interglacial variations of sedimentation on the West Australian continental margin: constraints from excess 230Th

Marine Geology 166 (2000) 11–30 www.elsevier.nl/locate/margo

Glacial/interglacial variations of sedimentation on the West Australian continental margin: constraints from excess 230Th H.H. Veeh a,*, D.C. McCorkle b, D.T. Heggie c b

a School of Earth Sciences, Flinders University, Bedford Park, SA 5042, Australia Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA c Petroleum and Marine Division, Australian Geological Survey Organization, Canberra, ACT 0200, Australia

Received 12 August 1999; accepted 26 January 2000

Abstract We determined the mass accumulation rates of biogenic (carbonate, organic carbon, Ba) and terrigenous (Al, Th) sediment components, together with excess 230Th in gravity cores from the Exmouth Plateau and the Perth Basin on the West Australian continental margin, in order to reconstruct paleoceanography and paleoclimate in this critical area during the Late Quaternary. We found the mass accumulation rates of these components, based on normalization to age-corrected initial excess 230Th in the sediments (constant flux model), to be more reliable for this purpose than the site-specific burial rates, since the latter appear to be affected by lateral sediment redistribution from the shelf and slope. Under the constraints provided by the constant 230Th flux model, there is little convincing support for enhanced productivity off Western Australia during the last glacial maximum and hence no compelling evidence in the sediment record for strong coastal upwelling comparable to that in the modern ocean off the west coasts of Africa and South America. The hypothesis of a major reorganization of ocean circulation in the southeastern Indian Ocean involving the replacement of the south flowing Leeuwin Current with the north flowing West Australian Current during the last glacial maximum should therefore be viewed with reservations. The flux pattern of Al and Th indicates a significantly lower input of terrigenous components during the last glacial maximum than during the Holocene, suggesting that the supply of terrigenous components to the deep sea floor in this area have been influenced primarily by variable runoff from monsoon-controlled northwestern Australia. The results are consistent with expanded aridity during glacial times and a somewhat wetter climate in NW Australia during interglacials. 䉷 2000 Elsevier Science B.V. All rights reserved. Keywords: Continental margin; Paleo-oceanography; Sedimentation; Th-230; Western Australia

1. Introduction Hemipelagic sediments near continental margins are under the influence of both oceanographic and land-based processes and as such should provide a detailed record of paleoclimatic variations during * Corresponding author. Tel.: ⫹ 61-8-201-2212; fax: ⫹ 61-8201-2212. E-mail address: [email protected] (H.H. Veeh).

the Quaternary. The West Australian continental margin is of particular interest in this regard for several reasons. Whereas the western continental margins of the world are characterized by equatorward eastern boundary currents driven by equatorward wind stress which is generally associated with strong coastal upwelling, there is no such equatorward coastal current off Western Australia and no evidence of significant coastal upwelling at the present time, in spite of equatorward winds throughout much of the

0025-3227/00/$ - see front matter 䉷 2000 Elsevier Science B.V. All rights reserved. PII: S0025-322 7(00)00011-6

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Fig. 1. Surface currents in the Southern Indian Ocean (after Tomczak and Godfrey, 1994); STF ˆ Subtropical Front.

year (Smith, 1992). Instead, a current of warm, low salinity water, known as the Leeuwin Current (Fig. 1) flows poleward along the shelf edge against the prevailing wind (Cresswell and Golding, 1980; Thompson, 1984; Godfrey and Ridgeway, 1985; Godfrey, 1996). The driving force for the Leeuwin Current is thought to be the longshore pressure gradient in the upper ocean between the warm (low density) tropical waters south of Indonesia and the cool (high density) water of the Southern Ocean which overrides the opposing equatorward wind stress and effectively blocks coastal upwelling along Western Australia (Thompson, 1984; Godfrey, 1996). There is some evidence in the sediment record that the ocean circulation off Western Australia may have been different during the last glacial maximum (LGM). Population statistics of planktonic foraminifera indicate that sea surface temperatures (SST) off Western Australia were significantly lower than now during the LGM, suggesting that the Leeuwin Current may have ceased to flow at that time (Prell et al., 1980; Wells and Wells, 1994; Wells et al., 1994). Such a change in ocean circulation during the LGM could have arisen (Prell et al., 1980) due to a northward shift of the Subtropical Front (STF) and an associated

strengthening of the West Australian Current (Fig. 1), which could have replaced the Leeuwin Current to form a cold equatorward current similar to eastern boundary currents now occurring along the west coasts of South America and Africa. Without the Leeuwin Current, the prevailing equatorward wind along Western Australia would almost certainly have produced strong coastal upwelling due to offshore Ekman transport, in analogy to the present day ocean off the coasts of Peru and Namibia. Inasmuch as coastal upwelling is usually associated with enhanced productivity (Suess and Thiede, 1983), the sediment record off Western Australia should contain some evidence for this. In addition to this postulated paleoceanographic scenario, the glacial to Holocene decrease in aridity of northwestern Australia (Williams et al., 1993) should have left some imprint in the distribution of terrigenous sediment components in the marine sediment record for some distance offshore. The present study is an attempt to unscramble the productivity and terrigenous input signals from the sediment record off Western Australia, in order to aid in the reconstruction of the paleoceanography and paleoclimate in this region during the Late

H.H. Veeh et al. / Marine Geology 166 (2000) 11–30 Table 1 Core locations Core Number

Latitude

Longitude

Depth (m)

Exmouth Plateau 53GC04 53GC07

19⬚35.1 0 S 18⬚54.5 0 S

113⬚32.1 0 E 112⬚37.9 0 E

956 2256

Perth Basin 57GC15 57GC19

29⬚22.9 0 S 27⬚19.2 0 S

113⬚13.0 0 E 111⬚37.6 0 E

2750 2755

Quaternary. We will rely on organic carbon (Mu¨ller and Suess, 1979; Pedersen, 1983; Sarnthein et al., 1988; Berger and Herguera, 1992) and Ba (Goldberg and Arrhenius, 1958; Dehairs et al., 1980; Bishop, 1988; Dymond et al., 1992; Gingele and Dahmke, 1994; Francois et al., 1995; Rutsch et al., 1995) as proxies for productivity, and on Al and Th (Bostrom et al., 1973; Chester, 1990; Cochran, 1992) as indicators of terrigenous sediment components. We will show that the mass accumulation rates of these various sediment components are constrained by the inventory of excess 230Th in the sediment cores, which can be compared to the known production rate of 230Th from its parent 234U in the overlying water column, thus permitting the recognition of sediment focusing and/or erosion at a given location (Cochran and Osmond, 1976; Suman and Bacon, 1989; Thomson et al., 1993; Scholten et al., 1994). In addition, the depth profiles of age-corrected initial excess 230Th concentrations in sediment cores (constant flux model) provide a means of point-bypoint estimation of regional sedimentation rates of any chosen sediment component out of the water column, which is relatively insensitive to lateral sediment redistribution on the sea floor (Suman and Bacon, 1989; Francois et al., 1990).

