SPECIATION IN ANCIENT LAKES 6 Review Paper

Dynamics of a Kalahari long-lived mega-lake system:hydromorphological and limnological changesin the Makgadikgadi Basin (Botswana)during the terminal 50 ka

Frank Riedel • Andrew C. G. Henderson • Karl-U. Heußner •

Georg Kaufmann • Annette Kossler • Christian Leipe •

Elisha Shemang • Linda Taft

Received: 7 December 2012 / Accepted: 26 July 2013

� Springer Science+Business Media Dordrecht 2013

Abstract The Kalahari features a long-lived lacus-

trine system which may exist since the Early Pleisto-

cene. The emergence of an extant cichlid fish radiation

from this (palaeo-) lake during the Middle Pleistocene

indicates an ancient lake character. The early history

of the system remains speculative, but it is established

that lake extensions matching modern Lake Victoria in

size have occurred during the Late Pleistocene. It has

been assumed that the hydrographical dynamics

chiefly depended on the inflow from the Okavango

River and thus on ITCZ-controlled precipitation. Our

studies, which focused the hydromorphological and

palaeolimnological development of the Makgadikgadi

Basin during the last 50 ka, suggest that from c.

46–16 ka it did not receive water from the Okavango

River but from palaeo-rivers located in the northern

and south-western catchment. A northward shift of the

winter rainfall zone during the Last Glacial Maximum

sustained a high lake level for a period of c. 6 ka.

During Heinrich Event 1 (17–16 ka) the lake probably

desiccated abruptly and completely. Higher lake

levels, controlled by water from the Okavango river

system, were reached again during the Holocene

before the lake dried up in the middle of the last

millennium.

Keywords Southern hemisphere � Kalahari

Desert � Long-lived lacustrine system � Basin

geomorphology � Palaeolimnology � Palaeolake

modelling

Guest editors: T. von Rintelen, R.M. Marwoto, G.D. Haffner &

F. Herder / Speciation in Ancient Lakes – Classic Concepts and

New Approaches

F. Riedel (&) � G. Kaufmann � A. Kossler �C. Leipe � L. Taft

Institute of Geological Sciences, Freie Universitat Berlin,

Malteserstr. 74-100, 12249 Berlin, Germany

e-mail: paleobio@zedat.fu-berlin.de

F. Riedel

Key Laboratory of Plateau Lake Ecology and Global

Change, Yunnan Normal University, No. 1 Yuhua

District, Chenggong, Kunming, China

A. C. G. Henderson

School of Geography, Politics and Sociology, Newcastle

University, Newcastle Upon Tyne NE1 7RU, UK

K.-U. Heußner

Scientific Department of the Head Office, Deutsches

Archaologisches Institut, Im Dol 2-6, 14195 Berlin,

Germany

E. Shemang

Department of Earth and Environmental Sciences,

Botswana International University of Science and

Technology, Private Bag 041, Gaborone, Botswana

123

Hydrobiologia

DOI 10.1007/s10750-013-1647-x

Introduction

When David Livingstone crossed the Kalahari in the

middle of the nineteenth century he recognised the

wide distribution of fossil shells of molluscs that he

interpreted to have lived in the same vast freshwater

lake. He related the vanishing of the lake to climate

change and to tectonic events which had changed the

course of the Zambezi and created an outflow across

the Victoria Falls, subsequently draining the lake

(Livingstone, 1857). Since then many studies

have identified the existence of a large Quaternary

lacustrine system comprised of several sub-basins

(Passarge, 1904; Grove, 1969; Baillieul, 1975; Ebert

& Hitchcock, 1978; Cooke, 1979; Mallick et al., 1981;

Heine, 1982, 1987; Butzer, 1984; Cooke & Verstap-

pen, 1984; Lancaster, 1989; Thomas & Shaw, 1991,

2002; Partridge & Scott, 2000; Ringrose et al., 2005,

2009; White & Eckardt, 2006; Burrough et al., 2009a,

b). Termed Lake Palaeo-Makgadikgadi by Grey &

Cooke (1977; see also Cooke & Verstappen, 1984), it

spans an area similar to extant Lake Victoria (White &

Eckardt, 2006).

The age of the lake system is unknown. An OSL

date of 288 ± 25 ka from palaeo-shoreline sediments

(Burrough et al., 2009a) indicates a minimum age of c.

300 ka. This age is at the limit of the OSL dating

technique (Cordier et al., 2012). Based upon Early

Stone Age tools McFarlane & Eckardt (2006) sug-

gested a minimum age of the lake floor of 500 ka. This

is in line with the molecular dating of a Pleistocene

cichlid fish radiation which emerged in Lake Palaeo-

Makgadikgadi (Joyce et al., 2005), the oldest clade

dated to c. 600 ka (Genner et al., 2007). Recently

Moore et al. (2012) speculated that the lake system

could have been initiated as early as 1.4 Ma, during

the Early Pleistocene. Geological data indicate that the

Neogene drainage from the west towards the Limpopo

was interrupted during Late Pliocene to Early Pleis-

tocene by the uplift of epeirogenetic axes resulting in

the creation of a huge internal drainage (Du Toit,

1933; Cooke, 1980; Haddon & McCarthy, 2005), in

which Lake Palaeo-Makgadikgadi could develop.

While it can only be speculated about the Early to

Middle Pleistocene history of this long-lived lake

system, geomorphological data on Late Pleistocene

and Holocene development are comprehensive. The

hydrological dynamics and forcings and feedbacks,

however, are still contentious (Burrough et al., 2009a, b;

Riedel et al., 2009, 2012). Burrough et al. (2009b)

reviewed geomorphological features related to lacus-

trine development and, using a large set of luminescence

dates (Burrough et al., 2009a), inferred seven so-called

mega-lake phases during the last c. 100 ka, four of these

occurred since 40 ka. Burrough et al. (2009b) argued the

Okavango River is the most important source of water

for Lake Palaeo-Makgadikgadi and suggested cooler

temperatures that reduced evapotranspiration drove

mega-lake expansion. Consequently more water flowed

from the Okavango Delta into the Kalahari depressions

leading to lake highstands. For the period from 19 to

15 ka, however, Burrough et al. (2009b) proposed a

southward shift of the ITCZ resulting in higher precip-

itation in the catchment area of the Okavango River and

thus driving a mega-lake highstand.

These interpretations are based upon the presump-

tions different lacustrine sub-basins were intercon-

nected and thus approximately equally high palaeo-

shorelines developed more or less synchronously.

Moreover, it also assumes the water budget of the

lacustrine system was dependent on the Okavango

River and therefore on ITCZ-controlled precipitation.

In order to test these hypotheses, we re-examined

geomorphological, geological and palaeontological

features of the lacustrine system with a focus on the

development of the Makgadikgadi Basin during the

last c. 50 ka to establish what evidence there is for so-

called mega-lake phases and occurrence of an inter-

connected lacustrine system since MIS 3, but also

to examine the possible causes of lake level

fluctuations.

Regional setting

Morphology of lacustrine basins and drainage

system

Lake Palaeo-Makgadikgadi is centred on the Makga-

dikgadi Basin, the largest and deepest of five major

lacustrine sub-basins within the Makgadikgadi–Ok-

avango–Zambezi basin (MOZB; Ringrose et al., 2005,

2008). The MOZB is a structural depression of the

south-western branch of the East African Rift System

(Reeves, 1972; Scholz et al., 1976; Thomas & Shaw,

1991; Carney et al., 1994; Modisi, 2000; Ringrose

et al., 2005; Kinabo et al., 2007) spanning c.

120,000 km2 (Thomas & Shaw, 1991), and which is

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123

mainly controlled by a series of northeast-southwest

trending faults (Baillieul, 1979; Cooke, 1980; Nugent,

1990, 1992; Shaw & Thomas, 1992; Haddon &

McCarthy, 2005; Ringrose et al., 2005; Kinabo

et al., 2008; Shemang & Molwalefhe, 2011). Two

phases of enhanced tectonic activity during the last

50 ka have been suggested, around 40 ka (Kinabo

et al., 2007; Ringrose et al., 2008) and from 36 to 28 ka

(Thomas et al., 2000, 2003; Wanke, 2005). McCarthy

et al. (1993) discussed the influence of neo-tectonics

on the water dispersal of the Okavango wetlands and

suggested that creation of interconnected graben

systems diverted water flow.

The lacustrine basins occupy c. 66,000 km2 of the

MOZB (Eckardt et al., 2008; Burrough et al., 2009a, b;

Figs. 1, 2). The Makgadikgadi Basin occupies the

eastern area of the MOZB spanning about 37,000 km2

(Ebert & Hitchcock, 1978; Cooke & Verstappen, 1984;

Ringrose et al., 1999, 2005; Burrough et al., 2009b;

Fig. 2). The Makalamabedi Basin (c. 1,200 km2, this

study, Fig. 2) is located between the Makgadikgadi

Basin and the Okavango Delta, a large alluvial fan and

fluvio-lacustrine wetland (c. 22,000 km2, Cooke,

1980; Shaw & Thomas, 1992; Andersson et al.,

2003). The Ngami Basin (c. 2,600 km2, Burrough

et al., 2009b) lies south of the Okavango Delta and the

Fig. 1 SRTM-3 Digital Elevation Model of southern Africa showing the Kalahari drainage system; with geographical references

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123

Mababe Basin (c. 2,300 km2, this study, Fig. 2)

stretches northeast from the Okavango Delta. The

shallow Caprivi Depression (c. 2,000 km2, Shaw &

Thomas, 1988; Fig. 2) represents the northernmost of

the MOZB basins.

The basins exhibit different generations of palaeo-

shorelines. The Makgadikgadi Basin lies below an

altitude of 950 m a.s.l. (Grove, 1969; Grey & Cooke,

1977; Cooke, 1980; Mallick et al., 1981) and is most

clearly demarcated to the west by the Gidikwe Ridge

which was created by aeolian sands during palaeolake

lowstands and was shaped by water activity during

palaeolake highstands. Shorelines at Gidikwe Ridge

were reported to be most prominent at c. 936 and

945 m a.s.l. (Cooke & Verstappen, 1984; Thomas &

Shaw, 1991; Burrough et al., 2009a, b). Sua Pan, in the

east, has an average elevation of 900 m a.s.l. (mini-

mum elevation of 890 m a.s.l.) and demarcates the

deepest area of the Makgadikgadi Basin (Cooke, 1980;

Thomas & Shaw, 1991; Eckardt et al., 2008; Burrough

et al., 2009b; Fig. 2). Evidence for lake levels at

920 m a.s.l. and around 912 m a.s.l. has been reported

from several locations of the Makgadikgadi Basin

(Cooke, 1980; Cooke & Verstappen, 1984; Shaw &

Cooke, 1986; Thomas & Shaw, 1991).

The Makgadikgadi Basin is interconnected with the

Makalamabedi Basin through the Boteti River valley

which forms a gorge at Gidikwe Ridge and the

attached Moremaoto Ridge which represents the

eastern demarcation of the Makalamabedi Basin

(Fig. 2). The Moremaoto Ridge rises to 955 m a.s.l.

and was considered by Gumbricht et al. (2001) to have

Fig. 2 SRTM-3 Digital Elevation Model exhibiting the struc-

tural depressions of Northern and Middle Kalahari; with

geographical references. Triangles and corresponding numbers

refer to sample locations as indicated in Table 2. A–D settle-

ments, A Khumaga, B Sukwane, C Rakops, D Nata

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123

a different origin than Gidikwe Ridge. Curvilinear

features at c. 940 m a.s.l. (Gumbricht et al., 2001;

Fig. 2) on both banks of the Boteti River were

interpreted as old shorelines indicating a lake which

occupied the Makalamabedi Basin in the past (Shaw

et al., 1988). Only the Boteti River valley proper is

lower than 930 m a.s.l. (Fig. 2).

