kenya Rift by 14–11 Ma and the surge of rift faulting in Kenya and Ethiopia at 14–11 Ma. However, as noted by Pik, Marty, Carigman, and Lave (2004), the Ethiopian plateau started to be uplifted by 30 Ma as a result of the widespread volcanism. Sembroni, Faccenna, Becker, Molin, and Abebe (2016) speculated that by that time an about 800-m-high small dome existed in the area that is now the Ethiopian plateau.

These findings make it likely that topography in East Africa indeed influenced regional climate and a gradual shift toward colder and more arid conditions by 15–10 Ma. However, numerical model simulations are no proof, and data constraining the timing of topography are needed to better evaluate tectonic drivers of climate change in Africa in the middle to late Miocene. Furthermore, a deeper understanding of the cause(s) for climate shift(s) in Africa also requires considering global drivers like the Asian monsoon and East African jet. The combination of isotopic shifts, expansion of grassland at the expense of forests, and increased aridification concur with a shift near 10 Ma in Asian climate toward the present monsoonal pattern (Molnar, Boos, & Battisti, 2010) largely concurrently with the increased abundance of C4 grasses in Africa (Cerling, 2014). One of the strongest arguments for a strengthening of the Indian monsoon near 10 Ma comes from ocean drilling in the Arabian Sea of eastern Indian Ocean (Kroon, Steens, & Troelstra, 1991). If monsoon winds did strengthen, sea surface temperatures would have dropped if the thermocline at ~10 Ma was deeper than it is today, as Philander and Fedorov (2003) suggested. Colder sea surface temperature in the Indian Ocean off the East Africa coast would have deflected moisture transport away from East Africa toward southeast Asia, which collectively may have cooled East Africa and caused a shift toward arid conditions. If so, the East African jet would have developed in the late Miocene. A final element in the global climate mix might be the glaciers that started to grow in the Northern Hemisphere by 10–6 Ma.

The Pliocene warm period appears to be associated with the initial development of strong Walker circulation. The youngest profound change in African climate took place in the late Pliocene. Geo- and biochemical datasets as well as climate models have shown that East African aridification is primarily controlled by Indian Ocean sea surface temperature (Cane & Molnar, 2001) and that precessional variations in C3 and C4 plants have been induced by changes in monsoonal precipitation driven by changes in low-latitude insolation (Philander & Fedorov, 2003). These findings imply that East African climate change since ~3–2.8 Ma has largely been governed by ocean–atmosphere interactions in the low latitudes and the East African Jet intensified in the late Pliocene. Cane and Molnar (2001) argued that changes in surface–ocean circulation, driven by the final closing of the Indonesian seaway at ~4–3 Ma, were responsible for the increased aridification of East Africa. According to their model, the northward motion of Australia and the associated northward displacement of New Guinea switched the source of flow through Indonesia from warm South Pacific to relatively cold North Pacific waters. This in turn decreased sea surface temperatures in the Indian Ocean, leading to reduced rainfall over East Africa (Cane & Molnar, 2001). The Cane and Molnar hypothesis can be seen as a novel alternative to the commonly held opinion—that the onset of significant Northern Hemisphere glaciation was the main driver for aridification in East Africa (deMenocal, 2004).

The second-order climate changes in the Pleistocene, namely the wet phases at 1.9–1.7 Ma and 1.1–0.9 Ma coincide with the time for intensification of the Walker circulation (1.9–1.7 Ma) and the mid-Pleistocene revolution (1.0–0.7 Ma) (Trauth et al., 2007). High-latitude forcing is required to compress the Intertropical Convergence Zone so that East Africa becomes locally sensitive to precessional forcing, resulting in rapid shifts from wet to dry conditions (Maslin, Shultz, & Trauth, 2015; Trauth et al., 2007).

In general, this review emphasizes the importance of the ~10 Ma time window for overall East African tectonics and climate. The Western Branch started to form by ~10 Ma, and sea floor spreading in the Gulf of Aden caused a surge of extensional deformation in the Ethiopian and Kenya Rift. Our discussion also suggests that rift structure and architecture is probably less important for regional East African climate change. The latter appears primarily controlled by plume-related surface uplift and the strengthening of the Asian monsoon and East African jet, both of which made East Africa drier. Regional surface uplift also largely controlled the development of the EARS and the pronounced differences in lake size, depth, and biodiversity between the Eastern and Western Branch lakes. However, rift structure is important for local climate. The commencing uplift of the Rwenzori Mountains at ~2.5 Ma caused reversals in river flow, capturing of catchments, and biogeographic turnover (Brachert et al., 2010). However, a growing rain shadow is not obvious in δ‎18O signatures of hippopotami teeth. This may be the result of the overriding effect of evaporation on δ‎ 18O responding to aridification of the basin floor by a valley air circulation system through relative deepening of the valley. On the other hand, a synchronous arid pulse is not clearly recorded in palaeosol data and mammalian fauna of the Eastern Branch. This discrepancy indicates that rift mesoclimate may represent an underestimated aspect in previous paleoclimate reconstructions from rift valley data.

An important point about climate drivers is that we often tend to think of one single driver for explaining climate change (i.e., Northern Hemisphere glaciation, the Indonesian throughflow, East African topography, or the Asian monsoon). However, it might well be that a mix of all or some of these processes controlled East African climate.


We have argued that increased surface uplift above the African Superswell between 15 and 10 Ma can be broadly linked with pronounced changes in East African climate causing a general trend from C4 to C3 environments. Important for East African climate at this time is also the strengthening of the Indian monsoon. A first significant step in the development of Walker circulation seems responsible for the Pliocene warm (wet) period in East Africa. There is the possibility that tectonic processes restricting the Indonesian and Panamanian seaways were important for this Pliocene stage in climate change. Significant cooling commencing at 3–2.8 Ma again seems to be connected to changes in the Indian Ocean and a drop in sea surface temperatures caused by the closing of the Indonesian throughflow and significant Northern hemisphere glaciation. At about 2.5 Ma, major tectonism and volcanism in Kenya and Ethiopia as well as the commencing uplift of the Rwenzori Mountains in the northern Western Branch can be linked to sea floor spreading in the Red Sea and were important processes controlling local rift mesoclimate. A significant tectonic reorganization at ~0.4 Ma coincides with the formation of Lake Victoria. Because the Lake Victoria basin is tectonically connected to rift shoulder uplift of the two EARS branches, its formation suggests enhanced uplift by 0.4 Ma. In the EARS, hominin fossils are largely found in the more arid Eastern Branch. In contrast, lake biodiversity is much more pronounced in the Western Branch with its great subsidence and wetter climate. The tectonic differences between the Western and Eastern branches reflect different plate strength, largely controlled by mantle plume activity.


An anonymous reviewer and Martin Claussen provided valuable suggestions for improving this article. Reuben Hansman helped with designing some figures.


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                                                                                                                                                                                                                                                                                                                    (1.) East Africa is geographically largely defined as in Johnson (Quaternary Climate Variation in Eastern Africa).