2. Material and methods

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archived in a cold room at the Australian Geological Survey Organisation (AGSO). The sediments were sampled at closely spaced intervals ranging from 5 to 20 cm. Each sample represents a homogenized 2 to 3 cm section, the mid-point of which is used in graphing analytical data. The sediments consist predominantly of foraminifera and calcareous nanoplankton, with minor terrigenous components. Opaline silica (mostly radiolaria) does not exceed 3% (Corlis, 1990). 2.2. Dry bulk density and salt correction Dry bulk density (DRB) was measured by determining the weight after drying (60⬚C) of a known volume of wet sediment. The weight loss (in percent of wet weight) upon drying, together with the known salinity of bottom water in the area, has been used to correct the analytical data for salt content. The salt content in these sediments varies between 1.3 and 1.7%. 2.3. Geochemistry CaCO3 and organic carbon (Corg) concentrations were determined by coulometric titration (Engleman et al., 1985; Lee and Macalady, 1989) of the CO2 released upon acidification with phosphoric acid (for CaCO3), and upon oxidation with dichromate/sulfuric acid (for Corg), in an adaptation of the method of Weliky et al. (1983). Both steps were carried out sequentially on the same sample and in the same digestion bottle. The reproducibility for CaCO3 and Corg at the concentration levels of the present study was ^0.5 and ^6%, respectively. The accuracy of the method was checked against reagent grade calcite and sucrose, as well as against a phosphate rock standard containing a known concentration of Corg at about the same concentration level as in the sediments of this study. Al and Ba were measured by X-ray fluorescence (XRF) at AGSO as discussed by Cruikshank and Pyke (1993), with a precision of ^0.5 and ^3%, respectively.

2.1. Sample description and processing 2.4. U and Th isotopes The gravity cores used were collected on the crest and western slope of the Exmouth Plateau, and on the lower continental slope facing the Perth Basin (Table 1, Fig. 2) during Rig Seismic cruises 53 and 57. The cores were split on board, sealed in plastic and

The U and Th isotopic analyses involved complete dissolution of the samples, followed by separation on anion exchange resins as described by Anderson and Fleer (1982). A calibrated 228Th/ 232U spike was used

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3. Results and discussion 3.1. Chronology

Fig. 2. Bathymetry of study are and core locations. Depth contours in meters.

to determine chemical yield. The U and Th isotopes were electroplated on to stainless steel discs, and the isotope ratios measured by alpha spectrometry on a Canberra Model S100 MCA coupled to a Canberra Quad-Alpha multiple input system with surface barrier detectors. Analytical errors based on counting statistics (^1s ) are 3% for U, 7% for Th ( ˆ 232Th), and 1–2% for 230Th.

2.5. Radiocarbon dates AMS radiocarbon age determinations (carbonate and organic carbon fractions, see Table 2) were carried out by the Australian Nuclear Science and Technology Organisation (ANSTO), Menai, NSW, and by the Institute of Geological and Nuclear Studies, Lower Hutt, New Zealand.

Oxygen isotope stratigraphy, together with linear sedimentation rates previously used (McCorkle et al., 1994) are reproduced in Table 3. The ages assigned to individual oxygen isotope stage boundaries had been obtained by correlation of d 18O records from the planktonic foraminifera Globigerinoides sacculifer (250– 300m, w/o sac) with the SPECMAP stacked d 18O record (Pisias et al., 1984; Martinson et al., 1987). Because of the low resolution of the isotopic data available in our cores, only major stage boundaries have been identified. Sedimentation rates are assumed to be constant between these stage boundaries, and the sediment surface is taken to have zero age. Radiocarbon ages (Table 2) are reported as conventional ages as defined by Stuiver and Polach (1977), but with a reservoir correction of 400 years (Bard, 1988). Conventional 14C ages greater than 9000 years BP have been converted to calendar years, using the calibration of Bard et al. (1998), and are shown in Fig. 3, together with oxygen isotope event markers 2 and 3 in terms of the SPECMAP age model. Most of the 14C ages of the carbonate fraction are significantly older than the ages inferred from oxygen isotope stratigraphy at corresponding depths in the cores, and the age difference appears to increase with depth. A notable exception are the 14C ages based on the organic carbon fraction (core 53GC07), which are in close agreement with SPECMAP model ages. The surface ages are comparable to 14C ages measured at the surface of other deep sea cores with similar sedimentation rates and can be attributed to bioturbation (Berger and Johnson, 1978; Erlenkeuser, 1980), which would introduce a positive age bias into the surface mixed layer prior to burial of the sediment. However, the divergence between 14C ages (carbonate fraction) and SPECMAP model ages below the surface mixed layer cannot be due entirely to bioturbation and is of some concern, as the respective chronologies would produce significantly different sedimentation rates. For instance, the sedimentation rate for stage 2 in core 53GC07, based on linear regression of the 14C ages, would be only 2.69 cm/ ky (Fig. 3), as compared to the sedimentation rate of 3.52 cm/ky based on oxygen isotope stratigraphy.

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Table 2 Radiocarbon (AMS) ages Laboratory reference

Age a (years B.P.)

Age b (years B.P.)