The Makalamabedi Basin is interconnected with the

Ngami Basin through the Boteti and Nhabe river

valleys (Fig. 2). The Ngami Basin is demarcated in the

east and in the north by several ridges to which palaeo-

shorelines from 930 to 940 m a.s.l. are related (Thomas

& Shaw, 1991; Burrough et al., 2007, 2009b) such as

Magotlawanen Ridge (936 m a.s.l.), the Dautsa Ridge

complex (930–936 m a.s.l.) which bifurcates at its

northern terminus, probably due to tectonic activity

(Huntsman-Mapila et al., 2006), and Kerang Ridge

(938–940 m a.s.l.). A 945 m a.s.l. palaeo-shoreline

was identified at the south-eastern margin of the basin

(Thomas & Shaw, 1991; Shaw et al., 2003; Burrough

et al., 2007). The lowest elevation in the Ngami Basin

is 919 m a.s.l. (Shaw, 1985a).

The Mababe Basin is heart-shaped (Fig. 2) and

most clearly demarcated in the west and in the north by

the prominent Magikwe Ridge, a palaeo-shoreline

system (Grove, 1969; Mallick et al., 1981; Shaw,

1985a; Thomas & Shaw, 1991; Burrough & Thomas,

2008). The Magikwe Ridge is lower in the south than

in the north, rising from 930 to 945 m a.s.l., which is

explained by tectonic activity (Gumbricht et al.,

2001). Burrough & Thomas (2008) reported eleva-

tions from 930 to 954 m a.s.l. referring to Gumbricht

et al. (2001). The 954 m a.s.l. of Gumbricht et al.

(2001, p. 260) apparently represents a typing error.

The authors wrote ‘‘… rises in elevation from south to

north by 15 m, from a low of 930 m in the south to

954 m at the northern end …’’. Later in the paper

(Gumbricht et al., 2001, p. 261) it is written ‘‘… lies

between the 930 and 945 m contours’’, which repre-

sents exactly the 15 m difference mentioned before.

The Magikwe Ridge bifurcates where it is cut by the

Savuti Channel (Shaw, 1985a). Burrough & Thomas

(2008) related a 945 m a.s.l. lake stand to the northern

termination of Magikwe Ridge. Mallick et al. (1981)

identified relics of palaeo-shorelines at a similar

elevation in a dune field west of Magikwe Ridge.

Palaeo-shorelines at an elevation of 936 m a.s.l. were

described from the northern Magikwe Ridge, across

delta sediments in the northeast of the basin and along

the south-eastern boundary of the basin (Shaw,

1985a). The minimum elevation of the Mababe Basin

is c. 920 m a.s.l. (Shaw, 1985a; Gumbricht et al., 2001;

Burrough & Thomas, 2008).

The Mababe Basin is interconnected with the

Caprivi Depression through the Savuti Channel

(Fig. 2). The central east of the Caprivi Depression

is characterised by the ephemeral shallow Lake

Liambezi (United Nations, 2000; personal observa-

tion). Between the lake and the Chobe escarpment

southeast of it, stretches a floodplain at an elevation of

928–930 m a.s.l. which exhibits sandy ridges with

heights of 932–936 m a.s.l. (Shaw & Thomas, 1988).

These ridges and alluvial terraces at similar elevations

along the Chobe escarpment have been interpreted to

represent relics of the shore of a palaeolake Caprivi

(Shaw & Thomas, 1988).

Based upon palaeo-shoreline evidence White &

Eckardt (2006) modelled four different lake system

stages. The 912 and 920 m a.s.l. lake levels were

effective in the Makgadikgadi Basin only. The lake

did not approach Gidikwe Ridge except for its

southern part. The 936 m a.s.l. lake level was effective

in the Makgadikgadi, Ngami and Mababe basins and

in the Caprivi Depression and in a small area of the

Makalamabedi Depression. Gidikwe Ridge (Makga-

dikgadi Basin), Magotlawanen Ridge (Ngami Basin)

and Magikwe Ridge (Mababe Basin) were active

beach zones. The basins and depressions supposedly

were interconnected by rivers and channels (White &

Eckardt, 2006). The 945 m a.s.l. lake level scenario

corresponds with the mega-lake phases suggested by

Burrough et al. (2009b). Water bodies in Ngami Basin,

Mababe Basin and Caprivi Depression formed a single

southwest-northeast stretching sheet of water (Lake

Thamalakane stage of Shaw, 1988) capturing the

Makalamabedi Basin and interconnected with the

Makgadikgadi Basin via a several kilometres broad

waterway (White & Eckardt, 2006).

The modern drainage system of the MOZB is

characterised by three major river systems, Cubango-

Okavango, Kwando-Linyanti-Chobe and Zambezi, the

Lundaschwelle representing the northern watershed

(Fig. 1). The upper Zambezi River has its catchment on

the Angolan and Zambian highlands and drains the

Caprivi Depression. It has been proposed that the upper

Zambezi formerly terminated in the MOZB, based on

the hypothesis that it was captured by the middle

Zambezi during Early or Middle Pleistocene (Du Toit,

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1933; Lister, 1979; Cooke, 1980; Thomas & Shaw,

1988, 1992; Nugent, 1990, 1992; Moore & Cotterill,

2010). However, delta sediments which could be

related to such an upper Palaeo-Zambezi have not yet

been identified in the MOZB. The Kwando-Linyanti-

Chobe river system actually represents part of the

upper Zambezi system. It originates on the Angolan

highlands and reaches the Zambezi at the north-eastern

edge of the Caprivi Depression (Shaw & Thomas,

1988; Fig. 1). The transition from Kwando to Linyanti

is fault controlled, the south-easterly flowing Kwando

bending sharply to become the north-easterly flowing

Linyanti which is named Chobe River closer to the

confluence with the Zambezi. The Linyanti receives

periodical overspill from the Okavango Delta through

the Selinda Spillway (Fig. 2). The Selinda Spillway

has formerly fed the Savuti Channel (Thomas & Shaw,

1991; Burrough & Thomas, 2008), which enters the

north-western Mababe Basin. There are no reports

about a lake in the Mababe Basin during historical

times (Gumbricht et al., 2001; Burrough & Thomas,

2008) although the largest of the north-eastern dry river

valleys, the Ngwezumba (Fig. 2), has been known to

flood sporadically (Shaw, 1985a).

The Cubango-Okavango system has its catchment

of c. 165,000 km2 (Wilk et al., 2006) on the Angolan

highlands (Fig. 1). The Okavango has formed a large

delta in the north-western MOZB, representing a

wetland with river branches and channels. The water

level of the delta wetland controls the overspill to river

beds particularly of the Kunyere, the Thamalakane or

the Selinda (Fig. 2). The Kunyere flows into the Nhabe

River, since several years filling up the Ngami Basin

which had been desiccated for decades. Early reports

about Lake Ngami came from Andersson (1857) and

Livingstone (1857) who both saw a wide sheet of

water. Andersson (1857) learned about rapid lake level

decrease from local people. Wilson & Dincer (1976)

emphasised that Lake Ngami of latest history desic-

cated within 2 years and filled up within 1 year. The

historical lake record was summarised by Shaw

(1985b) and Shaw et al. (2003). West and northwest

of Lake Ngami significant fossil river systems exist

such as the Eiseb and the Grootlaagte (Shaw et al.,

1992; Fig. 1). The Thamalakane River is fault con-

trolled and flows to the southwest in front of the central

Okavango Delta. It branches into the Boteti River

(Zouga River of Livingstone, 1857) and the Nhabe

River (Fig. 3). The interaction between these rivers has

not been described, however, is essential for under-

standing the dynamics of Lake Palaeo-Makgadikgadi

(see ‘‘Results’’ and ‘‘Discussion’’ sections). The Boteti

River runs east through the Makalamabedi Basin,

subsequently cutting deeply the Moremaoto-Gidikwe

ridges, entering the Makgadikgadi Basin and turning

sharply to the south some 30 km east of Gidikwe

Ridge. Approximately 70 km further south the Boteti

becomes directed easterly again, crossing the northern

part of the nowadays dry basin of Lake Xau, which had

been a freshwater lake still in the 1960s (Grove, 1969).

The Boteti river bed terminates in the southern

Ntwetwe Pan, which is the central pan of the Makga-

dikgadi Basin (Fig. 2). The neighbouring Sua Pan is

controlled by the Nata River, which has its catchment

northeast of the Makgadikgadi Basin and fills Sua Pan

periodically to shallow depths. The last hydrological

interaction between Ntwetwe Pan and Sua Pan

occurred probably during the first half of the last

millennium (Riedel et al., 2012). A significant fossil

river system, the Okwa Valley, formerly fed the

Makgadikgadi Basin from the southwest (Breyer,

1982; Shaw et al., 1992; Nash et al., 1994; Key &

Ayres, 2000; Nash & McLaren, 2003; Figs. 1, 2).

Chronological framework

Geological ages of lake system related features have

been determined with radiocarbon and luminescence

dates (Street & Grove, 1976; Heine, 1978, 1982, 1987,

1988; Cooke, 1980; Helgren & Brooks, 1983; Cooke

& Verstappen, 1984; Helgren, 1984; Shaw, 1985a;

Shaw & Cooke, 1986; Shaw, 1988; Thomas & Shaw,

1991, 2002; Shaw et al., 1992, 2003; Ringrose et al.,

2005; Huntsman-Mapila et al., 2006; Burrough et al.,

2007; Burrough & Thomas, 2008). Radiocarbon dates

have been obtained particularly from calcretes and

from molluscan shells. Calcretes are not formed

during lake highstands but rather under playa lake

conditions. On the other hand, it has not been tested

whether recent shells of regional molluscs contain old

carbon in order to calculate a potential hard water

reservoir effect for the fossil shells. Thomas et al.

(2003) proposed that a dating error resulting from a

possible hard water reservoir effect does not exceed

1 ka. It has been demonstrated from other lake

systems, however, that reservoir effects can produce

much larger dating errors (Fontes et al., 1996; Wu

et al., 2010; Wunnemann et al., 2010). Furthermore, it

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has not been established that the dated molluscs lived

in lacustrine environments, e.g., a fossil shell of the

gastropod Bellamya was used to infer lake conditions

for the Caprivi Depression (Shaw & Thomas, 1988)

but modern individuals of this genus are inhabitants of

regional river systems (Riedel et al., 2009). Riedel

et al. (2009) concluded that probably all fossil

molluscs listed in Thomas & Shaw (1991) are most

likely of riverine origin and thus cannot be used to

infer lacustrine environments (see next section).

Lake system-related luminescence data have

mainly been obtained from palaeo-beach sediments

(Burrough et al., 2009a), the intermediate bleaching of

which (resulting in a younger than actual age),

however, cannot be excluded. It is well possible that

during the suggested repeated mega-lake highstands

(Burrough et al., 2009b) older beach accumulations

were reworked by wave action during a later highstand.

Moreover many elevations of dated sediments were

measured with an error range of ±6–8 m a.s.l. (Bur-

rough et al., 2009a) making it difficult to clearly

differentiate between 936 and 945 m a.s.l. lake levels.

Burrough et al. (2009b) used the term mega-lake solely

in respect of a 945 m a.s.l. highstand, which is in

contrast to Grey & Cooke (1977) who introduced the

term in relation to both the 936 and 945 m a.s.l. palaeo-

shorelines. According to Burrough et al. (2009b)

mega-lakes which spanned about 66,000 km2 occurred

during the last 50 ka at 38.7 ± 1.8, 26.8 ± 1.2,

17.1 ± 1.6 and 8.5 ± 0.2 ka. The durations of the

mega-lake phases have not been inferred.

