Distribution of major, trace and rare-earth elements in surface sediments of the Wharton Basin, Indian Ocean

CHEMICAL GEOLOGY i x , m I)Z~f,

ISO TOPE GEO,S'('IEVCE

ELSEVIER

Chemical Geology 121 ( 1995 ) 201-215

Distribution of major, trace and rare-earth elements in surface sediments of the Wharton Basin, Indian Ocean J.N. Pattan a, Ch.M. Rao a, N.C. Higgs b, S. Colley b, G. Parthiban a aNational Institute of Oceanography, Dona Paula, Goa 403 004, India blnstitute of Oceanographic Sciences, Deacon Laboratory, Wormley, Godalming, GU8 5 UB, UK

Received 29 October 1993; revision accepted 1 September 1994

Abstract Bulk chemical analyses of eighteen surface sediments taken on an east-west transect (lat. 15 o40'S) in the Wharton Basin, Indian Ocean, were analysed for major, trace and rare-earth elements (REE) by inductively coupled plasma-atomic emission spectrometry (ICP-AES). The relationship between A1 and K suggests in situ formation ofauthigenic phillipsite. The relationship between P and Ca suggests accumulation of biogenic apatite/fish bone debris from high biological productivity waters, which is confirmed by the linear relationship with P and Cu. A strong positive correlation among redox sensitive metals (Mn, Fe, Cu, Ni and Zn) suggests that they are incorporated into a ferromanganese oxide phase, probably micronodules. Calcareous oozes from the transect have low REE abundance (La ~ 8 ppm ) with a very strong negative Ce/Ce* value (0.41), siliceous oozes have high REE abundance (La ~28 ppm) with moderate positive Ce/Ce" (1.24) and red clays have highest REE abundance (La ~ 66 ppm) with a weak positive (1.05) to moderate negative Ce/ Ce* (0.73). Shale-normalized REE patterns and (La/Yb)n ratios record similar enrichment of heavy REE (HREE) relative to light REE (LREE) in all sediment types. This might be due to retention of the bottom water REE pattern and the presence of fish bone debris. Correlation coefficient and R-mode factor analyses suggest a very strong positive association (R 2> 0.9) of REE with P, and thus most likely with biogenic apatite/fish bone debris. As P is directly associated with biological activity it is thus related to surface primary productivity. REE in the Wharton Basin sediments appear to reflect a combination of surface water effects and diagenetic processes. 1. Introduction

Deep-sea sediments are composed of varying amounts of calcium carbonate ( C a C O 3 ) , biogenic silica (opal), detrital material, Fe-Mnoxyhydroxides and hydrothermal material. The proportions of these components vary according to water depth, proximity to terrestrial and hydrothermal sources and the biological productiv[SBI

ity of the overlying waters. The abundance of chemical elements has been used in defining sediment sources, elucidating mechanisms of formation of the sediments, estimating abundance of different components, quantifying authigenic deposition rates and fluxes of various elements, and in understanding depositional environments (Goldberg and Arrhenius, 1958; Krishnaswami, 1976; Graybeal and Heath, 1984; Thomson et al., 1984; Toyoda and Masuda 1990; and references therein). In particular the rare-

0009-2541/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSD10009-2541 ( 94 )00112-X

202

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

earth elements (REE) form a coherent group with trivalent oxidation state, except for Ce, which oxidises to the tetravalent state and Eu, which reduces to the divalent state. REE are widely used as indicators of various geochemical processes including depositional environment (e.g., Piper, 1974; Taylor and McLennan 1985; Murray et al., 1992), redox conditions (Glasby et al., 1987; Liu et al., 1988 ), surface productivity (Toyoda et al., 1990), and to trace aeolian and hydrothermal input (Elderfield, 1988). Sediment from the Wharton Basin has not been studied to the same extent as sediment from other basins in the Indian Ocean and there are no previous reports on REE in sediment from this basin. In the present work sediments from an east-west transect located at 15 ° 4 0 " representing various lithologies (including calcareous ooze, siliceous and red clays) with variable productivity, aeolian input and influence of Antarctic Bottom Water have been studied to address the source, abundance, mode of incorporation and factors controlling the distribution of major elements, trace metals as well as REE fractionation.

I00° ---J-

90 °

I

io%

• I0"

Area of....~tudy~

L.

o'.

~

//~iff

"°~

WHARTON

BASIN

io'-

- IO*

2. Materials and methods

The uppermost surface sediments (0-5 cm) were sampled from gravity cores along the E - W transect of the Wharton Basin, Indian Ocean (Fig. 1 ), collected during a cruise by the Russian RV "Yuzhmorgeologia" under an I N D O - U S S R joint research programme. The sediments were washed with deionized water: as a result, some of the adsorbed fraction or loosely bound organics might have been lost. After washing, sediments were oven dried at 60 ° C, and powdered in an agate mortar. A salt correction was not applied to the sediment chemical analyses. Digestion of the sediment and cation-exchange separation of REE was carded out according to the procedure of Jarvis and Jarvis ( 1985 ). Major and trace elements and REE were analysed with simultaneous inductively coupled plasma-atomic emission spectrometry (ICP-AES) using a Philips ® PV-8060 at the Institute of Oceanographic

2C

~°

,go"

ff

Fig. 1. Schematic m a p showing the locations of surface sediments along the E - W transect in the W h a r t o n Basin. Contour lines are in meters.

Sciences, Deacon Laboratory, U.K. C a C O 3 w a s determined by evolution of C O 2 using orthophosphoric acid and subsequent determination of CO2 using a Coulometrics ® Coulometer. Si was analysed by ICP-AES following lithium metaborate fusion (Ingamells, 1966). The international sediment standards BCSS-1 and MESSI were used to assess the accuracy of the analyses: the accuracy agreed within the precision of the analysis (Si, 20.76%; AI, +_4%; Ti, +_4.6%; Fe, +2.1%; Mn, 20.86%; Mg, 21.2%; Ca, +-0.2%; Na, +-7.6%; K, +-9.8%; P, +_6.6%; Cu, +_2.1%; Ni, 22.6%; Pb, +_4%; Zn, +_8.1%; La,

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

___3.6%; Ce, _+0.6%; Nd, _+2.3%; Sm, _ 2.6%; Eu, ___1.6%; Gd, +2.3%; Dy, +0.2%; Er, +6.8%; Ho, +7.2%, Yb, + 1.5%; and Lu, < _+0.1%). Statistical analysis was carried out using an in-house program at the National Institute of Oceanography, Goa (Fernandez and Mahadevan, 1982).

3. Geology of the area

The Wharton Basin is part of the Indian Ocean and is bordered by Ninetyeast Ridge in the west, Broken Ridge in the south, the Australian continent in the east, and in the north by the Cocos Rise, Sumatra and the Indonesian archipelago. This basin lies beneath areas of higher biological productivity (Parsons et al., 1977 ) and low sedimentation rates (Lisitzin, 1972). The sea floor is carpeted by siliceous ooze, red clay and calcareous ooze (Udintsev, 1975). Smectite, which constitutes the bulk of the fine fraction along with moderate amounts of kaolinite, is probably derived from the alteration of local volcanic material and from atmospheric fallout (Venkatrathnam and Biscaye, 1973). Ferromanganese nodules recovered in this basin have Mn, Cu and Ni contents similar to Central Indian Basin nodule average but in general are depleted in Fe, Co and Pb (Cronan and Moorby, 1981 ). The primary productivity of the water column above the siliceous ooze and red clay is comparatively high ( 150-250 mg m -2 day-1 C) when compared to the calcareous ooze from the Ninetyeast Ridge ( < 100 mg m -2 day -1 C; FAO, 1981 ).

