The carbon isotopic record of the C37:2 alkenone in the South Atlantic: Last Glacial Maximum (LGM) vs. Holocene

Palaeogeography, Palaeoclimatology, Palaeoecology 221 (2005) 123 – 140 www.elsevier.com/locate/palaeo

The carbon isotopic record of the C37:2 alkenone in the South Atlantic: Last Glacial Maximum (LGM) vs. Holocene Albert Benthiena,T, Nils Andersenb, Sonja Schultec, Peter J. Mqllerd, Ralph R. Schneiderd,e, Gerold Weferd a

Alfred-Wegener-Institut fu¨r Polar- und Meeresforschung, 27515 Bremerhaven, Germany b Christian-Albrechts-Universita¨t zu Kiel, Leibniz-Labor, 24118 Kiel, Germany c Institut fu¨r Chemie und Biologie des Meeres (ICBM), Universita¨t Oldenburg, 26111 Oldenburg, Germany d DFG Forschungszentrum Ozeanra¨nder der Universita¨t Bremen, 28334 Bremen, Germany e Universite´ de Bordeaux, EPOC, Avenue des Faculte´s, 33405 Talence, France Received 29 March 2004; received in revised form 30 December 2004; accepted 18 February 2005

Abstract The carbon isotopic fractionation (e p) of the C37:2 alkenone was analysed for 19 South Atlantic sediment samples from the Last Glacial Maximum (LGM). Our study covers the equatorial and subtropical ocean including the coastal upwelling regions off South Africa, the equatorial upwelling, and the oligotrophic western basins. The results were compared to the Holocene e p records from the respective core locations (Andersen, N., Mu¨ller, P.J., Kirst, G., Schneider, R.R., 1999. Alkenone y13C as a proxy for past PCO2 in surface waters: results from the Late Quaternary Angola Current. In: Fischer, G., Wefer, G. (Eds.), Use of Proxies in Paleoceanography: Examples from the South Atlantic. Springer, Berlin, pp. 469–488; Benthien, A., Andersen, N., Schulte, S., Mu¨ller, P.J., Schneider, R.R., Wefer, G., 2002. Carbon isotopic composition of the C37:2 alkenone in core–top sediments of the South Atlantic Ocean: Effects of CO2 and nutrient concentrations. Glob. Biogeochem. Cycles 16, 10.1029/ 2001GB001433). Generally, alkenone e p was lower during the LGM compared to the Holocene. Higher glacial e p values were only found in sediments from the Angola Basin and in one sample from the eastern crest of the Walvis Ridge. Considering present understanding of LGM–Holocene changes in surface-water conditions (i.e. nutrient level, primary productivity, phytoplankton assemblages), the observed glacial/interglacial difference in e p indicates that multiple factors controlled the isotopic fractionation in alkenone producing algae depending on the regional setting. In the oligotrophic areas of the South Atlantic the lower than Holocene glacial e p values can be partly explained with a decrease in surface-water PCO2 during the LGM. In contrast, the Holocene to LGM decrease in e p values in the coastal upwelling areas as well as in the eastern tropical Atlantic most probably reflects much higher glacial haptophyte growth rates induced by an increase in surface-water nutrient concentrations. The exceptional opposite trend of the e p differences in the Angola Basin can be explained by a shift in the phytoplankton community towards a greater dominance of diatoms under glacial conditions, thus leaving less nutrients

T Corresponding author. Tel.: +49 471 4831 1039; fax: +49 471 4831 1149. E-mail address: [email protected] (A. Benthien). 0031-0182/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2005.02.008

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available for haptophytes. In this way, the isotopic record of alkenones indicate lower haptophyte growth rates during the LGM although other palaeoceanographic proxies point to enhanced productivity and higher nutrient levels. D 2005 Elsevier B.V. All rights reserved. Keywords: Alkenones; Stable carbon isotopes; Emiliania huxleyi; Growth rates; Carbon dioxide; Nutrients

1. Introduction Ice core studies demonstrated that the CO2 partial pressure in the atmosphere ( pCO2) was about 80–100 ppmv lower during glacial times than during interglacial periods (Barnola et al., 1987; Petit et al., 1999). Although there is yet no consensus as to the causes, it is agreed that these glacial–interglacial fluctuations are linked to changes in oceanic circulation, sea water carbonate system, and biological productivity (e.g., Broecker, 1982; Knox and McElroy, 1984; Sarmiento and Toggweiler, 1984; Boyle, 1988; Keir, 1990; Martin, 1990; Francois et al., 1997; Archer et al., 2000; Sigman and Boyle, 2000). One possible mechanism might be the photosynthetic uptake of CO2 in oceanic surface-waters and export to the deep ocean by sinking of phytoplankton particles (biological pump). Thus, the reconstruction of ancient surface-water conditions is crucial to understand the processes responsible for the CO2 exchange between ocean and atmosphere. The carbon isotopic composition (y13C) of phytoplankton organic matter can provide important information about the marine environment during carbon fixation. Based on a variety of experimental and field observations it has been suggested that the availability of aqueous CO2 (CO2aq) is a major factor controlling the y13C of phytoplankton (e.g., Popp et al., 1989; Rau et al., 1989; Freeman and Hayes, 1992). However, the results of subsequent laboratory and field experiments as well as model results indicate that a number of additional factors such as nutrient availability, algal growth rate, carbon acquisition mechanisms, irradiance, and cell size and geometry are also important factors regulating the carbon isotopic fractionation (e p) in marine algae (e.g., Francois et al., 1993; Goericke et al., 1994; Laws et al., 1995; Rau et al., 1996; Bidigare et al., 1997; Popp et al., 1998b, 1999; Burkhardt et al., 1999; Keller and Morel, 1999; Riebesell et al., 2000; Gervais and Riebesell, 2001; Rost et al., 2002).

