Response of trace metal redox proxies in continental shelf environment: The Eastern Arabian Sea scenario

Continental Shelf Research 106 (2015) 70–84

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Response of trace metal redox proxies in continental shelf environment: The Eastern Arabian Sea scenario Shiba Shankar Acharya, Mruganka K. Panigrahi n, Anil K. Gupta 1, Subhasish Tripathy 2 Department of Geology & Geophysics, IIT Kharagpur, WB 721302, India

art ic l e i nf o

a b s t r a c t

Article history: Received 13 December 2014 Received in revised form 7 July 2015 Accepted 10 July 2015 Available online 11 July 2015

Major and trace elements (viz. Al, Si, Ca, Na, Fe, Ti, K, S, Mn, V, Cu, Zn, Mo, Ni, Th and U) along with organic carbon and nitrogen content were analyzed in the sediment samples of a gravity core (GC-08/SK291; 12°34′N: 74°11.47′E) which lies within the upper edge of intense oxygen minimum zone (OMZ) of Eastern Arabian Sea. The aim was to determine how the geochemistry of the redox-sensitive elements (viz. V, Cu, Zn, Mo, Ni, and U) was influenced by the suboxic water column and bottom water condition and whether the sediments show a unique signature of such redox condition. The weathering intensity was also determined and the provenance of sediments has been inferred to discern the uniformity in sediment supply in the studied location. The chemical index of alteration (CIA), Chemical Index of Weathering (CIW) and Al–Ti–Zr ternary diagrams suggest low to moderate source area weathering of granodioritic to tonalitic source rock composition, which is similar to the regional geology of the continental mass nearby. The relative variations of major elements such as Si, Al and Ca suggest that the terrigenous fractions in the studied sediments are diluted by marine carbonates, unlike by biogenic opal, which is the case for most shelf regions of the world experiencing upwelling. The relative distributions of total organic carbon (TOC), total sulfur (TS) and total nitrogen (TN) suggest that the organic matter in the studied sediments is primarily of terrestrial origin irrespective of its presence in the deeper part of the shelf. We adopted the multi-proxy procedure for analyzing the redox condition prevalent during the deposition of sediments (e.g., enrichment of redox-sensitive elements, authigenic U, and ratios of trace metals such as U/Th, V/Cr, V/Mo, Ni/Co and (Cu þMo)/Zn). The most striking result is a fully oxic signature reflected consistently by all redox proxies despite the occurrence of studied location within the OMZ of Arabian Sea. The redox proxies estimated from the bulk composition as well as the marine fraction of trace metals, do not reveal any significant difference in the depositional condition. The seasonal variation of O2 concentration seems to be the dominant factor which controls the response of the trace metals to the water column anoxia in the studied location. The major implication of the present study is that the trace element redox proxies alone are not always suitable to discern the redox condition of deposition in continental shelf sediments which are characterized by high sediment input from continents and significant seasonal variation of O2 concentration in the water column. & 2015 Elsevier Ltd. All rights reserved.

Keywords: Arabian Sea Redox condition Trace elements Oxygen minimum zone Seasonal variation

1. Introduction The redox environment prevalent during the deposition of marine sediments is reconstructed using the geochemical behavior of redox-sensitive trace metals such as V, Cu, Zn, Mo, Ni, and U (Jones and Manning 1994; Tribovillard et al., 2006). Studies on n

Corresponding author. Fax: þ91 3222 282268. E-mail address: [email protected] (M.K. Panigrahi). 1 Present address: Wadia Institute of Himalayan Geology, Dehradun 248001, India. 2 Present address: Indian Institute of Technology, Bhubaneswar, Odisha 751013, India. http://dx.doi.org/10.1016/j.csr.2015.07.008 0278-4343/& 2015 Elsevier Ltd. All rights reserved.

behavior of these metals in suboxic and anoxic bottom water condition have been conducted in several sedimentary records, such as for the gulf of California (Brumsack, 1989), Black Sea (Anderson et al., 1989; Brumsack, 1989; Colodner et al., 1995), Cariaco Basin (Lyons et al., 2003; Piper and Dean, 2003), Gulf of Mexico (Swarzenski et al., 2008), Peru and Chilean Margins (Böning et al., 2004, 2005), Arabian Sea (van der Weijden et al., 2006; Pattan and Pearce, 2009) and Bay of Bengal (Pattan et al., 2013). A number of geochemical techniques have been used as indices of redox condition in recent sediments and sedimentary rocks. In general, they can be divided into three groups: (i) those based on formation of pyrite during early diagenesis (degree of pyritization (DOP)) and the relation between pyritic sulfur and

S.S. Acharya et al. / Continental Shelf Research 106 (2015) 70–84

organic carbon (e.g. Berner and Raiswell, 1983); (ii) those using the enrichment or depletion of authigenic trace metals in different depositional environment (e.g. Morford and Emerson, 1999); and (iii) those employing threshold values for various trace metal ratios such as U/Th, V/Cr, V/Mo, and Ni/Co (e.g. Riquier et al., 2006). Uncertainties still remain in the interpretation of these redox proxies in marine sedimentary records. The trace metal redox responses often get superimposed by the diagenetic signatures after sediment burial and also get affected by the changes in organic matter export to the sea floor. The growing evidence suggests that these trace metal redox proxies are not consistent in all environments of deposition. Nameroff et al. (2002) studied trace metal distribution in water column, settling particulate matters and sediment samples from eastern tropical north pacific and showed that redox-sensitive trace metals (Re, Mo, Cd, and U) signature in sediments does not reflect the bottom water redox condition. Similarly, McKay et al. (2007) studied the distribution of Re, U, Cd, Mn and Mo along with Iodine/organic carbon ratio in continental margin sediments of N.E. Pacific and showed that the response of these redox-sensitive proxies are not always reliable in environments characterized by high sedimentation rates and/or low sedimentation rates in combination with deep bioturbation. Therefore, a wide range of marine settings and environmental conditions need to be investigated to further develop our understanding in the responses of these trace metals to various redox conditions. Although the sediment columns from deeper basin of Arabian Sea were studied previously for trace metal responses (van der Weijden et al., 2006; Pattan and Pearce, 2009), the redox behavior of trace elements has not been studied in Eastern Arabian shelf sediments. Arabian Sea is characterized by intense upwelling which results in a high phytoplankton production that in return, enhances the consumption of dissolved oxygen. This high oxygen consumption in combination with a moderate rate of thermocline ventilation by Indian Ocean central water, Red Sea water and Persian Gulf water result in massive mid-water (150–1500 m) oxygen minimum zone (OMZ) (Wyrtki, 1971; Naqvi, 1987; Helly and Levin, 2004; Kurian et al., 2013). This prominent OMZ along with high organic input are likely to modify the geochemistry of the redox-sensitive metals in sediments. To address these problems, we studied the geochemistry of sediments of a gravity core retrieved from the eastern Arabian Shelf (12°34′N: 74°11.47′E), which is overlain by the suboxic water