4. Results and discussion

The results of bulk chemical analyses are normalized to a water- and C02-free basis and presented in Table 1. Interelemental correlations between major and trace elements and the REE of the entire data set are listed in Table 2. The average bulk chemical compositions of sediments from the Wharton Basin, Central Indian Basin (CIB), and Central Pacific Basin are presented in Table 3 for comparison.

203

4.1. M a j o r a n d trace elements

In deep-sea sedihaent, silica is derived mainly from lithogenous and biogenic sources. The siliceous oozes have higher SiO2 content (60.372.5%) than red clays (53.9-65.8%). These concentrations are nearly similar to those of Central Pacific basin sediments (Nohara and Kato, 1985) but are lower than those of the CIB (Table 3 ). The calcareous oozes have very low SiO2 contents (0.09-8.8%). The correlation matrix for siliceous ooze and red clay shows a strong negative relation between Si and AI, suggesting that these two elements have different sources, while there is a positive correlation between these two elements in the calcareous ooze. This positive correlation might be due to dilution by carbonate, which can seen in the factor analysis (Fig. 2). The "excess" silica computed for siliceous sediments and red clay using the method of Leinen (1977) showed that the former has more opal (21.5-49.6%) than the latter (5.5-12.55%) and is considered to be of biogenic origin. This biogenic silica most likely reflects the proximity of the subequatorial divergence with its associated high biological productivity (Parsons et al., 1977). Stations 73, 77 and 78 are located above the calcium carbonate depth (CCD) on the Ninetyeast Ridge. Sediments at these stations have very high carbonate contents ranging from 51.4% to 96.2% (Table 1). In contrast, the siliceous oozes and red clays are below the CCD and have a very low Ca content (Table 1 ). Average Sr content is low in siliceous ooze (85 ppm) and red clay ( 110 ppm) and highest in calcareous ooze (1017 ppm). Sr shows strong positive correlation with Ca (Table 2 ), reflecting its well-known association with calcite in the carbonate phase (e.g., Pingitore et al., 1992; Murray and Leinen, 1993). The association of Ca and Sr is clearly brought out in R-mode factor analysis (Fig. 2). Factor 1 accounts for 44% of the variance which reflects the carbonate phase. This factor has comparatively high negative loadings on Ti, Li, Cu, A1, Mg, Fe, Pb and Zn, as well as moderate negative loadings on Cr, Ni, K, Co, Mn, Si and Ce. This suggests that all these negatively loaded

204

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

Table 1 Bulk chemical analyses of major (%), trace and rare-earth elments (ppm) of surface sediments from the Wharton Basin, Indian Ocean (values are renormalised to water- and CO2-free basis) Sample Sediment (%) No. type Si02

A1203

Fe203

MnO MgO

CaO

Na20

K20

TiOe

P:O5

Ba

Co

Cu

46 47 48 51 52 55 56 57 58 60 62 63 70 73 74 75 77 78

8.78 19.10 21.75 20.15 13.91 16.75 19.53 16.98 16.96 7.64 19.26 16.23 15.19 1.74 18.19 15.44 0.93 0.09

5.30 8.04 9.40 10.76 6.24 7.20 10.29 7.50 7.72 3.12 10.08 7.31 6.74 0.75 8.04 6.52 0.37 0.09

0.28 0.99 1.61 2.09 1.06 1.38 2.42 1.28 1.44 0.73 3.26 1.81 1.10 0.21 1.51 1.22 0.18 0.04

51.44 0.06 0.09 0.07 0.08 0.58 0.27 0.18 0.16 6.82 0.38 0.10 1.23 84.63 0.02 0.04 93.19 96.25

0.83 2.20 2.95 2.80 2.83 3.75 3.02 4.05 5.67 5.26 3.72 6.09 3.43 3.72 4.33 2.88 0.51 2.13

1.15 2.18 2.22 2.28 2.24 2.91 3.16 2.15 2.34 1.39 2.70 2.39 2.34 0.33 2.32 2.50 0.12 0.09

0.50 0.92 1.01 0.97 0.55 0.72 0.81 0.79 0.79 0.30 0.93 0.75 0.64 0.07 0.80 0.64 0.04 0.02

0.35 0.30 0.38 0.52 0.25 1.11 2.64 0.40 0.33 0.18 0.88 0.42 0.44 0.31 0.48 0.32 0.26 0.20

503 448 299 331 2,778 548 597 1,378 475 1,206 422 612 1,025 609 1,220 1,046 693 52

12 88 130 153 60 120 203 101 111 39 223 124 106 9 133 107 0 2

60 27 359 75 304 82 435 68 268 30 236 49 243 54 304 49 256 53 110 24 529 68 358 52 268 46 35 4 372 63 256 48 17 3 8 2

co rc rc rc so rc rc rc rc so rc rc rc co rc so co co

27.95 62.87 56.96 56.70 69.68 62.72 53.92 63.44 60.98 72.54 54.65 61.50 65.79 7.02 60.36 67.85 3.90 0.39

( ppm )

3.41 3.33 3.62 3.94 3.17 2.90 3.93 3.23 3.61 2.03 4.35 3.39 3.09 1.23 3.94 2.88 0.51 0.71

co=calcareous ooze; so=siliceous ooze; r c = r e d clay. C e / C e ' = C % / [ 0 . 5 × ( L a n + N d n ) ] Yb ) n = ( Lasample/Lashale ) / ( Ybsample/Ybshale).

elements are diluted by the carbonate phase. Furthermore, this dilution effect can also be seen in factor 2 which comprises negatively loaded REE, P, and most of major elements (Fig. 2 ). A1 is an important component of particulate matter derived from continents (e.g., Taylor and McLennan, 1985) and is relatively immobile during transformation of sediments to rock (Stumm and Morgan, 1981 ). A1 is therefore frequently used to trace the abundance and accumulation of the terrestrial component (Murray et al., 1991, 1992). However, a small portion of AI in the deep-sea sediments is also attributed to submarine weathering of ridge rocks and hydrothermal activity (McMurtry and Yeh, 1981 ). The average A 1 2 0 3 content of siliceous and red clay is 16.9% and is similar to that of Pacific pelagic clay sediments but lower than that of Atlantic sediments (18%) (El-Wakeel and Riley, 1961 ) and higher than in adjacent CIB sediments (Table 3 ). Fe in deep-sea sediments is generally associated with lithogenous, hydrogenous and metalliferous components (Chester and Aston, 1976). The

Li

Ni 55 368 379 467 232 297 496 297 279 102 830 389 250 18 364 254 2 5

(Murray et al., 1993). (La/

average F e 2 0 3 content (7.78%) in the siliceous and red clay sediments is similar to that of the Central Pacific but higher than in CIB sediments (Table 3). The average F e 2 0 3 / A 1 2 0 3 ratio (0.45) is lower than that of Pelagic Clay (PC) (0.58; Taylor and McLennan, 1985 ), indicating there is no "excess Fe" in these sediments. Furthermore, the lack of "excess Fe" and the good positive correlation with A1 and Ti suggest that Fe is mainly derived from detrital clays. Ti is a better tracer of the terrigenous fraction than A1, because of the potentially large biologically affiliated A1 flux (Murray et al., 1993), but a minor quantity is also derived from oceanic weathering of basalt (Bostr/Sm et al., 1973). The average TiO2 content of the siliceous ooze and red clay (0.75%) is higher than in Pacific Pelagic sediments and nearly one and half times more concentrated than in CIB sediments (Table 3). The A1203/TiO2 ratio (av. 22.4) is higher than the ratios of Post Archean Average Shale (PAAS) (18.9) and PC (20.7). This ratio is lower in the average Pacific pelagic clay sediment (Bischoff

