Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls

Abstract

The summer rainfall zone (SRZ) in the South African interior experienced pronounced hydrological and vegetation changes during the Holocene inferred to be driven mainly by shifts in atmospheric and oceanic circulations systems. The exact mechanisms controlling these changes are still debated. To gain better insights into the Holocene environmental changes in the South African SRZ and their driving factors, we analysed compound-specific carbon and hydrogen isotopes of plant wax n-alkanes (δ¹³C_wax and δD_wax) from a marine sediment core covering the last 9900 years. The core has been recovered offshore the mouth of the Orange River, predominantly draining the South African summer rainfall region. Our data indicate a dry early Holocene and a gradual increase of wetter conditions with a higher abundance of C₄ vegetation towards the middle Holocene. Wettest conditions occurred around 3900 cal. yr BP. The last 3900 years were characterised by a gradual aridification overlain by variable wetter conditions. During the ‘Little Ice Age’ (LIA: ca. 640–310 cal. yr BP), relatively dry conditions with elevated C₄ plant contributions occurred. This opposite behaviour, that is, more C₄ plant contribution during drier conditions compared to the remainder of the Holocene, points towards an influence of winter rainfall in the lower Orange River catchment during the late-Holocene and a decline in summer rainfall. We emphasise the importance of changes in the latitudinal insolation gradient (LIG) as a potentially important controlling mechanism for hydrologic and vegetation changes in the SRZ.

Keywords

compound-specific isotopes Holocene insolation gradient palaeoclimate South Africa summer rainfall zone

Introduction

Climate in southern Africa is sensitive to changes in large-scale atmospheric and oceanic circulation systems inducing pronounced seasonal precipitation variability. Although palaeoenvironmental investigations do not yield homogeneous results neither for the summer rainfall zone (SRZ) nor the winter rainfall zone (WRZ) of South Africa (Figure 1), they generally show opposing Holocene climatic trends in both regions (e.g. Burrough and Thomas, 2013; Chase and Meadows, 2007; Hahn et al., 2016; Weldeab et al., 2013; Zhao et al., 2016). While the northeast and central parts of the SRZ were found to have experienced relatively dry conditions during the early Holocene (Chevalier and Chase, 2015; Holmgren et al., 2003; Kristen et al., 2010; Metwally et al., 2014; Zhao et al., 2016), the WRZ exhibited relatively wet climatic conditions (Chase et al., 2015a; Scott and Woodborne, 2007; Scott et al., 2012; Valsecchi et al., 2013; Weldeab et al., 2013; Zhao et al., 2016). During the transition from the early to middle Holocene, increasingly wetter conditions were found for the SRZ, particularly in the north-eastern area (e.g. Kristen et al., 2010; Metwally et al., 2014; Zhao et al., 2016). In contrast, the WRZ became drier (e.g. Chase and Thomas, 2006; Weldeab et al., 2013; Zhao et al., 2016). Records for the late Holocene, however, show diverging trends with partially wetter conditions in the SRZ as well as in the WRZ (e.g. Brook et al., 2010; Neumann et al., 2014; Valsecchi et al., 2013). Other studies inferred drier conditions in some parts of both regions during the Holocene (Breman et al., 2012; Chase et al., 2015b; Truc et al., 2013). Rainfall changes in the WRZ are supposed to be mainly driven by the latitudinal changes in position of the Southern Hemisphere Westerlies (SHW; Chase and Meadows, 2007; Chase and Thomas, 2007; Cockcroft et al., 1987; Reason and Rouault, 2005; Stager et al., 2012), partially controlled by changes in sea ice extent around Antarctica (Blamey and Reason, 2007; Gasse et al., 2008; Reason et al., 2002; Stuut et al., 2004) and also by the Benguela upwelling intensity (Lim et al., 2016; Shi et al., 2000). Changes in the SRZ were linked to latitudinal migrations of the tropical easterlies (Chevalier and Chase, 2015; Gasse et al., 2008; Norström et al., 2009), changes in Indian Ocean sea surface temperatures providing moisture for summer precipitation (Chevalier and Chase, 2015; Truc et al., 2013; Weldeab et al., 2014), and local insolation changes controlling atmospheric convection and moisture advection into the SRZ (Chevalier and Chase, 2015; Gasse et al., 2008; Partridge, 1997). There are, however, temporal offsets between the maximum local insolation and maximum rainfall as well as the most humid/warmest periods in palaeorecords (e.g. Chevalier and Chase, 2015; Kristen et al., 2010; Lim et al., 2016; Lyons et al., 2014; Norström et al., 2014; Truc et al., 2013; Zhao et al., 2016). In respect of this discrepancies, model simulations (e.g. Bosmans et al., 2012, 2015; Dallmeyer et al., 2013, 2015; Davis and Brewer, 2009, 2011; Loutre et al., 2004) and palaeoenvironmental studies (Chen et al., 2016; Kuechler et al., 2018; Li and Wu, 2010; Moossen et al., 2015; Risebrobakken et al., 2007; Thevenon et al., 2002; Vimeux et al., 2001) demonstrate the importance of changes in the latitudinal insolation gradient (LIG) rather than that of local insolation for past climatic changes. While the local summer insolation in low latitudes is mainly driven by precession, the driving forcing behind the LIG is the obliquity of the Earth’s axis (Davis and Brewer, 2009). The LIG leads to different heating between different latitudes. As a consequence, the LIG influences the latitudinal temperature gradient and, therefore, affects the latitudinal position and extent of atmospheric circulation patterns, such as the Hadley Cell, resulting in a displacement and change of strength of rain-bearing systems (Bosmans et al., 2012; Davis and Brewer, 2009, 2011; Jiang et al., 2015; Young and Bradley, 1984). Changes of the LIG are yet, however, often not considered in palaeoenvironmental reconstructions. The few palaeoenvironmental reconstructions considering the LIG have shown that, for instance, the strength of the Indian Summer Monsoon over Tibet closely follows the summer LIG between 30°N and 44°N, with highest Indian Summer Monsoon intensities during a LIG maximum rather than summer insolation maximum (Ramisch et al., 2016). Recently, it has been suggested that the West African climate was dominated by the LIG when precessional variability was low (Kuechler et al., 2018). Furthermore, Vimeux et al. (2001) demonstrated that a stronger annual LIG between low and high latitudes induced a shift of Antarctic precipitation regions. The LIG between middle and low latitudes was also been found to be responsible for the change of sea ice expansion in the East Antarctic realm (Denis et al., 2010). However, for the southern African region, the LIG has so far not been considered in palaeoenvironmental reconstructions.