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column. The objectives of the present study are (i) to work out the geochemistry of redox-sensitive elements and (ii) to identify whether this ‘known suboxic’ nature of water column can be traced from the redox proxies of the sediments. Besides, deciphering the provenance and weathering intensity through major element geochemistry has also been one of the objectives of the present work. We believe that the present study will put better constraints on the applicability of trace metal redox proxies, especially in the shelf sediments which are dominated by terrigenous inputs.

2. Samples and analytical methods A gravity core (SK-291/GC-8) was raised from the shelf region (water depth of 180 m; 12°34′N and 74°11.47′E) of southeastern Arabian Sea during ORV Sagar-Kanya cruise 291 in December 2011 (Fig. 1(a)). The dissolved oxygen concentration measured by the SBE 9plus CTD measuring instrument equipped with SBE 43 dissolved oxygen sensor, shows that the core site is overlain by a suboxic water column (Fig. 1(b)) (0.2o dissolved O2 o 2 ml/l, Tribovillard et al., 2006). The accuracy of the SBE 43 is specified as 2% of saturation, which is 0.14 ml/l for a solubility of 7.05 ml/l at 32 ppt and 6° C, or 0.11 ml/l for a solubility of 5.27 ml/l at 32 ppt and 20° C (Martini et al., 2007 and reference therein). The chemical composition of GC-8 core sediments (34 samples) were analyzed at 5 cm intervals up to 50 cm and 10 cm intervals beyond that up to 300 cm. Sediment samples were made salt free, dried at 60 °C for 36 h and pulverized for chemical analysis. Total carbon (TC), nitrogen (TN) and sulfur (TS) (CNS) were measured with a Eurovector elemental analyzer EA3000, using helium as a carrier gas. For CNS analysis, about 20 mg of pulverized sediment sample was combusted at a temperature of 1010 °C in a stream of high-purity oxygen. The gaseous combustion products were chemically reduced and were separated by gas chromatography and quantified with a thermal conductivity detector (TCD). Analytical precision as checked by parallel analysis of Sulfanilamide (CHNS grade standard), was within 2% for TC and TN and 4% for TS. Total inorganic carbon (TIC) was determined after treatment of the sediment with orthophosphoric acid in a Shimadzu TOC5000A analyzer with SSM-5000A Solid Sample Module. Total

Fig. 1. (a) Bathymetric map of SE Arabian sea having the sample location plotted by Ocean Data View software (ODV4. 6. 2, Schlitzer, 2015). (b) Dissolved oxygen in seawater measured through SBE 9plus CTD measuring instrument equipped with SBE43 dissolved oxygen sensor.

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Table 1 Results of bulk chemical and CNS analysis of GC-08 sediments. SiO2 wt%

Al2O3

Fe2O3

MnO

MgO

Na2O

CaO

0 5 10 15 20 25 30 35 40 45 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 230 240 250 260 270 280 290 300 Depth

19.071 18.512 18.229 18.486 17.804 18.332 20.062 21.895 29.218 26.931 27.701 38.702 42.204 42.84 40.149 40.707 41.786 41.24 33.691 40.608 36.757 37.815 39.221 34.711 35.001 30.196 19.208 18.333 19.978 32.978 37.845 44.612 38.605 34.926

5.739 5.48 5.392 4.752 5.156 5.276 5.799 6.547 9.047 8.269 8.468 12.289 13.656 13.721 12.374 12.743 13.048 12.806 10.344 12.711 11.465 12.097 12.438 10.647 10.718 8.997 5.571 5.36 5.771 9.851 11.6 13.785 11.886 10.56 TIC

3.564 3.743 3.815 5.179 4.635 5.023 5.671 4.325 4.663 4.52 4.361 5.638 6.185 6.133 5.915 5.931 5.834 5.804 5.102 5.54 5.517 5.309 5.987 5.703 5.679 5.782 5.543 5.568 4.956 5.715 5.783 6.037 5.852 5.805

0.032 0.032 0.032 0.035 0.034 0.035 0.036 0.035 0.044 0.041 0.042 0.061 0.07 0.069 0.069 0.071 0.069 0.069 0.056 0.068 0.065 0.073 0.065 0.06 0.062 0.051 0.037 0.036 0.039 0.054 0.067 0.075 0.068 0.059

2.571 2.562 2.606 3.121 2.312 2.648 3.004 2.73 2.95 2.861 2.798 3.424 3.622 3.678 3.273 3.429 3.493 3.408 3.251 3.408 3.371 5.105 3.295 2.995 3.104 2.918 2.365 2.347 2.258 3.049 3.313 3.775 3.328 3.108 TOC

1.548 2.455 2.277 1.865 1.239 1.935 1.802 2.204 1.454 1.551 1.319 3.102 3.114 1.837 9.332 1.806 1.941 1.71 1.589 1.798 1.872 1.814 1.765 1.473 1.559 1.651 1.449 1.002 1.107 1.502 1.682 1.894 1.729 2.119