J.N. Pattan et al. /Chemical Geology 121 (1995) 201-215

205

(ppm) Pb

Sr

Y

8 45 65 67 49 47 53 53 50 17 85 53 51 10 60 53 4 3

932 30 69 50 72 79 86 80 112 40 104 159 180 386 101 82 92 68 152 27 115 193 85 84 109 80 1,110 31 101 90 84 68 1,104 26 925 12

Zn

La

Ce

Nd

Sm

102 207 195 219 216 156 228 159 152 75 237 131 167 34 198 180 38 27

14.9 41.1 58.9 50.8 27.0 83.0 177.6 47.1 37.9 16.0 89.0 46.0 46.0 8.8 52.0 43.2 6.4 3.0

22.5 93.6 112.6 120.4 84.5 121.8 180.7 102.8 97.8 42.3 146.0 107.0 104.0 8.2 115.6 104.0 3.3 0.6

12.2 2.7 41.2 10.1 58.3 13.5 47.1 10.7 28.5 7.0 86.5 19.3 186.0 37.3 51.9 12.7 40.1 9.7 18.1 4.7 92.5 20.3 50.5 12.3 48.4 12.0 9.5 2.3 56.3 13.4 45.6 11.1 6.2 1.6 2.6 0.7

Eu

Gd

Dy

0.7 2.5 2.4 2.5 9.3 9.3 3.4 13.2 13.6 2.6 10.3 10.2 1.7 6.4 6.1 5.1 20.2 21.3 10.9 44.0 45.7 3.2 12.0 11.8 2.4 9.5 9.3 1.1 4.3 4.1 5.4 21.3 22.0 3.1 11.7 11.7 2.9 11.3 11.5 0.6 2.2 2.2 3.4 13.0 12.9 2.7 10.6 10.6 0.4 1.5 1.5 0.2 0.7 0.6

et al., 1979), indicating an insignificant contribution of Ti from submarine weathering. Because of the low sedimentation rate (Lisitzin, 1972 ) and insignificant contribution of Ti from submarine weathering, the high Ti and AI contents of the Wharton Basin sediments are probably due to the supply of lateritic soil from west Australia by southeasterly winds and northeasterly aeolian transport from the adjacent Indonesian archipelago (Venkatrathnam and Biscaye, 1973 ). TiO2 and A 1 2 0 3 s h o w g o o d positive correlation (Table 2) indicating their continental origin. The Mg, K and Na contents in the calcareous ooze are generally lower than those in red clay and siliceous ooze (Table 1 ). Mg shows strong positive correlation with A1, Fe and K (Table 2 ), suggesting its derivation from continental weathering and its negative association with Ca indicates that these two elements have different sources. Potassium and alumina show a strong linear relation (Table 2) suggesting lithogenous contribution, but the regression line intersects the

Ho

Er

Yb

0.5 1.5 1.5 1.8 5.3 5.2 2.8 7.9 7.5 2.1 6.2 5.8 1.2 3.3 3.2 4.6 13.1 11.9 10.0 29.3 23.1 2.4 6.5 6.0 1.8 5.2 5.1 0.8 2.2 2.1 4.7 13.5 12.2 2.4 6.7 6.3 2.3 6.4 6.1 1.4 1.4 1.2 2.6 7.3 6.6 2.1 6.1 5.7 0.3 1.1 1.0 0.1 0.4 0.5

Lu

SREE

0.2 0.8 1.2 0.9 0.5 1.8 4.0 0.9 0.8 0.3 1.9 1.0 0.9 0.2 1.0 0.9 0.2 0.1

61.5 220.3 292.8 267.0 169.2 388.8 748.6 257.3 219.7 95.9 428.9 259.1 251.8 37.1 283.9 242.9 23.6 9.7

Ce/Ce"

(La/Yb)n

0.79 1.07 0.67 1.17 1.43 0.68 0.47 0.98 1.18 1.18 0.76 1.05 1.04 0.39 1.01 1.11 0.25 0.19

0.86 0.68 0.68 0.75 0.73 0.60 0.66 0.67 0.64 0.66 0.63 0.63 0.65 0.62 0.68 0.65 0.57 0.50

K axis, indicating excess K, the K is most likely in the form of phillipsite (Fig. 3a). Microscopic examination of the coarse fraction of the sediments confirms the presence of phillipsite crystals with varying shapes and sizes. The higher concentration of Na compared with K in the red clay, siliceous sediments and calcareous oozes (Table 1) may indicate the presence of sodic feldspars in the clays (Nohara and Kato, 1985 ) or preferential biological removal of Na from seawater by certain calcareous organisms (EIWakeel and Riley, 1961 ). In deep-sea sediments, phosphorus is mainly present in the organic matter and phosphorite (apatite and fish bone debris) (Moody et al., 1988). A minor amount may also be provided by metalliferous sediments from mid-oceanic processes (Moody et al., 1988). The P content is low in both calcareous (0.11%) and siliceous sediments (0.1%) compared to red clays (0.32%). The high P content in red clays could be due to regeneration of P from the oxidation of organic matter: this oxidation refluxes P partly

