Geochimica et Cosmochimica Acta, Vol. 64, No. 20, pp. 3505–3513, 2000
Copyright © 2000 Elsevier Science Ltd
Printed in the USA. All rights reserved
0016-7037/00 $20.00 ⫹ .00
Pergamon
PII S0016-7037(00)00445-2
Mapping of C4 plant input from North West Africa into North East Atlantic sediments
YONGSONG HUANG,1,2* LYDIE DUPONT,3 MICHAEL SARNTHEIN,4 JOHN M. HAYES,5 and GEOFFREY EGLINTON5,6
1
Biogeochemistry Research Centre, Department of Earth Sciences, University of Bristol, Bristol, BS8 1RJ, UK
2
Department of Geological Sciences, Brown University, Providence, RI 02912-1846 USA
3
FB Geowissenschaften Bremen, Universitaet Bremen, D-28334 Bremen, Germany
4
Geologisch-Palaeontologisches Institut, Universitaet Kiel, D-24118 Kiel, Germany
5
Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA
6
Hanse-Wissenschaftskolleg, D-27753 Delmenhorst, Germany
(Received December 11, 1999; accepted in revised form April 30, 2000)
Abstract—Mapping the abundance of 13C in leaf-wax components in surface sediments recovered from the
seafloor off northwest Africa (0 –35°N) reveals a clear pattern of ␦13C distribution, indicating systematic
changes in the proportions of terrestrial C3 and C4 plant input. At 20°N latitude, we find that isotopically
enriched products characteristic of C4 plants account for more than 50% of the terrigenous inputs. This signal
extends westward beneath the path of the dust-laden Sahara Air Layer (SAL). High C4 contributions,
apparently carried by January trade winds, also extend far into the Gulf of Guinea. Similar distributions are
obtained if summed pollen counts for the Chenopodiaceae-Amaranthaceae and the Poaceae are used as an
independent C4 proxy. We conclude that the specificity of the latitudinal distribution of vegetation in North
West Africa and the pathways of the wind systems (trade winds and SAL) are responsible for the observed
isotopic patterns observed in the surface sediments. Molecular-isotopic maps on the marine-sedimentary time
horizons (e.g., during the last glacial maximum) are thus a robust tool for assessing the phytogeographic
changes on the tropical and sub-tropical continents, which have important implications for the changes in
climatic and atmospheric conditions. Copyright © 2000 Elsevier Science Ltd
(Pagani et al., 1999). In order to constrain these variables, there
is an urgent need to acquire data of C3 and C4 plant distributions during the critical sedimentary time horizons, such as the
last glacial maximum when pCO2 was ca. 30% lower than the
pre-industrial level of Holocene (Raynaud et al., 1993). Lake
sediment records are limited by their scope, since they record
only local vegetation changes (Talbot and Johannessen, 1992;
Giresse et al., 1994; Huang et al., 1999a;b). Individual marine
sediment cores, on the other hand, may record multiple factors,
such as changes in aeolian transport pathways and wind
strength (Sarnthein et al., 1981), in addition to the changes in
C3 and C4 plant on land (Collister et al., 1993). Accordingly,
mapping the C4 plant input over a wide geographic area at a
given marine-sedimentary time horizon may represent the best
approach to gain a comprehensive view of the ecosystem
changes on the adjacent continents.
In this study, we have tested the hypothesis that mapping
␦13C values of leaf-wax biomarkers, notably the n-C29 alkane
extracted from core-top sediments of the north east Atlantic,
can provide valuable information about the distribution of C3
and C4 plants on the adjacent west African continent. For this
purpose we have also used pollen distribution in marine sediments as an independent estimate for the input of C4 plants. We
have chosen this area because: 1) much of the terrigenous
organic carbon preserved in sediments from the Atlantic Ocean
around northwest Africa derives from atmospheric transport of
particles from the neighboring continent (Chester et al., 1972;
Simoneit et al., 1977; Sarnthein et al., 1981); 2) previous
organic geochemical studies have shown that both the aeolian
dusts and the marine sediments collected in the region contain
abundant terrigenous components (Simoneit et al., 1977; Cox et
al., 1982; Poynter et al., 1989); and 3) C4 plants are abundant
1. INTRODUCTION
Plants use two main carbon fixation pathways during photosynthesis: Calvin-Benson (C3), Hatch-Slack (C4) cycles (e.g.,
O’Leary, 1981). C3 plants, including trees, shrubs and coolclimate grasses, generally have ␦13C values in the range of ⫺22
to ⫺33‰. In contrast, C4 plants, including many tropical
grasses and sedges, usually exhibit ␦13C values in the range of
⫺9 to ⫺16 ‰. Thus, the ␦13C values of bulk organic matter
have been used to reconstruct the past changes in C3 and C4
plant abundance from paleosols (Krishnamurthy and DeNiro,
1982; Guillet et al., 1988), and in some cases from marine
sediments (France-Lanord and Derry, 1994). However, organic
matter in lake and marine sediments includes input from both
terrestrial and aquatic sources which often have different isotopic compositions (Westerhausen et al., 1992; Meyers and
Ishiwatari, 1993; Huang et al., 1999a). Isotopic study of sedimentary organic matter thus requires compound specific isotopic analysis (CSIA) (Hayes et al., 1990). CSIA provides a
means to reconstruct the changes in terrestrial C3 and C4 plant
distributions by measuring the ␦13C values of higher plant
biomarkers in aquatic sediments (Collister et al., 1993; Bird et
al., 1995; Huang et al., 1993; Huang et al., 1999a;b; Kuypers et
al., 1999).
Significant debate exists on the climatic and atmospheric
conditions controlling the balance of C3 and C4 plants in the
natural ecosystems. At the center of debate, there is a lack of
understanding on the relative impact of controlling factors,
notably atmospheric pCO2 (Cerling et al., 1997; Street-Perrott
et al., 1997; Kuypers et al., 1999), aridity and temperature
*Author to whom correspondence should be addressed (yongsong_
[email protected]).
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Y. Huang et al.
on the western African continent (White, 1983). Our data
provide the basis for studying the changes in C4 plant abundances on the west African continent during marine-sedimentary time horizons such as glacial/interglacial cycles.