Palaeontological record

Although MOZB fluvio-lacustrine sediments have

been studied to some extent (e.g., Passarge, 1904;

Fig. 3 Air photograph taken in May 2010 and documenting the confluence of Thamalakane, Boteti and Nhabe rivers, the two latter of

which exhibiting the phenomenon of flowing in opposite directions

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Cooke & Verstappen, 1984; Heine, 1987; Shaw, 1988;

Thomas & Shaw, 1991; Shaw et al., 1997; Gumbricht

et al., 2001; Thomas et al., 2003; Ringrose et al., 2005,

2009; Huntsman-Mapila et al., 2006; Riedel et al.,

2009), palaeontological records are still scarce.

Livingstone (1857) was the first to report a wide

distribution of molluscan shells which he considered

to represent fossil remains of a palaeolake, although

despite this Passarge (1904) concluded that most shells

Livingstone spotted were not fossil but of (sub-) recent

origin (see ‘‘Results’’ and ‘‘Discussion’’ sections). The

molluscs collected by Passarge were studied by

Martens (in Passarge, 1904). Beside terrestrial gastro-

pods, five aquatic gastropod and two bivalve species

were described from outcrops of the Boteti and Nhabe

river valleys (modern taxonomic names in parenthe-

ses, following Brown, 1994; Appleton, 2002): Pla-

norbis salinarum (=Biomphalaria salinarum), Physa

parietalis (=Bulinus parietalis), Melania tuberculata

(=Melanoides tuberculata), Ampullaria occidentalis

(=Pila occidentalis), a single shell of Vivipara

passargei (=Bellamya cf. capillata), Unio kunenensis

(=Coelatura kunenensis) and Corbicula africana

(=Corbicula fluminalis). All species are known from

the extant southern African fauna.

From a Boteti River section in the south-western

Makgadikgadi Basin Riedel et al. (2009) described a

very similar fossil molluscan community, which

additionally comprised a species of Potadoma, a

gastropod genus that no longer occurs in southern

Africa. Riedel et al. (2009) considered the molluscan

palaeo-community of riverine origin and not of

lacustrine origin. A Bellamya shell was dated to c.

46 ka cal. BP (Table 2). Riedel et al. (2009) reviewed

records of fossil aquatic molluscs from the MOZB and

concluded that probably all palaeo-communities had

lived in river systems. One exception is palaeolake

sediments at the Tsodilo Hills (northern Botswana)

from which Thomas et al. (2003) identified several

fossil species of gastropods and the bivalve Corbicula

cf. fluminalis. These sediments cover an age from c.

41.5 ka cal. BP to c. 17 ka cal. BP (Thomas et al.,

2003). The altitude of this palaeolake depression is c.

1,005 m a.s.l. and thus too high to be related to a

mega-lake Palaeo-Makgadikgadi.

Ostracods have not been treated in Passarge’s

monograph (1904). Possibly Grey & Cooke (1977,

p. 128) were the first to report fossil ostracods from the

Makgadikgadi Basin. They used their findings to

conclude lacustrine conditions in front of Gidikwe

Ridge supporting the interpretation that the ridge

represents a strandline feature. Shaw (1985a) men-

tioned estuarine calcretes from the Mababe Basin

which contained ostracods and were dated to c. 20.5

and c. 16 ka cal. BP. The ostracods, however, were

not investigated. A preliminary taxonomic record of

fossil ostracods from the Makgadikgadi Basin was

given by Riedel et al. (2012) who listed and figured

shells from lacustrine sediments deposited on ‘‘Kubu

Island’’. This granitic outcrop is located well ‘‘off-

shore’’ at the south-western edge of Sua Pan (Fig. 2)

and was examined to some extent because of its

interesting position. The maximum elevation of

‘‘Kubu Island’’ is 926 m a.s.l. and thus too low that

mega-lake high stands could have created palaeo-

shorelines. In sediment layers from elevations

between 913 and 918 m a.s.l., Riedel et al. (2012)

found valves of the following species: Limnocythere

thomasi-group, Sarscypridopsis glabrata, Limnocy-

there aff. inopinata, Potamocypris aff. variegata,

Ilyocypris sp. and Strandesia sp. This palaeo-commu-

nity reflects a saline lake phase. The total amount of

ostracod valves was too low for obtaining sufficient

calcium-carbonate for AMS-radiocarbon dating. The

‘‘Kubu Island’’‘‘ sediments contained a few fossil fish

remains, which however have not been studied (Riedel

et al., 2012).

Thomas & Shaw (1991) compiled literature records

of fossil diatoms using the term diatomaceous earths

but did not present taxonomic details and ecological

interpretations of the palaeo-communities. On the

other hand already Reichelt (in Passarge, 1904)

identified more than 40 different fossil species of

diatoms, most of which indicating brackish water

lacustrine conditions. The samples had been collected

from outcrops along the Nhabe, Thamalakane and

Boteti river valleys (Passarge, 1904). From the site

descriptions it can be concluded that these samples are

of late Quaternary age. Most of the fossil diatoms

analysed by Reichelt (in Passarge, 1904) came from

the locality Meno a kwena (c. 10 km southeast of

Moremaoto) where the Boteti River cuts Gidikwe

Ridge and thus can be attributed to the Makgadikgadi

Basin. From an outcrop in the Boteti Valley at

Moremaoto Shaw et al. (1997) identified diatoms at

the genus level. The sediments were dated to 32-27 ka

and related to a lake highstand in the Makgadikgadi

Basin. This site, however, is located at the backside of

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123

Gidikwe Ridge and thus lies within the eastern

Makalamabedi Basin (Gumbricht et al., 2001). Dia-

toms from sediments from the eastern part of the

Ngami Basin have been identified at the species level

by Robbins et al. (1998) and Shaw et al. (2003). Based

upon ecological interpretation of diatoms and dating,

shallow freshwater environments were concluded to

have prevailed during the Holocene. In addition,

Robbins et al. (1998) reported fossil fish remains. A

longer record from the central Ngami Basin covering

c. 42 ka was presented by Huntsman-Mapila et al.

(2006). Twelve samples from the 4.6-m long sediment

record were analysed to infer lake conditions by using

diatoms and geochemical parameters (Huntsman-

Mapila et al., 2006): From 42 to 40 ka brackish water

conditions prevailed in a relatively deep lake. At c.

19 ka the environment changed from shallow and

turbulent to deeper water conditions prevalent until c.

17 ka. Shallower alkaline conditions were predomi-

nant between 16 and 5 ka. Lake level was high around

4 ka, decreased until 2.4 ka, increased again until

0.8 ka and subsequently declined (Huntsman-Mapila

et al., 2006). Thomas et al. (2003) listed diatoms from

the palaeolake Tsodilo which, however, was not in

relation with Lake Palaeo-Makgadikgadi.

Climate dynamics since MIS 3

Contemporary climate

Using the Koppen–Geiger classification, the catch-

ment area of Cubango-Okavango, Kwando-Linyanti-

Chobe and Zambezi rivers on the Angolan and

Zambian highlands (Mazvimavi & Wolski, 2006;

Fig. 1) is characterised by temperate-dry-winter-

warm-summer climate (Cwb) in the west and

temperate-dry-winter-hot-summer climate (Cwa) in

the east, with a larger patch of tropical savannah

centred on the Barotse Plains (Kottek et al., 2006; Peel

et al., 2007). The average annual rainfall on the

headwaters is 1,300–1,400 mm (Mazvimavi & Wol-

ski, 2006; McSweeney et al., 2008a, b; Rodrıguez-

Fonseca & Xavier, 2009), most of it falling during the

summer, the rains starting in October/November and

terminating in April/May (Livingstone, 1857; Hughes,

2006; Wilk et al., 2006; Kampata et al., 2008;

McSweeney et al., 2008a, b; Rodrıguez-Fonseca &

Xavier, 2009). The moisture is mainly transported by

the East African Monsoon and the southward shift of

the ITCZ (Verschuren et al., 2009) and thus originates

from the Indian Ocean. The influence of the West

African Monsoon bringing moisture from the Atlantic

Ocean appears to be less pronounced which, however,

is under discussion (McHugh & Rogers, 2001; Barker

et al., 2007; Burrough et al., 2009b). The arid-steppe-

hot climate (BSh) further south is characteristic of the

Kalahari Desert and the MOZB. Precipitation values

over the MOZB range from 650 mm per year in the

northeast, to 500–300 mm annually in most other

parts, rain falling mainly from October to April

(Cooke, 1980; Connelly & Gibson, 1985; Thomas &

Shaw, 1991; Burney et al., 1994; Burrough & Thomas,

2008; Nash & Endfield, 2008). Southwest of the

MOZB, the Kalahari represents a true desert (arid-

desert-hot climate) with precipitation of less than

150 mm per year, falling from October to March

(Rautenbach & Smith, 2001). The winter rainfall zone

is limited to the Western Cape province of South

Africa and therefore has no effect on the hydrological

system of the MOZB (Burrough et al., 2009b). The

modern circulation system over Africa is figured, e.g.,

in Nicholson (1996), Gasse (2000), Gasse et al. (2008)

and Chase et al. (2012).

Past climates

Lake level fluctuations in the MOZB naturally reflect

climate variability but also may reflect tectonic events.

Moreover it has been summarised in the preceding

chapter that lake dynamics in the MOZB depends on

different sources of moisture and moisture pathways

may have changed in the past. We therefore give a

brief summary of palaeoclimate variability of southern

Africa s.l. to provide a framework for discussing

possible changes in the climate system [an outline

review of late Quaternary southern African climates

was provided by Burrough et al. (2009b)]. For clearer

geographic reference we use present time names of

countries.

MIS 3 (59–24 ka) Data from ODP Site 1078 were

used to infer that from c. 32 ka (beginning of the

record) desert and semi-desert extended rather far to

the north of Angola (Dupont et al., 2008). Sediment

records from off Namibia indicated that during the

period from 42 to 23 ka strong wind intensity

persisted over the South Atlantic which led to strong

upwelling (Shi et al., 2001) and humid conditions for

Hydrobiologia

123

the period 32–18 ka were suggested (Stuut & Lamy,

2004). Heine (1988) suggested humid conditions in

the central Namib Desert and the Namibian highlands

from c. 35–33 ka. Vogel (1982) inferred a humid

phase from 39 to 28 ka, rainfalls becoming less

accentuated from 28 to 23 ka. Brook et al. (2006)

found evidence for vigorous fluvial activity c. 25 ka in

south-western Namibia. In respect of the south-

western Kalahari, Heine (1982) postulated similar

conditions from 31 ka across the MIS 3 termination.

In southern South Africa humidity increased until

about 50 ka and rapidly decreased with subsequent

significant but less strong humidity changes (Partridge

et al., 2004; Chase & Meadows, 2007). Stalagmite

growth in Wolkberg Cave continued from 58 to 46 ka

(Holzkamper et al., 2009). Further north in South

Africa, a stalagmite from Wonderwerk Cave archived

signals which were interpreted to suggest wetter

conditions at c. 33 ka (Brook et al., 2010). Kristen

et al. (2007) reported humid intervals from 54 to 50 ka

and 37 to 35 ka inferred from sediments of the

Tswaing impact crater lake (north-eastern South

Africa; see also Partridge et al., 1997). Carr et al.

(2006) proposed enhanced humidity in the winter

rainfall zone of South Africa from [47 to c. 33 ka.

Further north in the Kalahari, major dune-building

occurred from about 46–41 ka and 36–29 ka (Stokes

et al., 1997, 1998; Partridge et al., 2004). Wetter

regional conditions than present occurred in the north-

western Kalahari (Tsodilo Hills) from 40 to 32 ka

(Thomas et al., 2003). For the area northeast of the

MOZB Thomas et al. (2009) inferred lake highstands,

from about 38.4 to 35.5 ka (44–33 ka) based on dated

sand ridges at Lake Chilwa (Malawi). In western

Zambia dune building occurred from 32 to 27 ka

(O’Connor & Thomas, 1999). Woltering et al. (2011)

provided a TEX86 temperature record for MIS 3 in a

range from 22.9 to 25.1�C using Lake Malawi

sediments.

