Using 10Be cosmogenic isotopes to estimate erosion rates and landscape changes during the Plio-Pleistocene in the Cradle of Humankind, South Africa

Journal of Human Evolution 96 (2016) 19e34

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Using 10Be cosmogenic isotopes to estimate erosion rates and landscape changes during the Plio-Pleistocene in the Cradle of Humankind, South Africa Paul H.G.M. Dirks a, b, *, Christa J. Placzek a, David Fink c, Anthony Dosseto d, Eric Roberts a a

College of Science and Engineering, James Cook University, Townsville, Qld 4811, Australia School of Geosciences, University of the Witwatersrand, Private Bag 3, Wits 2050, South Africa Australian Nuclear Science and Technology Organization, PMB1, Menai, NSW 2234, Australia d School of Earth and Environmental Science, University of Wollongong, Wollongong, NSW 2522, Australia b c

a r t i c l e i n f o

a b s t r a c t

Article history: Received 17 February 2015 Accepted 7 March 2016

Concentrations of cosmogenic 10Be, measured in quartz from chert and river sediment around the Cradle of Humankind (CoH), are used to determine basin-averaged erosion rates and estimate incision rates for local river valleys. This study focusses on the catchment area that hosts Malapa cave with Australopithecus sediba, in order to compare regional versus localized erosion rates, and better constrain the timing of cave formation and fossil entrapment. Basin-averaged erosion rates for six sub-catchments draining the CoH show a narrow range (3.00 ± 0.28 to 4.15 ± 0.37 m/Mega-annum [Ma]; ±1s) regardless of catchment size or underlying geology; e.g. the sub-catchment with Malapa Cave (3 km2) underlain by dolomite erodes at the same rate (3.30 ± 0.30 m/Ma) as the upper Skeerpoort River catchment (87 km2) underlain by shale, chert and conglomerate (3.23 ± 0.30 m/Ma). Likewise, the Skeerpoort River catchment (147 km2) draining the northern CoH erodes at a rate (3.00 ± 0.28 m/Ma) similar to the Bloubank-Crocodile River catchment (627 km2) that drains the southern CoH (at 3.62 ± 0.33 to 4.15 ± 0.37 m/Ma). Dolomite- and siliciclastic-dominated catchments erode at similar rates, consistent with physical weathering as the rate controlling process, and a relatively dry climate in more recent times. Erosion resistant chert dykes along the Grootvleispruit River below Malapa yield an incision rate of ~8 m/Ma at steady-state erosion rates for chert of 0.86 ± 0.54 m/Ma. Results provide better palaeodepth estimates for Malapa Cave of 7e16 m at the time of deposition of A. sediba. Low basin-averaged erosion rates and concave river profiles indicate that the landscape across the CoH is old, and eroding slowly; i.e. the physical character of the landscape changed little in the last 3e4 Ma, and dolomite was exposed on surface probably well into the Miocene. The apparent absence of early Pliocene- or Mioceneaged cave deposits and fossils in the CoH suggests that caves only started forming from 4 Ma onwards. Therefore, whilst the landscape in the CoH is old, cavities are a relatively young phenomenon, thus controlling the maximum age of fossils that can potentially be preserved in caves in the CoH. © 2016 Elsevier Ltd. All rights reserved.

Keywords: Sediba Cradle of Humankind Landscape Erosion Caves

1. Introduction The history of landscape change in southern (and eastern) Africa is important in relation to hominin evolution given that our species evolved in Africa and we are, at least in part, the product of our

* Corresponding author. E-mail addresses: [email protected] (P.H.G.M. Dirks), christa.placzek@jcu. edu.au (C.J. Placzek), fi[email protected] (D. Fink), [email protected] (A. Dosseto), [email protected] (E. Roberts). http://dx.doi.org/10.1016/j.jhevol.2016.03.002 0047-2484/© 2016 Elsevier Ltd. All rights reserved.

natural environment and dynamic changes therein (e.g., Bailey et al., 2011; deMenocal, 2011; Roberts et al., 2012; Dirks and Berger, 2013). Not only do landscape features, such as rifts or cave systems act as fossil depositories, but tectonic processes controlling the formation of these features could represent forcing mechanisms that indirectly influence hominin evolution itself (e.g., King and Bailey, 2006; Sepulchre et al., 2006). For instance, Bailey et al. (2000) argued that the rifts and volcanic landscapes of northeast Africa present dynamic, diverse and continually rejuvenating landscapes conducive to human occupation (see also King et al., 1994). This argument, in which active faulting and the associated

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landscape dynamics act as an evolutionary vector, has also been made for the high-plateau landscapes of southern Africa (Bailey et al., 2011; Reynolds et al., 2011), but is seemingly at odds with the general notion that southern Africa represents a relatively  and de Wit, 2003; de Wit, 2007; stable, static landscape (Doucoure Partridge, 2010; Decker et al., 2011; Scharf et al., 2013). In this paper, we attempt to estimate the rate of change of the landscape in the Cradle of Humankind (CoH) using the in situ production of the cosmogenic radionuclide 10Be in surface samples, to assess if the landscape is as tectonically dynamic and, therefore, variable in time and space as suggested by Bailey et al. (2011) and Reynolds et al. (2011). The principal driver of landscape change on the highveld in southern Africa is generally identified as climate variability (e.g., Hopley et al., 2007). This variability has been modelled for the high plateau landscapes of East Africa and has been linked to events such as the global advent of ice-ages, the onset of El Nino/Southern Oscillation (ENSO) patterns in the Pacific Ocean (e.g., Potts, 1998; Bobe et al., 2002; Maslin and Christensen, 2007; Trauth et al., 2007), and movement of the Inter-tropical Convergence Zone (Stager et al., 2011). Since the mid-Pliocene, southern Africa has experienced a general drying evidenced by fragmentation of forest cover and expansion of savannah grasslands (e.g., Vrba, 1995; galen et al., 2007). The assumption that Hopley et al., 2007; Se paleo-climate records from southern Africa mainly record global climate events implies that the high elevation landscapes of southern Africa are old; i.e. that no uplift of the landscape occurred that would have influenced local climate variability (e.g., de Wit, 2007; Erlanger et al., 2012). However, it has been suggested that significant uplift of the landscape has occurred since the Pliocene (e.g., Partridge and Maud, 1987), in which case regional climate variability in the last 5 Ma (Mega-annum) is not simply the result of global change, but also the result of uplift linked to the formation of the southern African plateau, which would have influenced regional temperature and rainfall patterns (e.g., Sepulchre et al., 2006). Thus, interpretation of changes in environmental variables since the advent of hominins on the landscape is complicated in that the interconnected tectonic and climate forcing processes caused variable impacts at two different scales: 1) the local to regional scale, where these processes control floral and faunal assemblages, as well as the development of local topography including cave formation, soil rejuvenation, and water availability (e.g., King and Bailey, 2006; Bailey et al., 2011; Dirks and Berger, 2013); and 2) the continental scale, in which overall uplift of the landscape influenced climate patterns across all of Africa (e.g., Sepulchre et al., 2006). Complicating this issue is the fact that the uplift history of southern Africa is poorly constrained in time. Some authors suggest that an old flat planar landscape experienced moderate uplift during the Miocene with renewed rapid uplift by as much as 900 m since the Pliocene (Partridge and Maud, 1987; Partridge et al., 2006, 2010; Roberts and White, 2010), with alternative models arguing for more gradual uplift or for topography to be old and largely Cretaceous in age (e.g., Tinker et al., 2008a,b; Flowers and Schoene, 2010; Erlanger et al., 2012; Guillocheau et al., 2012; Decker et al., 2013). Direct estimates of uplift rates and erosion, and, therefore, changes in the landscape, are hard to obtain for southern Africa (Saria et al., 2013) and rely mostly on indirect measurements of either temperature (Beukes et al., 1999), denudation (Brown et al., 2002; de Wit, 2007; Flowers and Schoene, 2010), erosion rates (e.g., Erlanger et al., 2012; Bierman et al., 2014), or the interpretation of geomorphological features such as knick points and raised marine terraces dated bio-stratigraphically (e.g., Partridge and Maud, 1987). Ancillary information on uplift in the region comes from sediment fill from off-shore alluvial deltas (e.g., McMillan,