53GC04 0–3 23–26 37–40 67–70 77–80

OZA216U NZA2986 OZA217U NZA2989 OZA218U

3340 ^ 90 9920 ^ 90 14; 500 ^ 80 19; 940 ^ 140 25; 400 ^ 200

11,490 17,060 23,510 29,800

53GC07 1–3 1–3 c 22–24 32–34 47–49 47–49 c 57–59 72–74

OZA211 OZA211 OZA213 OZA214 OZA212 OZA212 NZA1518 OZA215

2790 ^ 100 1790 ^ 90 9290 ^ 80 13; 400 ^ 150 17; 990 ^ 90 15; 400 ^ 130 21; 530 ^ 210 27; 100 ^ 250

10,720 15,730 21,220 18,140 25,360 31,730

57GC15 1–3 11–13 16–18 36–38

OZA219U OZA220U OAZ221U NZA2988

4160 ^ 70 10; 000 ^ 150 13,000 ^ 120 30; 300 ^ 320

11,590 15,250 35,300

57GC19 2–4 15–17 20–22 35–37 50–52 55–57

NZA1519 NZA1522 OZA222U NZA1523 OZA223U OZA224U

3550 ^ 60 9800 ^ 90 9400 ^ 80 19; 200 ^ 170 25; 900 ^ 250 26; 000 ^ 700

11,340 10,850 22,640 30,370 30,490

Sample interval (cm)

a b c

Conventional radiocarbon age, less reservoir correction of 400 years (Bard, 1988). Calendar age, based on calibration by Bard et al. (1998): Cal. age (years B:P:† ˆ ⫺3:0126 × 10⫺6 [ 14C age] 2 ⫹ 1.289 [ 14C age] ⫺ 1005. Analysis done on organic carbon fraction. All other analyses were done on total carbonate fraction.

A possible cause for the positive 14C age bias of the carbonate fraction in our cores could be lateral input of resuspended older sediment components (see below). This is not inconceivable, considering the proximity of the continental shelf, which would have been exposed during the LGM. More detailed 14 C ages, preferably of single species foraminifera rather than of bulk carbonates which includes the fine fraction most likely to be subject to redistribution, would be required to resolve this problem. For the sake of consistency with previous work in this area (McCorkle et al., 1994), we will proceed with the age model based on oxygen isotope stratigraphy (Table 3).

3.2. Geochemistry The concentrations of CaCO3, Corg, U, Ba, Al and Th in the sediment cores are shown in Fig. 4. An obvious feature is the down-core variations of carbonate which are in opposition to those of the terrigenous components. Dissolution of carbonate as a controlling factor is unlikely, since the cores come from water depths well above that of the modern lysocline at ⬃3700 m in this region (Peterson and Prell, 1985) and show little visible evidence of corrosive bottom water (McCorkle et al., 1994). We will show below that variations of the carbonate concentrations in the cores are primarily controlled by the

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Fig. 3. Radiocarbon (calendar) ages vs depth in cores, together with linear sedimentation rates (solid lines) for oxygen isotope stages 1 and 2 as shown in Table 3. Open diamonds ˆ carbonate fraction, closed circles ˆ organic carbon fraction. The dotted horizontal lines are the event markers 2.0 and 3.0 as defined by Pisias et al. (1984), corresponding to oxygen isotope stage boundaries 1/2 and 2/3, respectively. The dashed line in core 53GC07 is the sedimentation rate based on the radiocarbon (calendar) ages during oxygen isotope stage 2 (LGM). Note that sedimentation rates based on 14C (calendar) ages would reduce the difference in sedimentation rates between the LGM and the Holocene, as compared to sedimentation rates based on oxygen isotope stratigraphy.

flux of non-carbonate terrigenous material acting as a diluent. There is no discernable pattern in the Corg concentrations, except for a slight downward decrease below

the sediment surface, presumably reflecting continued decomposition of organic matter upon burial (Froelich et al., 1979; Heggie et al., 1987). This decay of organic matter within the sediment is not always

Table 3 Oxygen isotope stratigraphy and linear sedimentation rates (McCorkle et al., 1994) Core

Stage

Depth (cm)

Age (ky)

Sedimentation rate (cm/ky)

53GC04

1 2 3⫹4 5 1–5

0–26 26–71 71–169 169–294 0–294

0–12 12–24 24–74 74–130 0–130

2.16 3.73 1.96 2.23 2.26

53GC07

1 2 3⫹4 5 1–5

0–29 29–71 71–171 171–319 0–319

0–12 12–24 24–74 74–130 0–130

2.37 3.52 2.01 2.65 2.46

57GC15

1 2 3⫹4 5 1–5

0–14 14–37 37–70 70–100 0–100

0–12 12–24 24–74 74–130 0–130

1.16 1.91 0.65 0.55 0.77

57GC19

1 2 3⫹4 5 1–5

0–22 22–53 53–116 116–190 0–190

0–12 12–24 24–74 74–130 0–130

1.83 2.53 1.28 1.32 1.46

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Fig. 4. Depth profiles of biogenic and terrigenous sediment components in sediment cores from the Exmouth Plateau and the Perth Basin. The open circles in the depth profiles of U and Ba denote the concentrations of U and Ba in detrital components. The horizontal lines are event markers (stage boundaries) based on oxygen isotope stratigraphy (McCorkle et al., 1994).

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Fig. 5. Comparing concentrations of Ba vs Al for sediment samples in cores from the Exmouth Plateau and the Perth Basin. The solid line denotes the Ba/Al ratio (0.0075) used to correct for detrital aluminosilicates in these cores. The dashed line represents the Ba/ Al ratio (0.0065) in average terrigenous shale from Australia (Taylor and McLennan, 1985).