MIS 2 (24–12 ka) In Angola desert and semi-desert

conditions terminated around 22 ka and rain forest

started to expand in the northern lowlands. Between

18.8 and 15.4 ka conditions were cool but not arid.

After that period climate became wetter and warmer

(Dupont et al. 2008). Further south in Namibia

relatively humid climate changed to drier conditions

around 18 ka (Stuut & Lamy, 2004). Vogel (1982)

stated decreasing intensity of the rainfall over Namibia

from 23 to 19 ka while Shi et al. (2000) inferred arid

and cold conditions for south-western Africa from 21

to 17.5 ka. On the other hand, Brook et al. (2011)

suggested flooding of Etosha Pan at c. 19–16.7 ka and

Stone et al. (2010) dated moisture indicating mud units

from western Namibia to 16.9 ka. A phase of

aridification was postulated for the period

14.3–12.6 ka (Shi et al., 2000). Robbins et al. (1996)

suggested that MIS 2 was significantly wetter than MIS

1. In respect of southern South Africa Talma & Vogel,

(1992) calculated minimum temperatures 5–7�C lower

than today during the period 18.5–15.5 ka. Brook et al.

(2010) provided evidence for relatively wet conditions

in north-western South Africa from 23 to 17 ka. This is

in agreement with Chase & Meadows (2007) who

argued for a humid Last Glacial Maximum across

southern Africa suggesting a rainfall gradient along a

transect from the Western Cape province to the interior

of the Kalahari. Chase & Meadows (2007) related the

humid conditions with a shift of the winter rainfall

zone, which was proposed earlier by Van Zinderen

Bakker (1976) and Heine (1981). Lee-Thorp &

Beaumont (1995) and Hurkamp et al. (2011)

concluded that rainfall occurred in winter but also in

summer. From north-eastern South Africa Holmgren

et al. (2003) reported stalagmite growth from 24.4 to

12.7 ka. Postglacial warming was initiated around

17 ka. From the same region Kristen et al. (2007)

inferred from lake sediments a humid interval from 15

to 10 ka and thus during the transition from MIS 2 to

MIS 1. This is in contrast to the conditions in eastern

and in north-western South Africa where drier

conditions prevailed from c. 16 to 13.7 ka (Norstrom

et al., 2009) and from 17 to 13 ka (Brook et al., 2010),

respectively. In respect of the Kalahari Stokes et al.

(1997, 1998) reported dune building phases from 26 to

20 ka and 16 to 9 ka and Chase & Brewer (2009)

concluded that all of the major dune fields were active

at 21 ka. In this respect it is noteworthy that stalagmite

growth in Lobatse Cave (south-eastern Botswana)

terminated at 21.6 ka after a centennial growth of

3 mm since 26.7 ka (Holmgren et al., 1994). Heine

(1981, 1982) proposed a relatively humid climate until

19 ka for the south-western Kalahari. Thomas et al.

(2003) reported wetter than today conditions from the

north-western Kalahari during the period 27–12 ka

with a possible dry spell from 22 to 19 ka. Speleothem

growth at Drotzky’s Cave occurred between c. 19.4 and

15.6 ka (Shaw & Cooke, 1986). In western Zambia

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123

dune building occurred from 16 to 13 ka (O’Connor &

Thomas, 1999). Thomas et al. (2009) inferred Lake

Chilwa (Malawi) highstands for the periods 24.3–

22.3 ka (26.2–21 ka) and for 16.2–15.1 ka and

13.5–12.7 ka (17.9–12 ka). From Lake Malawi

sediments covering the last 25 ka Barker et al. (2007)

described wet-dry intervals spaced approximately every

2,3 ka with a pronounced wet phase around 13,5 ka.

Using the TEX86 index Barker et al. (2007) estimated

surface water temperatures compared to modern values

of -3.5�C during LGM, -1�C during YD and ?3�C at

c. 13.8 ka. Stager et al. (2011) reported a catastrophic

mega-drought during Heinrich Event 1 (17–16 ka)

having affected large areas of western (including

Angola), central, eastern and south-eastern Africa,

however, indicated wetter conditions of uncertain

origin, timing and geographic context for the MOZB.

MIS 1 (Holocene) Angola experienced a dry phase

until c. 8 ka, with subsequent increase of moisture until

c. 4 ka with another shift to more arid conditions

during the late Holocene (Dupont et al., 2008). Proxy-

data from the Okavango Panhandle which presumably

indicate climate conditions on the Angolan highlands

were interpreted to show relatively dry climate from 7

to 4 ka with a punctuated wet phase around 6 ka (Nash

et al., 2006). From 4 to 1 ka conditions became

progressively wetter before the present day conditions

were approached (Nash et al., 2006). Shi et al. (2000)

reported aridification of Namibia from 11 to 8.9 ka.

Based upon Etosha Pan studies Brook et al. (2007)

suggested drier than today conditions until 8 ka and

four subsequent periods of increased wetness: 7–5,

4.5–3.5, 2.5–1.7 and at c. 1 ka. Chase et al. (2010)

identified a series of rapid aridification events in north-

western Namibia since c. 3.8 ka. A pollen record from

Drotsky’s Cave (north-western Kalahari) was

interpreted to indicate mainly drier than today

climate during the early Holocene, with a shift to

wetter conditions than today between 7 and 6 ka which

prevailed until c. 3 ka (Burney et al., 1994). In north-

eastern South Africa the Early Holocene experienced

warm, evaporative conditions, which were initiated

during late MIS 2, around 13.5 ka (Holmgren et al.,

2003). Evident cooling occurred from c. 6 to 2.5 ka,

followed by warmer climate until 1.5 ka. Maximum

Holocene cooling occurred at 1700 A.D. (Holmgren

et al., 1999, 2003). Conditions were mainly dry, with

the exceptions of the millennia from 10.5 to 9.5 ka and

1.5 to 0.5 ka (Norstrom et al., 2009). Lake Chilwa

(Malawi) highstands were dated to *11 and 8.5 ka

(Thomas et al., 2009). Barker et al. (2007) inferred

surface water temperatures of Lake Malawi of -1�C at

8.2 ka and ?5�C at 5 ka, compared to modern values

(25–29�C). In western Zambia dune building occurred

from 10 to 8 ka and from 5 to 4 ka (O’Connor &

Thomas, 1999).

Materials and methods

Geomorphological and geological features of the

MOZB were studied during five periods of fieldwork,

August to October 2007, March and April 2008, May

and July 2010, June 2011 and May 2013. Except for

the Mababe Basin, we could collect data from all other

major depressions including part of the drainage

systems. The 2007, 2008 and 2013 field studies were

conducted with the aid of Differential GPS (D-GPS,

Ashtech). Several outcrops, exhibiting lacustrine or

fluvio-lacustrine palaeo-environments, were sampled.

Diatom extraction and slide preparation for micro-

scopic analysis were performed following Battarbee

et al. (2001). The detailed analyses of fossil remains of

molluscs, ostracods and diatoms with the aid of a

scanning electron microscope and geochemical anal-

yses are still ongoing and therefore some of the data

presented here are preliminary. Taxonomy is compiled

in Table 1. AMS-radiocarbon-dating has been con-

ducted at Poznan Radiocarbon Laboratory (Poland).

Elevation models

The modelling of the topography and of possible palaeo-

lake levels is based on SRTM 90 m Digital Elevation

Data Version 4 (Reuter et al., 2007; Jarvis et al., 2008).

To cover the study area, we used nine tiles (40-15,

40-16, 40-17, 41-15, 41-16, 41-17, 42-15, 42-16, and

42-17) available in GeoTiff format at http://srtm.csi.

cgiar.org/ (retrieved in March 2012). The data were

processed using ArcGIS Desktop 10 (ESRI, 2010).

Basin floor deformation modelling

During the wetter phases, the water of the palaeolakes

in the MOZB has depressed the surface because of the

gravitational weight of the water load. The limited size

and depth of the palaeolakes allows us to model the

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123

surface deformation with an elastic model in carthe-

sian geometry (Steffen & Kaufmann, 2006). We

employ the modelling software ABAQUS (Hibbitt

et al., 2005), and we describe the MOZB as an elastic

block with 2,000 km side length and 12 vertical layers,

simulating the crust and the upper mantle. Looking

onto the surface, the central grid located over the

MOZB with 800 km side length has 81 9 81 nodes,

and it is surrounded by a peripheral frame, thus the

entire grid has 101 9 101 nodes. The resolution in the

central grid is 10 km, while the peripheral grid has a

non-linear resolution, mimicking the horizontal

boundary conditions. The vertical layer thickness is

10 km, except for the lowermost layer, which is

400 km thick. The material properties for the crust and

the upper mantle, namely density, bulk and shear

modulus, are taken from the elastic model PREM

(Dziewonski & Anderson, 1981). The depth of the

Moho discontinuity at the location is assumed to be in

35 km depth, here the boundary between crust and

mantle is located. A linear, elastic rheology is used for

both the crust and upper mantle. The impounded water

of the palaeolakes is applied as surface pressure, thus

the water depth is multiplied by water density

q = 1,000 kg/m3 and gravitational acceleration

g = 9.81 m/s2. When the water load is applied and

the surface is depressed accordingly, we account for

the solid earth around our restricted block model by

applying several boundary conditions: (i) the nodes at

the model bottom cannot move in the vertical direction

Table 1 Classification of mollusc, ostracod and diatom taxa mentioned in this study

Gastropoda Viviparidae Gray, 1847

Ampullariidae Gray, 1824

Bellamya Jousseaume, 1886

Lanistes Montfort, 1810

Pila Roding, 1798

B. capillata (Frauenfeld, 1865)

–

P. occidentalis (Mousson, 1887)

Pachychilidae Troschel, 1857

Thiaridae Troschel, 1857

Potadoma Swainson, 1840

Melanoides Olivier, 1804

–

M. tuberculata (Muller, 1774)

M. victoriae (Dohrn, 1865)

Lymnaeidae Lamarck, 1812

Planorbidae Rafinesque, 1815

Radix Montfort, 1799

Biomphalaria Preston, 1910

Radix natalensis (Krauss, 1848)

B. pfeifferi (Krauss, 1848)

B. salinarum (Morelet, 1868)

Bulinus Muller, 1781 B. parietalis (Mousson, 1887)

Bivalvia Corbiculidae Gray, 1847

Unionidae Fleming, 1828

Corbicula Megerle von Muhlfeld, 1811

Coelatura Conrad, 1853

C. fluminalis (Muller, 1774)

C. kunenensis (Mousson, 1887)

Ostracoda Lymnocytheridae Klie, 1938 Limnocythere Brady, 1867 L. inopinata (Baird, 1843)

L. thomasi Martens, 1990

Ilyocyprididae Kaufmann, 1900

Cyprididae Baird, 1845

Ilyocypris Brady & Norman, 1889

Potamocypris Brady, 1870

Sarscypridopsis McKenzie, 1977

Strandesia Stuhlmann, 1888

–

P. variegata (Brady & Norman, 1889)

S. glabrata (Sars, 1924)

–

Diatomacea Achnnanthes Bory, 1822

Aulacoseira Thwaites, 1848

Campylodiscus Ehrenberg, 1840

Cocconeis Ehrenberg, 1838

Cyclotella (Kutzing) Brebisson, 1838

Epithemia Kutzing, 1844

Gomphonema Ehrenberg, 1832

Navicula Bory, 1822

Nitzschia Hassall, 1845

Rhopalodia O. Muller, 1895

Surirella Turpin, 1828

–

–

–

–

–

–

–

–

–

–

–

Hydrobiologia

123

(no slip); (ii) the nodes at the vertical model bound-

aries must not move in horizontal direction perpen-

dicular to the model sides (no slip). Here we implicitly

assume that the deformation caused by the water load

decays with depth and over large horizontal distances.

To ensure that the boundary conditions as well as the

model size have no effect on the modelling results, we

have carried out tests with different resolutions, and

found the used grid as appropriate for the model.