2003; Tinker et al., 2008a; Guillocheau et al., 2012) and modelling of stream profiles (e.g., Roberts and White, 2010). Because of this paucity of precise data it is difficult to reliably reconstruct the degree and rate of landscape change in the CoH since the early Pliocene, beyond broad inferences based on geomorphology (e.g., Partridge and Maud, 1987; Reynolds et al., 2011; Dirks and Berger, 2013). Surface exposure dating using cosmogenic radionuclides has provided estimates for the rate of landscape change in the vicinity of hominin-bearing cave sites (e.g., Partridge et al., 2003; Dirks et al., 2010; Granger et al., 2015), but work to date has been limited to a few locations, and has been mostly focussed on measuring the burial age of fossil-bearing deposits (e.g., Partridge et al., 2003; Granger et al., 2015), rather than attempting to reconstruct the landscape more broadly (e.g., Dirks and Berger, 2013). 1.1. An overview of cosmogenic nuclide methods Cosmogenic nuclides can be used to estimate the length of time that a rock or sediment has been exposed at, or near, the Earth's surface (Gosse and Phillips, 2001; Dunai, 2010; Granger et al., 2013, 2015). In situ cosmogenic nuclides are produced by secondary cosmic radiation that bombard the Earth's surface, and interact with the atomic nuclei in minerals, to form both stable (3He, 21Ne) and radioactive (10Be, 26Al, 36Cl) isotopes. In rock, most of the cosmic radiation is absorbed within the first few metres from the surface as the cosmic ray flux decreases exponentially with depth. As cosmic radiation is absorbed, cosmogenic nuclides are formed within the rock. The production rate for a particular cosmogenic nuclide varies with latitude, elevation, depth and rock density, and is further dependent on shielding effects, i.e. the amount of sky that can be seen from a sampling point. The concentration of cosmogenic nuclides in a mineral grain (from a rock sample or a fluvial sediment) is a function of the rate with which that grain has been brought to the surface, a process referred to as denudation, and is linked to the rate of erosion (Lal, 1991; Granger and Muzikar, 2001; Gosse and Phillips, 2001); i.e. higher concentrations of nuclides imply longer exposure times near the surface and, therefore, slower erosion rates (erosion rates in m/Ma [metres per million years or Mega-annum] or mm/ka [millimetres per thousand years or Kilo-annum]) and exposure ages for minerals are inversely proportional. Cosmogenic radionuclides decay at a rate given by their decay constant, and the rates of production and decay can be combined to give the total concentration of cosmogenic radionuclides in a sample, which is a function of age (see Material and methods section below for details). The concentration of cosmogenic nuclides initially increases linearly with time, allowing surface exposure ages to be determined, until a steady-state concentration is achieved where loss from both decay and surface erosion balances production. For surfaces at steady-state where erosion (and not decay) controls the cosmogenic concentration, the steady-state concentration is inversely proportional to the erosion rate. The two most frequently measured cosmogenic radionuclides are 10Be and 26Al, which form when cosmic rays interact with O and Si nuclei respectively. O and Si are the constituent atoms of quartz (SiO2), which is one of the most common and widely distributed minerals in the Earth's crust and has been shown to be a closed system following production, making 10Be and 26Al highly suitable nuclides for dating geological materials and estimating the rate of erosional processes (e.g., Lal, 1991). Each of these radionuclides is produced at a different rate, and can be used individually to provide exposure ages (e.g., Granger and Muzikar, 2001; Bierman et al., 2014). In combination the 26Al/10Be concentration ratio of both nuclides can provide an estimate of burial below the depth of

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production, such as the burial of sediment in Sterkfontein Cave (e.g., Partridge et al., 2003; Granger et al., 2015). Almost all estimates of erosion rates from southern Africa come from rock outcrops. These studies indicate that for at least the last one million years, average erosion rates on outcrops in the stable, flat areas along the southern African escarpment and inland plateau are generally between 1 and 4 m/Ma (e.g., Kounov et al., 2007 for 3He and 21Ne; Decker et al., 2011 for 3He; Decker et al., 2011 for 3He and 21Ne; Scharf et al., 2013 for 10Be; Bierman et al., 2014 for 10Be and 26Al). Erosion rates obtained from basalt (0.5e4.0 m/Ma [Kounov et al., 2007; Decker et al., 2011, 2013]) are generally higher than those obtained from more weathering resistant quartzite (1.0e2.1 m/Ma [Kounov et al., 2007; Scharf et al., 2013; Bierman et al., 2014]). Denudation rates along the dry western escarpment of southern Africa are significantly lower (0.4 m/Ma for 10Be and 26Al [Cockburn et al., 2000]; 1.5e3.0 m/Ma for 3He and 21Ne [Kounov et al., 2007]) than rates along the wetter eastern escarpment (6.6 m/Ma for 36Cl [Fleming et al., 1999]), suggesting that climate exerts a control on erosion across the region (Chadwick et al., 2013). The fact that rock outcrops protrude from the ground suggests that such exposed rock surfaces experienced lower erosion rates than the surrounding landscape. The rate of erosion of drainage basins is generally found to be significantly faster (by a factor 10 on average [Portenga and Bierman, 2011]) than that of rock outcrops, with the difference reflecting increased rates of rock weathering below soil. Thus, erosion rates determined from rock outcrops, in most cases, reflect minimum erosion rates, and may not reflect the regional rate of change in the landscape. The latter can be derived with basin-averaged erosion rates (Granger et al., 1996; Portenga and Bierman, 2011) using alluvial sediment samples from river catchments. Within a single drainage basin, different regions erode at different rates, meaning that sediments derived from different parts of the basin contain different nuclide concentrations. A stream that drains the catchment will mix sediment from these different source areas to reflect the average nuclide concentration, which can provide a basin-averaged erosion rate assuming that the different parts of the basin contribute sediment in proportion to their long-term erosion rate (Granger et al., 1996) and sediment is transported out of the catchment area (i.e., the catchment is a sediment-starved system [Riebe et al., 2004]). Because soils are typically well-mixed as a result of physical and biological processes, shallow, human-induced soil erosion does not typically affect basin-averaged erosion rates (Portenga and Bierman, 2011). Thus, the average nuclide concentration in stream sediment generally reflects the average erosion rate across the surface area of the basin above the sampling location (i.e. the basin-averaged erosion rate), and basin-averaged erosion rates are a better way to estimate longterm landscape change than estimates derived from rock outcrops. In southern Africa, few basin-averaged erosion rates have been obtained. Erlanger (2010) and Erlanger et al. (2012) report erosion rates of ~9e22 m/Ma for catchments along the coastal plain, increasing to ~86 m/Ma along the great escarpment, dropping back to ~5 m/Ma for catchments in the interior of South Africa. Scharf et al. (2013) record basin-averaged erosion rates of ~2.3e8 m/Ma for catchments draining the Cape Fold Belt to the south of the southern African plateau. Overall, available cosmogenic nuclide data suggest that the high plateau landscape of South Africa is relatively static and is eroding at rates that average <10 m/Ma (i.e. a low rate of erosion compared to the global mean of 5.4 m/Ma for rock outcrops and 54 m/Ma for sedimentary basins [Portenga and Bierman, 2011]), and thus it has been suggested that the topography is old and has been undergoing slow denudation for at least the past few million years (e.g., Decker

