Northeast African vegetation change over 12 m.y.

C₄ plants presently thrive in low- and mid-elevation sites in tropical climates with high temperatures and warm-season precipitation (Sage, 2004); however, this was not always the case. Today, the C₄ pathway appears in >19 families; most are grasses, followed by sedges and some dicotyledons (e.g., in the Amaranthaceae family). Evidence for C₄ evolution and expansion comes from rare plant macrofossils (Tidwell and Nambudiri, 1989), modern genetic population distributions (Edwards and Smith, 2010), and most commonly from stable carbon isotope records around the world (e.g., Cerling et al., 1997). These various lines of evidence indicate that C₄ plants appeared in the Oligocene but flourished only later in the Miocene. Reasons for this expansion are currently debated (see Edwards et al., 2010) but the ecological implications of C₄ plants and their rise to dominate some ecosystems is important for understanding terrestrial biospheric systems and the evolution of grasslands.

In northeast Africa (area shown in Fig. 1), there has been intense interest in reconstructing vegetation associated with the fossil record of faunal evolution preserved in the East African Rift system. In particular, questions center on the extent of forest versus grassland habitats (Cerling et al., 2011) and the incorporation of C₄ plants into diet (Uno et al., 2011). Palynological data indicate the presence of extensive grasslands during the Miocene and Pliocene (Bonnefille, 2010); however, because C₃ and C₄ grasses cannot be distinguished on the basis of pollen morphology, increases in C₄ biomass have to be assessed by isotopic means (e.g., Cerling et al., 2011).

C₄ photosynthesis has a distinct carbon isotopic fractionation relative to C₃ plants using the Calvin cycle alone, such that isotopic measurements can identify the appearance of C₄ vegetation in the geologic record. Carbon isotopic analyses have been applied to a range of substrates, including carbonates in paleosols (Wynn, 2000), the fossil teeth of mammalian herbivores (Uno et al., 2011), and plant leaf waxes in sedimentary deposits (Schefuss et al., 2011). The carbon isotopic compositions of pedogenic carbonates (δ¹³C_pc) capture local variations in the distribution of woody cover where paleosols form (Cerling et al., 2011), and have been used to show Late Miocene expansion of C₄ grasslands in some regions (e.g., Quade et al., 1989). The carbon isotopic composition of fossil teeth (δ¹³C_tooth) also provides perspective on C₄ distributions, but through the filter of animal dietary preferences. Early studies of δ¹³C_tooth records suggested that C₄ vegetation expanded in a step shift between 8 and 6 Ma (Cerling et al., 1997), but subsequent studies have shown that a shift toward a C₄-dominated diet occurred as early as 9.9–9.6 Ma in equids, followed by rhinocerotids at 9.6 Ma, hippopotamids and bovids at ca. 7.4 Ma, and suids between 6.5 and 4.2 Ma (Uno et al., 2011). The staggered timing of dietary shifts reveals ecological complexity but leaves open intriguing questions about the nature and timing of C₄ expansion in the landscape.

With the development of compound-specific isotopic analysis and its application to plant leaf waxes, plant leaf wax carbon isotopic composition (δ¹³C_wax) data have been reported from lake cores (e.g., Sinninghe Damsté et al., 2011) and marine sediments (e.g., Schefuss et al., 2011) spanning the last glacial cycle. In this paper, we extend this approach to much older marine sediments from Deep Sea Drilling Project (DSDP) Site 231 in the Gulf of Aden in northeast Africa (Fig. 1). Marine sediments downwind from northeast Africa provide a regionally integrative and continuous archive of the terrestrial environment. Terrigenous proxies including dust (deMenocal, 1995), pollen (Bonnefille, 2010), and plant leaf waxes (Feakins et al., 2005) have already provided evidence for environmental change from this site.

Here, we extend our initial report of δ¹³C_wax values from DSDP Site 231 across 11.3–1.4 Ma (Feakins et al., 2005) and provide a novel synthesis with pollen data from the same marine sediments. The combined isotopic and palynological approach enables differentiation of C₃ versus C₄ contributions to grass pollen and also allows us to evaluate how different types of C₃ ecosystems (rainforests versus seasonally dry woodlands) influenced the carbon isotope record. The continuous record of marine sedimentation permits direct comparison of the magnitude of Miocene and Pliocene C₄ expansions. We further compare the marine archive to soil carbonates in the East African Rift in order to understand the distribution of vegetation change.