50.124 49.668 49.671 47.153 43.111 44.334 42.409 41.287 32.431 35.439 34.412 16.8 11.66 12.299 11.347 10.939 12.953 12.648 22.702 12.018 14.363 16.444 13.275 18.487 17.305 25.812 37.963 40.67 34.586 23.844 14.907 10.868 13.971 19.256

cm 0

K2O

0.605 0.581 0.557 0.536 0.58 0.571 0.616 0.736 1.167 1.041 1.117 1.773 2.011 2.037 1.906 2.007 1.995 1.956 1.547 1.91 1.771 1.771 1.856 1.611 1.647 1.302 0.655 0.613 0.743 1.433 1.773 2.093 1.8 1.586 CaCO3

P2O5

TiO2

Cr ppm

Co

Cu

0.21 0.208 0.211 0.222 0.194 0.213 0.243 0.206 0.23 0.217 0.209 0.208 0.206 0.208 0.191 0.188 0.2 0.193 0.192 0.186 0.181 0.214 0.19 0.189 0.186 0.205 0.21 0.214 0.199 0.201 0.195 0.203 0.189 0.196

0.098 0.086 0.077 0.25 0.103 0.099 0.125 0.185 0.435 0.368 0.403 0.832 0.986 0.996 0.961 0.99 0.992 0.98 0.666 0.97 0.87 0.806 0.934 0.759 0.768 0.526 0.164 0.126 0.214 0.614 0.868 1.085 0.902 0.726

72.43 73.74 73.39 85.29 79.77 82.35 84.98 78.3 85.48 83.85 83.39 98.31 104.91 103.35 102.24 107.48 103.57 103.8 93.4 104.59 102.8 96.61 105.19 99.32 99.69 99.42 86.33 86.78 84.92 95.79 101.79 103.65 104.02 99.39 TS

19.94 20.69 20.63 24.9 25.37 26.44 28.57 23.8 26.17 25.8 25.16 33.12 36.96 36.45 36.37 36.19 35.56 35.15 30.36 34.33 33.99 32.55 36.05 33.86 33.52 32.31 29.25 29.27 27.69 32.66 34.48 36.66 35.31 33.36

17.44 17.52 17.08 18.53 17.6 17.55 16.98 18.29 23.52 22.4 24.36 34.44 38.84 37.15 37.01 39.25 37.33 37.75 29.99 37.88 36.58 35.21 37.61 32.44 32.75 26.22 18.85 19.07 22.08 27.77 34.97 40.05 36.58 29.8

V

35.17 35.92 35.24 38.19 45.74 46.6 50.62 49.9 74.02 64.05 69.3 116.31 133.88 129.8 126.34 135.05 131.14 129.19 98.8 129.93 118.34 114.35 127.75 107.66 113.31 91.49 56.02 53.85 59.31 97.96 122.05 139.56 127.09 108.92 TN

Zn

Rb

Sr

Zr

Mo

Th

U

Ni

45.97 45.65 45.61 50.07 48.58 48.53 47.49 46.43 57.25 54.61 57.86 76.98 88.28 86.4 86.04 89.27 85.43 86.51 70.91 84.73 81.09 79.22 84 73.88 74.8 62.35 66.81 54.59 56.65 67.47 81.13 89.13 83.89 71.99

9.36 9.04 8.34 2.02 7.62 6.58 5.74 10.46 28.45 24.06 28.86 66.42 83.83 80.65 83.88 90.45 83.14 82.72 51.71 83.16 72.74 70.67 73.29 54.05 56.86 31.65 8.08 6.47 14.38 38.29 67.12 88.44 70.18 48.74 LOI

1012.29 988.76 977.55 882.39 922.24 937.13 927.43 1074.22 1052.52 1089.06 1131.43 827.41 856.24 846.2 841.46 847.55 879.94 850.93 832.66 830.31 757.86 625.11 773.56 824.31 877.93 887.48 900.75 887.63 995.1 861.92 753.42 685.07 730.68 763.52

136.43 135.31 134.18 129.33 132.04 133.22 131.99 137.15 146.56 145.02 149.62 160.29 167.71 165.93 169.7 173.63 171.94 170.72 153.61 168.94 162.52 157.62 163.52 154.71 156.52 144.03 131.92 131.09 138.17 147.12 160.6 170.55 161.07 150.24

1.85 1.82 1.79 1.63 1.71 1.71 1.61 1.75 2.08 2.03 2.17 2.67 2.8 2.79 2.88 3.07 2.99 2.97 2.51 2.98 2.85 2.98 2.78 2.48 2.49 2.08 1.68 1.63 1.95 2.2 2.7 3.07 2.76 2.35

10.06 10.15 10.17 7.79 9.63 8.91 7.32 10.05 13.26 12.79 15.29 20.93 22.1 22.2 22.54 24.92 23.96 23.68 18.06 23.34 22.61 23.5 21.77 18.47 18.53 14.06 8.55 8.17 12.61 15.29 20.96 24.7 21.35 16.61

3.26 3.2 3.19 2.91 3.15 3.02 2.93 3.29 3.73 3.67 4.01 4.64 4.88 4.82 4.9 5.2 5.08 5.06 4.3 5.05 4.86 5.04 4.81 4.31 4.38 3.7 3.02 2.97 3.57 3.9 4.65 5.14 4.83 4.06

43.21 43.66 42.97 43.17 45.12 45.18 43.81 47.43 56.92 54.79 55.44 80.09 87.84 86.56 86.34 92.23 88.01 86.37 69.8 83.53 81.8 81.69 84.19 73.61 76.03 63.34 48.59 49.53 53.53 68.37 80.63 88.17 82.16 70.43 CIA

wt% 7.007

2.079

58.37

0.102

0.095

50 15

5

7.391

1.959

61.57

0.062

0.083

38.6 16

10

7.237

1.839

60.28

0.109

40

0.074 16

15

6.594

1.667

54.93

b.d.

0.078

41.4 17

20

6.546

1.86

54.53

0.068

0.069

52.3 24

25

6.56

1.585

54.64

b.d.

0.076

43 21

30

6.681

1.063

55.65

0.082

0.065

46.8 19

35

7.35

1.046

61.23

b.d.