21

et al., 2013; Scharf et al., 2013). In this context it is good to remember that across the varied landscapes of the Earth, cosmogenic nuclides indicate that erosion rates range from near zero to almost 10,000 m/Ma (Portenga and Bierman, 2011). This situation implies that our ancestors lived on a landscape with an overall form that resembles the landscape seen today (e.g., Dirks and Berger, 2013). Evidence from drainage channels and exposed caves in the CoH suggest that at least locally, channel incision and base level lowering has resulted in a few 10s of metres of erosion in the more recent past (Dirks et al., 2010; Dirks and Berger, 2013). An incision rate of 53 m/Ma was recorded for the Skeerpoort River west of Gladysvale (Dirks et al., 2010 using 10Be), and at Malapa, the type locality for Australopithecus sediba dated at 1.977 ± 0.003 Ma (Berger et al., 2010; Pickering et al., 2011a), a cave system deep enough to act as a death trap is now exposed on the surface, leaving only erosion remnants of sediment that during earlier times filled the deepest cave chambers (Dirks et al., 2010). This example and similar erosional remnants elsewhere in the CoH illustrate that, in some localities, significant landscape modifications occurred in the recent past, where they can be linked to particular geomorphological settings (Dirks and Berger, 2013). Bailey et al. (2011) and Reynolds et al. (2011) even go as far as to suggest that landscape features in the CoH, and especially in the Bloubank River valley, which hosts Sterkfontein, Swartkrans and Rising Star Caves, reflect frequent rejuvenation by active faults, a process normally associated with regional erosion rates of the order of 10 s of m/Ma (Portenga and Bierman, 2011). This paper makes use of the concentration of cosmogenic 10Be in quartz (Granger et al., 1996, 1997) from selected samples of rock and river sediment in and around the CoH. We report the first basin-averaged erosion rates from catchments across the CoH, as well as estimates of incision rates for valleys and valley slopes in the immediate vicinity of Malapa Cave. Placed together, these estimates better constrain our understanding of regional versus localized erosional processes that affect landscape modifications in the CoH. Our results question the suggestions made in Reynolds et al. (2011) that the landscape is dynamic and controlled by active faulting, and place critical constraints on landscape evolution in the CoH, including formation and opening of caves that potentially trap fossils. 1.2. Geology and geomorphology setting of the Cradle of Humankind The late Pliocene to Quaternary cave deposits in the Cradle of Humankind, UNESCO, World Heritage Area, South Africa are one of the world's most important geological settings, hosting hominin fossils and associated faunal and archaeological remains. Abundant hominin fossil remains have been recovered from a number of caves including Sterkfontein, Swartkrans, Malapa, Drimolen, Rising Star and Kromdraai (Fig. 1). These fossils have been ascribed to a range of species including Australopithecus africanus, Australopithecus robustus, A. sediba, Homo erectus and Homo naledi (e.g., Brain, 1993; Tobias, 2000; Berger et al., 2010, 2015). Hominin remains in the CoH caves are encased in clastic, cave-fill deposits situated in stromatolite-rich, dolomite sequences of the Malmani Subgroup (Martini, 2006), which was deposited on a late Archaean continental shelf (Eriksson et al., 2006). Geological descriptions of the most important fossil sites (e.g., de Ruiter et al., 2009; Pickering and Kramers, 2010; Dirks et al., 2010, 2015; Herries and Shaw, 2011; Pickering et al., 2011b) indicate that fossil deposition in caves involved a range of processes including death traps, scavengers, mud flows, predation and possibly even body disposal (Dirks et al., 2015). The geographic distribution of fossil-bearing caves in the CoH is not

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Figure 1. GoogleEarth image of southern Africa (a) showing the major rivers draining the Johannesburg area. (b) The area of Gauteng province where the CoH is located. (c) The CoH World Heritage Area shown on a digital terrain model (red >1600 m; blue <1200 m) with the upper reaches of the Crocodile River catchment. Centrally located in the CoH is the catchment of the Grootvleispruit with Malapa (8) at its core. Caves in the area are shown (blue circles) and the main hominin sites (yellow squares) are listed. MR ¼ Magaliesberg River; KR ¼ Crocodile River; SKR ¼ Skeerpoort River; BBS ¼ Bloubank River. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

random, indicating that the caves are not simply sedimentary depositories (Dirks and Berger, 2013), but also provided protection and possibly access to resources such as water. Their formation, location and access may, therefore, play a deterministic role in terms of fossil preservation. The area of the CoH to the west of the Johannesburg Dome (Figs. 1, 2), preserves a variety of landscape elements (duricrusts, terraces, paleo-soils) assigned to various African erosion surfaces, incised by the headwaters of the Crocodile River (Fig. 1; Partridge and Maud, 1987; Beukes et al., 1999; Dirks and Berger, 2013). The southern margin of the Johannesburg Dome is characterized by an escarpment reaching 1900 m in height, formed by quartzite of the 3.0e2.8 Ga (Giga-annum) Witwatersrand Supergroup that forms

the watershed between the Crocodile-Limpopo River system draining north and east, and the Orange-Vaal River system draining south and west (Figs. 1, 2). South of the escarpment is the start of a flat plateau mostly underlain by near-horizontal sedimentary rock sequences of Paleoproterozoic to Jurassic ages. To the east, north and west of the Johannesburg Dome outward dipping sediments of the Transvaal Supergroup form undulating hills accentuated by ridges underlain by quartzite and chert, bounded to the north by the Magaliesberg escarpment composed of coarse-grained quartzite, through which the Crocodile River has incised a narrow gorge, now dammed (Figs. 2, 3a; Partridge and Maud, 1987). In the past, sediments from the Karoo Supergroup would have covered most of the Johannesburg Dome as attested by

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Figure 2. Geological map of the Johannesburg Dome and surrounding lithologies. The boundary of the CoH is shown along the western margin of the dome.

erosional remnants of these sediments (e.g., Dirks and Berger, 2013). In the CoH several parallel ridges of sandstone, chert and dolomite crop out (Fig. 3), and the landscape rises from north to south, drained by two tributaries of the Crocodile River: the Skeerpoort River to the northwest, and the Bloubank River to the southeast (Figs. 1, 3a; Partridge et al., 2010). The catchment areas of these two tributaries host all major fossil-bearing caves (Fig. 1). The Skeerpoort River catchment is mostly underlain by chert-rich dolomite, sandstone, volcanics and shale, whilst the catchment of the Bloubank River includes large areas of granite and gneiss as well as chert-rich dolomite, volcanics and clastic sediment. Chert breccia of the Rooihoogte Formation is weathering resistant and underlies a 1e2 km wide, dissected dip slope that formed along the upper contact of the Malmani dolomite along sections of the Skeerpoort

River in the centre of the CoH (Dirks and Berger, 2013). A number of northwest draining tributaries to the Skeerpoort River, including the Grootvleispruit River, along which Malapa is situated, have cut narrow, steep-sided gorges through this dip slope. Prominent knick points characterized by waterfalls and, in places, spring sites have developed where creek beds cut across the Rooihoogte Formation; i.e. knick points in this area are lithologically controlled (Haviv et al., 2010; Dirks and Berger, 2013).

2. Material and methods Quartz-bearing samples were collected from chert-conglomeraterich erosion remnants belonging to the Rooihoogte Formation, and from modern stream-bed sediments from rivers that drain

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Figure 3. (a) Photo of the headwaters of the Grootvleispruit, looking NW across the centre of the CoH, with quartzite of the Magaliesberg forming the ridge in the background. (b) View of a typical erosion remnant of a silicified chert conglomerate dyke that has been preferentially weathered from the surrounding dolomite. The dyke in this example is 7.5 m tall and 1e2 m wide. (c) Example of a stream bed sediment sample to measure basin-averaged erosion rates. This sample (CHK110) was collected from the Grootvleispruit, 150 m NW of Malapa Cave. (d) Example of quartz-rich, rock chips collected from chert conglomerate dykes for cosmogenic age dating. This sample (CHK105) was collected from the prominent dyke shown in Figure 5b.

catchments of variable sizes and are underlain by varied geology, all of which drain parts of the CoH (Tables 1 and 2; Figs. 3c, d, 4, 5).

(>95%), to catchments dominated by granite and various clastic sedimentary units, with dolomite contributing as little as 15% (Table 2).

2.1. Stream sediment samples 2.2. Chert conglomerate samples Six sediment samples were taken from modern stream beds that drain the CoH (Fig. 4; Table 1) to the northwest and the southeast, within the sub-catchments of the Skeerpoort and Bloubank Rivers (Figs. 1, 4). All samples consist of several kilograms of sand and gravel (Fig. 3c), that were sieved to various grain sizes (Table 1). The 0.25 mm or larger grain size was processed in order to exclude windblown contributions. Two samples (CHK 110, 111) were taken from dry river beds near Malapa that drain sub-catchments of the Grootvleispruit River. A further sample (CHK112) was taken from the Grootvleispruit River below a series of waterfalls marking a prominent knick point. One sample (CHK113) was taken from the Skeerpoort River immediately above the confluence with the Grootvleispruit River, a spot marked by tufa deposits; a further sample (CHK118) from the Skeerpoort River was taken where it exits the CoH on Hartebeeshoek Farm. One sample (CHK207) was taken from the Crocodile River immediately below the confluence with the Bloubank River, which drains the southeast half of the CoH. Two different grain size fractions (0.25e0.5 mm, CHK207; and >1 mm, CHK207L) were analyzed for the Crocodile River sample. Sediment samples CHK 118 and 207 represent the full CoH catchment region. The surface areas of the (sub-)catchments above each sample site are presented in Table 2, together with estimates of the relative surface areas of various lithological units underlying each catchment. Sampled catchments vary in size from 3.5 to 627 km2, and range from catchments dominated by dolomite