We report a 12 m.y. leaf wax record of vegetation change from marine sediments of DSDP Site 231. In an earlier publication (Feakins et al., 2005), we presented 97 analyses from discrete time slices. Here we add 255 additional analyses extending upon these time windows with samples that integrate 2–3 k.y., as well as samples that integrate 300 yr during the Miocene for direct comparison with pollen analyses (Bonnefille, 2010). Molecular-level biomarker isotopic data were collected on fatty (alkanoic) acids as their corresponding fatty acid methyl esters using a gas chromatograph–isotope ratio mass spectrometer (GC-IRMS) and δ¹³C data are reported for the C₃₀n-alkanoic acid (Feakins et al., 2005). Age control is provided by an updated tephrostratigraphic and nannofossil age model for the core. Details of the method and age model may be obtained from the GSA Data Repository^1,2.

Plant proxy δ¹³C values are commonly interpreted in terms of percentages of C₄ vegetation using a mixing model approach, although the range of isotopic values in bulk plant tissues results in large propagated uncertainties. Working around these issues, Cerling et al. (2011) recently introduced an algorithm to calculate fraction of woody cover, trained on a modern survey of woody cover and area-weighted mean δ¹³C values of soil organic matter. While this approach is useful it does not differentiate between rainforest and seasonally dry trees, each of which may reach 100% canopy cover; here we explicitly test for the presence of rainforest taxa in the pollen record from the marine core. Secondly, this approach does not consider that open-canopy environments can be dominated by C₃ grasses, as may have been the case in the Miocene. For this reason we suggest that a paired pollen and carbon isotope approach is needed to reconstruct Middle to Late Miocene environments.

δ¹³C_wax values from DSDP Site 231 range between −31.1‰ and −22.3‰ (Fig. 2). The most negative values are reported at 11.1 and 4.3 Ma, and the most positive values at 1.4 Ma. There is a 4.4‰ positive shift in δ¹³C_wax between 11.1 and 8.7 Ma, and a gradual further increase of 1.9‰ by 4.9 Ma, followed by a 6.3‰ decrease in δ¹³C_wax values by 4.3 Ma. Higher sampling resolutions reveal variability within 10⁴–10⁵ yr cycles of up to 5‰, best sampled between 4.3 and 3.8 Ma. In addition to the cyclicity, between 4.3 and 1.4 Ma, δ¹³C_wax values display an overall trend toward more positive values.

In the terrestrial archives, Miocene data are sparse and the strongest signal is a trend to more positive δ¹³C values after 4.3 Ma where most pedogenic carbonate samples are reported (Fig. 2). This positive trend is evident amidst scatter of up to 11‰ in δ¹³C_pc values at any one time interval that reflects spatial variations in vegetation as fine as 1 m², including the continued presence of C₃ plants.

We view ecological change in northeast Africa as the primary driver of the δ¹³C_wax and δ¹³C_pc records and interpret these data as functions of the influence of C₄ vegetation on the landscape and the environmental conditions in which these plants grew. While δ¹³C values of atmospheric CO₂ (δ¹³C_CO2) can influence the δ¹³C values of plants, benthic foraminiferal constraints on δ¹³C_CO2 variability (Tipple et al., 2010) indicate that this effect is of minor importance before the Pleistocene glacial cycles in our record. Both δ¹³C_wax and δ¹³C_pc records show the presence of C₄ vegetation since at least 9 Ma, a minimum in δ¹³C values in the early Pliocene (ca. 4 Ma), and an increase in C₄ vegetation through the Pliocene and Pleistocene.

Palynology provides an important link to the species responsible for the carbon isotopic signals. We find that the diversity of plant species recorded in the marine sediments of DSDP Site 231 has remained largely unchanged across the last 12 m.y., but the proportions have shifted (Bonnefille, 2010; and this study). Of greatest relevance here, grass pollen ranges between 3% and 59% of total pollen, and displays a decline from 11.5 to 8 Ma, an increase around 8 Ma, and then an overall proportional decline from 8 to 1.6 Ma (Fig. 2). Declines in grass pollen occurred while δ¹³C_wax values increased toward greater C₄ biomass proportions, indicating that gains by C₄ grasses were outpaced by losses of C₃ grasses. Gains by C₄ grasslands are further minimized if we consider that Amaranthaceae (shrubs) may be responsible for some fraction of C₄ biomass expansion.