0.054

44.9 19

S.S. Acharya et al. / Continental Shelf Research 106 (2015) 70–84

Depth cm

40

5.685

1.346

47.36

0.086

0.054

59.9 17

45

5.335

1.892

44.44

0.104

57

0.072 16

50

5.71

2.348

47.56

0.041

b.d.

60.4 17

60

3.289

0.982

27.4

b.d.

0.077

50.3 15

70

2.449

0.696

20.4

b.d.

0.077

52.4 14

80

2.117

1.183

17.63

b.d.

0.045

62.4 14

90

2.059

1.027

17.15

0.087

0.078

62.3 12

100

2.202

0.871

18.34

1.285

b.d.

61.1 19

110

1.936

1.312

16.13

b.d.

0.078

60.4 15

120

2.257

1.024

18.8

b.d.

0.081

62.3

130

4.539

0.89

37.81

b.d.

0.068

60 19

140

2.389

0.83

19.9

0.075

0.086

61.4 19

150

3.049

0.816

25.4

b.d.

0.068

58.7 22

160

4.403

1.098

36.67

0.072

b.d.

60.5 18

170

2.364

1.379

19.69

0.17

0.074

61.4 19

180

3.788

1.224

31.55

0.127

0.075

61.8 22

190

3.737

0.983

31.13

0.216

0.057

60.8 22

200

5.033

1.325

41.92

0.234

0.075

56.8 21

230

6.34

1.558

52.81

0.118

0.061

50.4 25

240

6.444

1.803

53.68

0.071

b.d.

57.5 24

250

6.482

2.199

54

0.072

b.d.

56.5 24

260

4.229

2.518

35.23

0.069

b.d.

S.S. Acharya et al. / Continental Shelf Research 106 (2015) 70–84

18

60.3 19

270

3.4

0.561

28.32

b.d.

b.d.

60.9 20

280

1.866

0.979

15.54

b.d.

0.054

61.9 14

290

2.682

1.397

22.34

0.082

0.069

60.9 20

300

3.028

2.072

25.22

b.d.

b.d.

54.9 20

b.d.: Below detection, LOI: Loss on ignition, CIA: Chemical index of alteration, CaCO3 ¼TIC  8.33.

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organic carbon (TOC) was determined by the difference between TC and TIC. Carbonate content was calculated using the TIC concentration, assuming that calcite was the only carbonate-bearing mineral and according the formula: CaCO3 = TIC × 8.33. The bulk chemical compositions of samples were determined by using X-ray fluorescence spectrometry (PANalytical, AXIOS Model). Pulverized samples of 4 g each were made into pellets by mixing thoroughly with 2 g of boric acid and subjecting to a pressure of 15 t for 4 min in a hydraulic press. Major element oxides and trace elements were measured through quantitative program SUPERQ of PANalytical using 8 International reference standards for calibration. The precision and accuracy were checked by the parallel analysis of international reference materials (STSD1, Till-4, and MAG-1). The accuracy for the major elements were better than 3% variation (except Na, o5%) and for minor elements the variation was within 4%. The results of the above analysis are presented in Table 1.

3. Results 3.1. Bulk geochemistry The bulk geochemical signatures observed in the shelf sediments are the cumulative response of three types of fractions: 1) terrigenous fraction (fluvial, aeolian, and volcanogenic), 2) biogenic fraction, represented by carbonates, organic matter and silica, and 3) authigenic fraction, composed of insoluble oxyhydroxides and sulfides (Riquier et al., 2006). To identify the origin of the studied elements (detrital, biogenic or authigenic), correlation of aluminum and calcium carbonate with selected major and trace elements was attempted (Table 2), where aluminum is commonly used as a proxy for land-derived aluminosilicate fraction of the sediments, with very little affinity to move during diagenesis (e.g. Brumsack, 1989; Calvert and Pedersen, 1993) and calcium Table 2 Correlation coefficients (r) of Al and CaCO3 with selected major and trace elements and level of significance of correlations (p). Elements

Si Fe Mn Mg Na Ca K P Ti Cr Co Cu V Zn Rb Sr Zr Mo Th U Ni TOC CaCO3 TS TN Al

Al

CaCO3

r

p

r

p

0.998 0.699 0.974 0.771 0.248  0.969 0.995  0.521 0.986 0.909 0.879 0.980 0.980 0.949 0.982  0.626 0.974 0.967 0.971 0.972 0.980  0.603  0.961 0.345  0.107 1.000

o 0.001 o 0.001 o 0.001 o 0.001 0.157 o 0.001 o 0.001 0.002 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 0.126 0.603 n.a.

 0.969  0.788  0.961  0.667  0.282 0.967  0.972 0.592  0.980  0.951  0.930  0.972  0.979  0.967  0.968 0.666  0.954  0.939  0.940  0.937  0.973 0.546 1.000  0.383 0.048  0.961

o 0.001 o 0.001 o 0.001 o 0.001 0.106 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 o 0.001 0.001 n.a. 0.087 0.815 o 0.001

n.a. – Not applicable since the correlation is with the element itself.