easy to separate from changes in the supply of organic matter across the glacial to Holocene transition, especially in areas with slow sedimentation rate (Berger and Herguera, 1992), and can lead to erroneous interpretations of the sediment record in terms of changes in paleoproductivity. The only available d 13C values of Corg in these sediments are ⫺19.31 and ⫺18.68‰ at 2 and 48 cm, respectively, in core 53GC07 (obtained in connection with the AMS radiocarbon age determinations of the organic carbon fraction, E. Lawson, personal communication), which is typical of marine warm water plankton (Sackett, 1989) and well within the range previously reported for Corg in low-latitude open marine environments (Fontugne and Duplessy, 1986; Thunell et al., 1992). We therefore proceed with the assumption that the source for the Corg in our cores is predominantly marine. The U and Ba data are shown as both total and detrital concentrations (Fig. 4). Detrital U has been derived from the measured total U and Th concentrations by assuming that all Th resides in detrital phases, and that the average U/Th (w/w) ratio of detrital material is 0:27 ^ 0:07 (McCorkle et al., 1994). The large

excess of U over detrital U, starting a few centimeters below the sediment surface in cores 53GC04 and 53GC07, is authigenic U, a post-depositional feature which is related to the burial rate of organic matter and the development of anoxic conditions at depth in the sediment (McCorkle et al., 1994). Such uranium redox fronts close to the sediment surface are common in hemipelagic sediments (Thomson et al., 1990; Klinkhammer and Palmer, 1991; Veeh et al., 1999). The concentration of Ba in detrital phases has been based on the measured total Ba and Al concentrations and on the global average Ba/Al ratio of 0:0075 ^ 0:0025 in detrital aluminosilicates (Dymond et al., 1992). The total Ba concentrations lie well above those expected from association with aluminosilicates and show no relationship with Al (Figs. 4 and 5), indicating that most of the Ba in our cores is ‘biogenic’, i.e. it has been taken up by sinking biogenic particles in the water column (Dehairs et al., 1980; Bishop, 1988; Dymond et al., 1992). The concentration of biogenic Ba, or Ba[bio] in the present study, has been calculated as Ba[bio] ˆ Ba(total) ⫺ 0.0075Al. This normative method of deriving Ba[bio] is, of course, critically dependent on the true value of Ba/Al in detrital aluminosilicates in the area of study, which could well be different from the global average. However, the effect of possible variations in the Ba/Al ratio of detrital aluminosilicates on our normative Ba[bio] concentrations is minimized by the comparatively low (⬍30%) content of detrital components in our cores. 3.3. Excess

230

Th

Unsupported or excess 230Th in the deep sea sediments is routinely derived by correcting the measured total 230Th concentrations for 230Th associated with detrital U and assuming radioactive equilibrium between them. The occurrence of significant authigenic U in our cores requires an additional correction for in situ ingrowth of 230Th from this secondary U (see appendix). Correction for in situ ingrowth of 230Th involves assumptions about the source of the authigenic U as well as the time of its addition to the sediment (Francois et al., 1993). Although we believe (McCorkle et al., 1994) that the accumulation of authigenic U at depth in these cores, as in other

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Table 4 Radiochemical data Depth (cm)

Age (ky)

DBD (g/cm 3)

U (ppm)

Th (ppm)

230

Th (dpm/g)

230

Thx a (dpm/g)

230

Thxo b (dpm/g)

53GC04 1.5 11.5 24.5 38.5 48.5 58.5 68.5 78.5 88.5 98.5 123.5 143.5 163.5 183.5 203.5 223.5 243.5 263.5 283.5

0.7 5.3 11.3 15.4 18.0 20.7 23.4 27.8 32.9 38.0 50.8 61.0 71.2 80.7 89.7 98.7 107.6 117.1 125.6

0.84 0.80 0.82 0.89 0.89 0.89 0.90 0.92 0.88 0.88 0.84 0.83 0.85 0.86 0.86 0.87 0.87 0.84 0.85

1.61 2.59 4.44 4.70 4.23 3.58 2.89 3.55 3.41 3.63 4.23 3.57 3.72 2.82 4.30 4.24 4.35 4.42 3.81

3.04 2.54 1.52 1.02 1.12 1.28 0.92 1.13 1.09 1.01 1.42 1.66 1.65 1.71 1.49 1.09 1.26 2.49 3.27

2.57 1.95 1.87 2.14 2.12 2.01 1.76 2.23 2.12 2.08 2.31 2.17 2.32 2.07 2.79 2.62 2.68 3.04 2.60

1.97 1.45 1.60 1.94 1.90 1.76 1.58 1.94 1.73 1.56 1.34 1.10 1.04 0.94 1.03 0.79 0.66 0.76 0.48

1.98 1.52 1.78 2.24 2.24 2.13 1.96 2.51 2.34 2.21 2.14 1.93 2.00 1.98 2.35 1.96 1.78 2.24 1.53

53GC07 2 12 23 33 43 53 63 73 84 94 106 116 126 146 166 186 206 216 226 236 246 266 286 306

0.8 5.1 9.7 13.3 16.1 18.9 21.8 25.0 30.5 35.4 41.4 46.4 51.4 61.3 71.3 79.7 87.2 91.0 94.8 98.5 102.3 109.8 117.4 124.2

0.79 0.82 0.80 0.88 0.76 0.74 0.74 0.75 0.78 0.84 0.76 0.76 0.80 0.78 0.80 0.82 0.84 0.84 0.86 0.85 0.84 0.80 0.84 0.79

0.89 0.77 0.70 0.54 0.47 0.39 2.15 3.46 2.36 1.43 1.52 1.79 3.23 1.63 1.76 1.43 0.94 1.70 0.89 0.72 0.75 1.23 0.88 1.02

2.44 2.34 1.83 1.32 1.11 1.63 1.63 1.83 1.73 1.42 1.32 1.83 1.73 2.03 2.03 1.63 1.73 1.72 1.02 1.42 1.63 2.54 2.84 3.46

6.36 5.69 5.45 4.81 4.56 4.73 4.49 4.63 4.20 3.72 3.46 3.56 3.63 3.39 3.26 2.82 2.74 3.01 2.17 2.03 2.21 2.82 2.50 2.41

5.88 5.23 5.09 4.55 4.34 4.41 4.17 4.27 3.76 3.36 3.06 3.00 2.79 2.73 2.50 2.17 2.23 2.26 1.72 1.61 1.75 2.07 1.89 1.68

5.92 5.48 5.57 5.14 5.03 5.25 5.10 5.38 5.01 4.66 4.48 4.60 4.48 4.80 4.82 4.52 4.96 5.23 4.12 3.99 4.49 5.69 5.58 5.28

57GC15 2 7 12 17

1.7 6.0 10.3 13.5

0.89 0.93 0.93 1.03

0.84 1.01 1.06 1.04

2.26 2.91 1.90 0.93

7.59 7.41 6.72 6.81

7.14 6.83 6.31 6.54

7.25 7.23 6.98 7.50

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Table 4 (continued) Depth (cm)