Between 11.1 and 8.7 Ma, δ¹³C_wax values increased by 50% of their full range, which is broadly consistent with a 50% increase in C₄ biomass. The timing of C₄ increase in the marine record corroborates the timing of the Miocene C₄ expansion in the diets of equids, the first responders amongst the sampled fauna (Uno et al., 2011), and confirms a significant expansion of C₄ biomass in northeast Africa by 9 Ma which is not evident in the δ¹³C_pc record.

Pollen evidence has already suggested an absence of closed rainforest at East African hominin sites of Pliocene–Pleistocene age (Bonnefille, 1995) and from marine archives back to 12 Ma (Bonnefille, 2010; and this study). However this is in apparent contrast with molecular clock estimates for rainforest fragmentation and the isolation of West and East African rainforest taxa at ca. 8 Ma (Couvreur et al., 2008). In this study, we find δ¹³C_wax values vary between −30.2‰ and −25.4‰ from 12 to 5 Ma, inconsistent with modern rainforest values of ∼−38‰ (Vogts et al., 2007). Furthermore the pollen assemblages do not include humid Guineo-Congolian rainforest taxa, most likely indicating that a geographically restricted East African rainforest lay outside the source area for leaf wax and pollen.

Reconciling the data from multiple proxies requires consideration of how each proxy may record different aspects of the regional climate and environmental system. In the Gulf of Aden there are no significant fluvial inputs; terrigenous inputs are instead windblown (Fig. 1). Low-level winds follow the topographic barrier of the Ethiopian highlands, with high elevations in place since at least 12 Ma (Wichura et al., 2010) and similar-to-modern atmospheric circulation since at least 8 Ma (Sepulchre et al., 2006), although uplift and rifting continues (Trauth et al., 2007). Although both pollen and leaf waxes are produced by plants, their production and dispersal have limited functional connection in the plant and they are expected to have different (but uncertain) source integration characteristics across vegetated areas, which may vary both with plant type and climate regime. Studies off West Africa have shown distal transport is possible (Schefuss et al., 2003), although local sources dominate (Dupont, 2011). Here, in the Gulf of Aden, pollen taxa confirm integration of varied local and distal sources, thought to include the south Ethiopian escarpment, the Afar triangle, and to a lesser extent the Horn of Africa, but perhaps not sites as far south as the Turkana Basin (Fig. 1).

In comparison, the δ¹³C_pc values from East African Rift sites provide an indication of long-term trends within variable terrestrial landscapes within the rift valley, with regional climatic influences as well as local tectonic controls (Levin et al., 2011; Trauth et al., 2007). As pedogenic carbonates may form over 10³–10⁵ yr, δ¹³C_pc values do not capture the short-term variability and environmental extremes that may be recorded in the pollen and δ¹³C_wax records. However, the long-term vegetation trend in the Pliocene and Pleistocene is common to both the marine archive and the floodplain and lake-margin soils of the East African Rift system (Fig. 2).