Figure 1.

Modern mean monthly precipitation (1960–1990 avg., WorldClim 1.4 Global Climate data at 1 km² resolution; Hijmans et al., 2005) as well as atmospheric and oceanic circulation systems over South Africa during austral (a) summer (January) and (b) winter (July). Major atmospheric circulations systems, South East (SE) trade winds as well as easterlies and westerlies, and the oceanic Benguela and Agulhas currents are shown (Tyson and Preston-Whyte, 2000). The three rainfall zones of southern Africa are indicated as summer rainfall zone (SRZ), year-round rainfall zone (YRZ) and winter rainfall zone (WRZ).

The present-day transition zone between the SRZ and the WRZ in the central and western parts forms an important area to study the interaction between the SRZ and the WRZ on millennial timescales. However, arid conditions in this area hamper the formation of suitable archives. To date, only few high-resolution records exist from South Africa covering the entire Holocene period, such as from the Braamhoek wetlands, Mahwaqua, Makapansgat valley or De Rif rock hyrax middens in the Cederberg mountains (Holmgren et al., 2003; Neumann et al., 2014; Norström et al., 2009; Valsecchi et al., 2013). Most of the archives are discontinuous (e.g. Scott and Woodborne, 2007; Talma and Vogel, 1992; Weldeab et al., 2013), cover only short time intervals (Brook et al., 2010; Chase et al., 2015b; Kristen et al., 2010; Stager et al., 2012) or have a coarse temporal resolution (e.g. Lyons et al., 2014; Thackeray and Lee-Thorp, 1992). The gap in suitable archives can be filled by the marine sediment core GeoB8331-4 from the west-coast mudbelt, a Holocene sediment package, offshore western South Africa (Herbert and Compton, 2007; Mabote et al., 1997; Meadows et al., 2002). It has been shown that terrestrial organic material deposited in the northern mudbelt originates from the western and interior parts of the South African SRZ with a potential bias towards the western parts of the SRZ, reflecting a spatially integrated climatic signal (Herrmann et al., 2016; Zhao et al., 2016).

In this study, compound-specific carbon and hydrogen isotopes of leaf wax n-alkanes (δ¹³C_wax, δD_wax) are used to (1) reconstruct Holocene vegetation and hydrological changes in the central and western parts of the SRZ, (2) compare its climatic history to other palaeoenvironmental reconstructions from the region and (3) discuss potential forcing factors.

Regional setting

Climate and vegetation

South African’s climate is influenced by the tropical easterlies in the north and the SHW in the south as well as the effects of the cold Benguela and warm Agulhas currents in the Atlantic and Indian Ocean, respectively (Tyson and Preston-Whyte, 2000). The interplay of the atmospheric circulation systems leads to a pronounced seasonal precipitation variability in southern Africa. The SRZ (>66% precipitation between October and March) reaches from the interior to the east of southern Africa and receives precipitation mainly by tropical easterly winds from the Indian Ocean (Figure 1). In contrast, during austral summer, the WRZ (>66% precipitation between April and September) experiences low precipitation rates when the rain-bearing SHW are in their southernmost position. During austral winter, the SHW move equatorward to the southwestern Cape and lead to higher precipitation rates in the WRZ. In the SRZ, dry conditions are dominant during austral winter as a high pressure cell prevails over the interior and eastern southern Africa. A dynamic transition zone, the year-round rainfall zone (YRZ) receives equal amounts of precipitation during the summer and winter months (Chase and Meadows, 2007; Tyson and Preston-Whyte, 2000). Mean annual precipitation (MAP) and mean annual temperatures (MATs) are highly variable and range from ca. 20 to 1300 mm and 5°C to 23°C in the SRZ and from ca. 30 to 1200 mm and 10°C to 20°C in the WRZ (Hijmans et al., 2005), respectively.

These variable climatic conditions, together with other factors such as orography and soil diversity, generate a highly diverse vegetation distribution in southern Africa (Figure 2). It ranges from the Mediterranean Fynbos Biome in the southwestern Cape to the dry Succulent Karoo and arid desert in the northwestern parts as well as eastward from the Nama Karoo through Savanna, Grassland Thicket and Coastal Forest in the humid east (Cowling et al., 1997; Mucina and Rutherford, 2006). The vegetation of southern Africa consists of C₃, C₄ and CAM (Crassulacean acid metabolism) plants (Werger and Ellis, 1981). Woody C₃ plants and C₃ grasses are dominant in the high altitude grasslands of the Drakensberg Mountains and the Fynbos Biome. In contrast, C₄ grasses become more abundant with increasing aridity and growing season temperatures (Scott and Vogel, 2000; Vogel et al., 1978) and show highest abundance in the interior of southern Africa (Werger and Ellis, 1981). CAM plants exist throughout southern Africa but are most dominant in arid regions with strong precipitation seasonality like the Succulent Karoo and western Nama Karoo biomes (Mooney et al., 1977; Werger and Ellis, 1981).

Figure 2.

Overview of biomes in southern Africa and sample sites of this study. Colours indicate modern vegetation of southern Africa (after Mucina and Rutherford, 2006; Scott et al., 2012). Dots denote the locations of river samples (red) and marine surface sediments (blue). The location of the sediment core GeoB8331-4 (red star) is indicated. Black squares indicate terrestrial archives in South Africa discussed in the text. Rivers entering the Atlantic Ocean at the west coast are shown (VV = Verlorenvlei).