Age (ky)

DBD (g/cm 3)

U (ppm)

Th (ppm)

230

Th (dpm/g)

230

Thx a (dpm/g)

Thxo b (dpm/g)

22 27 32 37 42 47 52 57 62 67 72 77 82 87 92 97

16.2 18.8 21.4 24.0 31.7 39.4 47.1 54.8 62.5 70.5 78.5 87.6 96.7 105.8 114.9 124.0

1.08 1.06 1.02 1.06 1.06 1.03 1.02 1.00 0.99 1.01 1.03 1.04 1.05 1.03 1.04 0.95

1.11 0.76 0.76 0.69 0.67 0.69 0.90 0.60 0.64 0.72 0.86 0.61 0.65 0.78 0.86 1.10

0.76 0.95 1.07 0.75 0.89 0.94 1.12 1.25 1.27 0.89 1.19 1.19 1.33 1.45 1.92 2.95

6.37 6.77 6.97 5.14 4.93 4.43 4.14 4.22 3.53 3.33 3.32 2.96 2.94 3.04 3.09 2.76

6.11 6.51 6.69 4.91 4.66 4.13 3.74 3.89 3.15 2.96 2.85 2.59 2.53 2.55 2.52 2.00

7.22 7.82 8.23 6.23 6.36 6.10 6.05 6.60 5.64 6.03 6.35 6.10 6.53 7.32 7.81 6.84

57GC19 3 11 16 26 36 41 46 56 61 67 77 94 111 133 153 173

1.6 6.0 8.7 13.6 17.5 19.5 21.5 26.7 30.6 35.3 43.1 56.4 69.7 86.9 102.0 117.2

0.78 0.80 0.87 0.96 0.91 0.93 0.90 0.98 0.96 0.95 0.96 0.98 0.83 0.99 1.02 0.93

0.63 0.51 0.65 0.57 0.49 0.44 0.49 0.48 0.48 0.38 0.30 0.43 0.66 0.37 0.53 0.53

1.52 2.24 1.73 1.12 1.12 1.12 0.90 0.91 0.91 1.22 1.01 1.12 1.52 1.12 1.12 1.12

7.03 6.66 6.58 6.95 7.48 7.15 6.96 5.60 5.78 5.04 4.66 4.47 4.56 3.43 3.07 2.89

6.73 6.28 6.23 6.70 7.26 6.91 6.79 5.38 5.55 4.79 4.45 4.21 4.16 3.18 2.73 2.54

6.83 6.64 6.76 7.63 8.57 8.29 8.34 6.93 7.42 6.65 6.64 7.15 8.10 7.15 7.30 7.86

Thx ˆ excess 230Th, corrected for detrital 230Th and for in situ ingrowth of Thxo ˆ decay-corrected initial excess 230Th (see Appendix).

a 230

230

230

Th (see Appendix).

b 230

hemipelagic sediments, is most likely due to postdepositional diffusion into the sediment from overlying seawater along a pore water gradient, the timing of this process is not well defined. The time lag Dt between the age of the host sediment and the time of secondary U input should be approximately equal to the age of the sediment corresponding to the shallowest peak near the leading edge of the U redox front, assuming that the U redox fronts in cores 53GC04 and 53GC07 are still ‘active’ (McCorkle et al., 1994). Taking a value of 25 ky as

an upper limit for Dt, based on the approximate age of the sediment near the leading edge of the redox front in 53GC07, we can evaluate the effect of in situ ingrowth of 230Th on the 230Thx data (see Appendix). Except for core 53GC04, the contribution of in situ 230 Th ingrowth is comparatively small, and errors in the 230Thx data introduced by uncertainties in Dt would be no more than about 10%. For core 53GC04, however, such uncertainties could be quite large, because the 230Th generated by in situ decay of authigenic U is a significant proportion of the total

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Fig. 6. Excess 230Th vs depth in sediments. Closed circles ˆ excess 230Th ( 230Thx), open circles ˆ decay-corrected initial excess 230Th ( 230Thxo). The average sedimentation rates shown have been calculated from the exponential decrease (slope of regression line) of 230Thx with depth in sediment and are based on the assumption of constant 230Th flux from the water column. Note the compressed oxygen isotope stratigraphy and much slower sedimentation rate in core 57GC15, evidently due to post-depositional erosion by winnowing (see text).

excess 230Th due to the shallow depth of this core. However, even relatively large errors in Dt would not substantially alter the conclusions of this study which focus on the Glacial–Holocene transition, where the effect on the 230Thx data would be minimal. The authigenic U concentrations in cores 57GC15 and 57GC19 are so low that the correction for in situ ingrowth of 230Th is negligible. The 230Thx data together with the decay-corrected initial 230Thx data ( 230 Thxo) are shown in Table 4 and Fig. 6. The long term average sedimentation rates based on the 230Thx data are comparable to the average sedimentation rates (stage 1–5) based on oxygen isotope stratigraphy as previously defined (Table 3), giving us some confidence that the ingrowth-corrected 230 Thx data and our assignment of oxygen isotope stage 5/6 boundary (event 6.0) are mutually compatible. 3.4. Mass accumulation rates The mass accumulation rates of individual sediment components can be calculated according to: MAR i ˆ fi × DBD × S

…1†

where fi is the weight fraction of sediment component,

DBD the dry bulk density and S the linear sedimentation rate. The MARs determined by this method are, in effect, the mean burial fluxes of individual sediment components during the given oxygen isotope stages at the respective sites, making no allowance for possible variations of the MARs within a given isotope stage, nor for possible sediment redistribution prior to burial. Another, independent method of assessing mass accumulation rate is based on the well-supported assumption that 230Th, produced at a constant rate in the water column, is rapidly transferred to the sea floor (constant flux model). In this method, the 230Th flux to the sea floor is used as a constant flux tracer against which the paleoflux of other sediment components is evaluated by normalizing their concentration at a given depth in the sediment to 230Thxo (Bacon, 1984; Suman and Bacon, 1989). Thus: Fi ˆ fi × b × z= 230 Thxo

…2†

where Fi is the flux of the individual sediment components, fi the weight fraction of the sediment component, b the production rate of 230Th from 234U in the water column equal to 2:63 × 10⫺5 (dpm/cm 2/ky), z

22

H.H. Veeh et al. / Marine Geology 166 (2000) 11–30

Fig. 7. Mass accumulation rates of biogenic and terrigenous sediment components; Solid lines ˆ site-specific burial rates, based on mass accumulated between oxygen isotope stage boundaries, dotted lines ˆ regional rain rates, based on point-by-point normalization to 230Thxo under the constant 230Th flux model. Fa/Fp is a ‘focusing factor’, based on the inventory of 230Thxo in the sediment (Fa) and the known production of 230Th in the water column (Fp) above each site.

