High-resolution benthic foraminifer δ13C records in the South China Sea during the last 150 ka

High-resolution benthic foraminifer

δ

13

C records in the South China

Sea during the last 150 ka

Gang-Jian Wei

⁎

, Chi-Yue Huang

, Chia-Chun Wang

,

Meng-Yan Lee

, Kuo-Yen Wei

a

Key Laboratory of Isotope Geochronology and Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou, Guangdong 510640, China

b

National Taiwan University, Taipei, Taiwan

c

National Cheng Kung University, Tainan, Taiwan

Received 19 May 2005; received in revised form 16 July 2006; accepted 2 August 2006

Abstract

High-resolutionδ13C ratios of benthic foraminifer, Cibicidoides wuellerstorfi, in Core MD97-2151 from the southern South China Sea (SCS) and Core SO50-31KL from the northern SCS indicate deep-water variations in the SCS during the last 150 ka. Theδ13C records of both show a general variation pattern with higherδ13C during interglacial periods than glacial periods, and rapid increases during terminations. This variation pattern resembles those of the West Pacific intermediate/deep water, the sources of the SCS deep water and the general variation of global deep water, indicating that the variation of the SCS deep water is largely controlled by global factors, such as carbon cycles and global deep-water circulation. Spectra analysis exhibits robust Milankovitch cycles in the SCS deep-waterδ13C, again supporting global controls. Significant semi-precession cycles can also be seen in the spectra, implying that changes of local surface productivity also influence the SCS deep-waterδ13C.

© 2006 Elsevier B.V. All rights reserved.

Keywords: carbon isotopes; benthic foraminifer; deep-water ventilation; South China Sea

1. Introduction

The δ13C values of the dissolved inorganic carbon (DIC) in water is one of the best tracers for deep-water circulation (Curry and Lohmann, 1982; Shackle-ton et al., 1983; Duplessy et al., 1984; Duplessy and Shackleton, 1985; Boyle and Keigwin, 1985/86). When

deep-water flows, organic matter with more negative δ13

C sinks from surface water and decomposes, resulting in decreases of DICδ13C. Theδ13C of benthic foraminifer, Cibicidoides wuellerstorfi are very close to the dissolved CO2of deep-water, thus can well proxy

deep-waterδ13C (Belanger et al., 1981; Duplessy et al., 1984; Curry et al., 1988). Moreover, the relationship between theδ13C of C. wuellerstorfi and DIC is quite robust, and does not change with water depth ( McCor-kle and Keigwin, 1994). Therefore, theδ13C values of C. wuellerstorfi represent deep-water δ13C (Keigwin, 1998), and are generally used to reconstruct deep-water variations.

⁎ Corresponding author. Key Laboratory of Isotope Geochronology and Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou, Guangdong 510640, China. Tel.: +86 20 85290093; fax: +86 20 59290130.

E-mail address:gjwei@gig.ac.cn(G.-J. Wei).

0025-3227/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.margeo.2006.08.005

The South China Sea is the largest marginal sea in the west Pacific with several connections to the Pacific and Indian Ocean. The Bashi Strait, located in the northeast SCS with a sill depth of about 2500 m, is the only gateway that enables deep water to flow into the SCS from the west Pacific (Chao et al., 1996). At present, deep water from West Philippine Sea flows into the SCS through this gateway and fills the abyssal SCS from north to south (Wang, 1986). Two deep upwelling sys-tems off southwest Taiwan and Vietnam help to renew the SCS deep water (Chao et al., 1996). Paleoceano-graphic studies indicate that both the intermediate water mass (IWM) and deep-water mass (DWM) above 2500 m in the west Pacific contribute to the deep-water of the SCS (Jian and Wang, 1997), and the ven-tilation of SCS deep-water is closely related to that in

their sources (Jian and Wang, 1997). Because both the carbonate content in sediments and the sedimentary rates are low in the west Pacific, well-preserved high-resolution deep-water records are very rare. Therefore, records from the SCS are particularly valuable to un-derstand the deep-water variations of the west Pacific (Jian and Wang, 1997).