Marine sediments offshore southwestern Africa

The Namaqualand mudbelt (hereafter referred to as mudbelt) is a terrestrial Holocene sediment package along the west coast of South Africa. It stretches over 500 km and reaches a thickness of about 35 m near the Orange River mouth in the north (29°S) to about 2 m near the Berg River in the south (33°S; Birch, 1977; Meadows et al., 2002; Rogers and Bremner, 1991; Rogers and Rau, 2006; Schneider et al., 2003). This makes the mudbelt a high-resolution sedimentary archive attractive for Holocene environmental investigations (Birch, 1977; Grey, 2009; Hahn et al., 2016; Herbert and Compton, 2007; Leduc et al., 2010; Mabote et al., 1997; Weldeab et al., 2013; Zhao et al., 2016). The Orange River is the main sediment supplier for the mudbelt and comprises a catchment of almost 10⁶ km² with a Holocene mean annual mud flux of 5.1 million tonnes (Compton et al., 2010). The mud fraction of the suspended material is transported southward from the Orange River mouth by a poleward undercurrent (Birch et al., 1991; Mabote et al., 1997; Rogers and Bremner, 1991). In the central and southern parts, the mudbelt receives sediment contributions from the adjacent west coast biomes by berg-winds as well as the local ephemeral rivers and the Olifants and Berg rivers are the dominant sediment source to the southernmost mudbelt (Benito et al., 2011; Birch, 1977; Grey, 2009; Grey et al., 2000; Herrmann et al., 2016; Mabote et al., 1997; Zhao et al., 2015).

Material and methods

Suspension loads and flood deposits from the Orange River, collected in March 2014, were stored in combusted glass jars and plastic bags, respectively. Suspension loads were retrieved by centrifuging 100 L of pumped river water.

Marine surface sediments (first centimetre) of nine multi-cores from the mudbelt, recovered during RV Meteor cruise M57/1 in January–February 2003 (Schneider et al., 2003), were sampled for investigation of differential contribution to the mudbelt. ²¹⁰Pb_ex dating of the multi-cores GeoB8331-2 and GeoB8332-3 in the northern mudbelt revealed ages of less than 3 years for the upper centimetre (Leduc et al., 2010). Radiocarbon dating of the multi-cores GeoB8319-1 and GeoB8322-1 in the southern mudbelt show ages of 110–40 cal. yr BP between 11 and 14 cm (GeoB8319-1) and 630–370 cal. yr BP at 43 cm (GeoB8322-1), respectively (Taylor, 2004). In all studied, multi-cores pollen of neophytes (e.g. Pinus, Quercus), which were introduced during the 17th century in South Africa (Campbell and Moll, 1977; Richardson, 2000), were found in the upper three centimetre (Zhao et al., 2015). Thus, the upper first centimetre of the multi-cores can be considered to be of ‘modern’ age.

Gravity core GeoB8331-4 (887 cm, 29°08.12′S, 16°42.99′E) was retrieved from the mudbelt during RV Meteor cruise M57/1 at a water depth of 97 m. The sediment core consists of olive brown mud with a dark laminated layer of olive brown mud in the upper 15 cm (Schneider et al., 2003). The age model, based on seven radiocarbon ages of Nassarius vinctus, is described by Herbert and Compton (2007) and was extended with two additional radiocarbon ages by Hahn et al. (2015; Figure 3; Supplementary Table 1, available online). The core has a basal age of ca. 9900 cal. years BP. In total, 78 samples were collected with an average temporal resolution of ca. 40–240 years for the last 2000 cal. years and 2000–9900 cal. years BP, respectively.

Figure 3.

Calibrated ages with calculated uncertainty (2σ) for Geob8331-4 (Hahn et al., 2016; Herbert and Compton, 2007).

Preparation and lipid extraction

All samples were freeze-dried and grounded using a planetary mill (suspension loads) and an agate mortar and pestle (flood deposits and marine sediments). An internal standard (squalane) was added before extraction. Lipid extraction was carried out with a DIONEX Accelerated Solvent Extractor (ASE 200) at 100°C, 1000 psi for 5 min by using a solvent mixture of dichloromethane and methanol (DCM:MeOH, 9:1). The extraction was repeated three times for each sample, and the resultant total lipid extracts (TLEs) were concentrated by rotary evaporation. TLEs were desulphurised with activated copper. The hexane-soluble fractions were separated from the hexane-insoluble fraction by Na₂SO₄ column chromatography. Afterwards the hexane-soluble fractions were subsequently saponified at 85°C for 2 h using a solution of 0.1 M potassium hydroxide (KOH) in MeOH, and the neutral fractions were extracted with hexane. The hydrocarbons were separated with hexane from the neutral fraction using column chromatography with deactivated silica gel (1% H₂O, 60 mesh). The hydrocarbon fractions were further cleaned (separation of unsaturated from saturated hydrocarbons) by AgNO₃-Si column chromatography using hexane.

Instrumental analysis

Quantification of long-chain n-alkanes was carried out with a ThermoFischer Scientific Focus gas chromatograph equipped with a split/splitless injector operated at 340°C and a flame ionisation detector (GC-FID). n-Alkanes were quantified using an external standard which contains n-alkanes C₁₈ to C₃₄ in known concentrations. Precision of quantification is 5% based on repeated analyses of the external standard.

Compound-specific δ¹³C analyses of leaf wax n-alkanes were carried out using a ThermoFischer Scientific Trace GC Ultra coupled to a Finnigan MAT 252 irm-MS (GC-irm-MS) via a modified GC/C III interface operated at 1000°C. δ¹³C values were calibrated against CO₂ reference gas and are reported in ‰ notation against Vienna Pee Dee Belemnite (VPDB). Duplicate measurements of each sample yielded a reproducibility of 0.2‰ for n-C₂₉ and 0.1‰ for n-C₃₁ on average. The precision of the internal standard squalane was 0.3‰ (n = 154). Repeated analyses of the external n-alkane standard yielded a long-term accuracy and precision of <0.1‰ and 0.3‰, respectively.