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

206

Table 2 Correlation coefficient for entire data Si Si AI Fe Mn Mg Ca Na K Ti P

Ba Co U Li Ni Pb Sr Y Zn La Ce Nd Sm Gd

Dy Eu Ho Er Yb Lu

AI

Fe

Mn

Mg

Ca

Na

K

Ti

P

Ba

Co

Li

Ni

Pb

1.00 0.83 1.00 0.76 0.98 1.00 0.61 0.82 0.88 1.00 0.78 0.95 0.97 0.85 1.00 -0.98 -0.90 -0.85 -0.71 -0.86 1.00 0.67 0.49 0.44 0.50 0.50 - 0 . 6 7 1.00 0.87 0.93 0.92 0.82 0.90 - 0 . 9 2 0.56 1.00 0.77 0.99 0.92 0.81 0.95 - 0 . 8 6 0.46 0.88 1.00 0.21 0.42 0.51 0.59 0.43 - 0 . 2 9 0.10 0.57 0.37 1.00. 0.35 0.21 0.04 - 0 . 1 0 0.04 - 0 . 2 9 0.11 0.12 - 0 . 0 7 - 0 . 1 5 1.00 0.65 0.87 0.91 0.98 0.87 - 0 . 7 5 0.51 0.89 0.85 0.65 -0.11 1.00 0.73 0.88 0.88 0.87 0.88 - 0 . 8 7 0.51 0.81 0.89 0.24 0.06 0.86 1.00 0.75 0.97 0.94 0.77 0.91 - 0 . 8 3 0.43 0.84 0.98 0.30 - 0 . 1 4 0.82 0.88 1.00 0.59 0.83 0.88 0.97 0.85 - 0 . 7 0 0.43 0.80 0.84 0.50 - 0 . 1 0 0.95 0.92 0.81 1.00 0.76 0.92 0.92 0.90 0.90 - 0 . 8 5 0.52 0.88 0.91 0.35 0.09 0.91 0.95 0.88 0.91 -0.97 -0.88 -0.82 -0.82 -0.69 0.98 - 0 . 6 9 - 0 . 8 9 - 0 . 8 3 - 0 . 2 5 - 0 . 2 6 - 0 . 7 4 - 0 . 8 1 - 0 . 8 1 - 0 . 6 9 0.29 0.51 0.60 0.71 0.53 - 0 . 3 8 0.20 0.65 0.47 0.98 - 0 . 1 3 0.76 0.37 0.40 0.64 0.76 0.92 0.93 0.81 0.92 - 0 . 8 4 0.29 0.89 0.89 0.45 0.21 0.84 0.87 0.86 0.84 0.43 0.66 0.72 0.77 0.64 - 0 . 5 2 0.26 0.76 0.61 0.95 - 0 . 1 2 0.83 0.49 0.55 0.70 0.77 0.92 0.94 0.90 0.89 - 0 . 8 5 0.51 0.97 0.88 0.67 0.04 0.96 0.83 0.83 0.87 0.44 0.65 0.71 0.76 0.63 - 0 . 5 2 0.28 0.76 0.59 0.95 - 0 . 0 9 0.82 0.48 0.54 0.70 0.50 0.70 0.75 0.79 0.68 - 0 . 5 8 0.34 0.81 0.65 0.92 - 0 . 0 7 0.86 0.54 0.59 0.87 0.66 0.77 0.73 0.75 0.72 - 0 . 7 1 0.52 0.78 0.76 0.26 - 0 . 0 4 0.77 0.80 0.74 0.70 0.41 0.62 0.68 0.74 0.61 - 0 . 5 0 0.27 0.74 0.57 0.96 - 0 . 1 0 0.81 0.45 0.51 0.73 0.44 0.65 0.70 0.76 0.63 - 0 . 5 2 0.29 0.76 0.59 0.95 - 0 . 0 9 0.82 0.48 0.53 0.78 0.48 0.69 0.73 0.82 0.67 - 0 . 5 6 0.37 0.80 0.65 0.85 - 0 . 1 3 0.87 0.59 0.60 0.68 0.37 0.59 0.65 0.72 0.57 - 0 . 4 5 0.23 0.71 0.53 0.97 -0.12 0.78 0.42 0.48 0.69 0.42 0.64 0.70 0.76 0.62 - 0 . 5 1 0.27 0.76 0.59 0.95 - 0 . 1 3 0.82 0.47 0.53 0.78 0.39 0.61 0.67 0.74 0.59 - 0 . 4 8 0.25 0.73 0.56 0.96 - 0 . 1 3 0.80 0.44 0.50 0.67

Number of samples, n = 18. Level of significance at 99% is 0.59.

to seawater with a significant fraction renucleating and being buried as apatite (Moody et al., 1988). However, the average P205 content (0.54%) in these sediments is similar to Central Pacific pelagic sediments but two-fold enriched when compared with CIB sediments (Table 3 ). The highest PEOs content (2.64%), recorded in one of the red clays sediments (sample No. 56), is attributed to fish bone debris. Phosphatic fish debris is common in deep-water clays (Arrhenius et al., 1957) and it accumulates in the basinal regions below the CCD where the sedimentation rate is low. P205 and CaO (CaO values < 1%) show a moderate positive relation (Fig. 2b), suggesting the possible presence of Ca-phosphate perhaps as fish bone debris (Nohara and

Kato, 1985). A good linear relationship exists between P 2 0 5 ( P 2 0 5 values < 1%) and Cu (Fig. 2c) consistent with the mutual relationship observed between P and productivity (Froelich et al., 1982) and Cu with biogenic phases (Sawlan and Murray, 1983). The linear relationship between these two elements in the Wharton Basin may therefore reflect a higher primary productivity in the overlying water column which is in accordance with the observation made earlier

(FAO, 1981 ). The redox sensitive metals (Mn, Cu, Ni and Zn) in deep-sea sediments have either a hydrogenous, diagenetic or hydrothermal source. Mn in these sediments varies from 0.03% to 2.52%, with higher values associated with red clay and

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

207

Sr

Y

Zn

La

Ce

Nd

Sm

Gd

Dy

Eu

Ho

Er

Yb

Lu

1.00 -0.84 0.48 0.90 0.60 0.90 0.60 0.90 0.59 0.65 0.83 0.57 0.59 0.69 0.55

1.00 -0.35 -0.81 -0.49 -0.83 -0.49 -0.83 -0.49 -0.65 -0.71 -0.47 -0.50 -0.54 -0.45

1.00 0.54 0.98 0.76 0.98 0.76 0.98 0.96 0.39 0.98 0.98 0.91 0.98

1.00 0.66 0.89 0.65 0.89 0.65 0.69 0.67 0.62 0.64 0.66 0.61

1.00 0.85 0.99 0.85 0.99 0.99 0.48 0.99 0.99 0.94 0.99

1.00 0.89 0.85 0.89 0.59 0.75 0.83 0.85 0.87 0.82

1.00 0.99 0.49 0.99 0.99 0.94 0.99 0.99 0.99

1.00 0.55 0.90 0.99 0.95 0.98 0.99 0.98

1.00 0.47 0.49 0.71 0.43 0.51 0.46

1.00 0.99 0.94 0.99 0.99 0.99

1.00 0.94 0.99 0.99 0.93

1.00 0.93 0.95 0.93

1.00 0.99 0.99

1.00 0.99

siliceous ooze (1.16%) and low values with calcareous ooze (0.13%). Cu, Ni and Zn follow a similar trend to that of Mn in all sediment types. The average contents of Mn, Ni, Co and Zn are higher than in CIB and Pacific pelagic sediments, while the Cu content is lower than in Pacific sediments (Table 3 ). These redox sensitive metals show strong interelemental relationships between themselves (Table 2 ), suggesting coprecipitation with authigenic Mn-oxyhydroxide. Based on the presence of metallic micronodules in the coarse fraction, we infer that the Mn-oxyhydroxide phase in these sediments could be related to micronodules. Micronodules usually form in oxic environments by the supply of dissolved Mn 2+ via pore waters from the reducing

1.00

zone at subsurface depth and get oxidized to Mn 4+ a t the surface (Pattan, 1993). The moderate negative loadings (factor 2) of Ce, Mn, K, Co, Zn, Fe, A1, Pb and Mg may suggest association of Ce with micronodules and aluminosilicates, probably phillipsite (Fig. 3). The high content of these metals may also be due to postdepositional remobilization (Lynn and Bonatti, 1965; Graybeal and Heath, 1984; Pattan and Mudholkar, 1990; Banakar et al., 1991 ). Further, these metals might have been supplied from the dissolution of biogenic tests from the overlying high-productivity area. Ba is mainly associated with a biogenic source in marine sediment (e.g., Dymond et al., 1992 ) and is generally enriched in the sediments below

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

208

Table 3 Comparison of average bulk chemical composition of red clay ( R C ) , siliceous ooze ( S O ) and calcareous ooze ( C O ) from the Wharton Basin, Central Indian Basin and Central Pacific Basin Wharton Basin [ 1 ] RC

SO

CO

Central Indian Basin [2]

Central Pacific Basin [3]