Rock samples were collected from silicified chert conglomerate units in the Grootvleispruit River catchment (Fig. 1), down-valley (i.e. north) of Malapa Cave. The sampled units of silicified conglomerate occur at the base of the Rooihoogte Formation, and formed in the Paleoproterozoic as a result of fissure fill on paleokarst surfaces (Dirks and Berger, 2013). The fissure fill is oriented roughly vertically and consists of conglomerate dominated by laminated chert pebbles in a matrix of quartz sandstone. This fill has been strongly silicified by hydrothermal fluids, and weathers differentially from the surrounding dolomite to form prominent, dyke-like erosion remnants (Figs. 3b, 5a, b). The maximum height of these erosion remnants increases at a regular rate as one moves from the flat, African, paleo-erosion surface preserved along hill tops (at ~1580e1600 m), down the valley slopes into the valley bottom, where breccia walls locally reach heights of >20 m (Dirks and Berger, 2013). Samples of chert conglomerate were collected from two sites. The first site, site 1 (Table 1), is a 2e4 m wide and ~15 m high, walllike dyke that is oriented perpendicular to the modern river course and crosses the Grootvleispruit River ~1.9 km northwest of Malapa, with the valley floor at 1374 m (Fig. 5b, c). This site forms a prominent knick point within the Grootvleispruit River, and is representative of similar lithology-controlled knick points in the CoH (Dirks and Berger, 2013). Several distinct steps occur along the ridge of the silicified dyke (Fig. 5c), each forming a 2e3 m long sub-

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Table 1 Sample data and cosmogenic erosion rates and ages for samples considered in this study.a Sample name

Latitude (DD)

Longitude (DD)

Elevation (m)

Thickness (cm)

Shielding correction

[Be-10] atoms g1

Site 1: chert dyke crossing Grootvleispruit river (Fig. 5b,c) CHK101 25.879 27.793 1374 5.0 0.938 9.5907Eþ05 CHK105 25.879 27.793 1378 5.0 0.992 2.6181Eþ06 CHK102 25.879 27.792 1382 5.0 0.968 2.8746Eþ06 CHK103 25.879 27.792 1388 5.0 0.975 5.8745Eþ06 Site 2: chert dyke along valley slope (Fig. 5a) CHK115 25.893 27.801 1436 5.0 0.782 7.7182Eþ05 CHK116 25.893 27.801 1438 5.0 0.782 1.0819Eþ06 CHK117 25.893 27.801 1441 5.0 0.994 1.7857Eþ06 Grootvleispruit river sub-catchment upstream from Malapa CHK110 25.893 27.800 1398 0.5 e 2.0039Eþ06 CHK111 25.896 27.808 1423 0.5 e 2.2636Eþ06 Grootvleispruit river sub-catchment above chert knick points CHK112 25.875 27.789 1337 0.5 e 2.1110Eþ06 Skeerpoort river sub-catchment above confluence with Grootvleispruit river CHK113 25.875 27.780 1292 0.5 e 2.1084Eþ06 Skeerpoort river catchment above bridge CHK118 25.838 27.850 1248 0.5 e 2.1796Eþ06 Crocodile river catchment CHK207 25.933 27.905 1282 0.5 e 1.6912Eþ06 CHK207-L 25.933 27.905 1282 2.0 e 1.8834Eþ06 average basin-averaged erosion rate

þ/ atoms g1

Erosion rate (m/Ma)

External uncertainty (1s m/Ma)

Exposure age (ka)

External uncertainty (1s ka)

2.9797Eþ04 6.3279Eþ04 7.1528Eþ04 1.3867Eþ05

7.38 2.49 2.17 0.86

0.63 0.23 0.21 0.54

99,700 263,200 295,500 668,400

9,200 24,700 28,000 69,500

2.2243Eþ04 3.3710Eþ04 5.1028Eþ04

8.28 5.70 4.09

0.68 0.49 0.37

92,500 125,800 167,800

8,400 11,700 15,600

5.2645Eþ04 6.0704Eþ04

3.72 3.30

0.33 0.30

5.8753Eþ04

3.33

0.31

6.8646Eþ04

3.23

0.30

6.3227Eþ04

3.00

0.28

4.9431Eþ04 5.0067Eþ04

4.15 3.62 3.44 ± 0.31 (n ¼ 7)

0.37 0.33

a Samples CHK101, 102, 103, 105, 115, 116 and 117 are chert samples taken from chert conglomerate erosion remnants (Fig. 5). Elevations for these samples refer to the actual sample location; samples CHK 110, 111, 112, 113, 115, 118 and 207 are river sediment samples (Fig. 4). Elevations for these samples refer to the catchment-averaged elevations. Locations are given in Lat-Long decimal degrees (DD), WGS84. Sample 10Be content is provided relative to NIST SRM-4325 standard (10Be/9Be ¼ 2.790  1011; Fink and Smith [2007]). Uncertainties are given as 1s. The mean basin-averaged erosion rate for all catchments is an arithmetic mean (n ¼ 7) with standard deviation. The exposure ages (in ka) for chert conglomerate samples have been calculated with an assumed background erosion rate of zero.

Table 2 Basin-averaged erosion rates for each of the sampled catchments.a CHK110 Rock type (% total) Quartzite Chert conglomerate Dolerite Dolomite Shale Mafic volcanics Granite-gneiss Ultramafic schist Quaternary sediments Surface area (km2) Erosion rate (m/Ma) a

CHK111

CHK112

20.0

4.6

25.3

80.0

95.4

74.3 0.4

7 3.72 ± 0.33

3 3.30 ± 0.30

CHK113 4.6 25.4 1.4 15.1 44.3 9.1

15 3.33 ± 0.31

87 3.23 ± 0.30

CHK118 4.8 20.0 0.9 21.2 33.5 19.6

147 3.00 ± 0.28

CHK207 14.6 3.7 0.7 30.3 12.6 2.6 22.9 12.0 0.5 627 4.15 ± 0.37; 3.62 ± 0.33

Uncertainties are given as 1s. For each catchment the surface area (in km2) and underlying rock types (as % of total surface area) are indicated.

horizontal ledge, or terrace, separated by rises of 3e5 m. Each of these terraces can be interpreted to represent the base level of a paleo-river channel, which formed as the stream progressively incised into the dyke. Four chert samples (CHK102, 103, 104 and 105) were collected for 10Be analysis on the terrace surfaces to measure river incision rates. An additional chert sample (CHK101) was collected from a shelf approximately half a metre above the modern base level. The second site, site 2, is a prominent, 5.5 m high, dyke-like, chert conglomerate erosion remnant along the valley slope, located 280 m north-northwest of Malapa, and occurs at a height of 1435 m, i.e. at the same absolute height as Malapa Cave. The local valley bottom below the chert dyke is at 1398 m and the valley slope tops out at 1490 m (Fig. 5a). Three samples (CHK 115, 116, 117) were taken from along the crest of the chert dyke, at its highest elevation point, middle and lowest part, respectively (Fig. 5a).

2.3. Analytical methods Quartz was separated from samples, and 10Be isolated from the quartz at the Australian Nuclear Science and Technology Organization (ANSTO) laboratory, using a combination of standard HF/HNO3 methods (Child et al., 2000), and hot phosphoric acid (Mifsud et al., 2013). An additional leach in a warm 1% HF/HNO3 solution was performed in an ultrasonic bath to ensure complete removal of meteoric 10Be. All samples were purified to quartz with less than ~500 ppm Al, with final purified quartz masses ranging from 20 to 50 g. All samples were spiked with a 10Be-free solution of approximately 300 mg 9Be prepared from a beryl crystal solution at 1120 mg/g. Samples were loaded into quartz crucibles, dried, and calcined at 950  C for 3 h. The oxides were then crushed and mixed with niobium powder and loaded into aluminium holders for analysis by accelerator mass spectrometry (AMS) on the ANTARES accelerator against the NIST SRM-4325

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Figure 4. Sample location of river sediment taken for 10Be, basin-averaged, erosion rate measurements. The CoH is indicated with a white outline. The catchment perimeters (yellow (4a) and black (4b) lines) and sample locations are overlain on a sun-shaded digital terrain model (4a) and the geology (4b). The key to the geology is shown in Figure 2. The lake to the north of the figures is the Hartbeespoort Dam, and constitutes the local base level at 1164 m (water level, with base of the dam at 1125 m). The two principal catchments are those of the Skeerpoort River and its tributaries (sample CHK118), and the upper Crocodile River and Bloubank River (sample CHK207). Note the close correlation between topography and geology. North is towards the top of the figure.