2. SAMPLES AND EXPERIMENTAL
2.1. Samples
We have analyzed 41 marine surface sediments and 11 aeolian dust
samples (locations of the samples are shown in Fig. 1A,B). Marine
surface sediment samples (0 –2 cm) are from Meteor, Discovery, and
bofs cruises. Aeolian dust was collected during CYCLOPS (1974) and
TAF cruises (1972) using nylon mesh for ship-board sampling (Chester
and Johnson, 1971).
2.2. Compound Isolation
Marine surface sediments (4 –7 g) were freeze-dried overnight. The
dry sediments and aeolian dusts were spiked with known amounts of
internal standards (C36 n-alkane) and then ultrasonically extracted for
15 min each time using solvents of sequentially decreasing polarities
(MeOH ⫻ 2; MeOH:CH2Cl2 1:1 ⫻ 2; and CH2Cl2 ⫻ 2) (Huang et al.,
1999a). The total extracts were rotary-evaporated to near dryness at
30°C, then partitioned between 25 ml of 0.1 M KCl solution and
dichloromethane in a separatory funnel in order to remove salts.
CH2Cl2 extracts were collected, dried over pre-combusted anhydrous
Na2SO4. CH2Cl2 was then removed by rotary-evaporation at 30°C.
2.6. Gas Chromatography-Isotope Ratio-Mass Spectrometry
(GC-IRMS)
Analyses were conducted in Biogeochemistry laboratories, Bloomington, Indiana University, on a Finnigan-MAT delta S mass spectrometer interfaced onto a HP 5890 gas chromatography. The interface was
of an alumina reactor (0.5 mm ID) containing nickel and platinum
wires (0.1 mm OD). GC column was a 60 m, 0.32 mm ⫻ 0.5 ␮m fused
silica capillary column (Restek). Temperature program was from 60 to
200°C at 12°C/min, from 200°C to 310°C at 4°C/min, and then kept
isothermal for 40 mins. Analyses were triplicated (standard deviation
␴ ⫽ 0.1– 0.4 ‰). CO2 with pre-calibrated isotopic composition was
used as standard. ␦13C values are expressed versus PDB. For FAMEs
and alcohols, an isotopic correction using mass balance approach is
performed to remove the isotopic contribution from the derivatizing
reagents (i.e., methyl group for FAMEs and TMS group for alcohols)
(Huang et al., 1995; Huang et al., 1999a;b). We have used a simple
mass balance approach to remove/reduce the contribution of n-alkanes
from marine sources (including petroleum) to the C29 n-alkane (␦w29).
We write,
␦ 29 ⫻ A 29 ⫽ ␦ w 29 ⫻ 共 A 29 ⫺ Ax 29兲 ⫹ ␦ x 29 ⫻ Ax 29 共1兲
where 29 refers to C29 n-alkane in marine sediments; A ⫽ abundance,
w ⫽ higher-plant wax input; x ⫽ contamination. In order obtain ␦w29,
we need to estimate the Ax29 and ␦x29 (␦29 and A29 are measured
values). For this purpose, we have calculated the mean ratios of C29/C30
n-alkanes for all marine sediments (5.4 ⫾ 1.7) and aeolian dusts (7.6 ⫾
2.1) (Table 1). The lower mean C29/C30 ratio in marine sediments is
consistent with small input of long chain n-alkanes from marine
sources. Assuming that contamination to C29 and C30 n-alkanes in
sediments is equal (i.e., Ax29 ⫽ Ax30 ⫽ Ax), we can write,
2.3. Derivatization and Chromatographic Separation
The total extracts were transferred into 3.5 ml vials, dried under a
stream of N2, and solublized in 0.5 ml of toluene. 2 ml of freshly
prepared anhydrous 5% HCl/MeOH (by adding acetic chloride into
anhydrous methanol) was then added to each sample to transesterify the
lipids at 50°C over night (Christie, 1982). Samples were transferred to
20 ml test tubes, and were added with 5 ml of 0.1M KCl aqueous
solution. Lipids were recovered by extraction using hexane/ether (5:1)
and dried under a stream of N2. The total lipids were then separated
using thin layer chromatography (TLC) into aliphatic hydrocarbon,
fatty acid methyl esters (FAMEs) and alcohol fractions, using a solvent
system of hexane: ethyl acetate 7:1. Whenever necessary, aliphatic
hydrocarbons, FAMEs and alcohol fractions were further urea-adducted to obtain n-alkanes, n-acids and n-alkanols with greater purity
(Huang et al., 1995). Alcohols were derivatized to trimethylsilyl ethers
using bis(trimethylsilyl)-trifluoro-acetamide (BSTFA) at 60°C for 3 h
(Huang et al., 1995).
2.4. Gas Chromatography (GC)
Analyses were carried out on a Carlo Erba 5300 gas chromatography
fitted with an on-column injector and a fused silica capillary column
(HP1, 50 m ⫻ 0.32 mm, coated with OV-1). Hydrogen carrier was used
as carrier gas with a flow rate of about 2 ml/min. Typical temperature
program was: 40°C (isothermal for 1 min), 15°C/min to 150, then to
310°C at 5°C/min, isothermal for 30 min. GC is connected to a
Minichrom data processing system. The internal standard added prior
to the sample extraction is used as reference for quantification.
2.5. Gas chromatography-mass spectrometry (GC-MS)
70 eV EI GC-MS analyses were performed on a Carlo Erba Mega
gas chromatograph (on-column injection) interfaced directly with a
Finnigan 4500 mass spectrometer. Data acquisition and processing
were conducted using a INCOS data system. Typical conditions were:
column (CPSil-5CB, 50 m ⫻ 0.32 mm, film thickness 0.12 ␮m, fused
silica capillary; CHROMPACK), helium as carrier gas. Temperature
program was identical to that of GC analyses. The compounds were
identified by comparison of their mass spectra and retention time with
published data.
7.6 ⫽ Aw 29/Aw 30 ⫽ 共 A 29 ⫺ Ax兲/共 A 30 ⫺ Ax兲
or
(2)
Ax ⫽ 共7.6 * A 30 ⫺ A 29兲/6.6
where A30 and A29 are the measured abundances for C30 and C29
n-alkanes in marine sediments. The ␦13C values for contaminants
(␦x29) can be estimated from an even carbon, short-chain n-alkane
which is generally absent from the plant waxes. We have used C24
n-alkane because even carbon n-alkanes of shorter chain-length (i.e.,
⬍C24) are often absent from samples or co-eluting with other compounds.