The use of the isotopic composition of the C37 alkenone provides a way to reduce the number of factors influencing the y13C signal (Jasper and Hayes, 1990; Popp et al., 1998b; Andersen et al., 1999). Alkenones are exclusively produced by certain haptophyte algae, especially by Emiliania huxleyi and the closely related species Gephyrocapsa oceanica (Volkman et al., 1980; Marlowe et al., 1990; Conte et al., 1995; Sawada et al., 1996). Both species have a limited range in cell size and geometry. In addition, it is suggested that E. huxleyi preferentially uses CO2aq as the substrate for photosynthesis (Sikes et al., 1980). In contrast to many other phytoplankton species E. huxleyi has probably not developed an efficient carbon-concentrating mechanism (CCM) and thus is more sensitive to CO2 (e.g., Raven and Johnston, 1991; Badger et al., 1998; Rost et al., 2002). The present paper expands on a recent calibration study by Benthien et al. (2002) based on 29 core–top samples from different oceanic regions in the equatorial and subtropical South Atlantic. The results of this calibration indicated that the sedimentary record of the C37:2 alkenone reflects surface-water nutrient concentrations rather than the concentration of CO2aq. This is in agreement with findings based on laboratory experiments and field studies on particulate organic matter (Bidigare et al., 1997, 1999; Popp et al., 1999; Laws et al., 2001). Despite these limitations, Pagani et al. (2002) recently demonstrated in a study based on Holocene-aged, alkenone-based CO2aq estimates from the central Pacific Ocean that the use of alkenone y13C is a robust approach in the reconstruction of palaeoPCO2 if nutrient concentrations can be reasonably estimated. On the other hand, Schulte et al. (2004) have shown that the alkenone isotopic record in a deep-sea sediment core off Angola can be used as a proxy to estimate surface-water nutrient levels for the past 200,000 years. The two studies mentioned above were carried out in different oceanographic regions with different levels in primary productivity, nutrient regimes, and phytoplankton assemblages. For the use

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of alkenone y13C as a proxy for either palaeo-PCO2 or nutrient reconstructions, it is therefore very important to know how the various surface-water conditions in different oceanographic regions affect the carbon isotopic signal of alkenones over geological time. Here, we present the carbon isotopic record of the C37:2 alkenone from various regions in the South Atlantic during the Last Glacial Maximum (LGM). We compare the glacial values with those from the respective core– top sediments (Andersen et al., 1999; Benthien et al., 2002) and discuss the different environmental factors which can account for the observed Holocene to LGM shifts in the carbon isotopic fractionation e p.

2. Oceanographic setting The study area encompasses the equatorial and subtropical South Atlantic Ocean between 28N and 368S. The hydrography of surface and subsurface waters of the South Atlantic has been described in detail by Peterson and Stramma (1991) and is summarised in Fig. 1. The dominating feature of the

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surface-water circulation is a subtropical anticyclonic gyre. The eastern limb of this gyre is formed by the Benguela Current which is fed primarily by the South Atlantic Current (SAC) and the Agulhas Current (AGC). Near 308S, the main part of the Benguela Current separates from the coast and forms the northwestward flowing Benguela Ocean Current (BOC). Along the continental margin, the coastal branch of the Benguela Current (Benguela Coastal Current, BCC) transports cold and nutrient-rich water northward across the Walvis Ridge. At about 158N, it converges with the southward flowing Angola Current (AC), forming the NW–SE oriented Angola–Benguela Front (ABF). In the area of the BCC, the prevailing southerly and southeasterly winds force the coastal upwelling of cold, nutrient- and CO2-rich South Atlantic Central Water (SACW). Filaments of these water masses extend well offshore where they mix with waters from the BOC. Further offshore, the BOC feeds into the broad South Equatorial Current (SEC) that flows northwestward towards Brazil. Near 108S, the South Equatorial Current bifurcates into the northward-

Fig. 1. Bathymetric map of the South Atlantic indicating the locations of the surface sediments analysed in this study (Table 1). The general surface-water circulation pattern is modified after Peterson and Stramma (1991). Core locations were grouped in accordance with the biogeochemical provinces of Longhurst (1995) and marked by different symbols: Angola Coastal Current (ANG), Benguela Coastal Current (BENG), South Atlantic Tropical Gyre (SATL), Eastern Tropical Atlantic (ETRA), Western Tropical Atlantic (WTRA), Brazil Current (BRAZ).

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flowing North Brazil Current (NBC) and the southward-flowing Brazil Current (BC). In the equatorial area, the surface-circulation pattern is marked by a complex array of westward currents and eastward countercurrents. The SEC is separated by the eastward moving South Equatorial Countercurrent (SECC). The northern branch flows at 28 to 48S and interacts with the seasonally appearing North Equatorial Countercurrent (NECC). The convergence of these water masses results in downwelling of surface-waters. Between 58N and 58S, the Equatorial Undercurrent (EUC) flows eastward parallel to the equator at a depth of 50 to 125 m. The contact zone of the SEC and the EUC at 08 to 28S forms the equatorial divergence where upwelling of colder water masses from the thermocline occurs. Equatorial upwelling is most intense throughout boreal summer when the SEtrades are strongest. During this time the velocity of the South Equatorial Current is enhanced. The associated transport of warm tropical surface-waters results in deepening of the thermocline in the western equatorial Atlantic whereas in the eastern equatorial Atlantic the thermocline rises to a shallow depth of 20 to 30 m. The Angola Current is the southeasterly limb of the SEC which transports warm equatorial water with high salinity and low nutrient content into the Angola Basin. The complex interaction of SECC, AC, and BOC form gyres and fronts which cause oceanic upwelling. This, together with the shallow thermocline and nutricline during boreal summer, leads to a transport of nutrient-rich subsurface waters into the euphotic zone. Nearshore upwelling occurs by the piling up of the eastward-flowing Equatorial Undercurrent at the continental margin off Congo and Gabon (Servain et al., 1982). In the area off the Congo River mouth, upwelling is reinforced by the rapid outflow of freshwater which can be detected as far as 800 km offshore (van Bennekom and Berger, 1984).