Results

Dating

It has been pointed out that other authors (i.e., Burrough

et al., 2009a) produced a large set of OSL dates, on the

one hand to overcome problems of radiocarbon dating

(suitable material and hard water effect) and on the

other hand because OSL techniques allow much older

sediments to be dated. In order to calculate the potential

hard water error of fossil shells, ten modern gastropods

(Bellamya, Lanistes) from four rivers, Thamalakane,

Chobe, Zambezi and Nata, were radiocarbon dated. All

modern shells show modern ages (see Table 2 and

‘‘Discussion’’ section). Three fossil shells (Melanoides,

Corbicula) from the Okwa river valley (south-western

Makgadikgadi Basin) in a range of c. 17–16.2 ka cal.

BP and three fossil shells (Melanoides, Corbicula) from

the Nata river valley (north-eastern Makgadikgadi

Basin) in a range of c. 46.4–43.7 ka cal. BP were

dated. In addition to our dates we compiled in Table 2

another seventeen radiocarbon dates on regional mol-

luscs from the literature.

Lacustrine sediments from Sua Pan (eastern Mak-

gadikgadi Basin) were cored by us to a depth of 3 m

and organic rich samples radiocarbon dated (Table 2).

The lowest section (300–280 cm) was dated to c.

37 ka cal. BP, 3 samples from 200 to 120 cm to c.

4.6–2.1 ka cal. BP and 3 samples from 120 to 20 cm

to ages around 2 ka cal. BP.

Riedel et al. (2012) inferred historical lake levels in

the Makgadikgadi Basin using the zonation of living

Baobab trees on ‘‘Kubu Island’’. One size class was

calculated to represent ages of c. 1 ka. In 2010, one of

these trees died and samples for radiocarbon dating

could be taken in 2011. The four dates range from c.

840–900 years cal. BP (Table 2) and the tree thus

became 900–960 years old.

Fossil record and geomorphology

Previous studies suggested that lake levels within

MOZB sub-basins were coupled, similar levels having

occurred synchronously, due to the assumption that

the major water bodies were interconnected (see

‘‘Regional setting’’ section), and that the Okavango

river system was controlling the lake development at

least during the last 300 ka (e.g., Burrough et al.,

2009b). It has not been considered that lacustrine

phases in the Makgadikgadi Basin could have been

independent from lake development in the depressions

along the Okavango-Zambezi Rift Zone and thus from

the Okavango river system.

We surveyed all MOZB sub-basins except for the

Mababe Basin for fossil remains of riverine or

lacustrine environments, related palaeo-shorelines

and water dynamics, mainly with the aid of Differen-

tial GPS and satellite images. Particular focus lay on

the Makgadikgadi Basin. We present data from south-

western Sua Pan (eastern Makgadikgadi Basin), from

the Nata river valley (north-eastern Makgadikgadi

Basin), from the Okwa Valley (south-western Mak-

gadikgadi Basin), from the Boteti river valley (western

Makgadikgadi Basin, Makalamabedi Basin) and from

the Okavango-Zambezi Rift Zone (see Fig. 2).

South-western Sua Pan

C. 1.5 km east of ‘‘Kubu Island’’ we cored Sua Pan

sediments in order to obtain information about Holo-

cene lake development. We cored to 3.0 m depth. The

basal 20 cm of sediments were composed of medium

sand, mixed with fine sand and silt. The 2.8 m

sediments on top were composed of silt and fine sand

and were more organic rich. We studied the section in

20 cm sub-sections. In all 15 sub-sections ostracod

valves were abundant (c. 6–10 different species).

Remains of molluscs and diatoms were not contained.

The basal sandy sediments were dated to c. 37 ka cal.

BP. The finer grained sediments above the sandy sub-

section are of Holocene age, c. 8–2 ka cal. BP

(Table 2). The age of c. 8 ka has not been obtained

by dating but speculatively inferred by simple inter-

polation of thickness of the organic rich sediments on

top of the sandy layer which was dated to c. 37 ka cal.

BP and below the sediments which were dated to c.

4 ka.

Hydrobiologia

123

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0.3

3p

MC

Mo

der

n

Bel

lam

yasp

.sh

ell,

Zam

bez

iR

iver

(5)

S1

7.4

83

88

E2

4.2

55

03

,9

37

Po

z-3

67

21

10

2±

0.3

2p

MC

Mo

der

n

La

nis

tes

sp.

shel

l,N

ata

Riv

er(6

)S

20

.23

04

9E

26

.17

05

4,

91

3P

oz-

36

72

51

06

.63

±0

.33

pM

CM

od

ern

La

nis

tes

sp.

shel

l,N

ata

Riv

ersi

de

arm

(7)

S1

9.9

84

12

E2

6.4

40

18

,9

35

Po

z-3

67

24

10

6.3

3±

0.3

3p

MC

Mo

der

n

Ad

an

son

iad

igit

ata

1,

‘‘K

ub

uIs

lan

d’’

(8)

S2

0.8

93

00

E2

5.8

22

75

,9

11

Po

z-4

35

33

10

00

±3

0B

P9

01

±4

6ca

l.B

P

Ad

an

son

iad

igit

ata

2,

‘‘K

ub

uIs

lan

d’’

(8)

S2

0.8

93

00

E2

5.8

22

75

,9

11

Po

z-4

35

34

97

0±

30

BP

87

7±

47

cal.

BP

Ad

an

son

iad

igit

ata

3,

‘‘K

ub

uIs

lan

d’’

(8)

S2

0.8

93

00

E2

5.8

22

75

,9

11

Po

z-4

35

35

99

0±

30

BP

89

1±

47

cal.

BP

Ad

an

son

iad

igit

ata

4,

‘‘K

ub

uIs

lan

d’’

(8)

S2

0.8

93

00

E2

5.8

22

75

,9

11

Po

z-4

35

36

90

0±

30

BP

83

6±

55

cal.

BP

Su

aP

anse

dim

ents

,2

0–

40

cmb

elo

w

surf

ace

(9)

S2

0.8

88

65

E2

5.8

43

40

,9

03

Po

z-4

35

37

22

65

±3

0B

P2

26

7±

65

cal.

BP

Su

aP

anse

dim

ents

,6

0–

80

cmb

elo

w

surf

ace

(9)

S2

0.8

88

65

E2

5.8

43

40

,9

02

Po

z-4

35

38

19

20

±3

5B

P1

87

3±

37

cal.

BP

Su

aP

anse

dim

ents

,1

00

–1

20

cmb

elo

w

surf

ace

(9)

S2

0.8

88

65

E2

5.8

43

40

,9

02

Po

z-4

35

39

20

65

±3

5B

P2

04

6±

50

cal.

BP

Su

aP

anse

dim

ents

,1

20

–1

40

cmb

elo

w

surf

ace

(9)

S2

0.8

88

65

E2

5.8

43

40

,9

02

Po

z-4

35

40

21

20

±3

5B

P2

09

7±

49

cal.

BP

Su

aP

anse

dim

ents

,1

60

–1

80

cmb

elo

w

surf

ace

(9)

S2

0.8

88

65

E2

5.8

43

40

,9

01

Po

z-4

35

42

39

75

±3

5B

P4

46

4±

41

cal.

BP

Su

aP

anse

dim

ents

,1

80

–2

00

cmb

elo

w

surf

ace

(Rie

del

etal

.,2

01

2)

(9)

S2

0.8

91

36

E2

5.8

53

78

,9

01

Po

z-3

76

96

40

75

±3

5B

P4

64

3±

12

0ca

l.B

P

Su

aP

anse

dim

ents

,2

80

–3

00

cmb

elo

w

surf

ace

(9)

S2

0.8

88

65

E2

5.8

43

40

,9

00

Po

z-4

35

44

32

61

0±

43

0B

Po

nly

0.8

mg

C

37

12

9±

80

1ca

l.B

P

Bel

lam

ya,

Ch

ob

eR

iver

terr

ace

(Sh

awan

d

Th

om

as,

19

88

)(1

0)

Ch

ob

eG

ame

Lo

dg

e,9

31

Har

82

00

26

20

±1

40

BP

26

87

±1

86

cal.

BP

Mo

llu

scs,

Mo

lop

oR

iver

val

ley

,ea

sto

f

Ko

pp

iesk

raal

Pan

(Hei

ne,

19

81

,1

98

2)

S2

7�0

20

E2

0�3

20 ,*

85

0a

Hv

83

72

12

48

0±

22

0B

P1

47

37

±4

69

cal.

BP

Hydrobiologia

123

Ta

ble

2co

nti

nu

ed

Dat

edm

ater

ial

(lit

erat

ure

refe

ren

ce)

(lo

cati

on

nu

mb

erin

Fig

.2)

Geo

gra

ph

ical

coo

rdin

ates

,

elev

atio

nin

met

rea.

s.l.

Lab

.n

o.

or

sam

ple

ID

Rad

ioca

rbo

ny

ears

Cal

ibra

ted

yea

rs

Co

rbic

ula

sp.

shel

l,O

kw

ari

ver

bed

(11

)S

21

.34

85

9E

24

.39

80

2,

93

7P

oz-

43

59

81

32

90

±6

0B

P1

62

20

±4

16

cal.

BP

Bu

lin

us,

Tso

dil

oH

ills

(Th

om

aset

al.,

20

03

)

(12

)

S1

8.7

73

31

E2

1.7

31

12

,*

10

05

aB

eta,

TH

91

-21

13

74

2±

14

0B

P1

68

35

±2

70

cal.

BP

Mel

an

oid

estu

ber

cula

tash

ell

1,

Ok

wa

Val

ley

(11

)

S2

1.3

48

59

E2

4.3

98

02

,9

37

Po

z-4

35

97

13

79

0±

60

BP

16

97

1±

15

6ca

l.B

P

Mel

an

oid

estu

ber

cula

tash

ell

2,

Ok

wa

Val

ley

(11

)

S2

1.3

48

59

E2

4.3

98

02

,9

37

Po

z-4

35

96

13

85

0±

60

BP

17

07

2±

17

8ca

l.B

P

Mel

an

oid

es,

Gid

ikw

eR

idg

eat

Ok

wa

Val

ley

(Th

om

asan

dS

haw

,1

99

1,

Sh

awet

al.,

19

92

)(1

3)

S2

1�2

30 ,

E2

4�2

70 ,

92

0(?

)G

rN1

47

86

14

07

0±

15

0B

P1

73

29

±2

60

cal.

BP

Mel

an

oid

es,

term

inal

Ok

wa

Val

ley

(Th

om

as

and

Sh

aw,

19

91

;S

haw

etal

.,1

99

2)

(14

)

S2

1�2

40

E2

4�2

60 ,

92

0(*

93

6a)

GrN

14

78

71

44

90

±1

50

BP

17

63

2±

28

7ca

l.B

P

Lym

na

ea(=

Ra

dix

),C

ho

be

Riv

erte

rrac

e

(Sh

awan

dT

ho

mas

,1

98

8)

(15

)

Ser

on

del

laC

amp

,9

33

GrN

13

19

21

53

80

±1

40

BP

18

46

2±

32

3ca

l.B

P

Mo

llu

scs,

east

of

Mab

abe

Bas

in,

Ng

wez

um

ba

Riv

erte

rrac

e

(Th

om

asan

dS

haw

,1

99

1)

(16

)

Ng

wez

um

ba

Riv

erv

alle

y,

93

6G

rN1

47

88

15

57

0±

22

0B

P1

88

36

±3

10

cal.

BP

Lym

na

ea(=

Ra

dix

)/B

iom

ph

ala

ria

/Bu

lin

us,

at

Tso

dil

oH

ills

(Th

om

aset

al.,

20

03

)(1

2)

S1

8.7

73

31

E2

1.7

31

12

,*

10

05

aB

eta,

TH

91

-22

(1)

17

18

3±

14

0B

P2

06

19

±3

38

cal.