3. RESULTS AND DISCUSSION
3.1. Lipid Distributions and Abundances
Concentration ranges of total organic carbon (TOC) and
results of lipid analyses are summarized in Table 1 and Figure
2. The concentrations of TOC in the marine sediments and in
the aeolian dusts are similar. Long-chain n-alkanes with odd
carbon numbers (C25 to C35), and the corresponding evencarbon-number n-alkanols and n-alkanoic acids, are characteristic products of higher plants (Eglinton and Hamilton, 1963;
Tulloch, 1976). Distributions of long-chain n-alkanes, n-alkanoic acids, and n-alcohols, expressed in terms of standard
measures, occur in the sediments and dusts in similar carbon
number range (Fig. 2). These compounds derive mainly from
higher-plant leaf-waxes (Simoneit et al., 1977; Simoneit, 1997;
Cox et al., 1982; Eglinton and Hamilton, 1963; Westerhausen
et al., 1992; Gagosian et al., 1981). On a weight basis, they are
10 –100 fold more concentrated in the aeolian dusts than in the
sediments, indicating both dilution of the sediments by marine
input and degradation of the lipids during exposure in the water
column and at the sea floor. This contrast is attenuated when the
lipid concentrations are expressed relative to TOC, indicating
C4 plant input into North East Atlantic
Fig. 1. Northwest Africa and the Northeast Atlantic Ocean. A. Continent: The phytogeographical regions (White, 1983)
from North to South: Med, Mediterranean vegetation zone; MST, Mediterranean—Saharan transitional steppes; Sahara,
absolute desert, desert and semi-desert; Sahel, semi-desert grassland to Acacia wooded grassland; Savanna, dry savannas
and woodland; SDF, semi-deciduous forest; RF, tropical rain forest. The major wind systems are: the North East Trades,
low altitude winds stronger in winter; the Saharan Air Layer, medium altitude, mid-tropospheric winds, strongest in
summer; January Trades, low altitude winds carrying dust in winter from Northern Nigeria and Lake Chad areas, when
ITCZ is in southern-most position. Ocean: shipboard sampling locations for dust: each set of two stars connected by a line
denotes a dust sample which was collected during the course of the sail by the ship (Chester and Johnson, 1971); the shaded
areas over the North East Atlantic are those for which high occurrences of atmospheric haze are recorded; northern zone
during summer and southern zone during winter. Bathymetry: 200 and 2000m, shelf edge and break, respectively. B.
Continent: Source areas of Chenopodiaceae and Amaranthaceae (ChenoAms) pollen, mainly C4. Light; MST with some C3
grasses and C4 halophytes. Dark; sparse herb and shrub. Desert, extremely vegetation poor, hyper-arid desert area. Ocean:
isopol contours (percentage of total pollen) of ChenoAms pollen in surface sediment samples (Dupont and Agwu, 1991;
Hooghiemstra et al., 1986). Sampling sites are ⫹. C. Continent: source areas of grass pollen (Poaceae), C4. Light, Savanna
with C3 trees and shrubs. Dark, Sahel which is C4 grass dominated. Ocean: isopol contours of Poaceae pollen in surface
sediments (Dupont and Agwu, 1991; Hooghiemstra et al., 1986). D. Continent: Source areas of C4 pollen. The C4 biomass,
mostly made up by Cheno-Am and grasses, very approximately indicated by the intensity of shading. Ocean: Isopol
contours of main C4 components in surface sediments. Percentage plotted is Poaceae ⫹ 0.5 * Cheno-Am pollen percentage.
E. Continent: Circles, major source areas for dust deflation, such as Holocene lake deposits (Petit-Marie, 1991) (N.B. not
shown but to east of map is another major source area around Lake Chad, centred on ca. l4°N 11°E). Arrows mark major
river inputs of particulate terrigenous organic carbon (104 tonnes䡠y⫺1) (Ludwig et al., 1996). Ocean: Distribution of C4 plant
wax (% of total C3 and C4) in the surface sediment samples, calculated from values of ␦29 using a two-component mixing
equation with end member values of ⫺34 and ⫺19 ‰, respectively (Rieley et al., 1993; Collister et al., 1994). Sampling
sites are ⫹.
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Y. Huang et al.
Table 1. Compositional and distributional parameters for n-alkyl lipids of marine surface sediments and aeolian dusts.a
Parameters
TOC (%)b
n-Alkyl lipids
Chain length range
Cmaxc
CPId
ACLe
Concentration
(␮g/g.d.w)
Concentration
(␮g/g TOC)g
C29/C30 n-alkane
␦13C valuesh
Marine sediments (n ⫽ 41)
n-Alkanes
Major C23–C35
Minor C37–C43
Most C31
Some C29
1.7–6.9
(4.2 ⫾ 1.3)
28.3–29.9
(29.3 ⫾ 0.4)
0.2–2.2
(0.7 ⫾ 0.4)
0.2–2.7
(1.0 ⫾ 0.7)
1.6–8.9
(5.4 ⫾ 1.7)
⫺35.4 to ⫺25.3
(⫺27.8 ⫾ 2.3)
n⬘ ⫽ 41
Dusts (n ⫽ 11)
0.19–2.33 (0.86 ⫾ 0.58)
n-Acids
n-Alkanols
C14–C34
C22–C34
Most C26
Some C24
3.3–5.0
(4.1 ⫾ 0.5)
25.3–27.9
(26.2 ⫾ 0.5)
0.16–4.7
(1.1 ⫾ 1.0)
0.3–2.2
(1.2 ⫾ 0.5)
—
Most C28
Some C26 or C30
0.8–6.4
(3.8 ⫾ 1.5)
24.9–28.6
(27.2 ⫾ 0.9)
0.05–2.1
(0.5 ⫾ 0.5)
0.1–2.2
(0.7 ⫾ 0.5)
—
⫺29.5 to ⫺22.8
(⫺26.1 ⫾ 2.6)
n⬘ ⫽ 11
⫺30.7 to ⫺25.7
(27.9 ⫾ 2.0)
n⬘ ⫽ 11
n-Alkanes
C23–C35
Most C31
Some C29
2.8–8.2
(5.7 ⫾ 1.3)
29.1–29.7
(29.4 ⫾ 0.2)
2.9–49.9
(19.4 ⫾ 15.7)
5.3–33.4
(14.8 ⫾ 9.2)
4.4–11.7
(7.6 ⫾ 2.1)
⫺31.0 to ⫺25.3
(⫺28.0 ⫾ 1.8)
n⬘ ⫽ 8
0.38–2.54 (1.27 ⫾ 0.76)
n-Acids
n-Alkanols
C14–C34
C22–C34
Most C28
Some C24
4.4–6.2
(5.5 ⫾ 0.7)
25.2–26.8
(25.7 ⫾ 0.7)
0.82–103.5
(26.9 ⫾ 32.7)
2.1–48.4
(16.8 ⫾ 14.6)
—
Most C30
Some C28
5.3 ⫾ 10.4
(8.0 ⫾ 1.9)
27.6–29.1
(28.3 ⫾ 0.4)
5.3–36.0
(13.8 ⫾ 11.2)
4.4–16.8
(7.7 ⫾ 4.8)
—
⫺29.5 to ⫺24.2
(⫺27.0 ⫾ 2.7)
n⬘ ⫽ 4
⫺29.6 to ⫺25.2
(⫺27.3 ⫾ 1.4)
n⬘ ⫽ 8
The full data set containing abundance and ␦13C values of individual compounds is available at www.pangaea.de.