Studies of deep-water ventilation in the SCS, based on benthic foraminifers, were carried out by Oppo and Fairbanks (1987),Winn et al. (1992),Miao and Thunell (1996), Jian and Wang (1997) and Jian et al. (1999, 2000).δ13C values in benthic foraminifers were used to track deep-water variations (Oppo and Fairbanks, 1987; Winn et al., 1992; Jian and Wang, 1997; Wang et al., 1999). Previous studies showed that deep-waterδ13C in the SCS is generally lower than in the west Pacific (Winn

et al., 1992), and benthic δ13C in the southern SCS is generally more negative than in the northern SCS during glacials (Jian and Wang, 1997; Wang et al., 1999). High-resolution deep-water records in the SCS, however, are scarce, and deep-water variations are not well known. Moreover, spatial differences in organic matter fluxes from surface water are substantial in the SCS. Because higher organic fluxes provide more organic carbon to the deep water, the released carbon from decomposed organic matter may result in more negative δ13C in deep water. Such spatial differences may result in abnormal deep-water δ13C values (Wang et al., 1999), and obscure the deep-water ventilation signal in δ13C records. Nonetheless, to what extent they influence deep-waterδ13C is not well known in the SCS.

We present two high-resolution benthic foraminiferal δ13

C records from the southern and northern SCS, which document deep-water variations during the last 150 ka. The northern record was obtained from sediment Core SO50-31KL, recovered at water depth of 3360 m, far below the sill depth of the Bashi Strait. The southern record was obtained from sediment Core MD97-2151, recovered at water depth of 1598 m, above the sill depth of the Bashi Strait. Comparisons of δ13C in these two records and in records from other locations provide new insights into deep-water ventilation in the SCS within the last glacial/interglacial cycle.

2. Materials and methods

Sediment samples from Core SO50-31KL (18°45.4′N, 115°52.4′E, water depth 3360 m) and Core MD97-2151 (8°43.73′N, 109°52.17′E, water depth 1598 m) (Fig. 1) were used in this study. The recovered length was 2560 cm and 800 cm for Core MD97-2151 and Core SO50-31KL, respectively. Samples were taken every 4 cm in Core MD97-2151, and every 2 cm in Core SO50-31KL, corresponding to time resolution of about 0.2 kyr and 0.3 kyr for Core MD97-2151 and SO50-31KL, respectively. The benthic foraminifer, C. wuellerstorfi, was picked from theN150 μm size fraction of the samples for isotope analysis.

The foraminiferal tests were first ultrasonically washed in distilled water to remove adhering particles, and then soaked in 10% NaOCl solution for 24 h to remove organic materials. δ13C and δ18O were mea-sured on a Finnegan MAT Delta Plus mass spectrometer coupled with a Kiel automatic carbonate device in the Department of Geology, National Taiwan University. On average, about 3 C. wuellerstorfi tests were mea-sured for each sample. The oxygen and carbon isotopic ratios were calibrated to PDB standard via USGS

standard NBS 19. Analysis on NBS 19 and a lab marble standard during this study provided an external pre-cision of 0.08‰ and 0.05‰ for δ18

O and δ13C, re-spectively. About 50 samples from the two cores that had sufficient tests were double-checked, and the repro-ducibility is better than 0.08‰ and 0.05‰ for δ18

O and δ13

C, respectively, which agreed with that of NBS 19 and the marble standard. The oxygen and carbon isotope results are presented in a Supplementary Data in the Appendix.

3. Age models

The age models for these two cores were initially generated from the δ18O curve of the planktonic foraminifer Globigerinoides sacculifer, coupled with several AMS 14C dates of planktonic foraminifers (Chen and Huang, 1998; Lee et al., 1999). Because of significant carbonate dissolution, the G. sacculifer tests in Core SO50-31KL were lack, resulting in relatively low resolution for theδ18O curve (Chen and Huang, 1998). Here, we generated a new age model for Core SO50-31KL, based on theδ18O curve of C. wuellerstorfi and the previous AMS14C ages. The age control points are listed in Table 1. In contrast to G. sacculifer, C. wuellerstorfi is more resistant to dissolution, and the resolution of theδ18O curve of C. wuellerstorfi is higher, providing better age controls (Fig. 2).