SO

RC

CO

SO

RC

(%): SiO2 A1203 Fe203 MnO MgO CaO Na20 K20 TiO2 P205

51.4 14.4 7.4 1.4 2.9 0.2 3.2 2.1 0.6 0.6

53.4 13.7 5.8 0.9 2.5 1.8 2.9 1.6 0.6 0.3

131 321 378 56 186 948 52 52 106 128 66 0.66 0.73

86 258 283 42 159 651 60 27 98 52 18 0.68 1.24

2.20 0.53 0.23 0.08 0.47 51.89 1.20 0.10 0.02 0.15

61.20 9.91 4.50 0.86 2.18 1.00 2.24 1.60 0.38 0.22

n.d. 4.06 1.72 0.36 1.26 41.18 1.36 0.68 0.14 0.21

55.75 14.20 7.40 0.77 3.09 1.63 1.23 2.29 0.55 0.51

53.52 13.31 6.75 0.60 2.81 2.39 1.27 2.26 0.57 0.42

1.78 1.15 0.19 0.03 0.18 53.48 0.15 0.09 0.02 0.05

(ppm): Co Cu Ni Pb Zn Ba Li Pb Sr Y La (La/Yb)n Ce/Ce"

4 20 8 6 33 464 9 6 1,018 25 8 0.61 0.41

69 286 103 48 250 2,277 32 18 125 61 30 0.63 1.19

26 108 93 20 126 218 14 20 706 43 21 0.66 0.82

92 392 177 35 130 -

77 341 137 31 113 -

15 20 17 25 10 -

[ 1 ] = present study; [ 2] J.N. Pattan [ unpublished data, 1992; SO n = 14, CO n = 3 (n = n u m b e r o f samples) ]; [ 3] = N o h a r a

and Kato ( 1985 ). n.d. = not determined; - = data not available.

the surface high-productivity waters (Fisher et al., 1991). Dehairs et al. (1980) and Bishop (1988 ) found that Ba in suspended particles is predominantly in the form of barite and concluded that it has been formed in the upper water column by the breakdown of organic matter, release of Ba, and formation barite in microenvironments. The average Ba content is low in calcareous ooze (464 ppm) and red clays (668 ppm) but higher concentrations are recorded in siliceous ooze (1676 ppm) (Table 1 ). Estimation of biobarite, calculated from the formula of Dymond et al. (1992), suggests that Ba in the calcareous and siliceous ooze is mostly of biogenic origin, while in red clays it is not only associated with the biogenic phase but to a lesser

extent also includes an aluminosilicate controlled source. The association of Ba with the biogenic source is seen in factor 3 (Fig. 2). This factor has strong positive loadings on Ba with moderate positive loadings on Si, Na and K. The moderate negative loading of Ca and Sr suggests there are two biogenic sources like carbonate and opal. 4.2. Rare-earth elements (REE) REE abundances are presented in Table 1 and their shale-normalized (North American Shale Composition, NASC) patterns are shown in Fig. 4. Calcareous ooze is present on either side of the transect where the sediment contains extremely

J.N. Pattan et aL / Chemical Geology 121 (1995) 201-215

209

f~x

+

t

t

i

,

f

,

i

i

,

i

,

¢ ,

i

r

,

~

i

I

~

,

,

]@

i

i

i

I.(44.1%)

,

FACTOR

l °I'll

-I

-

llll 'il-llllll I't'""""

- I

~" '

FACTOR

,

-

,

2.(

3•

%)

,'i 'l 'ltlilllllll 01._ i_l...

÷1~

,

,

i

0 v

I l"l'i

r

r

r

,

i

,

,

,

r

i

r

,

,

~

l

,

r

FACTOR

,

i

-3.(

, _ . . ,

,

,

•

2-

i 'L

59.4

el

®

0 L•

ib

0 ,

,~

ab

25

AJzo3 % 0.6

7. 8 * / * )

. . . . .

0-40 0

SI

Fe AI

MQ Mn

Co

No

BO

TI K

P

Co

Cr

LI Cu

Pb NI

Sr

Zn

Ce Lo

Sm Nd

Gd Eu

Ho Dy

O

Yb Et

Lu

0-2e II

Fig. 2. Varimax factor loadings for major and trace elements, and REE in the surface sediments.

0

low REE abundances of 9-61 ppm (La ~ 8 ppm), which are ~ 8-50% lower than the average shale value. These samples have strong negative Ce/Ce* values (0.41, Table 1 ) and are similar to the seawater REE pattern. Similar distribution patterns have been reported in other parts of the world oceans (Spirn, 1965; Wang et al., 1986; Liu et al., 1988, Toyoda et al., 1990; Murray and Leinen, 1993). Because there appears to be very low REE abundance ( ~ 5 ppm) in ultrasonically cleaned foraminifera tests (Elderfield et al., 1981 ), the relatively high REE content in these calcareous oozes could be attributed to the ferromanganese oxide coatings (Boyle, 1981; Palmer, 1985). The strong negative Ce anomalies are associated with higher carbonate content (Table 1; Fig. 4a), contrary to the findings of Courtois and Hoffert (1977) in the Pacific Ocean. The siliceous ooze/sediments have higher REE concentrations ranging from 95 to 242 ppm (La ,,,28 ppm, Table 1 ) with moderate positive Ce/Ce" (av. 1.24). Because siliceous ooze typically has low REE content (Piper, 1974; Elderfield et al., 1981 ), the higher abundance in these sediments most likely reflects authigenic

@

V:

0

~"5

1.10

1"~5

2 '10

2'5

Pz 05 */* E

¢.,, 40 qp,,

82oot • 0/

0

0!2

© 014

0]6

0!8

1.0

P20~ % Fig. 3. Plots for all sediments: (a) A1203 vs. I(20; (b) P205 vs. CaO (P205 values < 1.2%, CaO values < 1%); and (c) P205 vs. Cu (P20 s values < 1.0%).. Cu Symbols: triangles=calcareous ooze; squares=siliceous ooze/sediment; circles = red clay.

minerals such as manganese micronodules or broken fragments thereof. Indeed, the shale-normalized patterns (Fig. 4b) are somewhat similar to that of manganese nodules (J.N. Pattan, Ch.M. Rao, S. Colley and N.C. Higgs, unpublished data, 1992 ). Toyoda et al. (1990) noticed a wide range of Ce/Ce* values from moderate

210

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

@ CALCAREOUS

.01

OOZE

'

RED

I)y

CLAY

~

~o ~2r

~)b llu

_

<~ Z

O-

Lo

C/

{~

I

N'd

i

i

i

Sm Eu

SILICEOUS

Gd

'

O~y I~o Er

i

Yb

llu

Yb

Lu

OOZE

z

I1_

LO

•

S

Eu

a

Dy

Ha E r

Fig. 4. Shale-normalized (NASC) REE patterns for: (a) calcareous ooze; (b) siliceous ooze; and (c) red clay.