Total organic carbon content by weight.
c
The carbon number for homologues with the highest abundance.
d
Carbon preference index ⫽ sum of odd C25 to C33/sum of even C24 to C32 for n-alkanes, or sum of even C24 to C32/sum of odd C23 to C31 for
n-acids and n-alkanols.
e
Average chain length ⫽ 兺[Ci] ⫻ i/兺[Ci], where C is concentration and i ranges from 23 to 33 for n-alkanes, or 22 to 34 for n-acids and n-alkanols.
f
Concentration relative to the dry weight of the sediment or dust.
g
Concentration relative to the weight of organic carbon of the sediment or dust.
h
Weighted average ␦13C values for C27, C29, C31 n-alkanes, or C26, C28, C30 n-acids and n-alkanols. n⬘ refers to the number of samples for the
specific type of compounds. Small amounts of C37 to C43 extended n-alkanes with no odd/even predominance were found in most of the marine
sediments and were attributed to petroleum sources (Huang et al., 1993). The ␦13C values for the n-C29 alkane have been corrected for the marine
n-alkane contamination, using the mass balance approach described in the “Samples and Experimental”.
a
b
that the degradation of terrestrial organic carbon is partly
compensated by the addition of marine organic matter. Degradation of leaf waxes must have also taken place during the
transport to the ocean sediment. Gagosian et al (1986) calculated a C29 n-alkane flux into the sediments of oligotrophic
Western North Atlantic and compared their data with the surface hydrocarbon data of Farrington and Tripp (1977). They
found that the atmospheric flux for C29 n-alkane was 100 –500
times of the sediment accumulation rate and concluded only
0.2% of the atmospheric flux survived degradation reactions in
the water column. Unfortunately, we do not have the time
series data required to calculate the flux of n-alkanes to the
surface sediments in the current study.
For the C25 to C33 n-alkanes, the abundance of odd- relative
to even-carbon homologues, expressed in terms of the CarbonPreference Index (CPI), is higher in the dusts than in the
sediments (Table 1; Fig. 2). The difference is of marginal
significance, but parallel differences (i. e., a preference for the
terrestrial biogenic homologues in the dust) are observed for
the C23 to C31 n-acids and n-alkanols. For instance, the relative
abundance of C20 to C24 n-alkanols is higher in marine sediments than in dusts (Fig. 2). It follows that some portion of the
long-chain, n-alkyl lipids in the sediments does not derive from
leaf waxes but instead from marine organisms and/or petroleum
(Uzaki et al., 1993; Poynter et al., 1989).
There are minor geographical variations in the concentrations of leaf wax n-alkyl lipids in marine surface sediments
(Huang et al., 1993). As in previous investigations (Poynter et
al., 1989), concentrations are not related to distance off the
coast or to water depth. Most samples contain 0.4 to 0.7 ␮g/g
dry wt. C23 to C35 n-alkanes, 0.3 to 0.6 ␮g/g n-alkanols, and 0.5
to 0.8 ␮g/g n-acids. The results are consistent with aeolian
transport as the primary mechanism for delivery of continental
material over long distances (Simoneit et al., 1977; Cox et al.,
1982; Gagosian et al., 1981). Deviations are observed only at a
few sites. Near river mouths, concentrations have been enhanced by rapid sedimentation and greater inputs of terrestrial
organic carbon. Concentrations of C22 to C34 n-acids are also
3– 4 times higher in the intense upwelling zone between 15 and
20°N (Huang et al., 1993), indicating enhanced input and/or
selective preservation of organic matter.
3.2. Carbon Isotopic Compositions of Leaf Waxes
In contrast to the relatively invariant abundance data, the
isotopic compositions of long-chain, n-alkyl lipids vary significantly and smoothly over large geographic areas (Fig. 3; Fig.
4). Table 1 and Figure 3 show that the total ranges approach
10‰ and that the standard deviations of the populations are
ten-fold greater than the errors of measurement. Small contributions from petroleum or fossil hydrocarbons and from marine
organisms to the marine sediments were also apparent. These
were well reflected in the lower CPI values of long-chain
n-alkyl lipids in marine sediments than in aeolian dusts (Table
1), as well as the previously reported C37 to C43 n-alkanes
(Huang et al., 1993). Smooth curves based on the abundances
of the even-carbon alkanes were used to estimate contributions
and isotopic compositions of the fossil hydrocarbons (see
C4 plant input into North East Atlantic
3509
Fig. 2. The normalized distributions of long chain n-alkyl lipids in marine surface sediments and aeolian dust samples.
“Samples and Experimental” for details). Subtraction of this
component yielded ␦ values for the non-fossil hydrocarbons
(Fig. 3). These corrections were small for the n-C29 alkane,
averaging less than 1‰ (Fig. 4a). Cross plots of the ␦13C values
for the n-C29 alkane (hereafter, ␦29) against those of weighted
averages of the C27 and C31 n-alkanes, C26, C28, and C30 acids
and alcohols yielded r2 ⫽ 0.88, 0.91 and 0.87, respectively, and
slopes that did not differ significantly from 1.0 (Fig. 4). The
acids were consistently enriched in 13C by 2‰, a difference
readily attributable to input from marine sources. We conclude
that ␦13C values of C29 n-alkane provide valid and sufficient
indications of inputs from terrestrial higher plants.