The long perspective afforded by the extended records allows us to constrain the nature of the Miocene C₄ expansion in direct comparison with the later Pliocene–Pleistocene expansions. The 12 m.y. leaf wax carbon isotope record suggests that the Miocene C₄ expansion was achieved in two directional shifts occurring 11–9 Ma and 6–5 Ma, separated by a period of apparent stability. On land, peak values for the Late Miocene are recorded at 5.72 Ma (δ¹³C_pc = −1.3‰) in the Lothagam Formation in Turkana (Cerling et al., 2003), and at 4.4 Ma in the Sagantole Formation (−0.4‰) in the Middle Awash study area of the Afar region (WoldeGabriel et al., 2009; Fig. 2). In the marine core, δ¹³C_wax values are high across 5.8–4.4 Ma, with a single point defining a peak (−23.9‰) at 4.9 Ma. These data show that during the Late Miocene, C₄ vegetation in northeast Africa reached proportions almost as great as the later peak at 1.4 Ma, but there was a return to some of the lowest δ¹³C_wax values reported (negligible C₄ proportions) by 4.3 Ma. Similar patterns are seen on land where there is a shift to more negative δ¹³C_pc values in both the Awash and Turkana Basins after 4.4 Ma (Cerling et al., 2011). However, despite the return to late Middle Miocene δ¹³C values, grass biomass and woody cover proportions are very different with only 10%–30% grass pollen in the earliest Pliocene versus 20%–50% grass pollen in the Middle Miocene, indicating dramatically different ecosystems. We find that C₃ grasslands and woodlands in the Miocene were replaced by a mix of C₄ grassland, C₃/C₄ shrubland, and woodland in the Pliocene. Overall, the Miocene and Pliocene trend was toward a reduction of grass pollen (presumably indicating a contraction of grasslands or a decline in grassland productivity) at the same time that C₄ grasses increased as a proportion of biomass (δ¹³C positive shift).

In the Pliocene and Pleistocene, a consistent trend of positive shifts in both δ¹³C_wax and δ¹³C_pc values and decline in the proportion of grass pollen begins after 4.3 Ma (Fig. 2). The C₄ expansion across 4.3–1.4 Ma was of equal magnitude to the Miocene expansion, but the nature of the Pliocene–Pleistocene expansion was very different with progressive C₄ biomass increases across numerous 10⁴–10⁵ yr cycles, corroborating the variability that was previously reported (Feakins et al., 2005). These cycles in δ¹³C_wax values are presumed to be obliquity- and precession-paced based on analysis of dust percentages in cores with continuous recovery collected around Africa (deMenocal, 1995). During this time a long-term reduction in the proportion of grass pollen suggests that absolute grassland productivity may have decreased or that grasslands in general declined relative to other habitats such as xeric shrubland (marked by abundant Amaranthaceae taxa), presumably related to increased aridity. These broad trends align with the progressive expansion of Northern Hemisphere glaciations after 2.7 Ma (Lisiecki and Raymo, 2005), supporting a role for the high latitudes in driving the tropical vegetation changes during the Pleistocene (Dupont, 2011; Schefuss et al., 2011). However, local insolation forcing (precession dominated) has also been implicated in tropical African vegetation changes (Denison et al., 2005; Dupont, 2011). We here confirm our earlier report (Feakins et al., 2005) that northeast African environmental variability began well before the onset of Northern Hemisphere glaciations, as early as 4.3 Ma.

This study synthesizes δ¹³C_wax, δ¹³C_pc, and pollen evidence to provide new insights into northeast African vegetation change during the past 12 m.y. Our enhanced δ¹³C_wax record reveals that the regional C₄ biomass expansion between 11 and 5 Ma (Late Miocene) was reversed by 4.3 Ma and then followed by an equally large C₄ expansion from 4.3 to 1.4 Ma (Pliocene–Pleistocene). The combination of δ¹³C and pollen data yields insights that neither the isotope nor pollen records could achieve alone. Firstly, extensive and/or productive Middle Miocene grasslands must have been C₃ dominated. Secondly, C₄ biomass increased substantially by ca. 9 Ma although grasslands were reduced in overall extent or productivity. Thirdly, the nature of the two C₄-biomass expansions differed: the Miocene expansion occurred amidst extensive C₃ grasslands, whereas the Pliocene–Pleistocene C₄ expansion occurred in a context of reduced grasslands. Consistencies between δ¹³C_wax and δ¹³C_tooth in the Late Miocene and δ¹³C_wax and δ¹³C_pc in the Pliocene and Pleistocene give confidence that these are robust features of both the East African Rift and the wider northeast African environment.

This research used samples provided by the Integrated Ocean Drilling Program. The manuscript was improved with the comments of the editor and three anonymous reviewers. Thanks to D. Montluçon and C. Johnson at Woods Hole Oceanographic Institution; M. Rincon, Z. Zhang, S. Oh, L. Petrie, and V. Nguyen at the University of Southern California (USC); and G. Buchet at the Centre National de la Recherche Scientifique for laboratory assistance. Funding was provided by the U.S. National Science Foundation HOMINID Grant 0218511 to Eglinton, and by USC to Feakins.