positive to moderate negative (0.5-1.6) in Pacific pelagic sediments comprising multiple lithologies. Elderfield et al. (1981) noticed positive Ce anomalies for siliceous sediments in the Pacific Ocean. The present deep-sea water shows negative Ce/Ce" (De Baar et al., 1985; Piepgras and Jacobsen, 1992). However, the siliceous ooze exhibits positive Ce/Ce', suggesting preferential removal of Ce from seawater by Fe-Mn-oxide. Red clays are greatly enriched in REE content, with concentrations varying from 220 to 749 ppm (Table 1 ). One of the red clay samples (Stn. 56) has the highest total REE abundance of 749 ppm, and is higher than some of the manganese nodules and micronodules from the CIB (Pattan et al., 1994). This value is nearly four times higher compared to shale. The red clay exhibits weak positive (av. 1.05) to moderate negative (av. 0.73) Ce/Ce* (Table 1 ). Higher Ce/Ce* in the red clays of the Atlantic are attributed to postdepositional diagenetic changes (Addy, 1979). In the Pacific, hydrothermal oxyhydroxides of Fe and Mn and fish bone debris are responsible for negative Ce/Ce" (Matsumoto et al., 1985; Toyoda et al., 1990). The negative Ce/Ce" in these red clays could be due to fish bone debris. Indeed, red clays which have high REE abundance and negative Ce/Ce* are associated with high P content (Table 1 ). In the absence of any hydrothermal influence in the Wharton Basin the negative Ce/Ce* most likely reflects the presence of fish bone debris which has strong negative Ce/ Ce* (Goldberg et al., 1963; Bernat, 1975; Elderfield and Pagett, 1986; Toyoda and Tokonami, 1990), as well as phillipsite and smectite, which also have negative Ce/Ce* (Piper, 1974). From these observations it appears that REE abundances and Ce/Ce* are related to general lithology and the abundance of authigenic components like micronodules, phillipsite and fish bone debris in the sediments. Shale-normalized REE patterns appear to exhibit weak positive Eu anomalies (Fig. 4) in these sediments. The calcareous ooze near the Ninetyeast Ridge has slightly higher Eu anomalies compared to siliceous and red clay sediment. This observation is interesting since positive Eu anomalies are reported in areas affected by either

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

aeolian or hydrothermal input (Elderfield, 1988), or strong reducing conditions (McLennan, 1989), or by the presence of detrital feldspar (Murray et al., 1991 ). In the absence of any hydrothermal influence and reducing conditions, the positive Eu anomalies could be due to alkali feldspars or aeolian input. The high Na/ A1 ratio suggests the presence of feldspar in the calcareous oozes from the Ninetyeast Ridge. Furthermore, siliceous and red clays have low Na/A1 ratios, indicating Eu anomalies in these sediments are due to aeolian input from the Australian continent. The degree of relative light REE (LREE) and heavy REE (HREE) enrichment can be estimated by the (La/Yb)n ratio (Calvert et al., 1987; Murray et al., 1992). The ratio (Table 1 ) showed consistent enrichment of HREE over LREE, suggesting retention of the seawater REE pattern in these sediments without any large-scale fractionation. This ratio is almost the same for all sediments (Table l). Although calcareous ooze, and siliceous and red clays have distinct REE abundance and Ce/Ce', their similar fractionation patterns suggest that they are independent of sediment type. These Wharton Basin sediments are within the realm of Antarctic Bottom Water (AABW). These waters may have a higher benthic flux of HREE because of progressive removal of LREE from the bottom water. Calvert et al. (1987) found similar fractionations in the valley nodules from the Pacific Ocean and attributed them to the bottom waters with high HREE content. Further, these Wharton Basin sediments have fish bone debris, phillipsite and smectite, which are also enriched in HREE over LREE (Piper, 1974; Bernat, 1975 ). REE patterns of ichthyoliths and bottom waters are similar, implying that ichthyoliths record REE patterns of overlying water at the sediment-water interface (Elderfield and Pagett, 1986). A very strong positive correlation (0.92) is noticed between Sm (intermediate REE) and P2Os (Fig. 5a), and an inverse relation between Ce/Ce* and P205 content (Fig. 5b). A similar relationship was observed for Pacific pelagic sediments (Toyoda et al., 1990). Fish bone de-

211

40

~O-

E t~ o.

20-

E

-L

O3 IO-

Io 0 •

®

A&

O O

I

I

Pz05 % 1.5

-)(-

1.0

CD

tJ

0,5

® 0 0

P205 % 1.5

I I.O-

°

°

C.D 0 ' 5 -

© 0

0.5

i I-5

I 2.5

3-5

MnO % Fig. 5. Plots of: ( a ) P 2 0 s vs. Sin; ( b ) P205 vs. C e / C e ' ; a n d (c) M n O vs. C e / C e " (symbols are the s a m e as in Fig. 3).

bris is known to be important in scavenging REE from seawater and contains high REE abundance with negative Ce/Ce* (Arrhenius et al., 1957; Bernat, 1975; Elderfield and PageR, 1986; Toyoda et al., 1990). A moderate inverse relation between Ce/Ce" and MnO was also observed (Fig. 5c). A similar relation was observed for Pacific pelagic sediments and was attributed

212

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

to hydrothermal influence (Matsumoto et al., 1985 ) because hydrothermal hydroxides rapidly scavenge the REE in seawater without preferential scavenging of Ce/Ce" (Ruhlin and Owen, 1986) or probably, with slight fractionation. In the absence of any hydrothermal influence in the Wharton Basin, the high Mn content could be due to the higher primary productivity of the water column (Martin and Knauer, 1984; Toyoda et al., 1990) and the presence of micronodules as well as upward diagenetic remobilization process. Therefore the strong negative Ce/Ce" with higher Mn content in these sediments might be due to adsorption/deposition of Mn-oxide on smectite and phillipsite which basically have negative Ce/Ce*. All the REE (La to Lu) show strong positive interelemental relationships ( > 0.8 ), indicating their coherent nature (Table 2 ). Ce behaves differently from the trivalent REE because of its oxidation to the insoluble tetravalent state. A strong positive relationship of REE with P is observed (Tables 2-5; Fig. 2). Ichthyoliths or fish teeth can contain very high REE concentrations nearly 50-200 times higher than shale (Arrhenius et al., 1957; Bernat, 1975; Wright et al., 1984; Elderfield and Pagett, 1986; Dymond and Eklund, 1978; Toyoda et al., 1990). Low sedimentation rates in the Wharton Basin results in high accumulation and better preservation of fish bone debris (Elderfield and Pagett, 1986). Based on the REE diffusion coefficient in fish teeth, Toyoda and Tokonami (1990) suggested that diffusion of REE from seawater into fish teeth before burial is too slow to produce such a high REE in fish bone and its high REE abundance must derive from the surrounding sediment pore fluids. This indicates the importance of diagenetic processes in transferring REE from sediments to fish bone debris (Shaw and Wasserburg, 1985 ). Biogenic apatite has large lattice defects and impurities, and as a result, the fish teeth enamel adsorbs REE. whereas dentine easily accepts REE into lattice (Toyoda and Tokonami, 1990). As P is related to biological activity, the strong association of REE with P in the Wharton Basin sediments indicates that REE are associated with surface primary productivity.

A small portion of REE in these sediments is also associated with authigenic components and is evidenced by the moderate positive correlation (~0.5) between REE and Mn-Fe-oxyhydroxides (probably manganese micronodules, Fig. 2). These Fe-Mn-oxyhydroxides scavenge the REE from the water column and are incorporated into micronodules (Ruhlin and Owen, 1986). Similarly, a weak positive correlation (~0.3-0.5) between REE and AI, Mg and K suggests their association within aluminosilicate which is probably phillipsite (Fig. 2) because aluminosilicates are not the REE carrier phase (Elderfield et al., 1981 ). A high REE abundance (300-400 ppm), and a Ce depletion is usually shown by phillipsites (Piper, 1974; Courtois and Hoffert, 1977 ). These samples have average Alex of 1.3 which constitutes ~ 10% of the total AI [Alex=Altot-(Tisap×(A1/Ti)sha], and only a few samples have higher Mg~x. The weak positive association between REE and aluminosilicate suggests that clays also might be carrying some of the REE as adsorbed phase. A very close association of yttrium with the REE is due to substitution into the lattice of calcium phosphate phase.