We also established a new age model based on the δ18

O of C. wuellerstorfi for Core MD97-2151, al-though the resolution of the initial age model, based on

Table 1

Age control points for Core SO50-31KL Depth (cm) Age (ka) Interpretation 5 0.924 C-14a 35 6.045 C-14a 41 7.441 C-14a 87 12.05 MIS 2.0 161 17.85 MIS 2.2 233 24.11 MIS 3.0 293 28 MIS 3.1 399 43.88 MIS 3.13 450 50.21 MIS 3.3 541 55.45 MIS 3.31 576 58.96 MIS 4.0 609 64.09 MIS 4.22 651 73.91 MIS 5.0 700 79.25 MIS 5.1

MIS represents oxygen isotope events correlated to SPECMAP. Ages are fromMartinson et al. (1987)except for MIS 3.1, which is from (Grootes et al., 1993).

G. sacculifer, is high enough. Because theδ18O curve of C. wuellerstorfi follows more closely the SPECMAP than that of G. sacculifer, which exhibits nearly the sameδ18O at MIS 5.1, MIS 5.3 and MIS 5.5 (Lee et al., 1999), the new age model may provide better age controls. Previous AMS14C ages younger than 10 ka (Lee et al., 1999) and the volcanic glass layer at 1558 cm were used as additional age control points. This glass is considered to be related to the latest eruption of the Toba volcano in Sumatra at∼71 ka (Lee et al., 1999). The age control points are listed in Table 2 and the oxygen isotope curves for Core SO50-31KL and MD97-2151 are shown inFig. 2.

The new age models for the two cores agree well with the former age models within errors of the SPECMAP ages (Martinson et al., 1987).

4. Carbon isotope results

Theδ13C curves of C. wuellerstorfi in Cores SO50-31KL and MD97-2151 are somehow similar both in

variation patterns and values (Fig. 3).δ13C is generally higher during interglacial periods than during glacial periods. The maximumδ13C (about 0.2∼0.3‰) in Core MD97-2151 occurs during Stage 1, and the minimum (about −0.9‰), showing an increasing trend towards the present since 150 ka (Fig. 3). Such trend is not so clear in Core SO50-31KL, where δ13C values are similar during Stages 3 and 5, and during Stages 2 and 4. However, the maximum δ13C (about 0.3‰) occurs during Stage 1, implying similar increasing trend to-wards the present like that in Core MD97-2151 (Fig. 3). There are several negative spikes in theδ13C curve of Core MD97-2151, centering at ∼33 ka, ∼38 ka and ∼44 ka during Stage 3. The decreases in δ13

C are about −0.2‰ to −0.4‰, and lasted for less than 1000 years in general. Such negative peaks were not seen in Core SO50-31KL, although the resolution of the two cores is similar.

During Terminations I and II,δ13C shows stepwise increases in both cores (Fig. 3). A positive δ13C excursions with amplitude of about 0.1–0.15‰ in

Fig. 2. Age correlation points for Core MD97-2151 (upper) and Core SO50-31KL (lower). Crosses indicate AMS14_{C dating points. Arrows indicate}

Core MD97-2151 and about 0.18‰ in Core SO50-31KL occurred from∼14 to ∼11 ka during Termination I. The positiveδ13C excursion from∼130 to ∼126 ka during Termination II in Core MD97-2151 is even larger, with amplitude of about 0.25∼0.40‰. Such stepwise increases in δ13C resemble the Younger Dryas-style events recorded in benthic δ13C from the tropical Atlantic during Terminations I to IV (Sarnthein and Tiedemann, 1990). Similar events were also detected in high-resolution benthic δ13C records from the Pacific (Lund and Mix, 1998) and Indian Ocean (Duplessy et al., 1988).

5. Fluctuations of the SCS deep-water

In general, deep-water δ13C variations in glacial– interglacial time scale are controlled by changes in the global organic carbon inventory and oceanic deep-water circulation (Shackleton, 1977; Duplessy et al., 1984). Climatically induced changes in the global organic

car-bon inventory may result in carcar-bon transfer between oceanic DIC, the terrestrial biomass (Shackleton, 1977) and shelf sediments (Broecker, 1982), which lead to changes in the mean ocean δ13C. Modeling indicates that carbon transfer may account for about two thirds of theδ13C shift within glacial/interglacial cycles in east-ern Pacific, and about one third in north Atlantic, and such significant changes of benthic foraminiferδ13C are synchronous on global scale (Ku and Luo, 1992). The variation patterns of Cores SO50-31KL and MD97-2151 with increasingδ13C towards the present are also evident in deep-water δ13C records from the eastern Pacific (Shackleton et al., 1983; Raymo et al., 1990), easternmost Indian Ocean (Holbourn et al., 2005) and Atlantic (Ruddiman et al., 1989; Sarnthein and Tiede-mann, 1990). This increasing trend seems to represent a general pattern for global deep-water, which appear to stem from carbon transfer between different reservoirs (Ku and Luo, 1992). Also, the increase inδ13C seems to agree with the increasing trend of organic carbon accumulation in the ocean toward the present (Ku and Luo, 1992). In addition to the carbon cycle, changes in deep-water circulation may influence global deep-water δ13