3. Materials and methods 3.1. Sediment samples The sediment samples were collected using a gravity corer during several cruises of R/V Meteor,

R/V Sonne, and R/V Victor Hensen (Table 1; Fig. 1). After core recovery, 10 ml subsamples were taken on board at 5-cm intervals using plastic syringes and stored at a temperature of 4 8C. In the home laboratory, all samples were freeze dried and ground in an agate mortar. The stratigraphy of the sediment cores is based on foraminiferal y18O. The samples for alkenone analyses were taken within the core section defined as LGM (19–23 kyr BP, http://www.pangaea.de/Institutes/ GeoB/Cores/LGM.html). In this context it is important to note that recently radiocarbon age offsets were detected between foraminifer tests and alkenones in sediments from the Benguela upwelling area (Mollenhauer et al., 2003). The age differences observed are between 1000 and 4500 years and in all cases alkenones were older than the foraminifer tests at the same core depth. Consequently, it is possible that in the present study the alkenone data from the Benguela upwelling region do not represent the LGM as defined by foraminiferal y18O but certainly reflect glacial environmental conditions. 3.2. Alkenone analysis Alkenones were extracted from 0.25 to 15 g aliquots of freeze-dried and homogenised sediment. We used UP 200H ultrasonic disrupter probes (200 W, amplitude 0.5, pulse 0.5) and three successively less polar mixtures of methanol and dichloromethane (CH3OH, CH3OH:CH2Cl2 1:1, CH2Cl2), each for 3 min. The three extracts were combined, desalted, dried over Na2SO4, and concentrated under N2. Additionally, the extracts were purified by passing them over a silica gel cartridge (Varian Bond Elut; 1 cm3/100 mg) and then saponified to remove possibly interfering esters. For the saponification, we added 0.3 ml of 0.1 M KOH in CH3OH:H2O (90:10) to the extract which was then heated at 80 8C in a capped vial for 2 h. After cooling, the alkenone-containing fraction was obtained by partitioning into hexane, evaporated, and finally taken up in 15–150 Al of a 1:1 CH3OH:CH2Cl2 mixture. Alkenone unsaturation ratios were determined using a HP 5890A gas chromatograph either equipped with a HP Ultra 1 fused silica column (50 m
0.32 mm
0.52 Am), split/splitless injection (1:10 split modus) or fitted with an on-column injector to a

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Table 1 Station data and analytical results for LGM sediments of the equatorial and South Atlantic Core (GeoB)

Cruise

Latitude

Angola Coastal Current (ANG) 1008-6 M6/6 6834.9S 1016-3 M6/6 11846.2S

Longitude

Water depth (m)

SST (8C) (LGM)

y13C37:2 (x) (LGM)

y13C37:2 (x) (Holocene)

e p (x) (LGM)

e p (x) (Holocene)

De p LGM–Hol (x)

10819.1E 11840.9E

3124 3411

22.3 21.7

24.9 24.3

23.1a 23.5a

14.0 13.3

12.9 13.2

+1.1 +0.1

Benguela Coastal Current (BENG) 1023-5 M6/6 17809.4S (1703-5) 1028-4 M6/6 20806.2S 1706-1 M20/2 19833.7S 1710-2 M20/2 23825.9S 1711-5 M20/2 23818.9S 1712-2 M20/2 23815.4S 3603-1 M34/1 35807.5S

11800.7E

1978

18.1

23.8

23.8a,b

12.3

13.0b

0.7

9811.1E 11810.5E 11841.9E 12822.6E 12848.5E 17832.6E

2209 980 2987 1967 998 2840

16.6 17.0 14.9 15.8 17.5 18.1

24.3 22.8 22.7 22.5 22.6 23.6

22.9 23.5a 23.0 23.9a 22.5 23.6

12.7 11.2 10.8 10.7 11.1 12.2

12.1 12.5 12.1 12.9 11.5 12.8

+0.6 1.2 1.3 2.2 0.4 0.6

South Atlantic 1032-2 1214-2 1413-1

6802.2E 7814.4E 9827.3W

2505 3210 3789

18.0 16.0 22.1

23.1 23.8 24.1

24.3 23.4 23.6

11.6 12.1 13.1

13.5 12.7 13.4

1.9 0.6 0.2

(ETRA) 3828.5S 1839.9S 3848.9S 8840.6S

7836.0W 12825.7W 14853.8W 11850.6W

4033 3225 3984 3161

23.1 23.0 24.2 24.1

22.7 23.2 23.1 24.1

23.6 23.4 23.6 25.7a

11.8 12.3 12.4 13.4

13.5 13.2 13.5 15.6

1.7 0.9 1.1 2.2

Western Tropical Atlantic (WTRA) 3117-3 JOPS II 4817.7S

37805.5W

800

25.7

24.1

24.2

13.5

14.3

0.8

Brazil Current (BRAZ) 2109-3 M23/2

45852.9W

2504

22.9

24.1

24.0

13.2

13.6

0.4

Tropical Gyre (SATL) M6/6 22854.9S M12/1 24841.4S M16/1 15840.8S

Eastern Tropical Atlantic 1041-1 M6/6 1105-3 M9/4 1117-3 M9/4 1903-1 SO84

27854.7S

KV KV Sea-surface temperature (SST) was obtained from the U37 index using the E. huxleyi calibration (U37 = 0.034T + 0.039) of Prahl et al. (1988). 13 Also shown are the Holocene y C and e p values from the respective core–top sediments (cf. Benthien et al., 2002). a Values are from Andersen et al. (1999). b Holocene values are taken from core GeoB 1703-5 (17827.1S; 11801.0E).

DB5ms (60 m
0.32 mm
0.1 Am). The oven temperature was programmed from 50 to 150 8C at 30 8C min1, from 150 to 230 8C at 8 8C min1, and from 230 to 320 8C at 6 8C min1 with a 45 min hold at 320 8C for split/splitless injection. For on-column injection, the oven temperature was programmed from 40 to 200 8C at 15 8C min1 (5 min initial time), from 200 to 250 8C at 5 8C min1 and from 250 to 300 8C at 3 8C min1 with a 35 min hold at 300 8C. The carbon isotopic analyses of the C37:2 alkenone were performed using a gas chromatograph (HP 5890A) connected via a combustion interface to a Finnigan MAT 252 mass spectrometer (irm-GCMS).