5. Conclusions

(1) The chemical composition of Wharton Basin sediments appears to be similar to that of Central Pacific Basin sediments and devoid of any hydrothermal influence. (2) Interelemental associations suggest the presence of micronodules, fish bone debris and phillipsite in these sediments. (3) The REE abundance and Ce/Ce" are distinct for each sediment type and are mostly controlled by the authigenic components like micronodules, phillipsite and fish bone debris. (4) The REE fractionation is independent of sediment type and is mainly governed by the overlying seawater REE pattern. (5) The REE abundances in these sediments are related to surface primary productivity and diagenetic processes.

J.N. Pattan et aL / Chemical Geology 121 (1995) 201-215

Acknowledgements The authors are grateful to the Director (N.I.O) for the permission to publish this paper and to Mr. R.R. Nair, Head Marine Geology Division, for encouragement. This work was carded out at the Institute of Oceanographic Sciences, Deacon Laboratory (IOSDL), Surray, U.K., during the stay of the first author under a British Council Fellowship. The authors thank Dr. Richard Murray, Bostan University, for his critical review and valuable suggestions. Mr. R. Uchil is thanked for the line drawings.

References Addy, S.K., 1979. Rare earth element pattern in manganese nodules and micronodules from the northwest Atlantic. Geochim. Cosmochim. Acta, 43:1105-1115. Arrhenius, G., Bramlette, M.N. and Piciotto, E., 1957. Localization of radioactive and stable heavy nuclides in ocean sediments. Nature (London), 180: 85-86. Banakar, V.K., Gupta, S.M. and Padmavati, V.K., 1991. Abyssal sediment erosion in the Central Indian Basin: Evidence from radiochemicat and radiolarian studies. Mar. Geol., 97: 423-431. Bernat, M., 1975. Les isotopes de l'uranium et du thorium et les terres rares dans l'environnement matin. Cah. ORSTOM (Off. Rech. Sci. Tech. Outre-mer), Srr. G6ol., 7: 65-83. Bischoff, J.L., Heath, G.R. and Leinen, M., 1979. Geochemistry of deep-sea sediments from the Pacific Manganese Nodule Province. DOMES sites A, B and C. In: J.L. Bischoff and D.Z. Piper (Editors), Marine Geology and Oceanography of the Pacific Manganese Nodule Province. Plenum, New York, N.Y., pp. 397-436. Bishop, J.K.B., 1988. The barite-opal-organic carbon association in oceanic particulate matter. Nature (London), 331: 341-343. BostriSm, K., Kraemer, T. and Gartner, S., 1973. Province and accumulation rates of opaline silica, AI, Ti, Fe, Mn, Cu, Ni and Co in Pacific pelagic sediments. Chem. Geol., 11: 128-148. Boyle, E.A., 1981.1 Cadmium, zinc, copper and barium in foraminifera tests. Earth Planet. Sci. Lett., 53:11-35. Calvert, S.E., Piper, D.Z. and Baedecker, P.A., 1987. Geochemistry of the rare earth elements in the ferromanganese nodules from DOMES site A, northern equatorial Pacific. Geochim. Cosmochim. Acta, 51: 2331-2338. Chester, R. and Aston, S.R., 1976. The geochemistry of deepsea sediments. In: J.P. Riley and R. Chester (Editors), Chemical Oceanography, Vol. 6. Academic Press, New York, N.Y., pp. 281-383.

213

Courtois, C. and Hoffert, M., 1977. Distribution des terres rares dans les srdiments superficiels du Pacifique sud-est. Bull. Soc. Grol. Fr., 19: 1245-1251. Cronan, D.S. and Moorby, S.A., 1981. Manganese nodules and other ferromanganese oxide deposits from the Indian Ocean. J. Geol. Soc. London, 138: 527-539. De Baar, H.J.W., Bacon, M.P., Brewer, P.G. and Bruland, K.W., 1985. Rare earth elements in the Pacific and Atlantic Oceans. Geochim. Cosmochim. Acta, 49: 1943-1957. Dehairs, F., Chesselet, R. and Jedwab, J., 1980. Discrete suspended particles of barite and the barium cycle in the open ocean. Earth Planet. Sci. Lett., 49: 529-550. Dymond, J. and Eklund, W., 1978. A microprobe study of metalliferous sediment components. Earth Planet. Sci. Lett., 40: 243-251. Dymond, J., Suess, E. and Lyle, M., 1992. Barium in deep sea sediments: A geochemical proxy to paleoproductivity. Paleoceanography, 7(2): 163-182. Elderfield, H., 1988. The oceanic chemistry of the rare earth elements. Philos. Trans. R. Soc. London, Ser. A, 325: 105126. Elderfield, H. and Pagett, R., 1986. Rare earth elements in icthyoliths: variations with redox conditions and deposition environment. Sci. Total Environ., 49:175-197. Elderfield, H., Hawkesworth, C.J., Greaves, M.J. and Calvert, S.E., 1981. Rare earth element geochemistry of oceanic ferromanganese nodules and associated sediments. Geochim. Cosmochim. Acta, 45:513-528. E1-Wakeel, S.K. and Riley, I.P., 1961. Chemistry and mineralogical studies of deep sea sediments. Geochim. Cosmochim. Acta. 25:110-146. FAO (U.N. Food and Agricultural Organization), 1981. Atlas of the Living Resources of the Seas. FAO Fish. Ser., Rome. Fernandez, A.A. and Mahadevan, R., 1982. Factor analysis of multivariate data. Natl. Inst. Oceanogr., Goa, Tech. Rep. 2/82. Fisher, N.S., Guillard, R.R.L. and Bankston, D.C., 1991. The accumalation of barium in marine phytoplankton grown in culture. J. Mar. Res., 49: 339-354. Froelich, P.N., Bender, M.L., Luedtke, N.A., Heath, G.R. and De Vties, T., 1982. The marine phosphorus cycle. Am. J. Sci., 282:474-511. Glasby, G.P., Gwozdz, R., Kunzendorf, H., Friedrich, G. and Thijssen, T., 1987. The distribution of rare earth and minor elements in manganese nodules and sediments from the equatorial and SW Pacific. Lithos, 20: 97-113. Goldberg, E.D. and Arrhenius, G.O.S., 1958. Chemistry of Pacific pelagic sediments. Geochim. Cosmochim. Acta, 13: 153-212. Goldberg, E.D., Koide, M., Schmitt, R.A. and Smith, R.H., 1963. Rare earth distributions in the marine environment. J. Geophys. Res., 68:4209-4217. Graybeal, A.L. and Heath, G.R., 1984. Remobilization of transition metals in surfacial pelagic sediments from the eastern Pacific. Geochim. Cosmochim. Acta, 48: 965-975.