BP

Bu

lin

us/

Mel

an

oid

es/?

Co

ela

tura

/Co

rbic

ula

,

Mo

lop

oR

iver

val

ley

(Hu

rkam

pet

al.,

20

11

)

S2

7�0

20 5

500

E2

0�3

00 1

400 ,

83

3B

eta-

25

93

20

18

94

0±

90

BP

22

82

6±

29

6ca

l.B

P

Mo

llu

scs,

Mo

lop

oR

iver

val

ley

,n

ear

Ko

op

an

Su

id(H

ein

e,1

98

1,

19

82

)

S2

7�1

40 0

000

E2

0�2

20 3

000 ,

*8

35

aH

v9

49

51

90

85

±1

12

52

28

27

±1

35

7ca

l.B

P

Mel

an

oid

es/C

orb

icu

la/u

nio

nid

s,n

ort

h-

wes

tern

Mak

gad

ikg

adi

Bas

in(H

ein

e,1

98

7)

(17

)

S2

0�1

60 3

300

E2

4�5

70 0

000 ,

*9

19

bH

v8

36

71

91

70

±6

60

BP

23

00

0±

83

6ca

l.B

P

Bel

lam

ya/M

ela

no

ides

,at

Tso

dil

oH

ills

(Th

om

aset

al.,

20

03

)(1

2)

S1

8.7

73

31

E2

1.7

31

12

,*

10

05

aB

eta,

TH

91

-12

26

99

±3

05

BP

27

34

4±

48

2ca

l.B

P

Mel

an

oid

es/B

uli

nu

s/C

orb

icu

la/u

nio

nid

s,

no

rth

-eas

tern

Mak

gad

ikg

adi

Bas

in(H

ein

e,

19

87

)(1

8)

S2

0�1

30 3

000

E2

6�1

50 5

700 ,

*9

17

bH

v8

37

12

59

10

±1

21

0B

P3

05

58

±1

13

1ca

l.B

P

Hydrobiologia

123

Ta

ble

2co

nti

nu

ed

Dat

edm

ater

ial

(lit

erat

ure

refe

ren

ce)

(lo

cati

on

nu

mb

erin

Fig

.2)

Geo

gra

ph

ical

coo

rdin

ates

,

elev

atio

nin

met

rea.

s.l.

Lab

.n

o.

or

sam

ple

ID

Rad

ioca

rbo

ny

ears

Cal

ibra

ted

yea

rs

Mel

an

oid

es/B

uli

nu

s/C

orb

icu

la/u

nio

nid

s,

no

rth

-eas

tern

Mak

gad

ikg

adi

Bas

in,

(Hei

ne,

19

87

)(1

8)

S2

0�1

30 3

000

E2

6�1

50 5

700 ,

*9

17

bH

v8

37

02

73

50

±5

50

BP

32

08

6±

48

6ca

l.B

P

Bel

lam

ya/M

ela

no

ides

/Bu

lin

us/

Co

rbic

ula

,at

Tso

dil

oH

ills

(Th

om

aset

al.,

20

03

)(1

2)

S1

8.7

73

31

E2

1.7

31

12

,*

10

05

aB

eta,

TH

91

-8(1

)3

21

65

±4

23

BP

36

63

3±

86

1ca

l.B

P

Bel

lam

ya/M

ela

no

ides

,at

Tso

dil

oH

ills

(Th

om

aset

al.,

20

03

)(1

2)

S1

8.7

73

31

E2

1.7

31

12

,*

10

05

aB

eta,

TS

OD

92

-11

36

08

2±

18

8B

P4

14

01

±2

55

cal.

BP

Mel

an

oid

estu

ber

cula

tash

ell

2,

Nat

aR

iver

sect

ion

(19

)

S2

0.2

20

05

E2

6.1

82

28

,9

16

Po

z-4

35

95

39

90

0±

60

0B

P4

37

01

±6

71

cal.

BP

Co

rbic

ula

aff.

flu

min

ali

ssh

ell,

Nat

aR

iver

sect

ion

(19

)

S2

0.2

20

05

E2

6.1

82

28

,9

16

Po

z-4

35

93

40

50

0±

70

0B

P4

41

15

±8

39

cal.

BP

Bel

lam

yaca

pil

lata

shel

l,B

ote

tiR

iver

sect

ion

(Rie

del

etal

.,2

00

9)

(20

)

S2

0.8

24

09

E2

4.3

69

43

,9

16

Po

z-2

74

26

42

00

0±

10

00

BP

45

56

0±

11

53

cal.

BP

Mel

an

oid

estu

ber

cula

tash

ell

1,

Nat

aR

iver

sect

ion

(19

)

S2

0.2

20

05

E2

6.1

82

28

,9

16

Po

z-4

35

94

42

80

0±

80

0B

P4

63

98

±1

37

9ca

l.B

P

Inb

old

:th

isst

ud

y;

oth

erd

ates

fro

mli

tera

ture

aE

lev

atio

no

fsi

tes

add

edo

rco

rrec

ted

bP

erso

nal

com

mu

nic

atio

nw

ith

Hei

ne

(20

12

);(?

)q

ues

tio

nab

le

Hydrobiologia

123

Nata river valley

We surveyed the river valley at several sites, from the

terminal river mouth area at Sua Pan to about 60 km

upstream (Fig. 2). In order to get upstream, we

followed a road which branches off the main road to

Kasane, about 6 km north of Nata village, accompa-

nying the Nata River in some distance to the east and

northeast. This road was under construction and on

both sides numerous quarries had been dug and the

material used as base for the pavement. We entered

several of these quarries and all of them exhibited

exclusively series of massive hardpan calcretes (e.g.,

Lionjanga et al., 1987). Fossil remains of aquatic

organisms could not be detected. Geological profiles

exposed in the river valley from the upper part at c.

950 m a.s.l., downstream to the bridge in Nata village

(river bed at c. 915 m a.s.l.) did also not exhibit fossils.

From about 1 km south of Nata bridge towards Sua

Pan, the river had cut 5–6 m into the sediments, in

which we locally found, at a level of c. 916 m a.s.l.

(river bed at c. 914 m a.s.l.), fluvio-lacustrine deposits

exhibiting fossil shells from the gastropods Melano-

ides tuberculata, Bellamya capillata, Biomphalaria

aff. pfeifferi and Radix sp., and the bivalve Corbicula

aff. fluminalis. The fauna has been dated to c.

46.4–43.7 ka cal. BP (Table 2).

Okwa river valley

The hydromorphologic relation of Gidikwe Ridge,

which bounds the Makgadikgadi Basin to the west, to

the two river valleys crossing it, Okwa and Boteti

(Fig. 2), has not been properly investigated. In our

opinion, understanding this relationship is crucial for

answering the question whether the Makgadikgadi

Basin was interconnected to other MOZB sub-basins

during lake highstands.

The considerable size of the Okwa river valley and

its fossil feeders in the main palaeo-catchment area,

which lies about 250–500 km WSW of Gidikwe Ridge

(Fig. 1), suggests that the Okwa was a major contrib-

utor to the water budget of the Makgadikgadi Basin.

Nowadays, people living near the Okwa Valley are not

aware that once a river flowed there. This is probably

due to the fact that the river is not even periodically

active. On satellite images the fossil river valley is

clearly demarcated, its relation to Gidikwe Ridge,

however, cannot be assessed properly by remote

sensing. Aerial survey confirmed that no cliffs are

exposed but the valley flanks are generally gently

sloping. The Okwa thus does not form a gorge at

Gidikwe Ridge as is the case with the Boteti (see next

section). Ground control along the terminal 7–8 km

exhibited that the densely vegetated river bed mean-

ders through the ridge at a level of c. 936–938 m a.s.l.

The river mouth cuts only the uppermost sediments of

the Makgadikgadi strandlines.

In the surface sediments of the terminal river bed

we found a palaeo-community of molluscs, the

gastropods Melanoides tuberculata, Radix cf. natal-

ensis, Bulinus cf. natalensis and the bivalves Corbic-

ula sp. and Coelatura cf. kunenensis. Melanoides and

Corbicula were the dominant taxa. Two Melanoides

shells were dated to c. 17 ka cal. BP and one

Corbicula shell to c. 16.2 ka cal. BP (Table 2).

Boteti river valley

The course of the Boteti river valley has been outlined

in the ‘‘Regional setting’’ section. We here focus two

areas of the Boteti, the west-east directed transition

from the Makalamabedi Basin through the gorge at

Gidikwe Ridge to the Makgadikgadi Basin, and the

north–south directed part between the villages of

Khumaga and Rakops (Fig. 2).

During our field work in 2007 and 2008 the river

valley was completely dry, which had been so since

two decades (personal communication with local

people). At numerous sites groundwater was pumped

from a few metres below the surface of the river bed in

order to water cattle. Cattle and wind had triggered

irregular undulation of the sandy river bed. In May

2010 the river had returned down to the village of

Sukwane (Fig. 2). In May 2011 the river had reached

(but not filled) the Lake Xau depression which is still

the terminal position. During a survey flight in May

2010, we could observe that the river was not

completely continuous, but blocked at different posi-

tions by higher elevated undulations through the sand

of which the water penetrated only very slowly to

lower elevations. 8 km south of Sukwane, where the

river bed was still dry in May 2010, we sampled an

outcrop of silty to sandy fluvio-lacustrine sediments of

1.6 m thickness in 10 cm steps. All samples contained

valves of ostracods. Thin layers comprising diatom

valves were intercalated. The sediments were not

dated but are considered older than 46 ka because they

Hydrobiologia

123

underlie mollusc bearing sediments at 916 m a.s.l.,

which were studied from a site 8 km further upstream

and which revealed an age of 46 ka cal. BP (Riedel

et al., 2009).

The north–south trending part of the Boteti is

connected to a fossil (partly periodically active), fault

controlled river valley which has its catchment area

north of the Makgadikgadi Basin (Palaeo–Boteti).

Altogether three fault controlled fossil river valleys

enter the Makgadikgadi Basin from the north (Figs. 2,

5A).

Upstream, the Boteti River has cut through Gidikwe

Ridge and formed an approximately 20–22 m deep

gorge (Fig. 2). D-GPS measurements yielded a max-

imum elevation of the ridge (a few hundred metres off

the gorge) of about 947 m a.s.l. whereas its northern

terminus is about 6 m higher. The bottom of the river

was measured when it was dry. It lies at c. 920 m a.s.l.

This elevation coincides with our observations during

May 2010 when we could study the water flow from

the Thamalakane River (overspill of Okavango Delta,

Fig. 2) into the Boteti as well as into the Nhabe River

and Lake Ngami (see next section).

Exposed in the gorge, in front of the ridge, we

recognised three different sediment units (S20.28672,

E24.26822). The base, at c. 925 m a.s.l., exhibited

0.4 m thick, whitish, comparatively light, silty sedi-

ments, usually typical of diatom content. Laboratory

analysis, however, exhibited only very few fragments

of diatom valves which could not be assigned to

certain taxa. The whitish sediments are overlain by a

9.5 m thick greyish to yellowish sequence character-

ised by large amounts of calcretes, silcretes and

rhizoliths. During this period playa lake conditions

prevailed. On top, at c. 935 m a.s.l., lies a 0.3 m thick,

diatom valves bearing, lacustrine sequence. Contained

diatom genera include the benthic Rhopalodia, Suri-

rella, Navicula, Achnanthes, Cocconeis, Nitzschia,

Epithemia and the planktic Cyclotella and Aulacose-

ira. OSL dating suggests that these sediments were

most likely deposited during MIS 5 (Schmidt et al.,

2012). The profile is overlain by greyish sandy soil,

which ranges in thickness between few decimetres and

few metres at different sites. The sand originates from

Gidikwe Ridge.