Compound-specific δD analyses were carried out on a ThermoFischer Scientific Trace GC coupled via a pyrolysis reactor operated at 1420°C to a ThermoFischer Scientific MAT 253 isotope ratio mass spectrometer (GC-irms). δD values were calibrated against H₂ reference gas and are reported in ‰ notation against Vienna Standard Ocean Mean Water (VSMOW). Duplicate measurements of each sample result in a reproducibility of 1.5‰ for n-C₂₉ and 0.5‰ for n-C₃₁ on average. The precision of the internal standard squalene was 3.5‰ (n = 150). Repeated analysis of an external alkane standard between every six analyses yielded a long-term accuracy and precision of <1‰ and 2.8‰, respectively.

Results

Individual compound-specific δ¹³C and δD values of n-C₂₉ and n-C₃₁ for the river samples and the marine surface sediments are given in Supplementary Tables 2 and 3, available online, respectively. The compound-specific δ¹³C values of the river samples and the marine surface sediments have been earlier reported and discussed by Herrmann et al. (2016). In the following, the compound-specific δ¹³C and δD values are presented as amount weighted mean of n-C₂₉ and n-C₃₁ and are hereafter referred to as δ¹³C_wax and δD_wax, respectively.

River samples

The δ¹³C_wax values of the Orange River flood deposits and suspension loads range from −31.4‰ to −26.9‰ and from −29.9‰ to −24.5‰, respectively. They show a general isotopic enrichment from the Vaal–Orange confluence towards the Orange River mouth, which is most pronounced in the suspension loads but less in the flood deposits (Figure 4). The δD_wax values range from −141‰ to −121‰ and from −148‰ to −121‰ for suspension loads and flood deposits, respectively (Supplementary Table 2, available online; Figure 4). The suspension loads tend to show isotopic enrichment from the confluence towards the west, except for the suspension load closest to the Orange River mouth. The flood deposits of the Vaal and Orange rivers exhibit highly variable δD_wax values (−121 ± 0‰ and −148 ± 1‰, respectively) before the confluence but are less variable in the further course of the river (Supplementary Table 2, available online; Figure 4).

Figure 4.

Amount-weighted compound-specific carbon and hydrogen isotope compositions of plant waxes in flood deposits (grey diamonds) and suspension loads (black dots) along a transect from the Vaal–Orange confluence (right) to the Orange River mouth (left). Letters (a–c) represent the ages of dated flood deposits given in the figure.

Marine surface sediments

δ¹³C_wax in marine surface sediments ranges from −27.3‰ to −26.2‰. From the northernmost to the central part of the mudbelt, δ¹³C_wax becomes enriched by about 1.1‰. The sediments in the southernmost part show a depletion of δ¹³C_wax by 0.7‰ compared with the central part (Supplementary Table 3, available online; Figure 5). The δD_wax values of the marine surface sediments range between −131‰ and −142‰ and reveal a similar pattern as for δ¹³C_wax showing an increase of 11‰ and 10‰ in the central part of the mudbelt relative to the northern and southern parts, respectively. Between 31.5°S and 32°S a decrease of ca. 9‰ from −132 and −141‰ for δD_wax is observed (Figure 5).

Figure 5.

Amount-weighted compound-specific stable carbon and hydrogen isotope compositions of plant waxes in surface sediments along the west coast of South Africa. Position of the poleward countercurrent (CC) is indicated by the grey dotted line and arrow, and the mudbelt is indicated by the hatched area. Colours indicate modern vegetation of southern Africa (after Mucina and Rutherford, 2006; Scott et al., 2012).

Sediment core GeoB8331-4

We found long-chain n-alkanes in all samples from core GeoB8331-4 with concentrations (sum of n-C₂₅ to n-C₃₃) between 2.0 and 7.2 µg/g (Supplementary Figure 1, available online). We calculated the average chain length (ACL) of the homologues n-C₂₇ to n-C₃₃ as follows:

AC L_{27 - 33} = \frac{27 \times C_{27} + 29 \times C_{29} + 31 \times C_{31} + 33 \times C_{33}}{C_{27} + C_{29} + C_{31} + C_{33}}

(1)

where, C_x is the concentration of the n-alkane with x carbon atoms. The ACL_27–33 is relatively constant ranging from 30.6 to 30.9 (Supplementary Table 4, available online). All n-alkane distributions are dominated by the n-C₃₁. The carbon preference index (CPI) of the homologues between n-C₂₇ and n-C₃₃ is calculated as follows:

CP I_{27 - 33} = 0.5 \times [\begin{array}{l} (\frac{C_{27} + C_{29} + C_{31} + C_{33}}{C_{26} + C_{28} + C_{30} + C_{32}}) + \\ (\frac{C_{27} + C_{29} + C_{31} + C_{33}}{C_{28} + C_{30} + C_{32} + C_{34}}) \end{array}]

(2)

where, C_x is the concentration of the n-alkane with x carbon atoms. The CPI_27–33 for all samples from GeoB8331-4 ranges from 6.6 to 10.4 (Supplementary Table 4, available online) indicating an origin from terrestrial higher plants and a relatively non-degraded state (Eglinton and Hamilton, 1967).

The δ¹³C_wax values of GeoB8331-4 show a range of about 1.6‰ (−27.6‰ to −26.0‰) for the entire Holocene (Figure 6a). Depleted δ¹³C_wax values (−27.6‰ to −27.2‰) occurred from ca. 9800 to 9500 cal. yr BP during the early Holocene with a gradual increase to more enriched δ¹³C_wax towards the middle to late Holocene, peaking with most enriched values, that is, −26.2‰ and −26.0‰, at ca. 3100 and 200 cal. yr BP, respectively. The time period from ca. 3100 to 500 cal. yr BP is characterised by fluctuating δ¹³C_wax compositions (−26.9‰ to −26.4‰), whereas a gradual enrichment of δ¹³C_wax values (−26.9‰ to −26.0‰) is observed from ca. 500 to 200 cal. yr BP. Afterwards, a gradual depletion from −26.0‰ to −27.4‰ occurred during the last 200 years.

Figure 6.