Core-top sediments were analyzed at the 41 sites marked in
Figure 3 and Figure 1E. Values of ␦29 are consistently more
positive between 15 and 20°N (⫺26.0 to ⫺26.5‰), but trend to
more negative values northward (down to ⫺32.2‰). Values of
␦29 increase with distance from the coast south of 15°N, ranging from ⫺34.7‰ at the mouth of the River Cess to ⫺27‰ in
abyssal sediments. In Figure 1E, such variations are expressed
in terms of percentage contributions from C4 plants to total
plant waxes. Contributions have been evaluated using a twocomponent mixing equation with C3 and C4 inputs represented
by ␦29 values of ⫺34 and ⫺19‰, respectively (Rieley et al.,
1993; Collister et al., 1994). The resulting contours indicate
smooth variations in the abundance of C4 inputs around the
northwest African coast, with a maximum around 20°N. This
pattern must derive both from contemporary vegetation and
from dust sources far inland. The supply of hydrocarbons from
both sources is dependent on local aridity, wind-strengths and
directions and, to a much lesser extent, fluvial transport. The
increase in C3 percentage near the river mouths may also in part
result from the bias to the C3 vegetation in fluvial transport as
shown by Bird et al. (1995), Bird et al. (1998).
The ␦13C values of leaf wax n-alkanes from the dust samples
are within the range shown in marine surface sediments, but the
overall variability is relatively small because of the limited in
sample locations (Table 1, Fig. 1A). However, a recent study
on a transect of dust samples collected along the Africa west
coast (covering the latitudinal range of 0 –35°N) shows a clear
latitudinal ␦13C variation of TOC (G. Lavik, private communication), and the pattern is similar to that of the alkane isotopic
data in the marine surface sediments.
Compound-specific isotopic analyses of terrigenous components in marine sediments have been previously related to
3510
Y. Huang et al.
recent vegetation (Hooghiemstra, 1986; Hooghiemstra, 1988).
The C4 plant distribution north of the Sahel is reflected in the
percentage distribution of pollen from the Chenopodiaceae and
Amaranthaceae (“Cheno-Am,” mainly C4) in the marine surface sediments of the eastern Atlantic (Fig. 1B). That of the C4
plants south of the Sahara is seen in that of the Poaceae (C4
grasses, Fig. 1C). These views are combined in Figure 1D,
which provides an overview of C4 inputs to sediments, judged
from pollen abundances. To attempt a single assessment of the
C4 pollen signal, we have summed percentages of Poaceae and
Cheno-Am, reducing those of the Cheno-Am by half. This
loading allows approximately for the distribution of C3 species
among the Cheno-Am, specifically among the non-halophytic
species of these families. Within the Poaceae, the amount of C3
species among the tropical grasses is negligible. Overall, the
pattern of C4 inputs inferred from values of ␦29 (Fig. 1E) is in
general agreement with the pollen distribution (Dupont and
Agwu, 1991; Hooghiemstra et al., 1986) (Fig. 1D).
3.4. Wind Regimes and Transport Pathways
Fig. 3. The ␦ C values of the higher-plant derived C29 n-alkanes
isolated from the marine surface sediments in the north east Atlantic.
13
vegetational sources for the area studied here (Huang et al.,
1993), in a core at a river mouth in North Queensland, Australia
(Bird et al., 1995) and (using lignin phenols) in the Gulf of
Mexico (Goñi et al., 1997). Here, we compare our isotope
mapping with pollen counts in marine surface sediments which
provide an alternative and completely independent proxy for C4
plant input (Dupont and Agwu, 1991; Hooghiemstra et al., 1986).
3.3. Phytogeographical Zones in Northwest Africa and
Pollen Distributions in Sediments
Phytogeographical zones in northwest Africa range from the
Mediterranean forest (Med in Fig. 1A) in the north to the
tropical rain forest (RF) in the south (White, 1983). North of
the Saharan desert (Med and MST), the vegetation consists
predominantly of C3 plants, with the exception of some halophytic species of the Chenopodiaceae (Dowton, 1975; Raghavendra and Das, 1978). Grasses of the Mediterranean-Saharan
transitional steppes (MST), just south of the Atlas Mountains,
are mostly C3 plants. In the sparse vegetation of the Sahara,
halophytes and other herbs count among the C4 plants (Dowton, 1975; Raghavendra and Das, 1978; Ehleringer et al., 1977).
The grasses of the Sahel, savanna including woodland, and
semi-deciduous forest (SDF) are C4 plants, while the woody
species are C3. In the tropical rain forest, woody C3 species
vastly dominate over C4 grasses (White, 1983). Thus, C4 plants
are mainly found in the Sahel and the savanna, but also in the
Sahara, a distribution that has been attributed to high temperatures during the growth season in conjunction with moisture
stress (Vogel and Fuls, 1978; Teeri and Stowe, 1976). CAMplants do not form a significant constituent of the vegetation of
northwestern Africa (Winter and Smith, 1996).
The distribution of pollen in seafloor sediments provides
independent information about inputs from contemporary and
Specific sources of the wind-blown dust have been inferred
from various proxies including lithology, the major- and traceelement contents of the silt and clay components, fresh-water
diatom distributions and pollen counts (Sarnthein et al., 1981;
Chiapello et al., 1997; Gasse et al., 1989; Hooghiemstra, 1988;
Pokas, 1991). The origin of the dust load is commonly assigned
to three main deflation regions (Pye, 1987) (Fig. 1A).
1. The Atlas Mountains and coastal plain, from which dust is
carried almost parallel to the coast by the low level North
East Trade (NET) winds.
2. The southern Sahara and Sahel, from which dust is raised in
the northern hemisphere summer by easterly winds into the
higher-level flow of the SAL (African Easterly Jet, AEJ) and
is carried beyond the continental margin between 10 and
25°N.
3. The southern edge of the Sahara (alluvial plains of Niger,
Faya Largeau and Chad), from which dust is uplifted by the
NET in the northern hemisphere winter and carried to a wide
area between 2 and 15°N off Africa.
The two major wind systems (Sarnthein et al., 1981; Dupont
and Agwu, 1991; Hooghiemstra et al., 1986), i.e., the NE
(January) trade winds at low altitudes and the SAL at mid
altitudes, are important transport agents for pollen and plant
waxes from northwest Africa to the Atlantic. The highest C4
input, as registered by both pollen abundances and ␦29 values,
is in regions directly influenced by the SAL (15 to 22°N, Fig.