We studied the main lithological units of the gorge

to the west, across Gidikwe Ridge and along the Boteti

Valley. The diatom beds of the gorge profile are

confined to the western flank of Gidikwe Ridge. There

are no continuous sediment units exposed intercon-

necting the Makgadikgadi and Makalamabedi basins

lithologically.

Okavango–Zambezi Rift Zone

‘‘If the lake (=Lake Ngami) ever becomes lower than

the bed of the Zouga (=Boteti), a little of the water of

the Tamunak’le (=Thamalakane) might flow into it

instead of down the Zouga; we should have the

phenomenon of a river flowing two ways; but this has

never been observed to take place here, and it is

doubtful if it ever can occur in this locality’’ (Living-

stone, 1857). We were fortunate enough to observe

this phenomenon of a river flowing in two directions in

May 2010 (Fig. 3). At the Toteng bridge, which is

close to the confluence of Kunyere and Nhabe (Fig. 2),

we could establish that the Nhabe River was flowing to

the southwest to Lake Ngami at an estimated speed of

3–4 km per hour. The Boteti River, on the other hand,

was flowing to the east and thus to the opposite

direction, with a speed of c. 2 km per hour near the

Thamalakane confluence and of c. 0.5–1 km per hour

at Moremaoto Ridge.

During our studies along the Zambezi Valley, at

Lake Liambezi (Fig. 2) and particularly at the Chobe

escarpment, we could not discover sediments indicat-

ing former stable lake phases but we found playa lake

and flood plain deposits. On the banks of the Chobe

River we sampled accumulations of sub-recent shells

of gastropods such as Melanoides, Lanistes and

Bellamya up to a level of c. 935 m a.s.l., approxi-

mately 5–6 m above the river water level (at the end of

September 2007).

Basin floor deformation modelling

In order to assess the influence of water load on the

basin depth on the one hand and on tectonic stress fields

on the other hand, we modelled the resulting vertical

deformation of two different palaeo-water levels which

is shown in Fig. 4. Although we have argued that lake

levels in the different sub-basins must not have

occurred contemporaneously, we here use synchro-

nous highstands for modelling maximum water load.

For the c. 936 m a.s.l. beach ridges, the lake loads are

rather limited to the different sub-basins of the

Makgadikgadi, Ngami, and Mababe. Only in the

Makgadikgadi Basin the water depth was in some

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123

areas more than 30 m, while in the smaller lake basins

water levels were at most below 20 m. Therefore, the

resulting surface deformation is rather small, with

10 cm in the Mababe Basin, and reaching 20–25 cm in

the Makgadikgadi Basin, and the deformation pattern

is mainly restricted to the extent of the palaeolakes. In

the Makgadikgadi Basin, the surface depression

resembles a flat bowl following the palaeo-shoreline,

thus the deformation is responsible for a slight increase

in the lake basin volume. For the c. 945 m a.s.l. beach

ridges, the Mababe and Ngami basins form an inter-

connected palaeolake, while the Makgadikgadi and the

Makalamabedi basins are still separated by the Gidi-

kwe ridge. Water depths in the interconnected Ngami-

Mababe palaeolake (Lake Thamalakane stage) reached

a maximum of c. 25 m, while the Makgadikgadi Basin

comprised a 45–55 m deep lake. This larger water load

results in a more pronounced surface deformation

pattern, with a maximum subsidence of 35 cm in the

central Makgadikgadi Basin (Fig. 4). Due to the larger

size of the water load, the deformation also affects the

area beyond the palaeolake shores, which is a result of

the elastic crust: The crust bends regionally also

outside of the water load due to the finite elastic

stiffness of the material. The bending of the surface

below Lake Palaeo-Makgadikgadi, 35 cm in the centre

and 10–15 cm along the shorelines, results in tilted

palaeo-shorelines, with shoreline elevation differences

between the rim of the basin and along isles in the

centre of the basin nowadays observable with about

20 cm altitude difference due to the surface relaxation

after the evaporation of the water.

Discussion

Dating

Modern ages of modern shells demonstrate that there

are no significant amounts of old carbon in the fluvio-

lacustrine system which may produce a hard water

effect and thus a dating error. Radiocarbon dates of

fossil shells are therefore considered comparatively

precise.

Fossil record and geomorphology

South-western Sua Pan

The sandy character of basal sediments from the Sua

Pan core which were dated to c. 37 ka cal. BP

suggests a shallow water, near-shore environment

Fig. 4 Lake basin floor modelling. Shown are selected river courses in blue, the palaeolake extensions for lake levels 936 m a.s.l.

(A) and 945 m a.s.l. (B), and the resulting surface deformation colour-coded as shown in the legend

Hydrobiologia

123

and thus low lake level and reduced humidity can be

inferred. This is in line with major dune building in the

Kalahari from 36 to 29 ka (Stokes et al., 1998) while

conditions were more humid in South Africa (Kristen

et al., 2007) and Lake Chilwa (Malawi) experienced

highstands (Thomas et al., 2009).

The finer grained sediments above the sandy sub-

section are of Holocene age, c. 8–2 ka cal. BP

(Table 2). The stratigraphical gap of c. 29 ka can

possibly be explained by deflation of sediments during

the early Holocene. This interpretation corresponds

with dry conditions which prevailed until 8 ka in

Angola (Dupont et al., 2008), Namibia (Brook et al.,

2007), north-western Kalahari (Burney et al., 1994),

north-eastern South Africa (Norstrom et al., 2009) and

western Zambia (O’Connor & Thomas, 1999). The

sediments from 2.8 to 1.4 m (c. 8–2 ka cal. BP) reflect

relatively high lake levels. Burrough et al. (2009b)

suggested a mega-lake phase of 945 m a.s.l. around

8.5 ka. The 2.8–1.4 m sediments were possibly

deposited during such highstand (and the subsequent

lake level decline) but B936 m a.s.l. because there is

no evidence of a transgression of Lake Palaeo-

Makgadikgadi into the Okwa river valley during the

last c. 17 ka. The highstand must have been reached

within a short period, possibly within decades, because

the Makgadikgadi Basin likely was dry during the

early Holocene. The subsequent lake level fall towards

the next desiccation phase, however, appears to have

lasted until 2 ka. The upper 1.4 m of the section

exhibited 4 radiocarbon dates ranging around 2 ka cal.

BP. It is proposed that these sediments have been

reworked by a sequence of shallow lake and playa lake

phases and drying up events.

Riedel et al. (2012) used the zonation of Baobab

trees (Adansonia digitata) on ‘‘Kubu Island’’ to infer

lake levels during the last millennium. The Baobabs,

however, were not dated but age was interpolated

(Riedel et al., 2012). The here presented radiocarbon

dates from one of the specimens (Table 2) are in line

with the estimated ages. Wetter conditions than today

around 1 ka (to 0.5 ka), interpreted from sites in the

north-western Kalahari (related to precipitation on the

Angolan highlands; Nash et al., 2006), Namibia

(Brook et al., 2007) and South Africa (Norstrom

et al., 2009), thus did not trigger lake levels higher than

908 m a.s.l. in the Makgadikgadi Basin (Riedel et al.,

2012).

Nata river valley

The fossil shells assigned to the gastropods Melano-

ides tuberculata, Bellamya capillata, Biomphalaria

aff. pfeifferi and Radix sp., and to the bivalve

Corbicula aff. fluminalis and dated to c. 46–44 ka cal.

BP can be used to characterise the palaeo-environ-

ment. While Corbicula aff. fluminalis and M. tuber-

culata can cope with brackish water, the other three

gastropods are typical freshwater taxa. Biomphalaria

aff. pfeifferi defines this molluscan community as

characteristic of riverine low energy habitats or of

shallow sheltered lacustrine environments (e.g.,

Brown, 1994; personal observations FR). Heine

(1987) reported a very similar taphocoenosis from a

site approximately 9 km further east, at an elevation of

c. 917 m a.s.l. (personal communication with Heine,

2012) dated to c. 32–30.5 ka cal. BP. This site was

described in relation with palaeo-shoreline features

and sediments thus were likely deposited under

estuarine–lacustrine conditions during a lake level of

c. 917–920 m a.s.l.

Okwa river valley

The palaeo-community of molluscs found in the

terminal river bed, the gastropods Melanoides tubercu-

lata, Radix cf. natalensis, Bulinus cf. natalensis and the

bivalves Corbicula sp. and Coelatura cf. kunenensis, is

suitable for outlining the palaeolimnology. Extant

individuals of the three gastropods and of Corbicula

sp. may live in lentic or lotic waters (Morton, 1986;

Brown, 1994; personal observations FR). M. tubercu-

lata and Corbicula sp. can even live in ponds with

salinities of up to c. 25 psu (Morton & Tong, 1985;

Piscart et al., 2005; Bolaji et al., 2011). Corbicula sp. can

survive short-term seasonal desiccation of a pond.

Modern C. kunenensis is a riverine endobenthic species

of sandy habitats (personal observations FR). This

bivalve indicates that the fossil molluscs lived in the

Okwa River and not in a bay of Lake Palaeo-Makga-

dikgadi which may have transgressed into the valley.

Because the identified molluscs have lived along the

river banks and they were found in the middle part of the

river bed, it is indicated that during the period of

deposition the Okwa was a rivulet only.

The offset in dating between Melanoides shells (c.

17 ka cal. BP) and Corbicula (c. 16.2 ka cal. BP) may

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123

be due to the fact that Corbicula became the dominant

taxon only when the other molluscan species could not

cope with the changing environment anymore. This is

that around 16.2 ka BP the river had ceased to flow

and only periodical ponds existed in the Okwa Valley.

Shaw et al. (1992) reported from the same river section

a Melanoides radiocarbon date of c. 17.3 ka cal. BP.

These authors gave an elevation of 920 m a.s.l. while

our data indicate that the lowest part of the Okwa

Valley is actually 15 m higher. This difference in

elevation data highlights a general problem of relating

palaeo-shoreline features across the MOZB.

Considering the extent of the fossil river valley it is

possible that a lake highstand in the Makgadikgadi

Basin and a fully active Okwa were contemporaneous

(see Fig. 5B). Because several palaeoclimate records

from southern Africa (Heine, 1981, 1982; Holmgren

et al., 2003; Brook et al., 2010) hint at wet conditions

Fig. 5 Lake level scenarios with focus on the Makgadikgadi

Basin. A Around 46 ka; main inflow from northern river

systems such as Palaeo-Boteti and Palaeo-Nata; lake level at

912 m a.s.l. B During LGM; main inflow from the Okwa River;

lake level at 936 m a.s.l.; lake level in the Makalamabedi Basin

(?) unclear. C During mid-Holocene; main inflow from the

Okavango River/Boteti River; lake level at c. 920 m. D Max-

imum extension during early last millennium; main inflow from

the Okavango River/Boteti River; lake level at 908 m a.s.l

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123

during the LGM, the Okwa was probably a large river

during that period. Chase & Meadows (2007) related

the increased humidity with a northward shift of the

winter rainfall zone. Hurkamp et al. (2011) suggested

that rainfall became less seasonal. The interpretation

that the Okwa ceased to flow around 17 ka is in line

with Brook et al. (2011) who reported that flooding of

Etosha Pan (Namibia) stopped at c. 16.7 ka. It is

proposed that the Okwa ceased relatively abruptly

because it did not cut significantly through Gidikwe

Ridge when the level of Lake Palaeo-Makgadikgadi

fell.

Boteti river valley

Based upon the ecological traits of the fossil gastro-

pods retrieved at Sukwane, and which were dated to c.

46 ka cal. BP, Riedel et al. (2009) inferred a riverine

environment and concluded that a lake in the Makga-

dikgadi Basin must have been below 914 m a.s.l.