H.H. Veeh et al. / Marine Geology 166 (2000) 11–30 230

230

the depth of water (cm) and Thxo the excess Th corrected for decay since the time of deposition. The conceptual advantage of this method is that it should be relatively insensitive to lateral sediment redistribution on the sea floor, because it is a function of the regional ‘rain rate’ of sediment components rather than a measure of their site-specific burial rate at a given location (Francois et al., 1990). An added advantage of this point-by-point reconstruction of the MAR of different sediment components is that it provides superior time resolution than the more traditional method of determining the site-specific burial flux of sediment components between dated horizons in a sediment core. The difference between the mass accumulation rates determined by these two independent methods is striking (Fig. 7), in particular in relation to the LGM, and the implications for the reconstruction of paleoproductivity would be quite different, depending on which method is adopted. The MARs determined by the conventional method show the burial fluxes of biogenic sediment components to be substantially higher at all core locations during the LGM as compared to the Holocene, as in previous findings (McCorkle et al., 1994). By contrast, there is little significant difference between the Holocene and the LGM in terms of the 230Thxo normalized fluxes of biogenic sediment components. Another, more subtle difference between the two methods is revealed by the MARs of the terrigenous components. Whereas the normalized fluxes of Al and Th indicate a minimum of terrigenous input during the LGM, the burial fluxes of Al and Th would indicate a minimum input during stage 3. We can think a several reasons for these differences: 1. The age model based on oxygen isotope stratigraphy is incorrect, at least with respect to the oxygen isotope stage boundaries defining the LGM. 2. The increased burial rates during the LGM reflect increased lateral sediment input in addition to the settling rate of sediment components from the water column (focusing). 3. The assumption underlying the constant flux model of 230Th is invalid due to enhanced scavenging in the water column (boundary scavenging) during the LGM. Although the SPECMAP age model based on

23

orbital tuning of the oxygen isotope record in deep sea sediments is widely used, it is not well calibrated by reliable absolute dating and specified stage boundaries may incorporate large errors. Whilst event 2.0 is relatively well defined, this is not so for event 3.0 which does not always permit precise identification in cores with low stratigraphic resolution (McCorkle et al., 1994) and which has an assigned age uncertainty of ^4930 years (Martinson et al., 1987). An uncertainty of this magnitude would have a comparatively large effect on the mass accumulation rate during the LGM when calculated by conventional methods. We can consider a test case by recalculating the MARs (both methods) in core 53GC07 on the basis of the 14C (cal) ages. The revised age model in this test case mainly affects the burial rates, whilst the effect on 230Thxo-normalized fluxes is negligible (Fig. 8). Although the pattern of the recalculated burial rates is now more compatible with that of the 230 Thxo-normalized fluxes, in particular with respect to the LGM, the burial rates are still consistently higher than the 230Thxo-normalized fluxes. Similar results can be expected for the other cores with significant differences between 14C ages and oxygen isotope stratigraphy. We therefore have to look elsewhere to explain the discordance between the burial rates and the normalized fluxes. Focusing (alternative 2) can be recognized by comparing the predicted flux (Fp) of 230Th from the water column to the measured inventory (Fa) of 230 Thxo in the sediment at each core location (Cochran and Osmond, 1976; Suman and Bacon, 1989; Francois et al., 1993; Thomson et al., 1993; Scholten et al., 1994). Thus: Z  230 Thxo × DBD × dr =b × …t2 ⫺ t1 † …3† Fa =Fp ˆ where t1 and t2 are the ages of horizons r1 and r2 in the sediment, the other terms of this equation as previously defined. In the present context, a value of F a =F p ⬎ 1 would indicate net lateral input of resuspended sediment components containing excess 230Th in addition to the vertical flux of particulate 230Th out of the water column, whilst F a =F p ⬍ 1 would indicate periods of non-deposition, or even net erosion e.g. by winnowing due to bottom currents (Fig. 9a). The Fa/Fp ratios shown in Fig. 6 suggest that the sediments on the Exmouth Plateau and in the Perth

24

H.H. Veeh et al. / Marine Geology 166 (2000) 11–30

Fig. 8. Mass accumulation rates in core 53GC07, based on 14C (calendar) ages. Note that burial rates during the LGM are greatly reduced in comparison to the burial rates shown in Fig. 6, but that the effect of the revised age model on the 230Thxo-normalized rates would be negligible.

Fig. 9. Schematic diagram illustrating different interpretations of the focusing factor Fa/Fp in terms of lateral sediment transport (a) and boundary scavenging (b). Note that lateral sediment redistribution can take place either by bottom currents or as horizontal plumes in the water column, and that the magnitude of boundary scavenging should increase with water depth due to the increase in dissolved 230Th (and corresponding increase in lateral advection of 230Th) with depth in the water column (see text).