1D,E). Inputs from C4 sources gradually decrease toward the
north, mainly because of diminishing strength of the SAL. The
input from the NE trade winds carrying C3 organic matter
becomes greater and is indicated by a latitudinal decrease in
␦29. At lower latitudes, in the Gulf of Guinea, the January
trades blow across the savanna and semi-deciduous-forest
(SDF) areas, carrying C4 components over long distances to the
ocean, leading to more positive values of ␦29 far offshore.
Additionally, mangroves (C3) form thick swampy belts along
some regions of the tropical coast, notably the Gulf of Guinea.
The numerous river systems (Fig. 1B–D) in the equatorial
regions favor fluvial transport of plant material and pollen from
C4 plant input into North East Atlantic
3511
Fig. 4. Plots of the ␦13C values of C29 n-alkanes (after correcting for the contribution from petroleum) against the
measured ␦13C values of other higher-plant derived compounds in marine surface sediments. A, Plot against the measured
␦13C values of C29 n-alkanes. B, Plot against the ␦13C values of C27 and C29 n-alkanes (weighted average). C, Plot against
␦13C values of C26, C28 and C30 n-acids (weighted average). D, Plot against ␦13C values of C26, C28, C30 n-alkanols
(weighted average).
the tropical forest and riverside vegetation. This process probably accounts for the more negative values of ␦29 near the coast
(represented by decreased estimates of contributions from C4
sources, Fig. 1E).
Primary production differs widely from vegetation zone to
vegetation zone (Dupont, 1993 and references therein). Forested areas (Mediterranean forest, tropical seasonal forest and
tropical rain forest) have a primary production typically exceeding 25 kg/m2; open vegetation has a lower primary production (savanna woodland 2–20 kg/m2, Sahelian grassland
and desert shrub land less than 5 kg/m2). The production of
C4-plant material thus will be largest in the savanna woodland.
Production of C4 material in the Sahel and the desert will be
moderate, although the relative proportion of C4 plants to the
vegetation might be bigger.
Matters of timing require further investigation. The C4 distribution based on ␦29, which is essentially congruent with that
based on pollen, is controlled largely by dust-borne materials.
The dusts come from extensive dried-up lake beds of ca.
mid-Holocene age on the northern and southern fringes of the
Sahara and the Sahel (Petit-Marie, 1991; circles, Fig. 1E). The
waxes must be of the same age. Indeed, paleohydrological and
paleobotanical studies of the late Quaternary of northwest
Africa have shown that such areas were more vegetated, probably with C4 grasses, during the pluvial periods of the Holocene
optimum (e.g., Petit-Marie, 1991), 8500 –3500 years B.P. Soil
samples from such regions contain abundant leaf-wax components with the usual distributions (Huang, unpublished data).
The fine fractions of these soils, laden with leaf-wax components (Bird and Pousai, 1997), can be carried over long distances (Simoneit et al., 1977; Gagosian et al., 1981).
Leaf waxes can also be picked up from contemporary vegetation by the wind streams en route. Distinction between
modern and ancient (8500 –3500 years B.P.) wax components
should eventually become possible via 14C dating of individual
compounds (Eglinton et al., 1997) and pollen isolates, but the
majority of the pollen input probably derives from contemporary vegetation. There is no established, quantitative relation-
3512
Y. Huang et al.
ship between either pollen counts or wax production and plant
biomass. Hence, the comparisons we draw between the maps of
two very different proxies (pollen and ␦29) are mostly qualitative because the basic data derive from integration of inputs
from numerous plant species growing in a variety of environments. Conjoint use of pollen-, biomarker- and dust-proxy data,
however, will be an important tool for assessing the different
contributions of vegetation and dust sources in terms of aridity,
wind direction and strength.
4. CONCLUSIONS
We have revealed a systematic variation of C4 plant input to
north east Atlantic sediments from north west Africa by compound-specific carbon isotope studies of C29 n-alkanes extracted from marine core-top sediments. This is the first time
that a molecular isotopic mapping over a wide geographic area
has been conducted to assess the organic carbon input from C3
and C4 plant into the marine sediments. The distribution is
readily explained by the C3 and C4 plant distributions on the
north west African continent and the wind regimes. Our findings for the parallel distribution of pollen counts and ␦29
underscore the importance of C4 vegetation as a source of
terrigenous organic matter contributed to modern oceans. C4
plant distributions are closely related to climatic (Vogel and
Fuls, 1978) and atmospheric (Cerling et al., 1997; Street-Perrott
et al., 1997) conditions and phytogeographical modeling techniques utilize these relationships (Jolly and Haxeltine, 1997).
Climate variability in the past must have affected the export of
both the quantities and types of organic carbon from land to the
oceans. For instance, at the last glacial maximum (LGM), the
expansion of C4 plants led to more positive ␦13C values in the
lacustrine sediment records of Africa (Ehleringer et al., 1977;
Street-Perrott et al., 1997). Maps like those developed here, but
for LGM time slices, would provide new views of African
paleophytogeography which could be compared to those afforded by the pollen record. Compound-specific isotope mapping thus provides a means of paleo-environmental studies for
areas of the world where C4 plant biomass has been significant.
C4 plants are widely distributed around the globe, especially
around the tropics and subtropics. Consideration of the global
role of C4 plants is of increasing importance in view of the
likely future elevation of atmosphere CO2 concentrations and
associated global changes in temperature and aridity, all of
which are important factors controlling the competitiveness of
C4 and C3 plants (Jolly and Haxeltine, 1997).
Acknowledgments—We acknowledge the financial support provided
the Royal Society Queens Fellowship (Y. Huang), American Chemical
Society—Petroleum Research Fund (Y. Huang, ACS-PRF #35208G2), the Deutsche Forschungsgemeinschaft (L. Dupont, grant We992/
26), the National Aeronautics and Space Administration (J. M. Hayes,
grants nagw-1940 and nag6-6660), and the Natural Environment Research Council for the TIGER Program and for the use of the analytical
facilities (GR3/2951 and 3758). We also thank the masters and crews
of the Meteor and the Discovery and acknowledge the support of the
German national program of climate research. We thank J. Carter,
Andrew Gledhill, and Jon Fong for technical assistance, Profs. John
Hedges, Alayne Street-Perrott, and John Raven for helpful discussions,
and Prof. Roy Chester and Dr. Helene Cachier for providing aeolian
dust samples. Data are available in the Pangaea database (www.
pangaea.de).