The GC was equipped with an on-column injector, a 2.5 m retention gap, and a fused silica column (SGEBPX 5 and Optima 1; each 50 m
0.32 mm
0.52 Am). The temperature program was 50–150 8C at 30 8C min1 (with 5 min initial time), then 150–230 8C at 8 8C min1, and 230–320 8C at 6 8C min1 (isothermal at 320 8C for 48 min). The isotopic composition of the C37:2 alkenone was calculated relative to Peedee belemnite (PDB) by comparison with co-injected nalkane standards (n-C34, n-C36, n-C37, n-C38) and a laboratory internal standard gas (CO2). Generally, each sample was measured four times revealing an analytical uncertainty less than 0.3x.

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3.3. Calculation of e p Calculation of the isotopic fractionation (e p) associated with photosynthetic carbon fixation requires knowledge of the carbon isotopic compositions of dissolved CO2 (d d) and haptophyte biomass (d p). Measurements on benthic and planktic foraminifera from deep-sea sediments suggest that the average y13C in ocean ACO2 (d ACO2) was between 0.2 and 0.5x lower during the LGM than it was during the Holocene (Curry and Crowley, 1987; Curry et al., 1988; Duplessy et al., 1988; Bickert and Wefer, 1996). In order to estimate the LGM– Holocene difference in d ACO2 for the surface-waters of the South Atlantic, we used the carbon isotopic record of planktonic foraminifera (Globigerinoides ruber p., Globigerinoides ruber wh., Globorotalia inflata) from various cores presented in this study. The results suggest an average shift in y13C of about 0.4x between LGM and Holocene. Assuming a surface-water mean value of 2.5x for pre-industrial d ACO2 (Kroopnick, 1985), we thus applied a value of 2.1x for glacial d ACO2. In this context it is important to point out that the isotopic composition of ACO2 may vary regionally. This holds especially true in regions with upwelling of nutrient- and CO2rich subsurface water. The possible effect on the alkenone isotopic fractionation is discussed in further detail in Section 5.3. The carbon isotopic composition of CO2aq was calculated from d ACO2 following the equation from Rau et al. (1996) based on Mook et al. (1974): dd ¼ dRCO2 þ 23:644 

9701:5 T

ð1Þ

where T is the temperature in Kelvin. We used alkenone temperatures determined according to the KV widely used U37 calibration from Prahl et al. (1988; Table 1). The isotopic composition of the phytoplankton biomass (d p) was derived by correcting for the isotopic fractionation e alkenone between the C37:2 alkenone and the bulk biomass of alkenone-producing organisms. We took 4.2x as the value for e alkenone (Popp et al., 1998a):   dC37:2 dp ¼ dC37:2 þ ealkenone 1 þ : ð2Þ 1000

The isotopic fractionation (e p) associated with the photosynthetic fixation of carbon was calculated using the following equation (Freeman and Hayes, 1992):   dd þ 1000 ep ¼  1 1000: ð3Þ dp þ 1000

4. Results The alkenone temperatures, y13C values, and the calculated isotopic fractionation e p of the LGM are given in Table 1 together with the Holocene values of the respective core sites published by Andersen et al. (1999) and Benthien et al. (2002). In order to simplify the data evaluation, we separated the sites into six groups according to the biogeochemical provinces defined by Longhurst (1995; see Fig. 1). These provinces are the Angola and Benguela Coastal Currents (ANG and BENG), the South Atlantic Tropical Gyre (SATL), the Eastern and Western Tropical Atlantic (ETRA and WTRA), and the Brazil Current region (BRAZ). The measured alkenone y13C values range from 24.9x to 22.5x and the resulting isotopic fractionation range from 14.0x to 10.7x, respectively (Figs. 2 and 3a). Relatively high y13C and low e p values were found in the upwelling regions of the Benguela Coastal Current as well as at the sites GeoB 1032 (SATL), GeoB 1214 (SATL) and GeoB 1117 (ETRA). By contrast, the regions of the Angola Coastal Current, the Western Tropical Atlantic and the Brazil Current as well as core GeoB 1413 (SATL) and GeoB 1903 (ETRA) reveal relatively low y13C and high e p values. Intermediate values were found at the BENG province sites GeoB 1023, GeoB 1028, and GeoB 3603 as well as GeoB 1105 and GeoB 1117 (both ETRA). The general distribution pattern of the glacial e p signal with lower values in eutrophic and higher values in oligotrophic regions resembles the Holocene pattern (Fig. 3a and b; Andersen et al., 1999; Benthien et al., 2002). Considerable deviations were found in the region of the Angola Coastal Current and at GeoB 1028 (BENG) located at the Walvis Ridge. These differences are also illustrated comparing LGM to Holocene variations in e p (De p LGM–Hol; Fig. 4). Generally, the glacial values are lower relative to the

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Fig. 2. Distribution of alkenone y13C data for the Last Glacial Maximum. The values are expressed in x PDB.

Holocene (De p LGM–Hol: 0.4x to 2.2x). Exceptions to this are again the two sites of the Angola Coastal Current region (GeoB 1008, GeoB 1016), and site GeoB 1028. At these core locations, the e p values were higher during the LGM compared to the Holocene (De p LGM–Hol: +0.1x to +1.1x).