H.H. Veeh et al. / Marine Geology 166 (2000) 11–30

Basin have been subjected to focusing, in particular during the LGM. This should not be surprising, considering the location of the cores on or near the continental slope, an environment where sediment resuspension and lateral redistribution by bottom currents and/or in the water column as horizontal plumes along isopycnals can be expected (Brewer et al., 1980; Honjo et al., 1982; Heggie et al., 1987; Biscaye and Anderson, 1994; Laine et al., 1994). Indeed, the F a =F p ⬍ 1 prior to stage 2 in core 57GC15 is a clear indication of local sediment erosion by winnowing (‘negative focusing’), which also explains why the burial rates during stages 3 ⫹ 4 and 5 in this core are much lower than in core 57GC19 from the same general area and depth. It is worth noting here that the MARs based on normalization to 230Thxo in this core do not show the effect of winnowing and thus provide a more reliable insight into the flux of sediment components from the water column than the burial rates. Variable boundary scavenging (alternative 3) would be a different interpretation of Fa/Fp (Fig. 9b). Although there is little doubt that the accumulation of excess 230Th in the deep sea sediments is balanced by its production in the water column in the global ocean (constant flux model), such balance may not be strictly observed in a local context. There is mounting evidence for enhanced scavenging, or ‘boundary scavenging’ of 230Th, near continental margins, as well as for reduced scavenging of 230Th in the open oligotrophic ocean, although the short residence time of 230Th in seawater (⬍50 years) limits the distance of laterally advected 230Th and hence the magnitude of boundary scavenging even in areas of high particle flux (Anderson et al., 1983a,b; Shimmield et al., 1986; Bacon, 1988; Taguchi et al., 1989; Lao et al., 1992; Yu et al., 1996). Although we cannot rule it out entirely, it is unlikely that boundary scavenging of variable intensity has been the major cause for the observed variations of Fa/Fp in our cores for the following reasons. The Fa/Fp ratios during the LGM range from 2.2 to 2.8 (Fig. 7), implying enhanced scavenging of 230Th from 120 to 180%. Since boundary scavenging depends on the lateral import of dissolved 230Th by advection from areas of low particle flux, it is difficult to see where such a large excess of 230Th during the LGM could have come from in view of its short residence

25

time. Inasmuch as the concentration of dissolved 230 Th in the water column increases with water depth in the open ocean (Nozaki et al., 1981; Nozaki et al., 1987; Cochran, 1992), we would expect to see a systematic increase in Fa/Fp with water depth at our core locations, but this is not the case (Fig. 7). We therefore proceed with the preferred interpretation that lateral import of resuspended sediment components is the predominant process responsible for the enhanced burial fluxes during the LGM. This interpretation would not only account for the higher burial fluxes of biogenic components, mainly driven by higher sedimentation rates during the LGM, but also for the somewhat anomalous burial flux variations of the terrigenous components which would be difficult to explain otherwise. In addition, lateral import of recycled (and older) sediment components would also provide a ready explanation for the anomalous 14C ages of the carbonate fraction in our cores (Fig. 3). A similar offset between 14C ages of the carbonate and organic carbon fractions has been reported for slope sediments of the Middle Atlantic Bight, where sediment focusing is at least partly responsible for the abnormally high Fa/Fp ratios measured in both sediments and sediment traps (Anderson et al., 1994a). A plausible source of the imported sediment is the upper continental slope and/or shelf (Anderson et al., 1994b; Biscaye and Anderson, 1994; Summerhayes et al., 1995). This could include the redeposition of sediment initially deposited at shallower depths with correspondingly lower 230Thxo concentrations, which would effect the interpretation of the 230Thxo-normalized fluxes. Francois et al. (1990) have examined the influence of post-depositional redistribution of sediments on the interpretation of normalized fluxes in some detail. If there has been no fractionation between different sediment components during resuspension and downslope transport, the effect on normalization would be minor, and the normalized fluxes would only slightly overestimate the true pelagic fluxes. On the other hand, if there has been selective resuspension of the fine sediment components, the most likely carrier phase of the excess 230 Th in the sediment (Francois et al., 1990; Thomson et al., 1993), it would have the opposite effect on normalization, and the normalized fluxes could

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H.H. Veeh et al. / Marine Geology 166 (2000) 11–30

somewhat underestimate the true pelagic fluxes. Thus the effects of lateral sediment redistribution on the normalized fluxes tend to compensate each other. Based on the above discussion, and for the purpose of the present study, we will consider the 230Thxonormalized fluxes shown in Fig. 7 as a more reliable and self-consistent estimate of the actual settling fluxes of biogenic and terrigenous sediment components than the burial fluxes. 3.5. Implications for paleoceanography and paleoclimate The 230Thxo-normalized fluxes of the biogenic sediment components (Fig. 7) do not provide convincing support for greatly enhanced paleoproductivity and hence coastal upwelling off Western Australia during the LGM. Neither the CaCO3 nor the Corg data show much evidence for increased export production at that time, although the glacial to Holocene change in the paleoflux of Corg may be obscured by the continued decomposition of organic matter within the Holocene section of each core (see above). Only the Ba data at one core location (53GC07) appear to provide limited support for somewhat increased productivity during the LGM as compared to the Holocene. However, the reliability of biogenic Ba as a quantitative paleoproductivity proxy can be questioned, based on recent evidence of early diagenetic remobilization of authigenic Ba near continental margins (Kumar et al., 1996; McManus et al., 1999). Any support for increased productivity associated with coastal upwelling off Western Australia during the LGM would therefore be based on only two of several productivity proxies, namely the previously cited abundance of benthic foraminifera (McCorkle et al., 1994), and possibly the 230Thxo-normalized fluxes of biogenic Ba. Of these, the abundance of benthic foraminifera would be a more reliable productivity proxy, as it is robust and not subject to diagenetic modification. Although there is little doubt that the abundance of benthic foraminifera on the sea floor reflects the supply of organic detritus (i.e. food) settling out of the water column, the abundance record of benthic foraminifera as a paleoproductivity record in sediments near continental margins is more