Comparison of compound-specific (a) carbon and (b) hydrogen isotope compositions of plant waxes in marine sediment core GeoB8331-4 with (c) ratio of Poaceae over Asteraceae (Poac/Ast; Zhao et al., 2016) from GeoB8331-4, (d) precipitation stack of the wettest quarter in northern (blue line) as well as central and eastern (green dashed line) in South Africa (Chevalier and Chase, 2015), (e) reconstructed precipitation during the wettest quarter (PWetQ) at Wonderkrater (Truc et al., 2013), (f) austral summer (December) insolation at 30°S and (g) December latitudinal insolation gradient (LIG) between 30°S and equator (Berger and Loutre, 1991). Stars indicate ¹⁴C dates.

δD_wax in core GeoB8331-4 varies between -149‰ and −137‰ (Figure 6b). The most enriched δD_wax composition (−137‰) was detected at ca. 9500 cal. yr BP during the early Holocene. δD_wax values show a gradual depletion from the early to the middle Holocene, with most depleted values (−149‰) at ca. 3900 cal. yr BP. An enrichment of δD_wax is observed during the late Holocene from ca. 3900 cal. yr BP to the present with a steepening of the enrichment for the last 600 years (Figure 7b).

Figure 7.

Comparison of compound-specific (a) carbon and (b) hydrogen isotope compositions of plant waxes in the marine sediment core GeoB8331-4 for the last 2200 years with (c) Antarctic sea ice extension based on diene/triene ratio (Etourneau et al., 2013). Stars indicate ¹⁴C dates.

Discussion

Source of terrestrial organic material

The enrichment of δ¹³C_wax (Figure 5) and increasing ACL_27–33 (Herrmann et al., 2016) in the surface sediments from the northernmost towards the central mudbelt indicate enhanced input from the adjacent dry CAM-rich Succulent Karoo along the west coast. This is supported by δ¹³C_wax values from flood deposits of the Buffels and Holgat rivers in the WRZ, which are isotopically more enriched (amount weighted mean: −25.5 ± 1.6‰) than the suspension loads (amount weighted mean: −28.7 ± 2.0‰) and flood deposits (amount weighted mean: −29.4 ± 1.6‰) from the Orange River (Herrmann et al., 2016).

The WRZ of southern Africa has higher MAP in the southwestern part and lower MAP in its northern part (cf. Figure 1). The hydrogen isotope composition of precipitation (δD_p) is influenced by various environmental processes such as the intensity of precipitation (amount effect) or distance from the moisture source (rainout, continental effect; Dansgaard, 1964). Earlier studies have shown that δD of plant waxes are mainly correlated to local δD_p (e.g. Garcin et al., 2012; Luo et al., 2011; Polissar and Freeman, 2010; Sachse et al., 2004; Schwab et al., 2015), but potentially also influenced by secondary factors such as evapotranspiration (e.g. Kahmen et al., 2013b; Krull et al., 2006; Liu et al., 2006), plant functional type and photosynthetic pathway (e.g. Gamarra et al., 2016; Gao et al., 2014; Hou et al., 2007; Liu and Yang, 2008; McInerney et al., 2011; Smith and Freeman, 2006). In general a deuterium enrichment in plant waxes is observed in areas with high evapotranspiration and lower amount of precipitation (e.g. Feakins and Sessions, 2010; Herrmann et al., 2017; Kahmen et al., 2013a; Krull et al., 2006). The enriched δD_wax in the marine surface sediments observed in the central mudbelt (Figure 5) may reflect a combination of these processes indicating input of terrestrial organic material from the dry Karoo Biome on the adjacent west coast. In the southernmost mudbelt-depleted δD_wax, relative to the central mudbelt, may indicate terrestrial input from the Fynbos Biome experiencing higher MAP than the dry Karoo. This is consistent with earlier published δ¹³C_wax and pollen data from the same sediments (Herrmann et al., 2016; Zhao et al., 2015), investigation of bulk sediments in the southern mudbelt (Birch, 1977) as well as δD_wax in South African soils (Herrmann et al., 2017). The more depleted δD_wax values in the northernmost mudbelt, in contrast, point to a source region with higher amount of precipitation and/or less evapotranspiration. As the Orange River is the major source of terrestrial material to the northern mudbelt, the terrestrial organic material in these marine sediments should mainly reflect environmental conditions in the SRZ. It has been shown that the terrestrial organic material transported by the Orange River comprises a spatially heterogeneous integrated signal from along the river course including the western parts of the SRZ (Herrmann et al., 2016). This is detectable in the earlier published δ¹³C_wax data of the suspension loads and flood deposits (Herrmann et al., 2016) as well as in the δD_wax data in this study. The deuterium enrichment in the suspension loads from the confluence towards the river mouth indicates additional input from regions with higher evapotranspiration and/or less amount of precipitation in agreement with the precipitation and aridity pattern along the central and lower reaches of the Orange River.

The strong difference in the δD_wax composition of about 27‰ of the flood deposits from the Vaal River (ORF35) and the Orange River before the confluence with the Vaal (ORF37) reflects different sources of the plant organic material. We are aware that both samples are of different age (Herrmann et al., 2016) and thus represent different periods. However, considering that modern precipitation has more depleted δD values in the Drakensberg Mountains than in the lower altitudes of South Africa’s SRZ (OIPC; Bowen and Revenaugh, 2003) and that δD_wax in soils of the SRZ tracks well the δD of precipitation (Herrmann et al., 2017), it can be assumed that the Orange River flood deposit reflects a flood event with an origin in the Drakensberg Mountains. Differences in the δD_wax composition of both flood deposits at location ORF29 might be due to different ages or contribution of n-alkanes by different sources as they also show slightly different n-alkane compositions (Herrmann et al., 2016). The combination of enriched δ¹³C_wax and slightly depleted δD_wax of the suspension samples closest to the river mouth (ORF24S) compared with the suspension loads ORF27S further upstream may be due to a local overprint.

In summary, the terrestrial organic material, including plant waxes, in the northern mudbelt and at site GeoB8331-4 is mainly derived from fluvial transport by the Orange River and originates from the central and western parts of the SRZ.