carbonate is primarily considered as the biogenic product of marine origin. The correlation coefficients (r) values in Table 2 clearly indicates that most of the major elements (Si, K, Ti and Mn) and trace elements (Ni, Cu, Zn, Zr, Rb, U, Th, Mo and Cr) have a siliciclastic origin and their fluctuation in sediments can be explained in terms of the variation in the detrital influx (Tribovillard et al., 2006). The negative correlation between CaCO3 versus most major and trace elements (Table 2) can be interpreted as due to dilution by carbonate production. The vertical profiles of most major and trace elements (Fig. 2) also reveal the dilution effect of calcium carbonate. The cyclicity in the variation of major elements concentrations can be traced easily from this figure. Two distinct patterns are observable at near surface sediments (10–40 cm) and at a depth of 180–270 cm. Thus, to be able to compare trace element concentrations which are not affected by the variable content of carbonate, the studied elements were normalized to Al. In addition to the element/Al ratio, enrichment factors of selected trace metals (Table 3) and some redox indices (e.g. U/Th, V/ Mo, Ni/Co, Table 4) have been evaluated to determine the redox conditions during the sediment deposition. The enrichment factor (EF) for an element X has been determined by the formula EFX = (X/Al ) sample/(X/Al ) average shale , where the average shale values were taken from Wedepohl (1971). In practice, a detectable EF for an element X corresponds to EFx 43 and a substantial enrichment to EFx 410 (Algeo and Tribovillard, 2009). Therefore, the absence of significant metal enrichments (Table 3) in the studied samples can be the result of oxygenated bottom water condition of deposition. The redox indices of studied samples have been used to derive information about paleo-oxygen level during sediment deposition based on the traditional redox classification scheme of Tribovillard et al. (2006). The results of redox indices have been reported as cross-plots in Fig. 3. All the indices of Fig. 3 refer to an oxic depositional condition. To estimate the fraction of elements which are not derived from the terrigenous input, authigenic fraction was estimated for a trace metal X by the standard formula ⎛ ⎞ X Authigenic X = Total X − ⎜Tisample × Ti ⎟ . Average Reference material ⎠ ⎝ shale (Wedepohl, 1971) and Al are generally used as reference component to calculate the authigenic fraction of metals (e.g. Calvert and Pedersen, 1993; Morford and Emerson, 1999; Tribovillard et al., 2006). However, the use of Al in the above equation (instead of Ti) gave negative values for the authigenic fractions of most elements. Hence, Ti was used instead of Al, which helped to remove the negative authigenic fraction for trace metals. The strong positive correlation between Al and Ti (r ¼0.986, po 0.001) and nearly invariant distribution of Al and Ti in the studied core (Fig. 3) suggest that the fraction of Ti in the studied sediments is predominantly terrigenous in nature. Therefore, the use of Ti in above equation serves the same purpose as Al (Xiong et al., 2012) to decipher the authigenic fraction in studied samples. The redox proxies estimated from the authigenic fraction of trace metals (Table 5) also revealed the oxic condition of deposition like the bulk chemical redox proxies (Table 4). The above result is also in line with the redox ratios of authigenic fractions estimated from the upper continental crust (Mclennan, 2001), a local granodioritic rock (Jayananda et al., 2000) and composition of saprolite after granodioritic gneisses (sample A1, Tripathi and Rajamani, 2007) (Table 5). To summarize, in the above sections, the chemical compositions of the sediments are strongly influenced by the detrital supply, which are diluted by the marine carbonates. The depositional environment remains completely oxic throughout the deposition of studied samples.

()

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Fig. 2. Vertical profiles of major and trace elements in sediments. The shaded regions show two distinct patterns at near surface sediments (10–40 cm) and at bottom sediments (180–270 cm), which are depositional signature rather than diagenetic alteration.

Table 3 Average redox sensitive element/Al ratio for GC-08 sediments, UCC (McLennan, 2001) and Average shale (Wedepohl, 1971). The enrichment factors (EF) for various elements in studied samples are calculated w.r.t average shale. Trace elements

Element/Al ratio (  10  4) GC-08

Cr Co Cu V Zn Mo U Ni

Enrichment factor (EF) Average shale

Max

Min

33.98 10.31 7.38 20.21 22.65 0.65 1.17 17.55

14.2 20.21 5.02 6.55 4.91 5.76 11.57 17.66 11.84 14.24 0.38 0.49 0.66 0.86 11.88 13.74

UC

negative correlation between MgO and TIC (Fig. 4(c)) shows that Mg is mostly incorporated into clay components, although a small contribution of Mg-rich calcite or dolomite cannot be excluded.

4. Interpretation 4.1. Weathering indices and provenance

Average 10.2 2.1 5.1 15 11 0.15 0.42 7.7

10.32 2.11 3.11 13.31 8.83 0.19 0.35 5.47

1.98 3.12 1.13 1.18 1.29 3.26 2.05 1.78

Table 4 Redox classification of depositional environment, after Tribovillard et al. (2006) and ranges of trace metal ratios for various redox domains. Redox classes and Oxygen concentration (ml/l) Redox Proxies

Oxic ( 42) Suboxic (2– 0.2)

Anoxic ( o 0.2)

References

U/Th

o 0.75

0.75–1.25

41.25

V/Cr

o2

2–4.25

44.25

V/Mo

10–60

2–10

o2

Ni/Co

o5

5–7

47

Jones and Manning, 1994 Jones and Manning, 1994 Gallego-Torres et al., 2010 Jones and Manning, 1994

3.2. CNS analyses The TOC concentration in the studied samples varies from 0.5 to 2 wt%, TS varies from 0.04 to 1.29 wt% and TN varies from 0.05 to 0.095 wt%. The TOC: TS ratio in the studied samples ranges from 4.55 to 57.26 with an average value of 19.45, which is very high compared to the normal marine sediments (2.8; Berner and Raiswell, 1983). Similarly, the molar TOC:TN ratios in the studied sediments varies from 11 to 31 (average 22) which are higher than the ratio in freshly deposited organic matter of marine system (  10, Emerson and Hedges, 1988). These ratios indicate that the organic matters in the studied sediments are dominantly of terrigenous in nature. The poor correlation between TOC and TN in Fig. 4(a) also reflect the lack of marine organic matter in studied sediments. The strong positive correlation between TIC and CaO in Fig. 4(b) shows that almost all CaO are present as CaCO3. The