difficult to interpret because of the proximity of the glacially exposed shelf as a potential source of recycled organic matter (Berger and Herguera, 1992). Thus the enhanced MARs of benthic foraminifera during the LGM as compared to the Holocene previously reported (McCorkle et al., 1994) may not be solely a reflection of increased primary productivity, but could in part be due to recycled organic matter. In other words, the abundance pattern of benthic foraminifera in our cores may not be immune from problems associated with sediment focusing. To sum up, there is no compelling evidence for strong coastal upwelling off Western Australia during the LGM similar to that observed in the modern ocean off the west coasts of Africa and South America. Our interpretation of the sediment record off Western Australia therefore does not provide unambiguous support for the hypothesis that the Leeuwin current had ceased flowing southward along the West Australian coast during the LGM as part of a major reorganization of ocean circulation of the eastern Indian Ocean (Prell et al., 1980; Wells and Wells, 1994; Wells et al., 1994). The MARs of Al and Th (Fig. 7) show major variations which are out of phase with that of the biogenic sediment components, indicating a higher supply of terrigenous sediment components to the Indian Ocean off Western Australia during the Holocene as compared to the LGM. This is somewhat counterintuitive, as one would have expected an increased aridity signal (eolian dust) in this area during the LGM, in analogy to the clear correlation between eolian dust flux maxima and glacial stages recorded in pelagic sediments downwind from major deserts (Rea, 1994), and in view of the more extensive aridity in northwestern Australia during the LGM as compared to the Holocene (Williams et al., 1993). A concentration gradient of kaolinite in pelagic sediments extending westward from Australia towards the center of the Indian Ocean has been attributed to predominantly airborne transport from Australian deserts by southeasterly winds (Griffith et al., 1968). However, all of our core locations are on the continental margin and hence under the influence of hemipelagic processes which may well mask the eolian aridity signal. Horizontal advection of terrigenous material within the water column near

H.H. Veeh et al. / Marine Geology 166 (2000) 11–30

continental margins is well documented by sediment trap studies (Brewer et al., 1980; Honjo et al., 1982; Heggie et al., 1987), and a similar process can be envisioned for the West Australian continental margin. In general, hemipelagic fluxes are much larger than eolian fluxes and tend to dominate the supply of terrigenous components to the deep sea floor for several hundred kilometers offshore, even in the vicinity of arid continents (Rea, 1994). Although there are no major rivers in Western Australia, the northwestern part of Australia is under the influence of a monsoonal climate with frequent cyclones, and increased runoff can be expected towards the north during the austral summer (Gentilli, 1991). Inasmuch as the monsoonal circulation is more intense during interglacials, than during glacials (Prell et al., 1980; Williams et al., 1993), higher rainfall and/ or increased chemical weathering during interglacials is a likely cause for the increased supply of terrigenous components to hemipelagic sediments in this area. The increasing MARs with increasing proximity of the core locations to the wetter part of northwestern Australia probably reflect this potential source for the terrigenous components.

4. Conclusions The mass accumulation rates of biogenic and terrigenous sediment components based on normalization to the constant 230Th flux tracer in the water column provide a more realistic paleoceanographic and paleoclimatic sediment record for the West Australian continental margin than the site specific burial rates, because the latter appear to be compromised by lateral sediment redistribution (focusing) which can lead to misleading interpretations. There is no unambiguous support for enhanced productivity off Western Australia during the LGM, and hence no compelling evidence in the sediment record for strong coastal upwelling comparable to that in the modern ocean off the west coasts of Africa and South America. The hypothesis of a major reorganization of ocean circulation in the southeastern Indian Ocean which involves the replacement of the south flowing Leeuwin Current with the north flowing West Australian Current during the LGM should therefore be viewed with caution.

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The mass accumulation rates of terrigenous sediment components show a distinctive pattern, with significantly lower fluxes during the LGM as compared to the Holocene. This is at variance with the glacial/interglacial pattern of eolian mineral deposition in most pelagic sediments and suggests that the eolian dust record on the West Australian continental margin is masked by local processes reflecting increased wetness associated with the monsoonal climate of northwestern Australia during interglacials. Acknowledgements We thank the officers, crew and scientists of R/V Rig Seismic Survey 53 (H. Stagg, Chief Scientist) and Survey 57 (J. Marshall, Chief Scientist) for collecting the cores used in this study as part of the Continental Margin Program of the Australian Geological Survey Organisation. We are grateful to C. Tuniz and E. Lawson (Australian Nuclear Science and Technology Orgnaisation), and R. Sparks (New Zealand Institute of Geological and Nuclear Sciences) for the radiocarbon (AMS) age determinations. The XRF analyses were carried out by J. Pyke (AGSO). We thank E. Barnett and N. Corlis (Flinders University) and J. Lane (AGSO) for laboratory assistance. J. Thomson and J. Scholten provided constructive reviews and comments on this manuscript. H. Veeh received financial support from the Australian Research Council and the Australian Institute of Nuclear Science and Engineering. D. McCorkle was supported by NSF grant OCE-8817620. Appendix A Calculation of excess 230

230

Th:

Thx ˆ 230 Th ⫺ 230 Thd ⫺ 230 Thin

…A1†

230 Thd ˆ detrital 230 Th ˆ 0:74 × 0:27 × Th where (see note 1). 230 Thin ˆ in situ ingrowth of 230Th from authigenic U

ˆ 238 Ua …1 ⫺ exp…⫺l230 t† ⫹ …234 Ua ⫺238 Ua †  …l230 =…l230 ⫺ l234 ††…exp…⫺l234 t† ⫺ exp…⫺l230 t†† …A2†

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where Ua ˆ authigenic U ˆ 1:14 × 238 Ua (assuming authigenic uranium has been derived from seawater), 238Ua ˆ authigenic 238 U ˆ 0:74 …U ⫺ 0:27 × Th†; l 234 the decay constant of 234U, l 230 the decay constant of 230Th and t the time of addition of authigenic U (see note 2). Calculation of initial excess 230Th: 234

230

234

Thxo ˆ 230 Thx =exp…⫺l230 t†

…A3†

where t is the age of the host sediment. Note 1: In the derivation of detrital 230Th, it is assumed that the measured Th (ppm) is of detrital origin, that the U/Th (weight ratio) in average detrital material is 0.27 (McCorkle et al., 1994) and that 230Th, 234 U, and 238U in the detrital material are in secular equilibrium. The factor 0.74 is needed to convert U (ppm) to 238U (dpm/g). Note 2: The time of addition of authigenic U is bounded by t0 (present addition) and ta the age of sediment (addition at time of deposition of host sediment). The distribution pattern of authigenic U in the cores suggests that the secondary U continues to diffuse downward into the sediment for some time Dt after deposition of the host sediment and then stops. Thus the value of t for the ingrowth correction can be determined according to t ˆ ta ⫺ Dt: The value of Dt has been estimated at about 25 ky, based on the depth of the leading edge of the U redox front in core 53GC07 and should be an upper limit (McCorkle et al., 1994).

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