Holocene vegetation and hydrological variability in the SRZ

Early to middle Holocene (9900–3900 BP)

The relatively depleted δ¹³C_wax values during the early Holocene indicate a higher relative abundance of C₃ vegetation or decreased coverage by C₄ grasses during this period (Figure 6a). The pollen data of the same samples reveal low representation of Poaceae pollen interpreted as reduced extension of Grassland/Savanna during the early Holocene (Zhao et al., 2016). In the modern Savanna and Grassland biomes C₄ rather than C₃ plants dominate in the lower altitudes, whereas the highest altitudes in the Grassland Biome are dominated by woody C₃ plants and C₃ rather than C₄ grasses (Cordova, 2013; Cowling et al., 1999; Vogel et al., 1978; Werger and Ellis, 1981). At present, the abundance of C₄ grasses increases with higher aridity and higher growing season temperatures (Scott and Vogel, 2000; Vogel et al., 1978). Based on the investigation of controlling factors of δD_wax compositions in modern soils from the SRZ (Herrmann et al., 2017), relatively enriched δD_wax compositions during the early Holocene suggest less rainfall/more arid conditions (Figure 6b). This is confirmed by low Poac/Ast pollen ratios of the same record (Figure 6c), an indicator for summer rainfall, with low values indicating less summer rainfall and vice versa (Zhao et al., 2016). Less summer rainfall during the early Holocene, indicated by shrubby karroid vegetation and low Poac/Ast ratios, was also found at Blydefontein and Braamhoek wetlands, respectively (Bousman, 1991; Norström et al., 2009). Further evidence for drier conditions during the early Holocene are found in palaeosols at Erfkroon (Lyons et al., 2014), wetland deposits at Florisbad (Scott and Nyakale, 2002), stalagmites at Makapansgat (Holmgren et al., 2003) as well as in lacustrine sediment cores from the Tswaing Crater (Kristen et al., 2010; Metwally et al., 2014; Partridge et al., 1997). In addition, the early Holocene has been inferred to have been relatively cool (Chevalier and Chase, 2015; Metwally et al., 2014), which is advantageous for C₃ rather than C₄ plants/grasses (Cowling et al., 1997; Werger and Ellis, 1981; Yamori et al., 2014). We therefore interpret the higher relative abundance of C₃ plants associated to enriched δD_wax values as reduced C₄ grass cover due to drier and cooler conditions. A higher contribution of plant organic matter from the WRZ during the early Holocene can be ruled out as this would have led to enriched δ¹³C_wax values (Figure 5).

A general trend to wetter conditions (more depleted δD_wax values) is observed from about 9200 to ca. 3900 cal. yr BP and increased contribution by C₄ plants (more enriched δ¹³C_wax values) is observed from about 9200 to ca. 3000 cal. yr BP, respectively. The observed wetter conditions in our record are consistent with higher Poac/Ast ratio of the same record reflecting more summer rainfall (Zhao et al., 2016). An establishment of grassier vegetation during the middle Holocene were observed in the interior of South Africa (Cowling et al., 1997; Scott and Nyakale, 2002; Scott et al., 2005). The marked increase of rainfall/wetter conditions from about 9200 to ca 7000 cal. yr BP in our δD_wax record is consistent with other records from the SRZ of South Africa, suggesting increasing humidity for this period (Chevalier and Chase, 2015; Lyons et al., 2014; Truc et al., 2013). Chevalier and Chase (2015) used regional pollen sequences to obtain precipitation stacks for the northern SRZ and the central/eastern SRZ covering the Holocene. Our δD_wax record is similar, with some deviations, to the northern rather than their central/eastern precipitation stack during the early and middle Holocene. Whereas the used records for the northern precipitation stack are consistent, records for the central/eastern precipitation stack are more variable and cover a wider area (Chevalier and Chase, 2015). For instance, in central South Africa, a marked drying trend at Florisbad (Chevalier and Chase, 2015; Scott and Nyakale, 2002) is observed from the early to the middle Holocene. In contrast, the pollen record at Blydefontein shows variable but slight increasing precipitation (Chevalier and Chase, 2015; Scott et al., 2005). In addition, pollen records from the eastern part show either less precipitation or stable precipitation conditions during this period (Chevalier and Chase, 2015; Finch and Hill, 2008; Neumann et al., 2010; Norström et al., 2009). We infer highest rainfall/wettest conditions in our record around 3900 cal. yr BP. This corroborates evidence of a wet phase in parts of the interior of South Africa, for example, at Florisbad (Scott and Nyakale, 2002) and Wonderwerk Cave (Thackeray and Lee-Thorp, 1992), whereas the record at Erfkroon shows increasing rainfall during this time (Lyons et al., 2014). Furthermore, some records in eastern South Africa, for example, Wonderkrater (Truc et al., 2013), Lake Eteza (Neumann et al., 2010) and Cold Air Cave (Repinski et al., 1999; cf. Figure 2) show wet phases. Also, the central/eastern precipitation stack by Chevalier and Chase (2015) shows increased precipitation during this period (Figure 6d).