As the terrigenous influx is the major controlling factor in the supply of major and trace elements in the studied samples, the variation in the source rock composition and weathering intensity were determined to reveal the subtle changes in these factors during the deposition of sediments that may affect the distribution pattern of studied elements. The intensity of chemical weathering of source-area can be quantified by chemical index of alteration Nesbitt and (CIA = Al2 O3/(Al2 O3 + K2O + CaO* + Na2O) x100; Young, 1982) and Chemical Index of Weathering (CIW = Al2 O3/(Al2 O3 + CaO* + Na2O) x100; Harnois, 1988), where CaO* represents the CaO incorporated in the silicate fraction of the samples measured and excludes CaO combined in carbonate and phosphate minerals. In these indices, Al2O3 is used as immobile element and CaO, Na2O and K2O as mobile elements. Thus, these indices increase with the degree of depletion of Na, K and Ca in sediment relative to Al e.g. CIA o60 suggests low weathering, 60– 80 moderate weathering, and 480 extreme weathering (Fedo et al., 1995). The studied sediments are carbonate rich (CaO ranges from 10–50 wt%), where the estimation of CaO* based on correction for carbonate and phosphate minerals as suggested by Fedo et al. (1995) leads to unrealistic low CIA values (ranges from 9 to 43). Therefore in this study, a correction for CaO suggested by McLennan (1993) was applied, where CaO was initially corrected for phosphate using P2O5 data (corrected-CaO ¼mole CaO*  mole P2O5  10/3). If the remaining number of moles of CaO is less than that of Na2O, the corrected-CaO value was adopted as CaO*, else CaO* value was assumed to be equivalent of Na2O. This approach excludes the impact of high carbonate content in the estimation of CaO*. Shao and Yang (2012) determined the CaO* value analytically (leached the samples by HCl to remove carbonates) and showed that the difference in the CIA values based on McLennan (1993) correction and analytically determined CaO* is not significant. The lower mean values of both CIA and CIW of studied sediments (55.6 and 60.8, respectively) in comparison to upper continental crust (UC) (McLennan, 2001) (56.92, 65.24) and Post-Achaean Australian Shale (PAAS) (Taylor and McLennan, 1985) (75.30, 88.32), suggests low to moderate source area weathering. Nesbitt and Young (1989) and Fedo et al. (1995) used the A–CN–K ternary diagram to deduce weathering trends and to infer the source rock composition.

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Fig. 3. (a) Distribution of (Cu þ Mo)/Zn ratio versus depth, Bivariate diagrams of (b) U and Th ratio, (c) Ni and Co and (d) V/Mo and V/Cr in the sediments.

The ternary plot of Al2O3–(CaO*þ Na2O)–K2O (Fig. 5) shows that all the sediments fall on a trend sub-parallel to the Al2O3–CaO*þNa2O join and within the weathering trends of average Tonalite and Granodiorite (Bahlburg and Dobrzinski, 2011), which suggests that these sediments are derived from a composition which is granodioritic to tonalitic in composition. The studied shelf region which lies in the vicinity of Mangalore receives sediments transported by the Netravati river. This river, after its origin in the Western Ghats, flows over a terrain, which is characterized by peninsular gneissic complex (Trondhjemite–tonalite–granodiorite suite; Naqvi and

Rogers 1987) and in the lower reaches the river drains plateaus covered with laterite. Netravati river watershed is composed of about 83% migmatites and granodiorite to tonalitic gneiss, 5% of charnockites, about 6% metasediments and 2% amphibolites (Gurumurthy et al., 2012). The Chikmagalur granodiorite and tonalite (Taylor et al., 1984) and granodiorite from east of Hoskote (Jayananda et al., 2000), form part of the peninsular gneissic basement in southern India. The plot of these rock types in A–CN–K diagram (Fig. 5) further confirms that the inferred source rock composition of the studied sediments seems to be in line with the regional geology of nearby region.

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Aluminum, Ti and Zr are considered to be the least mobile elements during chemical weathering. Because of this inert nature of these elements along with their short residence time in ocean, their relative proportion in the marine sediments mimics the source rock characteristics without any significant alteration (Garcia et al., 1991, 1994). The relative distribution of these elements can be used to discern the provenance by using Al–Ti–Zr ternary diagram (Fig. 6; Garcia et al., 1991, 1994). The studied samples plot closer to the granodioritic field (Garcia et al., 1991) and nearer to the composition of Chikmagalur granodiorite (Taylor et al., 1984). This inferred parent rock is similar to the source rock inferred from A–CN–K diagram (Fig. 5). The Al–Ti–Zr ternary diagram can also be used to evaluate the mechanical sorting in the sediments. Garcia et al. (1991) suggested that sediments with a wide range of TiO2/Zr in the above ternary diagram indicate high compositional maturity, unlike the narrow range of variations for compositionally immature sediments. Interpretation of this diagram is based on the premise that zircon has a strong tendency to be concentrated in the coarse grained fraction of the sediments while TiO2, although present in heavy minerals such as titanite (CaTiSiO5), illmenite (FeTiO3) or rutile (TiO2), is mainly retained in the fine grained alteration product i.e. with the clay fraction of sediments, together with Al2O3 (Garcia et al., 1991). Thus, with an increase in maturity of sediments TiO2 and Zr shows wider variation in Al–Ti–Zr ternary diagram. The presently studied sediments show a wide range of TiO2/Zr in this ternary diagram (Fig. 6) and hence indicate a greater maturity in sedimentation. Similar vertical variations of Ti, Zr and Al (Fig. 2) in the studied samples also reveal lack of grain size gradation and hence greater maturity

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Fig. 5. Major-element composition of studied sediments plotted as molar proportions on an Al2O3–(CaO*þ Na2O)–K2O diagram. Weathering indices of (1) average tonalite and granodiorite (Bahlburg and Dobrzinski, 2011); (2) Chikmagalur granodiorite and tonalite (Taylor et al., 1984) (3) granodiorite from east of Hoskote (Jayananda et al., 2000), upper continental crust (UC) and Post-Archean Australian Shale (PAAS) value (Taylor and McLennan, 1985), are also shown for comparison with studied samples. The dotted lines indicate the weathering trend of average tonalite and granodiorite.