C (Duplessy et al., 1984; Duplessy et al., 1988). Most of the Younger Dryas-style stepwise increases in deep-waterδ13C in the Atlantic is believed to relate to the short term reversal of North Atlantic Deep-Water (NADW) circulation, induced by meltwater during Terminations (Sarnthein and Tiedemann, 1990). The Younger Dryas-style stepwise increases in δ13C of Cores SO50-31KL and MD97-2151, and in cores from the Indian and Pacific Oceans (Duplessy et al., 1988; Mix et al., 1991; Lund and Mix, 1998) may also relate to changes in global deep-water circulation. Therefore, the δ13

C variations in Cores SO50-31KL and MD97-2151 from the SCS agree with records from other oceans, indicating a response to global deep-water variations.

The sources of the SCS deep-water are the inter-mediate water mass and the deep-water mass above 2500 m in the west Pacific (Jian and Wang, 1997). Thus, Core V19-27 collected at 1373 m water depth in the east Pacific is representative of the source of the SCS deep-water (Mix et al., 1991; Talley, 1993). Theδ13C in Core V19-27 remains about 0.2‰∼0.8‰ higher during the last 150 ka than in Core MD97-2151, and the differences during Stage 4 and Stage 6 is somewhat higher than during interglacials (Fig. 4). The δ13C difference between the two cores may be attributed to the δ13C decline in deep-water, as it flows from the source in the west Pacific to the SCS. Despite of this difference, the variation patterns of the two records are similar (Fig. 4).

Table 2

Age control points for Core MD97-2151 Depth (cm) Age (ka) Interpretation 3 0.938 C-14a 35 1.353 C-14a 71 2.143 C-14a 91 2.743 C-14a 127 3.697 C-14a 171 4.499 C-14a 211 5.648 C-14a 251 7.276 C-14a 311 9.927 C-14a 363.5 12.05 MIS 2.0 575.5 17.85 MIS 2.2 763.5 28 MIS 3.1 939.5 43.88 MIS 3.13 1085.5 50.21 MIS 3.3 1313 58.96 MIS 4.0 1421 64.09 MIS 4.22 1556 71 Toba ashb 1601 73.91 MIS 5.0 1661 79.25 MIS 5.1 1797 96.21 MIS 5.31 1865 103.29 MIS 5.33 1949 110.79 MIS 5.4 2097 122.56 MIS 5.51 2133 125.19 MIS 5.53 2197 129.84 MIS 6.0 2321 135.1 MIS 6.2

MIS represents oxygen isotope events correlated to SPECMAP. Ages are fromMartinson et al. (1987)except for MIS 3.1, which is from (Grootes et al., 1993).

a

All calibrated to calendar years, data fromLee et al. (1999).

b

In order to understand the driving forces behind variations of the SCS deep-water, we performed a spec-tral analysis on the δ13C of Core SO50-31KL and MD97-2151 from the SCS, and Core V19-27 from the east Pacific, using the software SPECTRUM (Schulz and Stattegger, 1997). Harmonic analyses (Siegel's Test) of theseδ13C records all exhibit robust Milanko-vitch cycles (Fig. 5). The 100 ka-eccentricity cycles, 41 ka-obliquity cycles and 23 ka-precession cycles are well expressed in theδ13C records of Core MD97-2151 and V19-27 within the 6 dB bandwidth. No eccentricity cycle is apparent in the records of Core SO50-31KL due to the short time span, but the obliquity and precession cycles are well expressed (Fig. 5).