5. Discussion 5.1. Background: effects on carbon isotopic fractionation in alkenone-producing haptophytes The magnitude of isotopic fractionation by marine phytoplankton is a function of the isotope effects associated with (1) the fluxes of CO2 into and out of the cell and (2) enzymatic carbon fixation. Following the results of laboratory and field studies, several models of photosynthetic 13C fractionation have been developed to describe the interaction of these effects and how they influence e p (for a review see Laws et al., 2001). While earlier models allow only one carbon source (e.g., Sharkey and Berry, 1985) or consider only diffusive CO2 transport into the cell (e.g., Francois et al., 1993;

Jasper et al., 1994; Laws et al., 1995; Rau et al., 1996) more recent considerations allow also for active uptake of either CO2 and HCO3 (Burkhardt et al., 1999; Keller and Morel, 1999). The discussion about active CO2 uptake and/or bicarbonate utilization by coccolithophores is still controversial (e.g., Paasche, 2002). In contrast to many other phytoplankton species (e.g., diatoms), however, the coccolithophorid Emiliania huxleyi appears to employ an inefficient CO2 concentrating mechanism (CCM) and thus relies at least partly on diffusive CO2 uptake (Raven and Johnston, 1991; Badger et al., 1998; Rost et al., 2002). For pure diffusive CO2 uptake the model of Keller and Morel (1999) reduces to the model proposed by Jasper et al. (1994) and to a simplified version of the model of Rau et al. (1996), respectively which both predict the isotopic fractionation e p to be a linear function of b/[CO2aq]: ep ¼ ef  

b  CO2aq

ð4Þ

where e f is the enzymatic fractionation associated with photosynthetic carbon fixation and b is the sum of

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Fig. 3. Distribution of the alkenone based e p values for (a) the Last Glacial Maximum and (b) the Holocene. The values are expressed in x PDB.

physiological factors influencing the carbon isotopic discrimination. According to this, increasing CO2aq concentration causes higher e p values, whereas an increasing b-value lowers e p and vice versa.

As previously mentioned, the results of laboratory experiments, field studies, and modelling showed that the variable b reflects primarily the intracellular carbon demand which depends on microalgal growth

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Fig. 4. The difference of e p between LGM and Holocene (De p

as well as cell size/geometry and cell membrane permeability (Rau et al., 1996; Popp et al., 1999). Since alkenones are known to be produced mainly by Emiliania huxleyi and the closely related Gephyrocapsa oceanica, variations in cell size, shape and membrane properties are assumed to be relatively small and thus only have a minor effect on the carbon isotopic composition (Popp et al., 1998b, 1999). Bidigare et al. (1997) demonstrated that in modern alkenone-producing haptophytes most variations in e p result from variations in growth rate (l) which were correlated to the concentration of soluble reactive phosphate (SRP). More recently, Bidigare et al. (1999) showed theoretically that during the IronEx II iron fertilisation experiment the growth rate was 7 times more important than CO2aq in causing variations in e p of marine phytoplankton. These findings are in agreement with the isotopic record of the sedimentary C37:2 alkenone. In a recent comparison of Pacific Ocean water column [CO2aq] with sedimentary e p37:2 values, Pagani et al. (2002) found that the magnitude of the isotopic alkenone signal is predominantly controlled by nutrient concentration and not [CO2aq]. Similar results were found in a calibration study based on core–top sediments

LGM–Hol).

131

The values are expressed in x PDB.

from the equatorial and South Atlantic (Benthien et al., 2002) and in a deep-sea sediment core recovered off Angola (Schulte et al., 2004). Therefore, LGM to Holocene variations in the sedimentary e p signal should be interpreted primarily as variations in haptophyte growth rate controlled by the availability of nutrients in surface-waters. In this context two considerations should be taken into account. (1) The growth-limiting nutrients in haptophyte algae are not known. As pointed out by Bidigare et al. (1997), it is unlikely that growth rates were directly controlled by phosphate levels since E. huxleyi has a low phosphorus requirement (see also Riegman et al., 2000). Instead, the authors suggested that essential micronutrients such as iron, zinc and cobalt might be the growth-rate limiting factors and that the concentrations of these micronutrients are correlated to [PO4]. Moreover, variations in other environmental conditions such as irradiance or changes in the growth-limiting resource may also affect the carbon isotopic fractionation in microalgae (e.g., Beardall et al., 1998; Riebesell et al., 2000; Rost et al., 2002). (2) In many oceanic systems the surface-water concentrations of dissolved CO2 and nutrients are positively correlated. As shown in Eq. (4), changes in

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nutrient-controlled growth (here expressed as b) and CO2aq concentration would therefore affect e p in opposite directions. At the moment, however, there is no established proxy to separate the effects of these parameters quantitatively (Popp et al., 1999). Thus, the variation of one parameter can only be quantified if the other parameter remained constant which is unlikely in a natural system. A promising approach to solve this problem might be the use of the Sr/Ca ratio in coccoliths as an independent indicator of algal growth rate (Stoll and Schrag, 2000; Stoll et al., 2002, Rickaby et al., 2002). However, the applicability of this new proxy has still to be investigated in further studies. In this paper we present y13C values of sedimentary alkenones from different oceanic systems including the low productivity areas in the western tropical Atlantic and the subtropical gyre as well as the nutrient and CO2-rich upwelling areas off central and southwest Africa. Since in these regions the observed oceanographic LGM–Holocene variations are quite different as well we will discuss our data in the following on a regional scale for the respective biogeochemical provinces defined by Longhurst (1995; Fig. 1). 5.2. Oligotrophic regions The lower than Holocene e p values recorded in the glacial sediments from the regions of the SATL, WTRA, and BRAZ regions (negative De p LGM–Hol values; Fig. 4) are consistent with both a glacial increase in growth rates and a glacial decrease in [CO2aq] if we consider the relationship between e p, CO2aq, and the b-value as expressed in Eq. (4). Only few palaeoceanographic reconstructions from these low-productivity regions are available (e.g., Ru¨hlemann et al., 1996; Wefer et al., 1996; Mulitza et al., 1998; Mollenhauer et al., 2004). They are in general agreement with studies from other oligotrophic ocean areas, showing that palaeoproductivity and organic carbon accumulation rates remained constant or decreased slightly during the last glacial relative to the Holocene (Curry and Crowley, 1987; Sarnthein et al., 1988; Mix, 1989; Howard and Prell, 1994). A change in overall phytoplankton productivity does not necessarily involve a change in haptophyte productivity and vice versa. It has long been recognised that