P.H.G.M. Dirks et al. / Journal of Human Evolution 96 (2016) 19e34

27

Figure 5. Sample location and position of cosmogenic samples taken from erosion remnants of silicified, chert conglomerate dykes in dolomite in the Grootvleispruit catchment. Samples were taken from erosion remnants along the valley slope at GR 27.80054e25.889256 (a) and along a chert dyke traversing the river valley at a right angle, at GR 27.79272e25.87873 (b, c) where the dyke forms a knick point (Grid References in Long-Lat, WGS84). Terrace surfaces representing the position of former base levels of the river bed are marked with black dashed lines. Erosion corrected exposure ages for samples CHK 101, 102, and 105 are indicated (Table 3).

standard (10Be/9Be ¼ 2.790  1011; Fink and Smith, 2007). Two chemistry procedural blanks were also prepared with the same beryl spike and gave a mean 10Be/9Be value of 4.3 ± 0.2 1015 (n ¼ 7), which represents < 1% of the lowest 10Be/9Be ratio from the collected samples. Errors for final 10Be atoms/g concentrations are calculated by summing in quadrature the statistical error for the AMS measurement, 2% for reproducibility, and 1% for uncertainty in the beryl spike concentration. 2.4. Computational methods This paper makes use of the concentration of cosmogenic 10Be in quartz (Granger et al., 1996, 1997). At a given location, the 10Be concentration depends on the net result of a site-specific production rate that increases the number of atoms that can be measured in the sample, and losses due to decay and erosion. To first order, the concentration (N) of a cosmogenic nuclide i in an exposed surface rock depends on the erosion rate (E) and the exposure time (t) as follows (e.g., Lal, 1991):

 Ni ðtÞ ¼



Pi

rE∧1

þ li

1  eðrE∧

1

þli Þt



(1)

where Pi is the production rate, r the rock density, L the effective cosmic-ray nucleon attenuation length, and li the decay constant (lBe-10 ¼ 5.00  107 a1 [Chmeleff et al., 2010; Korschinek et al., 2010]). Equation (1) assumes that the irradiated rock surface is eroding at a constant rate, that there was no initial (or inherited) nuclide inventory and the surface has not been buried. Depending on the geomorphic setting, the cosmogenic nuclide concentration of a sample can be used to estimate either a maximum erosion rate (assuming steady-state or erosional equilibrium in which the production rate equals losses) or a minimum exposure age (assuming that the erosion rate of the surface rock equals zero). The site geomorphology and sample context can be used to assess whether erosion or exposure is the dominant process. Analytical results for the bedrock samples were processed and have been expressed as maximum erosion rates or minimum exposure ages (Table 1). Several major schemes exist for scaling cosmogenic nuclide production rates from one location on Earth to another (Balco et al., 2008). Here, we have used the CRONUS v2.2 online calculator (Balco et al., 2008) and the time-dependent scaling scheme of Lal (1991) updated by Stone (2000). Elevation averaged production rates were determined for each catchment from which stream sediment samples were collected, using elevation maps constructed from the South African digital geological

28

P.H.G.M. Dirks et al. / Journal of Human Evolution 96 (2016) 19e34

map series with a pixel resolution of 5 m. All bedrock samples were collected as a series of rock chips chiselled from a surface area ~100 cm2 in size (Fig. 3d); sample thicknesses were c. 5 cm. Horizon shielding by surrounding hills was measured at 45 intervals, where shielding effects are minor (<10 in polar angle), and by connecting points of shielding, where the effect is significant. Shielding corrections for bedrock samples were estimated using the CRONUS online calculator. To aid with the interpretation of the cosmogenic data, river profiles have been constructed for the Bloubank and Skeerpoort rivers draining the CoH using digitizing functions and the digital terrain model available in GoogleEarth (Fig. 6). The profiles were constructed to assess the overall maturity of the river profiles and evaluate the presence and nature of knick points in relation to erosion resistant lithologies such as chert units, or tectonic factors such as faults.

therefore, essentially the same across the CoH. The basin-averaged erosion rate for all samples taken together (by calculating an arithmetic mean with standard deviation) is 3.44 ± 0.31 m/Ma (Table 1). The principal trend observed in these data is that 10Be concentrations and inferred erosion rates are similar across the entire data set irrespective of catchment size or substrate (Table 2). In addition, these erosion rates are extremely slow when compared to global data sets of erosion rates based on 10Be data: Portenga and Bierman (2011), for example, report erosion rates that vary from near zero to almost 10,000 m/Ma, with a global median erosion rate for sedimentary basins of 54 m/Ma. The slightly different results between the different catchments reflect only small variations in the amount of material that would have been removed from each catchment in the past few Ma. Hence, catchment geometries and relative relief across the sub-catchments will have changed little on the timescales since hominins occupied the landscape in the CoH.

3. Results 3.2. Chert conglomerate samples 3.1. Stream sediment samples 10

Measured Be concentrations in the stream sediment samples correspond to a range of basin-averaged erosion rates of 3.00 m/Ma to 4.15 m/Ma (Tables 1 and 2). The six (sub-)catchments that drain parts of the CoH have basin-averaged erosion rates of 3.62 ± 0.33 m/Ma (CHK207L > 1 mm fraction) to 4.15 ± 0.37 m/Ma (CHK207 ¼ 0.25e0.5 mm fraction) for the Crocodile River catchment; 3.00 ± 0.28 m/Ma (CHK118) for the Skeerpoort River catchment above the bridge on Hartebeeshoek Farm; 3.23 ± 0.30 m/Ma (CHK113) for the Skeerpoort River sub-catchment above the confluence with the Grootvleispruit River; 3.33 ± 0.31 m/Ma (CHK112) for the Grootvleispruit River sub-catchment above a series of chert knick points (Fig. 6); and 3.72 ± 0.33 and 3.30 ± 0.30 m/ Ma (for CHK110 and CHK111 respectively) for sub-catchments of the Grootvleispruit River above Malapa. All uncertainties are given as 1s, i.e. the basin-averaged erosion rates are within 2s error, and,

Measured 10Be concentrations from the chert dyke at site 1 (Fig. 5b) correspond to maximum erosion rates of 0.86 ± 0.54 m/Ma to 7.38 ± 0.63 m/Ma, or minimum exposure ages of 668 ± 70 ka to 100 ± 9 ka respectively (Table 1; uncertainties provided as 1s). The sample with the highest 10Be concentration (CHK103, giving a minimum exposure age of 668 ± 70 ka and a maximum erosion rate of 0.86 ± 0.54 m/Ma) is at the top of the chert ridge, with systematically decreasing 10Be concentrations for samples lower down the ridge (Fig. 5c, Table 1). The sample with the lowest 10Be concentration (CHK101, giving a minimum exposure age of 100 ± 9 ka and a maximum erosion rate of 7.38 ± 0.63 m/Ma) is within the modern stream channel (Fig. 5; Table 1). Sample CHK104 (Fig. 5c) was found to be deeply weathered and altered, and was excluded from the analyses. Concentrations of 10Be from the chert dyke at site 2 also display an increase from the base to the top of the wall (Fig. 5a, Table 1).

1600

1600

1400

1400

1200

1200 1000

0

10 20 30 40 Distance from divide (km)

50

0 20.00

15.00

15.00

10.00 5.00

10.00

Chert

Slope (x1000)

20.00

5.00

0.00

20 40 60 80 Distance from divide (km)

Slope (x1000)

1000

100

Quartzite

1800

Height (m)

1800

Height (m)

b. Bloubank-Crocodile river: Profile & slope

a. Skeerpoort river: Profile & slope

0.00 0

10 20 30 40 Distance from divide (km)

50

0

20 40 60 80 Distance from divide (km)

100

Figure 6. River profiles and slope angles for the main rivers draining the CoH. (a) the Skeerpoort River up to the Hartebeestpoort Dam. (b) The Bloubank-Crocodile River system up to the Hartebeestpoort Dam.