In addition to Milankovitch periodicities, high frequency cycle of 8∼11 ka is apparent in the δ13

C records of Core SO50-31KL and MD97-2151 with spectral powers above the 95% false alarm levels

(Fig. 5). In contrast, Core V19-27 does not exhibit such a periodicity, which may be due to its lower res-olution. These high-frequency cycles are probably half-precession cycles that may relate to insolation variations in tropical regions (Berger and Loutre, 1997), and they are also detected in some high-resolution paleoclimate records, such as pollen (Sun et al., 2003; Luo and Sun, 2005; Luo et al., 2005) and TOC contents (Wang, 1999) in the SCS, and paleoproductivity records in the Timor Sea (Holbourn et al., 2005). The power spectrum for TOC in Core MD97-2151 also shows half-precession cycles centering at 9–9.5 ka (Fig. 5). TOC variations are generally related to changes of organic flux from surface water, which is controlled by productivity (Berger et al., 1989), and deep-waterδ13C is significantly influenced by organic matter flux from surface water (Kroopnick, 1985). The half-precession cycles in the SCS deep-water δ13

C therefore suggest that theδ13C changes in Cores

Fig. 3.δ13_{C ratios of the benthic foraminifer C. wuellerstorfi in Cores MD97-2151 and SO50-31KL. Theδ}13_{C records of Core MD97-2151 were}

smoothed by a 3-point running average.δ18_{O curves were also shown to identify isotope stages. The shaded areas indicate Younger Dryas-style}

SO50-31KL and MD97-2151 are partly controlled by surface productivity in the SCS.

In the SCS, surface productivity is heterogeneous and changes during glacial/interglacial cycles (Wang et al., 1999; Jian et al., 2001; Wei et al., 2003). The deep-water in the SCS generally flows from north to south (Wang, 1986), andδ13C values in the south are

generally more negative than in the north (Jian and Wang, 1997; Wang et al., 1999). Recent observations, however, reveal a much complicated pattern for deep waterδ13C, which does not always follows deep-water circulation in the SCS (Cheng et al., 2005). Local surface productivity might have a marked effect on the deep-waterδ13C signal in the SCS.

Fig. 5. Harmonic analysis (Siegel's test) of theδ13_{C of Cores MD97-2151, SO50-31KL, and V19-27 and TOC contents of Core MD97-2151 using}

SPECTRUM (Schulz and Stattegger, 1997). Setting: OFAC = 4, HIFAC = 1,α=0.05, λ=0.4; Horizontal bar marks 6-dB bandwidth, dashed line indicates false alarm level of 95% confidence intervals. Numbers represent the respective frequency of the peaks.

Fig. 4. Comparison ofδ13C records of South China Sea deep water and Pacific intermediate/deep water. Solid lines represent the SCS deep-water records from Core MD97-2151 after smoothing by a 3-point running average. Dashed line represents the Pacific intermediate water record from Core V19-27 (Mix et al., 1991). Horizontal lines indicate the boundaries of oxygen isotope stages.

6. Summary

High-resolutionδ13C ratios of the benthic foramin-ifer, C. wuellerstorfi in Core MD97-2151 from the southern SCS and Core SO50-31KL from the northern SCS allow investigating deep-water fluctuations in the SCS during the last 150 ka.

Theδ13C records of these two cores agree well with each other, with higherδ13C during interglacials than during glacials, increasingδ13C toward the present, and Younger Dryas-style stepwise increases during Termi-nations. Such variation patterns agree with global deep-water, implying that variation of the SCS deep-water is generally controlled by global factors, such as carbon cycles and global deep-water circulation. This is further supported by the robust Milankovitch cycles in the SCS deep-water δ13C records. Whereas significant semi-precession cycles in the spectra of theseδ13C records suggest the influence of local productivity.

Acknowledgement

We thank Dr Chen Y.G. of the Department of Geo-logy, National Taiwan University, for help with the isotope analysis, and Dr Chen M.T. of the National Taiwan Ocean University for valuable advice. Proof reading by Dr. Sun W.D. of Guangzhou Institute of Geochemistry, CAS, helped improve the manuscript. The authors thank Dr A. Holbourn of the Kiel University, Dr. De Lange and an anonymous reviewer for the constructive comments and suggestions. Thanks also go to Dr. Jian Z.M. of the Tongji University for comments on the former version of this manuscript. This study was supported by National Science Council Grants 006, NSC86-2611-M-002-04 and NSC88-2119-M-019-002 to C.Y. Huang, and Grant 40173015 of the Natural Science Foundation of China (NSFC) to G.J. Wei.

Appendix A. Supplementary data

Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.margeo. 2006.08.005.

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