Special handling: T. E. Cerling
REFERENCES
Bird M. I., Summons R. E., Gagan M. K., Roksandic Z., Dowling L.,
Head J., Fifield L. K., Cresswell R. G., and Johnson D. P. (1995)
Terrestrial vegetation change inferred from n-alkane ␦13C analyses
in the marine environment. Geochim. Cosmochim. Acta 59, 2853–
2857.
Bird M. I. and Pousai P. (1997) Variations of ␦13C in the surface soil
organic carbon pool. Global Biogeochem. Cyc. 11, 313–322.
Bird M. I., Giresse P., and Ngos S. (1998) A seasonal cycle in the
carbon-isotope composition of organic carbon in the Sanaga River,
Cameroon. Limnol. Oceanogr. 43, 143–146.
Cerling T. E., Harris J. M., MacFadden B. J., Leakey M. G., Quade J.,
Eisenmann V., and Ehleringer J. R. (1997) Global vegetation change
through the Miocene/Pliocene boundary. Nature 389, 153–158.
Chester R. and Johnson L. R. (1971) Atmospheric dusts collected off
the west African coast. Nature 229, 1105–107.
Chester R., Elderfield H., Griffin J. J., Johnson L. R., and Padgham
R. C. (1972) Eolian dust along the eastern margins of the Atlantic
Ocean. Mar. Geol. 13, 91–105.
Chiapello I., Berganetti G., Gomes L., Chatenet B., Dulac F., Pimenta
J., and Santos Suares E. (1997) Origins of African dust transported
over the northeastern tropical Atlantic. J. Geophysical Research 102,
13,701–13,709.
Christie W. W. (1982) Lipid Analysis. 2nd Ed., Pergamon Press,
Oxford, England.
Collister J. W., Huang Y., and Eglinton G. (1993) Assessment of
terrestrial input into marine sediments during glacial/interglacial
sequences in the Northeastern Atlantic by compound-specific isotopic analysis of biological markers. In Org. Geochem. (ed. K. Øygard), 430 – 433, Falch Hurtigtrykk, Oslo.
Collister J. W., Rieley G., Stern B., Eglinton G., and Fry B. (1994)
Compound specific 13C analyses of leaf lipids from plants with
differing carbon dioxide metabolism. Org. Geochem. 21, 619 – 628.
Cox R. E., Mazurek M. A., and Simoneit B. R. T. (1982) Lipids in
Harmattan aerosols of Nigeria. Nature 296, 848 – 849.
Dowton J. S. (1975) The Occurrence of C4 Photosynthesis Among
Plants. Photosynthetica 9, 96 –105.
Dupont L. M. and Agwu O. C. (1991) Environmental control of pollen
grain distribution patterns in the Gulf of Guinea and offshore NWAfrica. Geolische Rundschau 80, 567–589.
Dupont L. M. (1993) Vegetation zones in NW Africa during the
Brunhes Chron reconstructed from marine palynological data. Quat.
Sci. Rev. 12, 189 –202.
Eglinton G. and Hamilton R. J. (1963) The distribution of n-alkanes. In
Chemical Plant Taxonomy (ed. T. Swain), 187–217, Academic
Press, London-New York.
Eglinton T. I., Benitez-Nelson B. C., Pearson A., McNichol A. P.,
Bauer J. E., and Druffel E. R. M. (1997) Variability in radiocarbon
ages of individual organic compounds from marine sediments. Science 277, 796 –799.
Ehleringer J. R., Cerling T. E., and Helliker B. R. (1977) C4 photosynthesis, atmospheric CO2, and climate. Oecologia 112, 285–299.
Farrington J. W. and Tripp B. W. (1977) Hydrocarbons in western
North Atlantic surface sediments. Geochim. Cosmochim. Acta 41,
1627–1641.
France-Lanord C. and Derry L. A. (1994) ␦13C of organic carbon in the
Bengal Fan; source evolution and transport of C3 and C4 plant carbon
to marine sediments. Geochim. Cosmochim. Acta 58, 4809 – 4814.
Gagosian R. B., Peltzer E. T., and Zafiriou O. C. (1981) Atmospheric
transport of continentally-derived lipids to the tropical North Pacific.
Nature 291, 312–314.
Gagosian R. B. and Peltzer E. T. (1986) The importance of atmospheric
input of terrestrial organic material to deep sea sediments. In Adv.
Org. Geochem. 1985 (eds. D. Leythaeuser and J. Rullkotter), Org.
Geochem. 10, 661– 669.
Giresse P., Maley J., and Brenac P. (1994) Late Quaternary palaeoenvironments in the Lake Barombi Mbo (West Cameroon) deduced
from pollen and carbon isotopes of organic matter. Palaeogeogr.
Palaeoclimatol. Palaeoecol. 107, 65–78.
Gasse F., Stabell B., Fournatier E., and Iperen Y. V. (1989) Freshwater
diatom influx in intertropical Atlantic: Relationships with continental
records from Africa. Quat. Res. 32, 229 –243.
C4 plant input into North East Atlantic
Goñi M. A., Ruttenberg K. C., and Eglinton T. I. (1997) Sources and
contribution of terrigenous organic carbon to surface sediments in
the Gulf of Mexico. Nature 389, 275–278.
Guillet B., Faivre P., Mariotti A., and Khobzi J. (1988) The 14C dates
and 13C/12C ratios of soil organic matter as a means of studying the
past vegetation in intertropical regions; examples from Columbia
(South America). Palaeogeogr. Palaeoclimatol. Palaeoecol. 65, 51–
58.
Hayes J. M., Freeman K. H., Popp B. N., and Hoham C. H. (1990)
Compound-specific isotopic analysis: A novel tool for reconstruction
of ancient biogeochemical processes. Org. Geochem. 16, 1115–
1128.
Hooghiemstra H., Agwu C. O. C., and Beug H. J. (1986) Pollen and
spore distribution in recent marine sediments: A record of NWAfrican seasonal wind patterns and vegetation belts. “Meteor” Forschung-Ergebnisse C40, 87–135.
Hooghiemstra H. (1988) Palynological records from northwest African
marine sediments: A general outline of the interpretation of the
pollen signal. Phil. Trans. R. Soc. Lond. B. 318, 431– 449.