The local summer insolation is considered to be an important forcing mechanism for the climatic evolution during the early and middle Holocene in southern Africa (Figure 6e) as proposed by other studies (e.g. Badewien et al., 2015; Chevalier and Chase, 2015; Collins et al., 2014; Hahn et al., 2016; Partridge, 1997; Zhao et al., 2016). First, increased local insolation leads to stronger atmospheric convection due to enhancement of the atmospheric low pressure cell during austral summer (Zhang et al., 2015). Second, higher insolation leads to a stronger thermal land–sea gradient resulting into a stronger easterly flow and thus to more precipitation in South Africa (Zhang et al., 2015). Chevalier and Chase (2015) suggested that the timing of the maximum humidity shifts from early to mid-late Holocene towards the south due to direct insolation forcing. In our record, highest rainfall/wettest condition is detected around 3900 yr BP (Figure 6c), whereas maximum local summer insolation at 30°S peaks later at around 1500 yr BP, which makes the local summer insolation alone an unlikely forcing mechanism. Furthermore, temporal offsets in maximum humidity in palaeorecords relative to local summer insolation maxima have been detected by several studies (e.g. Lim et al., 2016; Lyons et al., 2014; Metwally et al., 2014; Partridge, 1997; Singarayer and Burrough, 2015; Truc et al., 2013). Other studies suggested that, instead of local insolation alone, the LIG plays an important role in changing atmospheric and ocean circulation patterns (Bosmans et al., 2015; Davis and Brewer, 2009; Loutre et al., 2004) albeit not yet for southern Africa. Therefore, we compared our δD_wax record with the LIG between 30°S and the equator (30°S–0°S), as these latitudes comprise approximately the extent of the southern Hadley Cell, and temporal changes in LIG should have influenced the extent and position of the Hadley Cell (Bosmans et al., 2012). The highest LIG during the Holocene occurred at ca. 3500 cal. yr BP, which approximately coincided with the wettest conditions in our record (Figure 6g). This implies that the LIG between low and middle latitudes is a potentially important driver for Holocene precipitation changes in the interior of southern Africa. Models show that stronger heating in the middle latitudes compared with the equator would induce a poleward extension of the Hadley Cell due to a more stable subtropical jet stream at its poleward edge (Lu et al., 2007). Consequently, such a shift of the Hadley Cell would result into a southward shift of the rain-bearing wind systems (Bosmans et al., 2012). However, considering the LIG and palaeorecords, especially, in the eastern/northeastern part of South Africa, discrepancies exist between strength of LIG and humidity/precipitation records. It was suggested that changes in SST and thermal land–sea contrast are important factors driving humidity/precipitation on these regional scales (Chevalier and Chase, 2015; Chevalier et al., 2017; Lim et al., 2016; Truc et al., 2013), which might overprint the influence of the LIG in southern Africa.

Nevertheless, our finding is congruent with other palaeorecords from the Southern and Northern Hemisphere, which underline the importance of the effects of the LIG on the climate system (Jiang et al., 1998; Kuechler et al., 2018; Masson-Delmotte et al., 2005; Vimeux et al., 2001), such as the strength and shift of the monsoon systems (Li and Wu, 2010; Ramisch et al., 2016; Verschuren et al., 2009) and the expansion of sea ice (Denis et al., 2010).

Late Holocene (3900 BP–present)

This time period is characterised by an overall drying trend from ca. 3900 BP to the present suggested by increasingly enriched δD_wax. This overall drying trend is also visible in the central/eastern precipitation stack by Chevalier and Chase (2015; Figure 6) and at the Pella site (Lim et al., 2016) but is less pronounced in the northern part of South Africa (Chevalier and Chase, 2015, 2016 and references therein). The drying trend at Pella is associated to intensification of Benguela upwelling around 3.5–1.5 ka BP (Lim et al., 2016) falling in the period of decreasing LIG. The intensification of the Benguela upwelling, therefore, cannot be excluded as potential influence of the observed drying trend in our record. Within this drying trend, phases of stable and more humid condition occurred from around 2700 to ca. 640 and from about 310 to 240 cal. yr BP, respectively (Figures 6 and 7). Note that the temporal sampling resolution for the last ca. 2000 years is higher enabling the detection of short-term variations, which cannot be studied on the earlier parts of our record. The variable and more humid condition from about 2700 to ca. 640 cal. yr BP coincides with humid conditions in north-eastern South Africa (Chevalier and Chase, 2016). The last phase of this period (ca. 1000 to 640 cal. yr BP) corresponds to the Medieval Climate Anomaly (MCA) lasting from ca. 1100 to 700 cal. yr BP, with conditions generally wetter and partially warmer in the SRZ (Chevalier and Chase, 2015 and references therein; Holmgren et al., 2003; Lodder, 2011; Lyons et al., 2014). For instance, at Erfkroon, seasonal rainfall was relatively high (ca. 600–700 mm/yr) around 830 cal. yr BP compared with present day (ca. 400–500 mm/yr; Lyons et al., 2014). Further evidence of moister conditions are indicated by increased fluvial activity in the Orange River which was probably caused by intense precipitation events (Hahn et al., 2016; Lyons et al., 2014; Zawada et al., 1996; Zhao et al., 2017). The dated flood deposits from the Vaal (ORF35) and the lower Orange River (ORF25) fall in this period (Supplementary Table 1, available online). In a broader, that is, Southern Hemisphere, perspective, additional evidence for warming in the Southern Hemisphere during the MCA are found in New Zealand (Cook et al., 2002; Schaefer et al., 2009) and South America (Bertrand et al., 2014; PAGES 2k Consortium, 2013; Strelin et al., 2008). A reduced sea ice extent has been detected around Antarctica for this period, which led to a southward shift of the SHW and relatively dry conditions in the WRZ (Etourneau et al., 2013; Stager et al., 2012; Zhao et al., 2016). This anti-phase/dipole behaviour of the climate in the SRZ versus the climate in the WRZ is thus consistent with a large-scale southward shift of climate zones during the MCA.

For the following period of the ‘Little Ice Age’ (LIA: 640–310 cal. yr BP), we observe drier conditions consistent with other records from the SRZ suggesting colder and drier conditions, such as from tree rings in the Karkloof Forest (Natal midlands) from 650 to 450 cal. yr BP (Vogel et al., 2001) or stalagmite records from around 500 to 200 cal. yr BP at Makapansgat Valley (Holmgren et al., 2003) and in north-eastern Namibia (Voarintsoa et al., 2016). The dry conditions (enriched δD_wax) during the LIA were accompanied by a gradual increase in C₄ plant contributions (enriched δ¹³C_wax) from about 510 to 330 cal. yr BP in our record (Figure 7). This correlation seems to contradict the earlier finding of increasing C₄ grass cover with wetter conditions during the early and middle Holocene in our study area. Furthermore, an abrupt increase of n-alkane concentration and accumulation rates are observed for this period (Supplementary Figure 1, available online) as a consequence of higher sedimentation rates (Hahn et al., 2016; Herbert and Compton, 2007), which may have been caused by flood events in the Orange River (Zawada et al., 1996) but also in the ephemeral Holgat and Buffels rivers (Benito et al., 2011). For the late Holocene and especially the LIA, it has, however, been suggested that large expansion of Antarctic sea ice led to a northward shift of the rain-bearing westerlies (Etourneau et al., 2013; Lamy et al., 2001; Reason and Rouault, 2005) and, therefore, expansion of the WRZ in southwestern Africa (Stager et al., 2012; Weldeab et al., 2013; Zhao et al., 2016) and flood events in the ephemeral rivers (Benito et al., 2011). We therefore infer that the northward shift of the westerlies during the LIA was accompanied with a northward expansion of the WRZ, which led to elevated contribution of C₄ plant material to the core site during the LIA and explains the observed behaviour of δ¹³C_wax and δD_wax during this period.