214

J.N. Pattan et al. /Chemical Geology 121 (1995) 201-215

Ingamells, C.O., 1966. Absorption methods in rapid silicate analysis. Anal. Chem., 38 (9): 1228-1234. Jarvis, I. and Jarvis, K.E., 1985. Rare earth element geochemistry of standard sediments: A study using inductively coupled plasma spectrometry. Chem. Geol., 53: 335-344. Krishnaswami, S., 1976. Authigenic transition elements in Pacific pelagic clays. Geochim. Cosmochim. Acta, 40: 425-435. Leinen, M., 1977. A normative calculation technique for determining opal in deep-sea sediments. Geochim. Cosmochim. Acta, 41: 671-676. Lisitzin, A.P., 1972. Sedimentation in the world ocean. Spec. Publ. Tulsa Soc. Econ. Paleontol. Mineral., 17:218. Liu, Y.G., Miah, M.R.U. and Schmitt, R.A., 1988. Cerium: a chemical tracer for paleo-oceanic redox conditions. Geochim. Cosmochim. Acta, 52:136 l - 1371. Lynn, D. and Bonatti, E., 1965. Mobility of manganese in the diagenesis of deep sea sediments. Mar. Geol., 3: 457-474. Martin, J.H. and Knauer, G.A., 1984. Vertex: manganese transport through oxygen minima. Earth Planet. Sci. Lett., 67: 35-47. Matsumoto, R., Minai, Y. and Iijima, A., 1985. Manganese content, cerium anomaly and rate of sedimentation as aids in the characterization and classification of deep-sea sediments. In: N. Nasu, K. Kobayashi, S. Uyeda, I. Kushiro and H. Kagami (Editors), Formation of Active Ocean Margins. TTRRAPUB, Tokyo, pp. 913-939. McLennan, S.M., 1989. Rare earth elements in sedimentary rocks: Influence of provenance and sedimentary processes. In: B.R. Lipin and G.A. Mckay (Editors), Geochemistry and Mineralogy, Vol. 21. pp. 169-200. McMurtry, G.M. and Yeh, H.W., 1981. Hydrothermal clay mineral formation of East Pacific Rise and Bauer Basin sediments. Chem. Geol., 32: 189-205. Moody, J.R., Chaboudy, Jr., L.R. and Worsley, T.R., 1988. Pacific pelagic phosphorus accumulation during the last l0 M.Y. Paleoceanography, 3( 1 ): 113-136. Murray, R.W and Leinen, M., 1993. Chemical transport to the sea floor of the equatorial Pacific Ocean across a latitude transect at 135 ° W: Tracing sedimentary major, trace, and rare earth element fluxes at the Equator and the Intertropical Convergence Zone. Geochim. Cosmochim. Acta, 57: 4141-4163. Murray, R.W., Brink, M.R.B., Gerlach, D.C., Russ III, G.P. and Jones, D.L., 1991. Rare earth, major and trace elements in chert from the Franciscan Complex and Monterey Group, California: Assessing REE sources to fine grained marine sediments. Geochim. Cosmochim. Acta, 55: 1875-1895. Murray, R.W., Brink, M.R.B., Gerlach, D.C., Russ III, G.P. and Jones, D.L., 1992. Interoceanic variation in the rare earth, major, and trace element deposition chemistry of chert: Perspectives gained from the DSDP and ODP record. Geochim. Cosmochim. Acta, 56:1892-1913. Murray, R.W., Leinen, M. and Isern, A.R., 1993. Biogenic flux of AI to sediment in the Central equatorial Pacific

Ocea: Evidence for increased productivity during glacial periods. Paleaoceanography, 8 ( 5 ): 651-670. Nohara, M. and Kato, K., 1985. Chemical compositions of pelagic deep-sea sediments its relation to the formation of authigenic mineral phase under the chemical control of sea water. In: N. Nasu et al. (Editors), Formation of Active Ocean Ridges. TTRRAPUB, Tokyo, pp. 893-912. Palmer, M.R., 1985. Rare earth elements in foraminifera tests. Earth Planet. Sci. Lett., 73: 75-82. Parsons, T.R., Takahashi, M. and Hargrove, B., 1977. Biological Oceanographic Processes. Pergamon, New York, N.Y., 2nd ed., 332 pp. Pattan, J.N., 1993. Manganese micronodules: a possible indicator of sedimentary environments. Mar. Geol., 113: 331-344. Pattan, J.N. and Mudholkar, A.V., 1990. The oxidation state of manganese in ferromanganese nodules and deep-sea sediments from the Central Indian Ocean. Chem. Geol., 85: 171-181. Pattan, J.N., Colley, S. and Higgs, N.C., 1994. Rare earth element behaviour in co-existing macronodules, micronodules and sediments from the Central Indian Basin. Mar. Georesour. Geotechnol. (in press). Piepgras, J. and Jacobsen, S.B., 1992. The behaviour of rare earth elements in seawater: Precise determination of variation in the North Pacific water column. Geochim. Cosmochim. Acta, 56: 1851-1862. Pingitore, Jr., N.E., Lytle, F.W., Daviers, B.M., Eastman, M.P., Eller, P.G. and Larson, E.M., 1992. Mode of incorporation of Sr in calcite: determination by X-ray absorption spectroscopy. Geochim. Cosmochim. Acta, 56:15311538. Piper, D.Z., 1974. Rare earth elements in the sedimentary cycle: a summary. Chem. Geol., 17: 287-297. Ruhlin, D.E. and Owen, R.M., 1986. The rare earth element geochemistry of hydrothermal sediments from the East Pacific Rise: examination of a sea water scavenging mechanism. Geochim. Cosmochim. Acta, 50: 393-400. Sawlan, J.J. and Murray, J.W., 1983. Trace metal remobilization in the interstiatial waters of red clay and hemipelagic marine sediments. Earth Planet. Sci. Lett., 64:213230 Shaw, H.F. and Wasserburg, G.J., 1985. Sm-Nd in marine carbonates and phosphates: implications for Nd-isotopes in sea water and crustal ages. Geochim. Cosmochim. Acta, 49:503-518. Spirn, R.V., 1965. Rare earth distributions in marine environment. Ph.D. Thesis, Department of Geology and Geophysics, Massachusetts Institute of Technology, Cambridge, Mass., 165 pp. Stumm, W. and Morgan, J.J., 1981. Aquatic Chemistry. Wiley, New York, N.Y., 780 pp. (see especially Introduction Emphasizing Chemical Equilibria in Natural Water). Taylor, S.R. and McLennan, S.C., 1985. The Continental Crust: Its Composition and Evolution. Blackwell, London, 300 pp. Thomson, J., Carpenter, M.S.N., Colley, S. and Willson,

J.N. Pattan et al. / Chemical Geology 121 (1995) 201-215

T.R.S., 1984. Metal accumulation rates in northwest Atlantic pelagic sediments. Geochim. Cosmochim. Acta, 48: 1935-1948.

Toyoda, K. and Masuda, A., 1990. Sedimentary environments and chemical composition of Pacific pelagic sediments. Chem. Geol., 88: 127-141. Toyoda, K. and Tokonami, M., 1990. Diffusion of REEs in fish teeth from deep-sea sediments. Nature (London), 345: 607-609. Toyoda, K., Nakumura, Y. and Masuda, A., 1990. Rare earth elements of Pacific pelagic sediments. Geochim. Cosmochim. Acta, 54:1093-1103. Udintsev, G.B. (Editor), 1975. Geological and Geophysical

215

Atlas of the Indian Ocean. Pergamon, London, 151 pp. Venkatrathnam, K. and Biscaye, P.E., 1973. Clay mineralogy and sedimentation on the eastern Indian Ocean. DeepSea Res., 20: 727-738. Wang, Y.L., Liu, Y.G. and Schmitt, R.A., 1986. Rare earth element geochemistry of south Atlantic deep-sea sediments: Ce anomaly change at 54 My. Geochim. Cosmochim. Acta, 50: 1337-1355. Wright, J., Seymour, R.S. and Shaw, H.F., 1984. REE and Nd isotopes in conodont apatite: variations with geological age and depositional environment. Geol. Soc. Am., Spec. Pap., 196: 335-340.