Cooke & Verstappen (1984) discussed a 911 m a.s.l.

lake level around 46 ka (see Fig. 5A). This interpre-

tation is supported by records suggesting contempo-

raneous dry climate over the region (Stokes et al.,

1997, 1998; Thomas et al., 2000), however, more

humid than modern climate. The fluvio-lacustrine

molluscan palaeo-community from the Nata river

valley (see above) has about the same age and is

located at a similar elevation. It is thus quite possible

that Boteti and Nata were active at the same time at c.

46 ka. The activating moisture was transported to the

Nata River catchment area, which was extended to the

north in the past (Palaeo–Nata, Figs. 2, 5A), during

southern hemisphere summer by the East African

monsoon and South Indian Ocean anticyclone sys-

tems, respectively. The water flow of the Boteti may

have been activated by the same moisture source and

possibly not by the Okavango system. The reason for

this assumption is that the north–south trending part of

the Boteti is connected to a fossil river valley which

has its catchment area north of the Makgadikgadi

Basin (Palaeo–Boteti, Figs. 2, 5A), and that we

propose the Boteti may not have been connected to

the Okavango system via Gidikwe Ridge during that

period. During periods of enhanced precipitation the

palaeo-river systems entering the Makgadikgadi Basin

from the north may have maintained a c. 910–912 m

a.s.l. lake level (Fig. 5A). It can be speculated that the

Okwa River, which has probably been controlled by

the winter rainfall zone, was not active around 46 ka.

Archives from South Africa indicate a rapid decrease

of humidity after 50 ka (Partridge et al., 2004; Chase

& Meadows, 2007) and e.g., stalagmite growth at

Wolkberg Cave terminated at 46 ka (Holzkamper

et al., 2009).

Although the sediments exposed in the Boteti

Gorge have not been deposited during the last 50 ka

the outcrop is important to understand the relation with

the neighbouring Makalamabedi Basin. C. 7.5 km

NNW of the gorge site, at the backside of Gidikwe

Ridge, diatomaceous sediments are exposed a few

metres above the banks of the Boteti, at c. 925–927 m

a.s.l. Shaw et al. (1997) dated the sediments to

32–27 ka, studied the diatom flora and reconstructed

lake-marginal water energy conditions in relation with

lake level highstands in the Makgadikgadi Basin. Our

studies show that these sediments are in no relation to

similar sediment units exposed in the gorge. On the

other hand, the diatom beds of the gorge profile are

confined to the western flank of Gidikwe Ridge.

Overall there are no continuous sediment units

exposed which could evidence that lacustrine phases

in the Makgadikgadi Basin were affective in the

Makalamabedi Basin too. This supports the interpre-

tation of Gumbricht et al. (2001) that the Makama-

labedi Basin is bound by a distinguished Moremaoto

Ridge to the east while the Gidikwe Ridge bounds the

Makgadikgadi Basin to the west. Tectonic deforma-

tion of sediment beds exposed at Moremaoto Ridge

indicates a west–east striking fault. Actually Gumb-

richt et al. (2001) suggested tectonic control of the

formation of the Boteti Gorge which assumingly

occurred as late as during the Holocene.

Okavango–Zambezi Rift Zone

Considering the 2010 situation when the Thamalakane

River distributed water in two opposite directions (see

Fig. 3), the historical observations of Passarge (1904,

p. 204) are noteworthy. He wrote about the confluence

(translation from German with use of modern names):

«The Thamalakane is only 80–100 m wide, but deep,

and flowed at that time, during the beginning of the

seasonal flood, at an estimated speed of 2 km per -

hour. The water was clearly flowing to the east, to the

Boteti. West of the Thamalakane, in the Nhabe River,

there was no current recognisable.» Passarge travelled

by boat on the Nhabe River in 1897 and thus his report

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123

is unambiguous. From one of Passarge’s maps

(Ubersichtskarte der Mittleren Kalahari) as well as

from his descriptions, and using information from

Shaw (1985b) and Shaw et al. (2003), it can be

calculated that the level of Lake Ngami stood close to

930 m a.s.l. when Passarge travelled there. Thus

backflow from Lake Ngami to the Boteti can occur

at levels higher than 930 m a.s.l. only.

Shaw et al. (2003) identified Lake Ngami levels

during 4–3 ka as high as 936 m a.s.l. Following e.g.,

White & Eckardt (2006) who modelled isohypsic

palaeo-shorelines of the different MOZB sub-basins to

have developed isochronously, a 936 m a.s.l. lake

level in the Makgadikgadi Basin should have devel-

oped. Modern fluvio-lacustrine dynamics show a

different picture. Starting in 2009 and until 2012 Lake

Ngami has been filled by Okavango overspill to

approximately 928 m a.s.l., however, there is still no

lake in the Makgadikgadi Basin except for the

northern Sua Pan (at c. 900 m a.s.l.) which, however,

is periodically fed by the Nata River and some rivulets

from the east. On the other hand, our data from cored

Sua Pan sediments (see above) demonstrate the

existence of an extended lake in the Makgadikgadi

Basin during 4–3 ka. The source of the water was most

likely the Okavango system, which means that the

Boteti Gorge had formed prior to that period.

In respect of the Caprivi Depression (Fig. 2), Shaw

& Thomas (1988) suggested the existence of a late

Quaternary Lake Caprivi which linked the Zambezi

and Middle Kalahari drainage systems. A Lymnaea

shell from a 933 m a.s.l. Chobe River terrace was

radiocarbon dated to c. 18.5 ka cal. BP, and a

Bellamya shell from 931 m a.s.l. deposits was radio-

carbon dated to c. 2.7 ka cal. BP (Shaw & Thomas,

1988; Table 2). Shaw & Thomas (1988) considered

Bellamya a still-water genus indicating lake condi-

tions. Riedel et al. (2009) objected to this interpreta-

tion because they found modern Bellamya living

particularly in river environments with high water

energy, e.g., at the Popa Falls (Okavango River),

Sedudu Rapids (Chobe River) and Mambova Rapids

(Zambezi River). Shaw & Thomas (1988) related the

findings of the fossil gastropods to 936 m a.s.l. levels.

A palaeolake Caprivi sensu Shaw & Thomas (1988)

is questioned. The sub-recent gastropod shells found

by us on the banks of the Chobe River at c. 935 m a.s.l.

can be related to flooding events which did not sustain

lakes during the dry seasons. Such periodical river

lakes in the Caprivi Depression are probably not only

triggered by enhanced water flow in the Chobe river

system but also by Zambezi floods which may partly

be redirected to the Chobe River due to a drainage

bottleneck at the Mambova Rapids (Fig. 2). Shaw and

Thomas (1988) argued in a similar way but came to the

conclusion that a permanent lake was maintained.

Basin floor deformation modelling

The influence of water load on the basin depth is

insignificant in respect of hydrological properties. We

can speculate, however, that during lake highstands

the impounded water load, which brings an additional

normal stress on the surface, results in a slight

stabilisation of active faults, which otherwise would

be close to failure and rupture. When removing the

additional surface water load, tectonic activity along

these faults can possibly be triggered due to the small

stress release from the vanishing water load.

Conclusions

There is evidence that Okavango–Zambezi Rift Zone

lakes, which have mainly been fed by ITCZ controlled

rainfall over the Angolan highlands, were not inter-

connected with contemporaneous lake phases in the

Makgadikgadi Basin during MIS 3 and MIS 2. It has

been proceeded on the assumption that highstands

were isochronous because both lacustrine systems

feature palaeo-shorelines at c. 936 and c. 945 m a.s.l.

Published elevation data, however, may exhibit large

errors, of c. 15 m in the case of the terminal Okwa

Valley, and thus have to be treated with great caution.

On the other hand, it has been shown that during a

period of interaction between the different lacustrine

systems, during the late Holocene (4–3 ka), a c. 936 m

a.s.l. level of Lake Ngami did not result in the same

lake level in the Makgadikgadi Basin.

Understanding the spatiotemporal context of lake

development in the depressions of the MOZB has

further been hampered by the problem of proper dating.

It has been shown, however, that radiocarbon dating of

molluscan shells, which do not exhibit hard water

reservoir effects, and of organic rich lake sediments

produce reliable ages with uncertainties of 0.5–3%. We

thus can add information on six periods of lake

development: (i) Around 46 ka river systems with

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123

catchment areas north of the Makgadikgadi Basin, i.e.,

the Palaeo–Boteti and the Palaeo–Nata, maintained a

lake level of c. 910–912 m a.s.l. in the depression

(Fig. 5A). The moisture originated from the Indian

Ocean. (ii) At c. 37 ka a lake level below c. 908 m a.s.l.

is indicated by fossil lacustrine ostracods from a sandy

facies. (iii) High lake levels, probably reaching c.

936 m a.s.l., developed during the Last Glacial Max-

imum when the Okwa River controlled the hydrology

of the Makgadikgadi Basin (Fig. 5B) during a period of

a northwards extended winter rainfall zone until c.

17 ka. The development of high lake levels during that

period was fostered by significantly colder climate and

thus by reduced evapotranspiration and by less pro-

nounced seasonality. Around 17 ka the Okwa River

ceased flowing, supposedly due to the southward move

of the winter rainfall zone, which indicates that the

Kalahari was also affected by a mega-drought prevail-

ing over large areas of Africa during Heinrich Event 1.

The lack of hydromorphologic features hinting at a

post-Heinrich Event 1 lacustrine transgression into the

Okwa river valley implies that lake levels higher than

936 m a.s.l. have not developed in the Makgadikgadi

Basin since c. 17 ka. (iv) From c. 16 to 8 ka, probably

centred on the early Holocene, periods of desiccation

occurred as is indicated by deflation of lake sediments.

(v) A connection between Okavango–Zambezi Rift

Zone and Makgadikgadi Basin lacustrine systems most

likely opened between late MIS 2 and mid-MIS 1,

probably controlled by tectonics and resulting in the

formation of the Boteti Gorge, cutting through Mor-

emaoto and Gidikwe ridges. Tectonic activity may have

been triggered by stress release, following an abrupt

desiccation of the Makgadikgadi Basin after a lake

highstand, which likely happened e.g., during Heinrich

Event 1. The contribution of Okavango–Zambezi Rift

Zone water to the Makgadikgadi Basin during the

Holocene is indicated by the existence of an extended

lake in the latter depression, evidenced by fossil

lacustrine ostracods, and the Okwa River not being

active during that period (Fig. 5C). (vi) A relatively

large but shallow lake, at a level of c. 908 m a.s.l.

(Fig. 5D), developed for the last time in the Makga-

dikgadi Basin during the first half of the last

millennium.

The development of Lake Palaeo-Makgadikgadi

during the last 50 ka is characterised by strong lake

level fluctuations and thus substantial limnological

changes at different time intervals. The durations of

lake highstands remain unclear but may lie in a range

of decades to millennia. It is likely e.g., that from c.

23–17 ka a mega-lake at c. 936 m a.s.l. prevailed.

Only during such environmentally relatively stable

periods lacustrine biodiversity may have increased,

however, corresponding palaeo-communities have not

yet been recorded.

Acknowledgments We appreciate the field assistance of Kai

Hartmann, Mareike Schmidt, Caroline Seidig, Franziska Slotta,

Imke Steinmoller and Shuping Zhang (all FU Berlin, Germany).

Sebastian Erhardt (Imperial College, London, UK) conducted

D-GPS measurements in 2007 and 2008. Michael Taft (Abenden,

Germany) arranged the aerial survey of the Middle Kalahari in

2011. Many thanks go to Klaus Heine (University of Regensburg,

Germany) for sharing information on the Makgadikgadi Basin and

to Tomasz Goslar (Poznan Radiocarbon Laboratory, Poland) for

calculating radiocarbon dates of modern shells in respect of hard

water effect. Maike Glos (FU Berlin) processed sediment samples

and Jan Evers (FU Berlin) improved Figs. 3 and 4. We are grateful

to many helpful people in Botswana providing particular

information, to the Ministry of Minerals, Energy and Water of

Botswana for granting a research permit, and to the Deutsche

Forschungsgemeinschaft for financial support.

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