A short wet phase occurred in the Orange River catchment from around 310 to 240 cal. yr BP. This finding is consistent with other studies, which showed that a wet phase occurred at the end of the LIA (e.g. Heine and Völkel, 2011; Nash et al., 2016; Voarintsoa et al., 2016). Besides the long-term variation in insolation and LIG, which seem responsible for the overall drying trend during the late Holocene in our record, short-scale variations are clearly influencing the climate of southern Africa during the last millennium. Such short-scale variations could be attributed to variations of the El Nino Southern Oscillation (ENSO), changes in the Benguela upwelling strength, Antarctic sea ice expansion, solar activity or volcanic eruptions (e.g. Hodell et al., 2001; Lean and Rind, 1999; Lim et al., 2016; Reason and Rouault, 2002; Scott et al., 2012; Wanner et al., 2008). For instance, Reason and Rouault (2002) showed that during ENSO events, high pressure anomalies occur over southern Africa and the Indian Ocean. As a consequence, decreasing advection of moist air over eastern South Africa results in drier conditions in the SRZ (Reason and Rouault, 2002; Zhang et al., 2015). A direct link, however, cannot be seen in our record, which might be attributed to an overlay of other short-scale variations, small-scale regional differences or age model uncertainties.

Evidence of human impact is detected for the last ca. 150 years of the record. The marked gradual depletion of about 1.5‰ in the δ¹³C record can be attributed to the Suess effect changing atmospheric δ¹³C towards more depleted values due to anthropogenic fossil fuel emissions (Keeling, 1979; Suess, 1955). The reflection of this signal in the GeoB8331-4 record is a clear evidence for a rapid transfer of climatic signals in the Orange River watershed and quick deposition in the mudbelt.

Conclusion

In this study, we reconstructed palaeoenvironmental changes in the SRZ of southern Africa covering the Holocene based on a compound-specific isotope record from marine sediment core GeoB8331-4 retrieved in the vicinity of the Orange River mouth at a decennial to centennial resolution. The core receives terrestrial sedimentary contributions from the middle and lower Orange River catchment. Our compound-specific hydrogen isotope record indicates relatively dry conditions during the early Holocene supporting earlier studies from the SRZ inferring colder and drier climate during this time. Towards the middle Holocene, we detect a gradual increase of wetter conditions and a shift to more C₄ grasses. Precipitation changes were not solely driven by increasing local insolation but potentially also due to a stronger LIG displacing the position of maximum atmospheric convection southward. After 3900 cal. yr BP, gradual aridification occurred in the Orange River catchment, with periods of stable to more humid conditions (ca. 2700–640 cal. yr BP, ca. 310–240 cal. yr BP). Within this gradual aridification trend, the MCA (1000–640 cal. yr BP) is evident as a more humid period in the Orange River catchment. During the LIA (640–310 cal. yr BP), the Orange River catchment experienced a drying accompanied by an increase of C₄ plant contribution. We suggest that the long-term influence of the LIG and local insolation forcing is superimposed by short-term impacts of other controls, such as Antarctic sea ice extent, Benguela upwelling changes, solar activity, volcanic eruptions, ENSO and human influences.

The observed long-term influence of the LIG in our record is a new finding for this region. This hypothesis calls for further investigation and should be considered and tested in future studies as well as comprehensive compilations of regional records in southern Africa.

Supplementary Material

Supplementary Material, Supplementary_Figure_1 – Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls

Supplementary Material, Supplementary_Figure_1 for Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls by Nicole Burdanowitz, Lydie Dupont, Matthias Zabel and Enno Schefuß in The Holocene

Supplementary Material

Supplementary Material, Supplementary_Table1_rev – Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls

Supplementary Material, Supplementary_Table1_rev for Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls by Nicole Burdanowitz, Lydie Dupont, Matthias Zabel and Enno Schefuß in The Holocene

Supplementary Material

Supplementary Material, Supplementary_Table2_rev – Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls

Supplementary Material, Supplementary_Table2_rev for Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls by Nicole Burdanowitz, Lydie Dupont, Matthias Zabel and Enno Schefuß in The Holocene

Supplementary Material

Supplementary Material, Supplementary_Table3_rev – Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls

Supplementary Material, Supplementary_Table3_rev for Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls by Nicole Burdanowitz, Lydie Dupont, Matthias Zabel and Enno Schefuß in The Holocene

Supplementary Material

Supplementary Material, Supplementary_Table4_rev – Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls

Supplementary Material, Supplementary_Table4_rev for Holocene hydrologic and vegetation developments in the Orange River catchment (South Africa) and their controls by Nicole Burdanowitz, Lydie Dupont, Matthias Zabel and Enno Schefuß in The Holocene

Footnotes

Acknowledgements

We thank the captain, crew and scientists of the Meteor cruise M57/1 for recovering the studied material. We thank Ralph Kreutz, Oliver Helten and Alexander Weinhart for lab assistance. We also thank the editor for their time and helpful remarks as well as Frank Neumann and an anonymous reviewer for their constructive and helpful comments to improve the manuscript.

Funding

Funding was provided by the Bundesministerium für Bildung und Forschung (BMBF) within in the project ‘Regional Archives for Integrated Investigation’ (RAiN, 03G0840A), which is embedded in the international research program SPACES (Science Partnership for the Assessment of Complex Earth System Processes). N.B. was supported by GLOMAR – Bremen International Graduate School for Marine Sciences.

ORCID iD

Nicole Burdanowitz

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