Table 5 Mean and standard deviation (SD) of redox proxies estimated from authigenic fractions of trace elements. Reference materials

Average Shalea UCb Saprolitec Granodiorited

Redox proxies

Mean SD Mean SD Mean SD Mean SD

U/Th

V/Cr

V/Mo

Ni/Co

(Cu þMo)/Zn

0.26 0.08 0.28 0.09

0.76 0.44 0.80 0.44 3.99 2.84 0.94 0.41

25.73 13.83 28.10 13.96

1.87 0.38 2.07 0.34 0.93 0.53 6.52 5.40

0.41 0.08 0.48 0.05

a

Wedepohl (1971). Mclennan (2001). Sample A1 (Tripathi and Rajamani, 2007). d Mean composition of Hoskote granodiorite (Jayananda et al., 2000). b c

Fig. 6. Al–Ti–Zr variation diagram to discern provenance (Garcia et al., 1991).

Fig. 4. Bivariate plot of (a) TOC–TN, (b) TIC–CaO and (c) TIC–MgO. (Note: The correlation between MgO and TIC is calculated excluding abnormal value at 180 cm depth.)

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Fig. 7. Major components of studied sediments in the 5Al2O3–SiO2–2CaO ternary diagram. Data point for average shale (Wedepohl, 1971) also shown for comparison.

in sedimentation. 4.2. Relative variations of major elements Marine sediments are composed of variable mixtures of SiO2 (quartz/opal), Al2O3 (clay minerals), and CaO (carbonates). To compare the relative contribution of these major elements to the sediments of the studied core, they are plotted on a triangular diagram (Fig. 7). Most of the samples plot on the mixing line between the average shale and marine carbonate. Opal contents in the studied samples are low, which can be inferred from Fig. 7 as samples with higher biogenic silica concentration than average shale would fall closer to the SiO2 corner. Therefore, significant dilution of the terrigenous fraction by biogenic silica can be ruled out, which implies that distribution of elements is mainly influenced by dilution with carbonate, which has also been revealed from the negative correlation of major and trace elements with CaCO3, and from the vertical profiles of these elements (Fig. 2, Section 3.1). This finding is in line with the results of Kolla et al. (1981) and Schnetger et al. (2000), who reported low opal content for the shelf and deep Sea sediments of Arabian Sea. The carbonate dominated dilution of marine sediments has also been reported in

other upwelling-dominated shelf systems. For example, Plewa et al. (2012) reported low opal content in the sediments of NW Africa, where the distribution of the terrigenous fraction is proposed to be diluted by the marine carbonates. The major element geochemistry of these sediments can be explained by a simple mixing model with average shale and calcium carbonate as end members. Ti is mostly associated with the heavy minerals whereas K is derived from feldspar and micaceous minerals. Recently, Cuven et al. (2010) reported that K/Ti ratio is sensitive to the grain size variation in sediments, where the ratio increases in clay rich layers and decreases in silty facies. Hence, the similar vertical distribution of these elements (r ¼0.99, p o0.001, standard deviation of K/ Ti¼2.1) implies lack of grain size gradation in the studied core sediments. The same can also be inferred from the indistinguishable Ti–Al and K–Rb distribution in samples (Fig. 8), which are commonly used to discern detrital matter provenance (Shimmield, 1992). Manganese and iron concentrations have been normalized with that of Al to account for the fraction of these elements which are not diluted by carbonates. The Mn/Al ratio varied from 0.007 to 0.01 whereas the Fe/Al ratio ranged from 0.5 to 1.4. These ratios are above the upper continental crustal ratios of 0.0075 and 0.44, respectively (McLennan, 2001), indicating the presence of structurally unsupported Mn and Fe in the form of Fe–Mn oxyhydroxides. These metal oxyhydroxides are the prime scavengers of heavy metals in marine systems. Koschinsky et al. (2003) showed through laboratory sorption experiments that hydrated cations and labile cationic chloro-complexes in seawater like Co2 þ , Ni2 þ , Cu2 þ , Zn2 þ , and PbCl þ are preferentially adsorbed or ion-exchanged on the negatively charged surfaces of Mn oxides. The strong positive correlation of these elements with Mn (Fig. 9(a)– (e)) in the studied sediments supports the adsorption phenomena. However, the strong positive correlation of Mn and trace metals with Al (Section 3.1 and Table 2) suggests negligible authigenic Mn-oxide contribution to the studied sediments. Therefore, the strong positive correlation of heavy metals with Mn (Fig. 9) should be considered as a terrestrial signature, where after release from the source rock, the trace metals (e.g. Cu, Co, Ni etc.) got adsorbed onto the surface of Fe–Mn oxides and transported to the marine system. The release of heavy metals during chemical weathering and erosion and their subsequent scavenging by Mn and Fe oxides have been documented for the River Netravati (Gurumurthy et al., 2014), which is the main transporting agent of studied sediments (Section 4.1).

Fig. 8. Scatter plots of (a) Al–Ti and (b) K–Rb. Strong linearity suggests the lack of grain size gradation in sediments (see text for details).

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matter. 2) Negligible diagenetic pyrite formed in organic-rich sediments supplied by rivers as compared to analogous marine sediments because of the much lower concentrations of dissolved sulfate found in most fresh waters as compared to seawater (Berner and Raiswell, 1984). These result in very high TOC:TS ratio in sediments. 3) Under oxic conditions of deposition, there is a lack of sufficient diagenetic Fe and dissolved sulfide in the pore waters, which result in lower TS and hence, higher TOC:TS ratio in sediments (Qi et al., 2010). This interpretation of the terrestrial organic matter is further supported by the molar TOC:TN ratio in the studied sediments, which varies from 11 to 31. These high values can be the result of the dominance of organic matter of terrestrial origin. High-molecular-weight plant organic matters (e.g. lignin) of terrestrial system contains lower amount of nitrogen and thus have a higher TOC:TN ratio unlike marine organic matter which are relatively rich in proteins and shows low TOC:TN ratio. The molar TOC:TN ratio of phytoplankton and zooplankton is  6.6 (the Redfield value); freshly deposited organic matter has a ratio of  10; and terrigenous derived materials has ratios ranging from 20 to 200 (Emerson and Hedges, 1988). The lack of significant correlation between TOC and TN (r¼0.28, p¼0.189) also ruled out the possibility of marine-derived organic matter. This finding is in line with the work of Prasad et al. (2007) and Sardessai (1994), who reported terrestrial dominated organic matter in the mid and outer eastern Arabian shelf sediments. Recently, Pradhan et al. (2014) documented the sources and distributions of organic matters of Netravati and Gurpur estuaries and reported the dominance of plant organic matter having TOC:TN ratio ranging from 10 to 30 (Pradhan et al., 2014, Fig. 3). The similarity of the aforementioned range of TOC:TN ratio with the present studied sediments confirms the transportation of these organic matters to the studied location. 4.4. Geochemistry of redox-sensitive elements Many trace elements are redox-sensitive and become moderately to highly enriched under anoxic bottom water conditions, making them useful as indicators of paleoredox conditions. In this section, the pattern of variation of redox-sensitive elements such as Mn, V, Cu, Zn, Mo, Ni, and U and their inter-elemental relationships are examined.