a negative relationship exists between productivity and the relative contribution of small cells to the phytoplankton biomass (Malone, 1980; Chisholm, 1992). However, this general statement holds true mainly if data from very diverse environments are pooled (Chisholm, 1992). In a recent study of the ecology and biogeochemistry in the North and South Atlantic Subtropical Gyres, Maran˜o´n et al. (2003) have shown that despite a 20-fold range in productivity rates, the size structure of the phytoplankton assemblages did not change significantly, nor did the partitioning of C fixation between small and large photoautotrophs. It seems therefore very unlikely that the presumed small or even nonexistent changes in palaeoproductivity during the LGM relative to the Holocene caused large variations in the fractional contribution of haptophytes to the phytoplankton assemblages. More or less constant LGM to Holocene haptophyte productivity does not necessarily equate with small or nonexistent changes in haptophyte growth rates. The inference can only be drawn if one assumes that haptophyte standing stock, grazing pressure, and export rate also remained similar over time. Based on the available palaeoceanographic information for these regions this assumption can neither be confirmed nor ruled out. However, for the following considerations we presume minor LGM to Holocene variations in haptophyte growth rate. Hence, our observed variations in e p should mainly reflect a glacial decrease in [CO2aq]. In contrast to upwelling areas, oligotrophic openocean regions are generally characterised by a small air-to-sea disequilibrium with respect to CO2. Thus, the LGM–Holocene difference in surface-water PCO2 (DPCO2 LGM–Hol) should be close to the glacial 80– 100 Aatm decrease of atmospheric pCO2 (Barnola et al., 1987; Petit et al., 1999). Assuming that sitespecific haptophyte growth rates remained constant through time (b = const.), the estimation of DPCO2 LGM–Hol indicated by the variations in e p yield values between 17 Aatm at site GeoB 2109 and 54 Aatm at site GeoB 1032 (Table 2). The deviation from the expected decrease of 80–100 Aatm can be explained either by a decrease in growth rate (affecting b towards lower values) or by the overall uncertainty of the method. Error estimations for calculated PCO2 values provide uncertainties between 11 and 15% (Andersen et al., 1999; Pagani et al., 1999; Popp et al.,

A. Benthien et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 221 (2005) 123–140 Table 2 Estimated LGM–Holocene difference in surface-water PCO2 (DPCO2 LGM–Hol) Core PCO2 (GeoB) (Aatm)

Holocene

PCO2

LGM

(Aatm)

DPCO2 (Aatm)

South Atlantic Tropical Gyre (SATL) 1032-2 302 248 54 1214-2 308 258 50 1413-1 309 283 26

LGM–Hol

b-value (Holocene) (x Amol/l) 117 124 105

Western Tropical Atlantic (WTRA) 3117-3 322 288

35

94

Brazil Current (BRAZ) 2109-3 282 265

17

96

PCO2 values were calculated from [CO2aq] using Henry’s law: PCO2 = [CO2aq]/a where a is the solubility coefficient of CO2 (Weiss, 1974; Skirrow, 1975). LGM CO2aq concentrations were estimated from LGM e p values using a rearrangement of Eq. (4), with b LGM = b Holocene and e f = 25x (e.g., Bidigare et al., 1997). Holocene CO2aq concentrations and b-values were taken from Benthien et al. (2002).

1999). This could only explain part of the observed deviations although it is very improbable that uncertainties in the DPCO2 LGM–Hol estimations always result in an underestimation of the expected LGM– Holocene difference of 80 to 100 Aatm. Consequently, two other considerations are more likely: (1) the observed glacial e p values in the SATL, WTRA, and BRAZ regions indicate a smaller glacial decrease of surface-water PCO2 compared to atmospheric pCO2. This would imply a weak CO2 source during the glacial in these regions. (2) The smaller glacial e p values are the results of both a glacial decrease in CO2aq concentrations (affecting e p towards lower values) and a glacial decrease in haptophyte growth rates (effecting e p towards higher values). On the basis of available evidence, however, we cannot rule out either of these hypotheses. 5.3. Upwelling systems The LGM–Holocene e p variations in both the BENG and ETRA regions (De p LGM–Hol: 0.4 to 2.2x; Fig. 4) indicate enhanced growth rates of alkenone producing algae and thus higher nutrient levels and/or lower CO2aq concentrations during the Last Glacial Maximum. The former is in good

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agreement with our present knowledge of the Late Quaternary history of these areas. For the Benguela region, various palaeoceanographic reconstructions demonstrated increased upwelling and higher productivity during the last glacial relative to the Holocene based on changes in organic carbon content, alkenone temperatures, and phytoplankton assemblages (e.g., Summerhayes et al., 1995; Little et al., 1997; Kirst et al., 1999; Abrantes, 2000). In addition, the y15N records from three sediment cores off Namibia (GeoB 1710–1712) indicate that although productivity was higher, the enhanced upwelling of subsurface waters resulted in elevated nutrient concentrations at the surface (Lavik, 2001). In ocean surface-waters, the y13C of ACO2 is variable due to air–sea mixing processes, biological productivity and nutrient cycles. For this reason, the y13CACO2 is in general inversely correlated to the nutrient concentration (e.g., Broecker and MaierReimer, 1992). Increased upwelling and elevated surface-water nutrient levels during the LGM would therefore yield y13CACO2 values lower than the estimated glacial average of 2.1x (cf. Section 3.3). While these variations are generally small, this in turn would result in a lower glacial alkenone isotopic fractionation and consequently in an even more pronounced difference between LGM and Holocene e p values (more negative De pLGM–Hol). Enhanced productivity due to a trade-wind forced shoaling of the nutricline and an increase in upwelling during MIS 2 is also reported for the ETRA (e.g., McIntyre et al., 1989; Schneider et al., 1996; Wefer et al., 1996; Wolff et al., 1999). In core GeoB 1105 from the equatorial divergence (Fig. 1), the y15N record indicates that surface-water nutrients were less depleted during the last glacial than during the Holocene (G. Lavik, unpublished data). Therefore, the higher glacial nutrient availability in both areas, BENG and ETRA, may have caused enhanced haptophyte growth rates which in turn resulted in the observed lower alkenone isotopic fractionation. As already discussed in the previous section, variations in palaeoproductivity are not equivalent to changes in haptophyte productivity or growth rates. This holds particularly true in high productivity upwelling areas where haptophyte blooms generally follow diatom blooms (Kilham and Kilham, 1980; Harris, 1986; Westbroek et al., 1994). For example, an