P.H.G.M. Dirks et al. / Journal of Human Evolution 96 (2016) 19e34

These concentrations correspond to maximum erosion rates of 4.09 ± 0.37 m/Ma, 5.70 ± 0.49 m/Ma and 8.28 ± 0.68 m/Ma, or minimum exposure ages of 168 ± 16 ka, 126 ± 12 ka and 92.5 ± 8 ka for CHK117, CHK116 and CHK115, respectively (Table 1). 3.3. Stream profiles Profiles for the rivers and their slopes draining the CoH are shown in Figure 6. All profiles have typical concave shapes characteristic of mature rivers (Kirby and Whipple, 2012), and preserve one or more knick points where the rivers pass erosion resistant units. The Skeerpoort River shows a step in its profile at ~1400 m (after 12 km from the divide; Fig. 6a). For the next 7 km gradients are higher. Small tributaries that drain dolomite to the southeast of the Skeerpoort River, such as the Grootvleispruit River, show a series of vertical-step knick points characterized by waterfalls 2e10 m in height, between 1400 and 1300 m as they pass dyke-like erosion remnants of Rooihoogte Formation chert (e.g., Fig. 5a, b). The Bloubank-Crocodile River displays a stable, concave profile (Fig. 6b). Minor knick points occur where the Bloubank River passes erosion resistant quartzite of the Witwatersrand Supergroup at 1610 m (6 km from source) and 1390 m (28 km from source [Dirks and Berger, 2013]). If the Hartebeestpoort Dam (Fig. 2) is removed from the profile, the prominent quartzite ridge of the Magaliesberg quartzite does not form a knick point or slope break in the river profile, and the quartzite does not present a local base level controlling drainage in the CoH. 4. Discussion 4.1. Basin-averaged erosion rates Our data illustrate that for (sub-)catchments that drain the CoH, basin-averaged erosion rates are similar (ranging from 3.00 ± 0.28 m/Ma to 4.15 ± 0.37 m/Ma) with all results within 2s error. The slightly higher rate of erosion for the Crocodile River (CHK207 at an average rate of ~3.89 m/Ma), which drains the western portion of the Johannesburg Dome (Table 2; Fig. 1) may reflect slightly higher erosion rates for feldspar-rich, granite-gneiss, as opposed to quartz-rich, volcaniclastic and sedimentary cover rocks. The marginally higher 10Be concentration and, therefore, slightly lower erosion rates for the larger size fraction from the Crocodile River (CHK207L at 3.62 m/Ma as opposed to 4.15 m/Ma for the finer grained size fraction) may suggest that weathering of rocks to finer grain sizes proceeded more rapidly than processes resulting in coarser grain sizes (generally composed of weathering resistant chert fragments). All the basin-averaged erosion rates obtained in this study (Tables 1 and 2) are exceptionally slow by global standards, and represent some of the lowest reported anywhere (note, the global mean 10Be, basin-averaged erosion rate is 54 m/Ma [Portenga and Bierman, 2011]). The difference between the minimum obtained value of 3.00 ± 0.28 m/Ma for the Skeerpoort river catchment and the maximum obtained value of 4.15 ± 0.37 m/Ma for the fine-grained sediment fraction in the upper Crocodile River catchment, is small when compared to global variability, and even if this difference is real, it will have resulted in little differential weathering between various catchments at time scales of 1e3 million years; i.e. on average, throughout the CoH the physical character of the landscape will have changed little in the last 3 Ma. These results are consistent with basin-averaged erosion rates obtained from other parts of the southern African plateau or plateau rim (e.g., Scharf et al., 2013). Basin-averaged erosion rates across the CoH are within 2s error of one another, irrespective of catchment size or underlying geology (Table 2). In the Grootvleispruit River catchment, basin-averaged

29

erosion rates for the sub-catchment above Malapa Cave (CHK111, characterized by a broad open valley), the upper reaches of the Grootvleispruit River above Malapa Cave (CHK110, characterized by open valleys with river incision), and below the main knick point (CHK112, where river incision is prominent), are identical within 1s error at 3.30e3.72 m/Ma (across areas of 3e15 km2). In each of these sub-catchments, the substrate is dominated by dolomite (Table 2). These erosion rates also overlap with the basin-averaged erosion rates obtained for the larger Skeerpoort River catchment areas above samples CHK113 (at 3.23 m/Ma across 87 km2) and CHK 118 (at 3.00 m/Ma across 147 km2), that are characterized by varied topography (Figs 1, 4) in which shale forms the dominant lithology and dolomite forms only a minor proportion of the substrate (15.1% for CHK113 and 21.2% for CHK118 respectively; Tables 1 and 2). The erosion rates from catchments draining the CoH to the north west are also similar to the value obtained for the Bloubank-Crocodile River catchment, again characterized by varied topography that drains the CoH to the south east (3.62 m/Ma and 4.15 m/Ma across 627 km2). This observation is inconsistent with inferences made by Bailey et al. (2011) and Reynolds et al. (2011), who suggest that the topography of the Bloubank Valley may have been shaped by active faulting as deduced from the interpretation of lineaments on satellite imagery, even though there is no direct evidence of active faulting in the field (e.g., Dirks and Berger, 2013), nor is it evident from stream profiles (Fig. 6). Normally, valleys shaped by active, neo-tectonic processes experience basin-averaged erosion rates of the order of 10 s of m/Ma (Portenga and Bierman, 2011). In this case it may be the slow pace of the erosional processes that sculpt the landscape, which resulted in a close mimicking of the underlying geology, including older lineaments or lines of weakness within the rock. The similarity in erosion rates across all catchments suggests that in terms of overall sediment budget, the sediment contribution from local down cutting and valley floor lowering, most prominent in parts of the Skeerpoort River catchment, is small as a percentage of total sediment load. It also indicates that erosion rates in areas away from recent channel erosion and valley slope retreat are even lower than the CoH mean, basin-averaged erosion rate of 3.44 m/ Ma, which is consistent with the presence of erosion remnants of wad, and paleo-soil attributed to the Cretaceous (Partridge and Maud, 1987) and Miocene (Beukes et al., 1999). Catchments largely underlain by dolomite (e.g., CHK110, 111, 112) have eroded at the same rate (within 2s error) as catchments largely underlain by siliciclastic lithologies (e.g., CHK 113, 118, 207 Table 2), indicating that regionally, in the past few hundred thousand years at least, chemical weathering has been as fast as physical weathering. Normally, erosion proceeds by the removal of rock material from the top of bedrock, via a combination of physical, chemical and biological processes. Carbonate-rich rocks like dolomite are special in that they have the capacity to also weather chemically via dissolution from below. When conditions are right, i.e. when the climate is wet and the water table is high, carbonaterich rock can be removed via dissolution, which would be expected to speed up erosion rates, as well as create prominent karst topography. This, however, has not been observed (Tables 1 and 2) and suggests that, in the recent past at least, the area was affected by a relatively dry climate similar to today. Despite the similarity of erosion rates across different substrates, the geology is expressed in the topography, and, therefore, the small differential erosion rates that exist between lithologies are being imprinted over long time scales. This is a clear indication that the landscape evolved over a long period of time as the geology and topography mimic each other on a regional scale (Fig. 3). For example, the Crocodile River catchment is represented by the Johannesburg Dome composed of granite and gneiss, flat-lying geologic units form plateaus, quartzite units and other

30

P.H.G.M. Dirks et al. / Journal of Human Evolution 96 (2016) 19e34

weathering resistant lithologies form ridges (see also Scharf et al., 2013), and folded areas show diverse topography. This is also consistent with the mature concave profiles for rivers draining the CoH (Kirby and Whipple, 2012; Fig. 6). 4.2. River incision and local erosion rates around Malapa Although regional basin-averaged erosion rates are within 2s error of each other (Table 1) across the entire CoH, there is clear evidence that, locally, erosion rates can vary considerably (e.g., Dirks et al., 2010; Dirks and Berger, 2013). We have tried to quantify this variability more precisely for the area around Malapa Cave to better constrain the rate caves are being exposed and de-roofed along the Grootvleispruit River valley. The basin-averaged erosion rates obtained from samples located immediately below Malapa (CHK110 and CHK111) indicate that ~7 m of material would have been removed from Malapa Cave in the past 2 Ma (Table 1) if the site is representative for erosion rates across the sub-catchments, providing a first order estimate for the depth of Malapa Cave at the time of burial of A. sediba. However, various geomorphological features (e.g., see Dirks and Berger, 2013) indicate that erosion rates along the valley floor of the Grootvleispruit and Skeerpoort Rivers are locally considerably higher, with an incision rate of 53 m/Ma recorded for the Skeerpoort River below Gladysvale Cave, where the river is undercutting cliffs of soft shale (Dirks et al., 2010). A measure for the incision rate of the Grootvleispruit River below Malapa can be obtained from site 1 using the exposure ages calculated for samples from the horizontal terraces (CHK101, 102, 105; Fig. 5c), which have been interpreted as earlier base levels for the stream bed. Table 1 provides estimates for the minimum exposure age for each terrace, which have been calculated by assuming an erosion rate of zero. The chert wall at site 1 is resistant to erosion, a situation evidenced by its great height (Fig. 5), with each successive younger terrace surface experiencing a lower exposure period as the river continued to incise. Given that the highest terrace sample, CHK103, gives the highest 10Be concentration, we infer that this site represents erosional steady-state, and, thus, can provide a good estimate for the maximum (possibly chemical) weathering rate of the chert to allow erosion-corrected exposure ages to be calculated for the lower terraces as per equation (1). This implies that the terrace surface at CHK103 has been exposed for a sufficiently long period of time (~1.5 Ma) to have arrived at steady-state for the chert lithology. The maximum erosion rate calculated for CHK103 of 0.86 m/Ma (Table 1) is used to estimate the minimum erosion corrected exposure ages for each of the terraces (Table 3). These ages combined with the height differences between terraces along the wall (Fig. 5c) provide estimates for differential incision rates of 14.2 m/ Ma (between CHK105 to CHK101) and 57 m/Ma (between CHK102 to CHK105; Table 3). Such high incision rates cannot reflect the long-term average incision rate of the Grootvleispruit River for the