Huang Y., Collister J. W., Chester J., and Eglinton G. (1993) Molecular
and ␦13C mapping of aeolian input of organic compounds into
marine sediments in the Northeastern Atlantic. In Org. Geochem.
(ed. K. Øygard), 523–528, Falch Hurtigtrykk, Oslo.
Huang Y., Lockheart M., Collister J. W., and Eglinton G. (1995)
Molecular and isotopic biogeochemistry of the Miocene Clarkia
Formation: Hydrocarbons and Alcohols. Org. Geochem. 23, 785– 801.
Huang Y., Street-Perrott F. A., Perrott F. A., Metzger, P., and Eglinton
G. (1999a) Glacial-interglacial environmental changes inferred from
the molecular and compound-specific ␦13C analyses of sediments
from Sacred Lake, Mt. Kenya. Geochim. Cosmochim. Acta 63,
1383–1404.
Huang Y., Freeman K. H., Eglinton T. I., and Street-Perrott F. A.
(1999b) ␦13C analyses of individual lignin phenols in the lacustrine
environment: A novel proxy for deciphering the past terrestrial
vegetation changes. Geol. 27, 471– 474.
Jolly D. and Haxeltine A. (1997) Effect of low glacial atmospheric CO2
on tropical African montane vegetation. Science 276, 786 –788 (1997).
Krishnamurthy R. V. and DeNiro M. (1982) Isotope evidence for
Pleistocene climate changes in Kashmir, India. Nature 298, 640 –
641.
Kuypers M. M. M., Pancost R. D., and Sinninghe Damste Jaap S.
(1999) A large and abrupt fall in atmospheric CO2 concentration
during Cretaceous times Nature 399, 342–345.
Ludwig W., Probst J.-L., and Kempe S. (1996) Predicting the oceanic
input of organic carbon by continental erosion. Global Biogeochem.
Cyc. 10, 23– 41.
Meyers P. A. and Ishiwatari R. (1993) Lacustrine organic geochemistry—an overview of indicators of organic matter sources and diagenesis in lake sediments. Org. Geochem. 20, 867–900.
O’Leary M. H. (1981) Carbon isotope fractionation in plants. Phytochem. 20, 553–568.
Pagani M., Freeman K. H., and Arthur M. A. (1999) Late Miocene
atmospheric CO2 concentrations and the expansion of C4 grasses.
Science 285, 876 – 879.
Petit-Marie N. (1991) Paléoenvironnements du Sahara. CNRS.
Pokras E. M. (1991) Source areas and transport mechanisms for fresh-
3513
water and brackish-water diatoms deposited in pelagic sediments of
the equatorial Atlantic. Quat. Res. 35, 144 –156.
Poynter J. G., Farrimond P., Brassel S. C., and Eglinton G. (1989)
Aeolian-derived higher-plant lipids in the marine sedimentary recordL Links with palaeoclimate. In Palaeoclimatology and palaeometeorology: modern and past patterns of global atmosphere transport (eds. M. Leinen and M. Sarnthein), pp. 435– 462. Kluwer
Dordrecht.
Pye K. (1987) Aeolian dust and dust deposits. Academic Press, Cambridge.
Raghavendra A. S. and Das V. S. R. (1978) The Occurrence of
C4-Photosynthesis: A Supplementary List of C4 Plants Reported
During Late 1974 to Mid 1977. Photosynthetica 12, 200 –208.
Raynaud D., Jouzel J., Barnola J. M., Chappellaz J., Delmas R. J., and
Lorius C. (1993) The ice record of greenhouse gases. Science 259,
926 –934.
Rieley G., Collister J. W., Stern B., and Eglinton G. (1993) Gas
Chromatography-Isotope Ratio Mass Spectrometry of leaf wax nalkanes from plants of differing carbon dioxide metabolisms. Rapid
Comm. in Mass Spec. 7, 488 – 491.
Sarnthein M., Tetzlaff G., Koopmann B., Wolter K., and Pflaumann U.
(1981) Glacial and interglacial wind regimes over the eastern subtropical Atlantic and North-West Africa. Nature 293, 193–196.
Simoneit B. R. T., Chester R., and Eglinton G. (1977) Biogenic lipids
in particulates from the lower atmosphere over the eastern Atlantic.
Nature 267, 682– 685.
Simoneit B. R. T. (1997) Compound-specific carbon isotope analyses
of individual long-chain alkanes and alkanoic acids in Harmattan
aerosols. Atmosph. Env. 31, 2225–2233.
Street-Perrott F. A., Huang Y., Perrott A., Eglinton, G., Baker P., Ben
Khelifa L., Harkness D. D., and Olago D. (1997) The impact of
lower atmospheric CO2 on tropical mountain ecosystems. Science
278, 1422–1426.
Talbot M. R. and Johannessen T. (1992) A high resolution palaeoclimatic record for the last 27,500 years in tropical West Africa from
the carbon and nitrogen isotopic composition of lacustrine organic
matter. Earth Planet. Sci. Lett. 110, 23–37.
Teeri J. A. and Stowe L. G. (1976) Climatic patterns and the distribution of C4 grasses in North America. Oecologia 23, 1–12.
Tulloch A. P. (1976) Chemistry of waxes of higher plants. In Chemistry
and Biochemistry of Natural Waxes (ed. P. E. Kolattukudy), pp.
236 –252. Elsevier, New York.
Uzaki M., Yamada K., and Ishiwatari R. (1993) Carbon isotope evidence for oil-pollution in long chain normal alkanes in Tokyo Bay
sediments. Geochem. J. 27, 385–389.
Vogel J. C. and Fuls A. (1978) The Geographical Distribution of Kranz
grasses in South Africa. S. Afr. J. Sci. 74, 209 –215.
Westerhausen L., Poynter J., Eglinton G., Erlenkeuser H., and Sarnthein M. (1992) Marine and terrigenous origin of organic matter in
modern sediments of the equatorial east Atlantic: The ␦13C and
molecular record. Deep Sea Res. 40, 1087–1121.
White F. (1983) The vegetation of Africa. Unesco.
Winter K. and Smith J. A. C. (1996) Crassulacean Acid Metabolism:
biochemistry, ecophysiology and evolution. Ecolog. Stud. 114,
Springer Berlin Heidelberg, 449 p.