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increase in productivity during the LGM due to enhanced upwelling and nutrient supply may have caused a shift in the phytoplankton community towards a higher fractional contribution of diatoms. Thus, although the annually integrated productivity and export fluxes were higher, haptophyte growth rates could have remained constant or were even lower during the LGM compared to the Holocene. A way to test whether there was a shift in the phytoplankton community is to examine the concentrations of different molecular biomarkers (Hinrichs et al., 1999). The authors demonstrated that at site GeoB 1710 off Namibia the ratio of alkenone to loliolide concentrations [loliolide/(loliolide + C37 alkenones)] shows similar or even higher values during the Holocene (~0.47) than during the LGM (~0.28– 0.46). Since loliolides are also derived from algae other than haptophytes, a lower ratio is interpreted to reflect periods when alkenone-producers were a more dominant group in algae communities whereas high values characterise situations where other algae such as diatoms became relatively more important (for a detailed discussion see Hinrichs et al., 1999). This clearly indicates that in the area of the Benguela Coastal upwelling algae other than haptophytes were the dominant primary producers, but the relative abundance of haptophytes increased during the last glacial compared to the Holocene. Higher productivity and the resulting increase in carbon export out of the photic zone might also have lowered the concentration of CO2 in surface-waters like it is assumed for the Southern Ocean (e.g., Sarmiento and Toggweiler, 1984; Martin, 1990; Moore et al., 2000). As already mentioned, a decrease in [CO2aq] has the same effect on the isotopic fractionation as an increase in growth rate. On the other hand, enhanced upwelling of CO2-rich subsurface water would probably increase [CO2aq] in the surface layer. This is presently observed in the upwelling regions of the Arabian Sea and the equatorial Pacific (Sabine et al., 1997; Feely et al., 1995, 1999). Despite of high productivity, these areas are characterised by high surface-water CO2 partial pressure ( PCO2). If we therefore assume that in the BENG and ETRA regions the interaction of enhanced upwelling and increased productivity during the LGM yielded higher or at least similar CO2-levels relative to the Holocene, the resulting effect on e p should be

opposite to the trend observed in our data. Consequently, we conclude that the observed lower glacial e p values in the BENG and ETRA regions are mainly a result of higher haptophyte growth rates during the LGM compared to the Holocene. In contrast to the general trend in the Benguela and equatorial upwelling areas, the sediments from the ANG region recorded higher (GeoB 1008) or at least similar (GeoB 1016) isotopic fractionation during the LGM, indicating lower haptophyte growth rates and/ or higher CO2aq concentration relative to the Holocene (positive De pLGM–Hol values; Fig. 4). Similar to the BENG and ETRA areas, however, thermocline shoaling, oceanic upwelling, and in turn productivity and surface-water nutrient concentrations in the ANG region are suggested to have been enhanced during the last glacial compared to the Holocene (e.g., Jansen and Van Iperen, 1991; Schneider et al., 1994, 1997; Rutsch et al., 1995; Holmes et al., 1997). In order to explain the unexpected record in the alkenone isotopic signal, two possible causes, individually or combined, can be assumed for the higher glacial e p: (1) The observed changes in e p at these core locations are mainly a function of CO2 supply which requires similar growth rates during the LGM and the Holocene. The higher glacial e p then indicates higher [CO2aq] which might be explained by enhanced upwelling of CO2-rich subsurface waters. Whether the glacial CO2aq levels in the ANG region were significantly enhanced is still a matter of debate. PalaeoPCO2 reconstructions at site GeoB 1016 based on the y13C record of bulk organic carbon (Mu¨ller et al., 1994) and alkenones (Andersen et al., 1999) show conflicting results. Whereas the bulk y13C signal indicates that the Angola Basin was a weaker CO2 source during the LGM, the alkenone isotopic signal yields similar PCO2 values for both the LGM and the Holocene which consequently implies a stronger glacial CO2 source due the lower atmospheric pCO2. However, in this context it should be mentioned that these former studies are subject to the same restrictions as the present one in terms of biological effects on isotopic fractionation (see Section 5.1).

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(2) The second possibility which can explain the observed De pLGM–Hol values in the ANG region is that enhanced productivity together with a change in phytoplankton community depleted the available nutrient pool for haptophytes. Such a shift might have caused lower haptophyte growth rates although annually integrated primary productivity and export fluxes were higher during the LGM. Again using additional biomarker records evidence is found for a shift in the phytoplankton community in the Angola Basin as indicated by changes in the ratio of alkenone to loliolide concentrations in Angola margin sediments (Hinrichs et al., 1999). At site GeoB 1016 the loliolide/(loliolide + C37 alkenones) ratio decreased significantly from a glacial value of about 0.26 to 0.08 during the Holocene. Compared with the biomarker results from the Benguela coastal upwelling this indicates a shift towards a dominant contribution of alkenone producers to organic carbon burial during the Holocene. Further evidence that the e p variations in the Angola Basin are related to a shift in the phytoplankton community is derived from the findings of Schneider et al. (1997). They reported significantly enhanced opal accumulation rates at site GeoB 1008 compared to GeoB 1016 during the last glacial. Since GeoB 1008 is situated off the Congo River, the enhanced opal production at this site is explained by increased fluvial silica input (Schneider et al., 1997). If silica was not a limiting factor diatoms would have been able to utilise more of the other available nutrients, thus leaving less nutrients available for the haptophytes. This might explain (1) why the alkenone e p changes indicate lower nutrient levels despite other palaeoceanographic proxies show higher nutrient levels and enhanced productivity during the LGM (e.g., Holmes et al., 1997; Schneider et al., 1997) and (2) the observed e p differences of 1x between GeoB 1008 and GeoB 1016 (+1.1x vs. +0.1x; Fig. 4). Similar to the sediments in the Angola Basin, the record of core GeoB 1028 also showed higher isotopic fractionation during the LGM compared to the Holocene (Fig. 4). The core site is situated well seawards of the coastal upwelling zone, within the mixing area between Benguela Ocean Current and Benguela Coastal Current (Fig. 1). The results of