simple reason that the top of the sampled chert wall (at CHK103) is currently 14.2 m above the river channel (CHK101; Fig. 5c) and eroding at a very slow rate. If the top of the wall is lowered at 0.86 m/Ma then, say, 3 Ma ago, this part of the wall would have been ~17 m above the current base level of the channel. At incision rates of 14e57 m/Ma, the channel would have been many 10s of metres above the current base level, i.e. well above the expected height of the wall at that time, and inconsistent with the steepsided, symmetrical channel morphology of the valley. We, therefore, interpret the high incision rates (or rather, young ages) obtained for samples CHK102 and CHK105 to reflect repeated catastrophic failure of the wall in the past 400 ka. This episodic failure is probably reflected in the stepped nature of the wall (Fig. 5c). Using this logic, since 370 ka (the exposure age of CHK102) the wall would have failed possibly twice for it to be lowered by 6 m, in two ~3 m steps. There is no guarantee that the wall would have collapsed at a similar rate further into the past. If a constant erosion rate of 0.86 m/Ma for CHK103 is assumed, at the time A. sediba was living in the valley 2 Ma ago (Pickering et al., 2011a), the chert wall would have been ~16 m high, and, therefore, the average maximum incision rate for the stream in the past 2 Ma would have been 8 m/Ma; i.e. about double the basin average erosion rate. For site 2 a similar approach can be taken assuming a steadystate erosion rate for the chert of 0.86 m/Ma. This results in calculated surface lowering rates of 51 m/Ma and 33 m/Ma for CHK115 and CHK116, respectively (Table 3). These high erosion rates again cannot reflect long-term average lowering rates, but rather suggest that the chert dyke at site 2 experienced recent partial collapse events; i.e. these surfaces are not stable, leading to an overestimation of the erosion rate along this slope. The calculated maximum erosion rate of 8.28 ± 0.68 m/Ma for CHK115 at the base of the erosion remnant (Table 1) appears to be the more consistent estimate for the rate of landscape change in the area. Taken together, the basin-averaged erosion rates and rates for sites 1 and 2 suggest that the nearby Malapa Cave site may have experienced unroofing of the order of no less than ~7 m and no more than ~16 m in the past 2 Ma, i.e. less than the earlier estimates of >30 m (Dirks et al., 2010; Dirks and Berger, 2013). These earlier estimates were based on a linear correlation between the erosion rate measured along the plateau above Malapa and a measurement from a chert surface along the Skeerpoort River below Gladysvale Cave, which returned an anomalously high erosion rate. The data presented here show that such anomalously high erosion rates are due to local geomorphological processes, and should be used with care. 4.3. Implications for cave formation, cave opening and fossil trapping Fossils are preserved in caves because caves were open to the surface to allow faunal remains to be trapped either by accident or

Table 3 Exposure ages (in ka) and incision rates (in m/Ma) calculated for chert conglomerate samples.a Sample name

Relative elevation (m)

Site 1: Chert dyke crossing Grootvleispruit river (Fig. 5b,c) CHK101 0 CHK105 3.05 CHK102 6.3 CHK103 14.2 Site 2: Chert dyke along valley slope (Fig. 5a) CHK115 1.75 CHK116 3.7 CHK117 5.35 a

Exposure age (ka)

External uncertainty (ka)

103,140 313,000 370,200 0.86 m/Ma

10,200 37,100 46,700 e

95,100 133,500 183,800

92,60 13,597 19,500

Maximum incision rate (m/Ma) 14.2 57.0 e e 50.8 32.8 e

These were calculated using a minimum erosion rate of 0.86 m/Ma for chert obtained from sample CHK103 (Table 1). Uncertainties are given as 1s.

P.H.G.M. Dirks et al. / Journal of Human Evolution 96 (2016) 19e34

with the help of physical and/or biological agents (e.g., Tobias, 2000; Partridge et al., 2003; de Ruiter et al., 2009; Pickering and Kramers, 2010; Herries and Shaw, 2011; Reynolds and Kibii, 2011; Pickering et al., 2011b; Dirks et al., 2015). In working out the relationship between fossils and caves it is of crucial importance to understand when caves formed and when they opened to the surface. It is generally assumed that caves within the dolomite of the Malmani Subgroup were present throughout the Pliocene and perhaps earlier (Partridge, 1973; Martini, 2006), and that their exposure to the surface was either due to river incision and knick point retreat (Partridge, 1973), or the removal of overburden including Time Ball Hill shale, Karoo sediment or wad (Dirks and Berger, 2013). It can be demonstrated that cave opening through river incision occurred locally (Dirks and Berger, 2013), however, when viewed regionally, many caves, including fossil-bearing caves, do not occur near strongly incised rivers (e.g., Dirks and Berger, 2013), but instead lie scattered across the catchments draining the CoH. The 10Be basin-averaged erosion rates of 3.00e4.15 m/Ma found in this study (Tables 1 and 2), indicate that, on average, the land surface across the CoH did not lower by much in the last few Ma, and, in fact, the observed erosion rates rank amongst the slowest recorded on Earth (Portenga and Bierman, 2011). Locally, things may have eroded a little faster along river channels (Dirks et al., 2010), but overall, the land surface would have been lowered very slowly, and in large areas, the change in the past 2e3 Ma would have been minimal. There is a very important implication to this with respect to the trapping of faunal remains in caves. Because basin-averaged erosion rates in the CoH are extremely slow (Table 1), only small amounts of overburden will have been removed in the past few million years. This means that 1 to 5 Ma ago there must have been significant areas of dolomite outcropping on the surface across the CoH, where dolomite would have been exposed to periods of climate, vegetation and soil cover change. The slow erosion rates also mean that if caves existed within the dolomite as far back as 4 or 5 Ma ago, or even further back in time into the Miocene, some of these caves were likely open to the surface to trap sediment and faunal remains, and some would have been preserved today, in the same way that it is possible to still find palaeo-caverns infilled with Karoo-aged sediment (Wilkins et al., 1987; Dirks and Berger, 2013). For instance, various caves in the CoH are currently well over 40 m deep (e.g., Sterkfontein). At landscape lowering rates of 3e4 m/Ma (see also Granger et al., 2015) it will take many millions of years for sediment accumulations at the bottom of these cave systems to be exposed to the surface. The slow basin-averaged erosion rates in the CoH, therefore, raise the very interesting question about why we don't find older, early Pliocene- or Miocene-aged cave deposits in the CoH e at least they have not been reliably described from any of the cave systems. Is it that caves did not exist in the Early Pliocene and Miocene, or that they did exist, but have since been eroded away, or that they were not exposed to the surface? The extremely slow erosion rates strongly suggest that caves could have survived for a very long time, which implies that the caves may not have formed until fairly recently. An estimate for the minimum age of the caves can be obtained from the age of sedimentary deposits within them, bearing in mind that the age of the cavity and the age of the cave sediments can vary greatly (e.g., Palmer, 1991; Sasowsky, 1998; Stock et al., 2005; White, 2015). Most cave deposits with macro fauna in the CoH are less than 2.0 Ma old (for a summary see O'Regan and Reynolds, 2009), with older assemblages of up to 2.8 Ma described from Sterkfontein Member 4 (Vrba, 1975, 1976; Reynolds and Kibii, 2011). The oldest reported assemblage in the CoH is from Bolt's Farm, where micro-faunal remains have been interpreted at 4e4.5 Ma (Gommery et al., 2014), but there are no absolute dates to confirm

31

this age. The exact age of STW573 at Sterkfontein also remains contested (e.g., Partridge et al., 2003; Walker et al., 2006; Pickering and Kramers, 2010; Bruxelles et al., 2014; Granger et al., 2015), but considering all available data, it is likely to be less than 3 Ma (Reynolds and Kibii, 2011). Direct age dating of a basal flowstone in drill core from Sterkfontein Cave provides a maximum age for the cave deposits of 2.80 ± 0.28 Ma (Pickering and Kramers, 2010). The oldest dated flowstone from Swartkrans is ~2.3 Ma (Pickering et al., 2011b). Cosmogenic burial ages from Sterkfontein Cave are as old as ~4 Ma (Partridge et al., 2003; Granger et al., 2015), providing evidence for some of the oldest cave sediments in the CoH. Using 10Be and 26Al, Granger et al. (2015) report a burial-isochron age of 3.67 ± 0.16 Ma for sediment encasing STW573. They make the assumption that this age also provides an age for STW573, although this cannot be absolutely proven, because the isochron age can also be replicated by re-deposition and mixing of surface and cave sediment that previously resided close to surface. This age does, however, indicate that some form of cavity was present at Sterkfontein as far back as 3.7 Ma, providing the oldest reliable age estimate for a cavity in the CoH (e.g., Stock et al., 2005). Thus, if all existing evidence is taken together, the oldest cave deposits in the CoH are no more than ~4 Ma, with all conclusively dated flow stone deposits being younger than 2.8 Ma. This strongly suggests that the caves in the CoH were not accumulating fossilbearing clastic sediments, until 4 Ma, but for most caves more likely ~3 Ma. This would suggest that caves either did not form until 3e4 Ma ago, or that they were not exposed until that time (e.g., Palmer, 1991; Sasowsky, 1998), although our 10Be results indicate that dolomite would have been at surface. Thus, the absence of early Pliocene or Miocene aged cave deposits in the CoH strongly suggests that caves only started forming from 4 Ma, and for most cave systems probably only from 3 Ma onwards; i.e. whilst the land surface is old, caves within the landscape may be relatively young. This, in turn, raises the question of how and why caves formed between 4 and 2 Ma, if the regional landscape was stable and did not change significantly. Answering this question fully is well outside the scope of this paper, but has relevance to understanding the overall landscape dynamics in the CoH since the mid-Pliocene, and we will therefore briefly mention some of the salient points. Caves in the CoH are maze caves (Kavalieris and Martini, 1976; Palmer, 1991; Martini, 2006; Dirks and Berger, 2013), consisting of networks of mostly narrow, vertical cavities that formed along intersecting vertical joints, bedding planes and layer-parallel shear zones (Dirks and Berger, 2013). Most of the cave systems are small with limited connectivity between caves, and their morphology is indicative of diffuse recharge along a multitude of dilatant fractures, with dissolution taking place mostly within the vadose zone (Palmer, 1991; White and Culver, 2012). Intersection points between fractures are typically characterized by larger dissolution chambers, with further enlargement of chambers taking place under phreatic conditions as groundwater levels rose during wetter periods. Additional sculpting of caves took place through collapse and the upward propagation of cavities along fractures and joints (Martini, 2006). Locally there is also evidence for hypogenic cave formation (Palmer, 1991; White, 2015), with evidence for hot spring activity recorded in some caves (e.g., as evidenced by quartz veins intruding Malapa Cave sediments [Dirks et al., 2010]), but the extent of this process in the CoH is unknown. Along the Grootvleispruit River in the area upslope from Malapa, caves deepen at approximately the same rate as the land surface rises along the valley slope, suggesting that caves in this area had formed down to a similar depth prior to incision and valley slope erosion of the Grootvleispruit River (Dirks et al., 2010; Dirks and Berger, 2013). This observation has been made more generally for

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caves in the COH and beyond (e.g., Martini, 2006) and suggests that caves formed more or less at the same time across the CoH, when the paleo-water table was considerably higher, and that in the last few million years new cave formation, and deepening of existing cave systems was slow (Dirks and Berger, 2013). A drop in the water table, presumably coinciding with an initial drying of the climate (deMenocal, 2004), could result in halting cave formation processes (Palmer, 1991; White and Culver, 2012) and at the same time make caves accessible for entry. The drying of southern Africa, as documented by a general shift from C3 to C4 grasslands, started around 2.8 Ma, but became especially pronounced from 1.7 Ma onward galen et al., 2007), (e.g., deMenocal, 2004; Hopley et al., 2007; Se and incursions of Kalahari sand in the CoH are noted possibly as early as 2.5e2.8 Ma (Beukes et al., 1999). This drying out of the climate is entirely consistent with our observation that basinaveraged erosion rates are the same across catchments underlain by siliciclastic lithologies (i.e. shale, sandstone, granite) and dolomite, which may have eroded differently under wet conditions (e.g., Chadwick et al., 2013). On a regional scale, caves are controlled by lithological contacts within the Malmani dolomite and north-northeast and eastsoutheast trending basement fractures (Kavalieris and Martini, 1976; Martini et al., 2003; Martini, 2006; Dirks and Berger, 2013). At the time of cave formation, these fractures must have been open to allow water ingress. An extensional environment (Bird et al., 2006; Viola et al., 2012; Dirks and Berger, 2013) in which multiple fractures are allowed to open simultaneously would facilitate a diffuse groundwater recharge system, necessary to form the distributed maze caves that characterize the CoH (Palmer, 1991). One possible explanation for upper crustal extension of rocks on the South African high plateau and consequent cave formation is uplift of the entire landscape in the early Pliocene (e.g., Partridge and Maud, 1987; Moore et al., 2009). Such uplift may not only have influenced cave formation, but could have also contributed to the onset of drier climates from 2.8 Ma onward.

last 3 Ma. This is consistent with the concave shape of the river profiles draining the CoH in which knick points are controlled by erosion resistant units. The slow basin-averaged erosion rates indicate that dolomite was exposed on surface for a long period into the past; i.e probably well into the Miocene. This fact, in combination with the slow erosion rate of the landscape, and therefore slow exposure of caves and slow erosional removal of fossil-bearing cave breccia that would have become exposed on the surface, raises the important question about why no early Pliocene- or Miocene-aged cave deposits and fossils have been found in the CoH. We conclude that the paucity of such older deposits, taken together with the slow erosion rate of the general land surface, indicate that caves only started forming from 4 Ma, and perhaps mostly only 3 Ma, onwards; i.e. whilst the land surface in the CoH is generally old, caves within the landscape are a relatively young phenomenon, and their formation and opening controlled the maximum age of fossils that can potentially be found in caves in this region.

5. Conclusion

References

Basin-averaged erosion rates for catchments draining the CoH, as determined by cosmogenic 10Be measured in quartz from samples of river sediment, are within 2s error of each other, at 3.00 ± 0.28 m/Ma to 4.15 ± 0.37 m/Ma (uncertainties in 1s), and do not vary significantly with catchment size, landscape morphology or underlying geology. Thus, catchments dominated by a substrate of dolomite erode at the same rate as catchments underlain by mostly siliciclastic rocks, and erosion rates for sub-catchments vary little with overall erosion rates for larger catchments in which the sub-catchments are located. The fact that dolomite- and siliciclastic-dominated catchment areas erode at similar rates indicates that physical weathering has been the rate controlling process, with limited contributions from chemical weathering of dolomite. This, in turn, suggests that, in the recent past at least, the area experienced a relatively dry climate, similar to today. Erosion resistant chert conglomerate units along valley floors provide information on incision rates of river valleys. Assuming a steady-state erosion rate for chert dykes of 0.86 ± 0.54 m/Ma, an average incision rate of ~8 m/Ma has been obtained for the Rietvleispruit River below Malapa. This estimate together with the basin-averaged erosion rate indicates that Malapa Cave was at least 7 m, but probably no more than 16 m deep at the time of deposition of A. sediba. The low basin-averaged erosion values indicate that the landscape across the CoH is old and eroding slowly, and throughout the CoH the physical character of the landscape has changed little in the

Acknowledgements We would like to thank the three anonymous reviewers for their detailed and valuable input, which has helped to improve this manuscript. We thank the South African Heritage Resources Agency and the Nash family for allowing generous access to sites. Funding was received from the Australian Research Council (DP140104282), Australian Institute of Nuclear Science and Technology (AINSE grant #ALNGRA13015), the University of the Witwatersrand and, James Cook University. Thanks also to the University of the Witwatersrand's Schools of Geosciences and Bernard Price Institute for Palaeontology for extensive logistical support. Samples around Malapa were collected with permission from the South African Heritage Resources Agency (SAHRA) under permit number 80/08/ 09/001/51.

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