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various palaeoceanographic reconstructions concerning this region reveal an inconsistent picture, often denoted as the Walvis Opal Paradox (for a recent review see Berger and Wefer, 2002). On the one hand, Diester-Haass (1985) and Diester-Haass et al. (1988) concluded on the basis of biogenic opal accumulation that upwelling and productivity was less intense during glacials and that the Benguela Coastal Current flowed further on shore into the Angola Basin. On the other hand, Oberha¨nsli (1991), Schneider et al. (1995), and Summerhayes et al. (1995) suggested that due to an increase in trade-wind zonality during glacial times the coastal upwelling filaments shifted westward, thereby increasing nutrient concentration and productivity at the outer Walvis Ridge. In core GeoB 1028, the concentrations of the C37 alkenone closely co-vary with the total organic carbon (TOC) content, with enhanced values during the last glacial (Mu¨ller et al., 1997). In addition, variations in coccolithophore assemblages show that during MIS 2 the relative abundances of Emiliania huxleyi and Gephyrocapsa spp. do not significantly differ from those in core GeoB 1710 off Namibia (Mu¨ller et al., 1997; Baumann et al., 1999). Together, this leads to the assumption that in the area of site 1028 the productivity of alkenone producing algae was enhanced during the LGM as it was the case in the nearshore upwelling area. On the basis of available evidence we cannot unambiguously determine the reasons for the observed e p variation at this core location. Possible causes might be the same as considered for the sites GeoB 1008 and 1016 in the Angola Basin. However, to test this hypothesis additional information about changes in nutrient levels, growth rates, and phytoplankton assemblages is required.

6. Summary and conclusions We analysed the carbon isotopic fractionation (e p) of sedimentary alkenones from various oceanographic regions in the equatorial and subtropical South Atlantic during the LGM. The results were compared to the Holocene e p signal from core–top sediments at the respective core locations (Andersen et al., 1999; Benthien et al., 2002). Generally, the e p signal showed lower values during the LGM compared to the

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Holocene. Exceptions are the cores from the Angola Basin and one from the eastern crest of the Walvis Ridge where higher glacial e p values were recorded. Taking into account the present knowledge of Late Quaternary variations in surface-water hydrography, nutrient conditions and palaeoproductivity of the South Atlantic, our results indicate that regionally different factors controlled the magnitude of isotopic fractionation of alkenones in the different biogeochemical provinces: (1) In the oligotrophic areas of the Brazil Current, the South Atlantic Tropical Gyre, and the Western Tropical Atlantic, less rather than more nutrients were available during the LGM. The observed trend towards lower glacial e p values could thus not be explained by variations in haptophyte growth rates. Therefore, it is assumed that the isotopic record is caused mainly by decreased surface-water CO2aq concentrations during the LGM. However, the estimated, e p-based decrease in surface-water PCO2 is lower than expected from the glacial/interglacial atmospheric pCO2 variations. This implies either a small glacial CO2 source in these regions or slightly lower glacial haptophyte growth rates have biased the e p signal towards higher values. (2) In the upwelling areas of the Benguela Coastal Current and the Eastern Tropical Atlantic the observed variations in e p seem to reflect mainly changes in haptophyte growth rates which were controlled by the availability of surface-water nutrient concentrations. Due to the missing evidence for the amount of additional carbon dioxide brought to the surface under intensified coastal upwelling it could not be determined to what extent [CO2aq] has influenced the alkenone isotopic record. (3) In the Angola Basin, the higher glacial e p values indicate lower or at least similar haptophyte growth rates during LGM although other palaeoceanographic proxies point to enhanced productivity and higher surface-water nutrient levels. It is most likely that due to a glacial increase in the fluvial silica input from the Congo River diatoms became more important in the phytoplankton community, thus leaving less nutrients available for haptophyte algae.

(4) This study demonstrates that the magnitude of the LGM to Holocene record of the sedimentary y13C37:2 signal in the equatorial and subtropical South Atlantic is mainly controlled by two factors. The by far dominant factor is the regional variation in glacial to Holocene surface-water nutrient levels at the time of alkenone production. To a minor extent variations in the CO2aq concentration could also affect the sedimentary signal. However, evidence for variations in e p due to changes in [CO2aq] could only be found in areas with presumably no or only small Holocene to glacial variations in surface-water nutrient concentrations. These findings severely complicate the approach of alkenone e p as a global palaeo-PCO2 proxy. A reliable application of this approach depends strongly on our ability to better quantify algal growth rates. Studies on Sr/Ca ratios in coccoliths (Stoll and Schrag, 2000; Rickaby et al., 2002) have provided very promising results for a combined approach of these parameters together with the alkenone e p signal. On the other hand, our results imply the potential of the sedimentary y13C37:2 signal to serve as a proxy for palaeo-nutrient reconstructions (Schulte et al., 2004). In this context it is important to note that the carbon isotopic alkenone signal reflects surface-water conditions at the time of alkenone production. This study demonstrates that in combination with other palaeoceanographic proxies, the y13C37:2 signal may provide important information about variations in nutrient levels and resultant changes in the phytoplankton community. Addressing this implication clearly requires further studies both in the sediment and under well-constrained experimental conditions.

Acknowledgements We wish to thank the officers, crews and scientists aboard R/V Meteor, R/V Sonne, and R/V Victor Hensen for their assistance with coring and sampling operations during cruises to the South Atlantic. We are grateful to Wolfgang Bevern, Birgit MeyerSchack, and Monika Segl for their assistance with

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the mass spectrometers and both Hella Buschhoff and Ralph Kreutz for their technical support in the laboratory. Brian Popp and one anonymous reviewer contributed with detailed and helpful criticism. We would also like to thank Gaute Lavik and Gesine Mollenhauer whose comments and critical reviews on an earlier version significantly improved the paper. This research was funded by the Deutsche Forschungsgemeinschaft (Sonderforschungsbereich 261 at Bremen University).

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