Fig. 9. Relative distribution of Co, Cu, Zn, Ni and Pb with respect to Mn in the sediments.

4.3. Nature of organic matters The nature of organic matter (i.e. terrigenous or marine) can be defined with the help of TOC:TS and TOC:TN ratio. The average TOC:TS ratio in the studied samples (19.45) is very high as compared to the normal marine sediments (2.8; Berner and Raiswell, 1983). This may be the result of dominance of terrigenous organic matter deposited under oxic conditions in the sediments. Such high ratio can be the result of following three factors. 1) The terrestrial organic matters are less labile and less susceptible to the microbial activity and hence have a greater rate of preservation than marine organic matters (Aller, 1998 and references therein). Thus, it results in higher TOC:TS ratio than normal marine organic

4.4.1. Molybdenum, Copper and Zinc In marine sediments, Mo is considered as a proxy of redox conditions because of its conservative behavior in oxygenated waters and enrichment in anoxic sediments. Mo distribution in the studied sediments closely follows that of Mn and Fe that is reflected by the strong correlation with Mn (r¼ 0.97, po0.001) and Fe (r¼0.63, po0.001). This is possibly due to the association of Mo with Mnoxyhydroxides and Fe-monosulfides (Shimmield and Price, 1986). Another alternative of Mo removal from seawater is the fixation of Mo by particulate or dissolved organic matter (Brumsack and Gieskes, 1983). As TOC contents in the sediments show a negative correlation with Mo (r¼  0.6, po0.001), its distribution cannot be attributed to organic matter. The lack of significant marine input of organic matter may be the reason for this relation. The ratio R = (Cu + Mo) /Zn has been proposed by Hallberg (1976), as an indicator of the oxygenation of bottom waters. The basic principle behind the use of this ratio lies in the fact that in reduced environment with H2S in bottom water, the precipitation of Cu is favored over Zn in the sediments, which is the result of differences in the solubility product of their sulfides in reduced environments (Hallberg, 1976 and references therein). Hence, this ratio is expected to increase under anoxic conditions and decrease under oxidizing conditions. The variation of R within a small range (0.3–0.48) in the studied sediments reflects the predominance of oxic condition

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during the deposition of sediments (Fig. 3(a)). A similar observation was also made in the sediments of Eastern Gotland Basin, where Hallberg (1976) reported the R values within 0.25–0.55 for the sediments ventilated by the inflows of oxygenated water. 4.4.2. Uranium and Thorium In oxic water, U is present as soluble U(VI) in the form of tricarbonate species (UO2(CO3)34-). In oxygenated marine settings, dissolved U(VI) is neither scavenged by particulates to the sediments nor is it reduced to thermodynamically favored U(IV) (Anderson et al., 1989). However, anoxic basins are the sites of authigenic U deposition where reduction of U(VI) to U(IV) occurs under conditions similar to those of Fe(III) to Fe(II) reduction (Tribovillard et al., 2006). The U concentration in the studied sediments (mean 4.1 ppm) is nearly similar to the average shale value (3.7 ppm, Wedepohl, 1971) without any significant enrichment (2.05, Table 3) which suggests an oxic depositional condition. Barnes and Cochran (1993) use laboratory experiments to suggest that under oxic conditions U gets associated with Fe–Mn oxides. In the studied sediments, the strong positive correlation of U with Mn (r¼0.97, po0.001) and moderate positive correlation with Fe (r¼0.62, po0.001) also reflects the oxic depositional condition. Jones and Manning (1994) suggested that authigenic U contents (Authigenic U¼Total U Th/3, Wignall and Myers, 1988) of o5 ppm are considered to be indicative of oxic environments. The absence of authigenic U in the present sediments further supports an oxic depositional environment throughout the core length. In contrast to U, This relatively immobile and is associated mostly with the detrital clay fraction of the sediment. Thorium concentration in the sediments varies between 8 ppm to 22 ppm with the mean concentration of 16.6 ppm. Because of the different behavior of U and Th, U/Th ratio can be used as a redox indicator. High U/Th ratio ( 41.25) suggests deposition of sediments in anoxic condition, whereas lower values ( o0.75) indicate the oxic environment of deposition (Jones and Manning 1994). The low and almost uniform U/Th ratio (average 0.27) in the core is indicative of prevalence of an oxic environment without any drastic change throughout the period of core deposition (Fig. 3(b)). 4.4.3. Nickel and Cobalt In oxic environments, Nickel behaves as a micronutrient which occurs as soluble Ni2 þ or NiCl þ ion (Tribovillard et al., 2006). Nickel complexation with organic matter accelerates scavenging in the water column and thus its enrichment in sediments (Calvert and Pedersen, 1993). Unlike Ni, Cobalt behaves similarly as Mn in seawater and sediments, i.e. it can diffuse out of sediments under

reducing conditions (Heggie and Lewis, 1984) and hence increase the Ni/Co ratio. Jones and Manning (1994) assigned Ni/Co ratios o5 for oxic conditions, 5–7 suboxic conditions, and 47 for anoxic conditions. The studied sediments reflect the oxic condition of deposition, as Ni/Co ratio varies within a narrow range of 1.5–2.5 with an average of 2.16 (Fig. 3(c)). 4.4.4. Vanadium and Chromium In oxic water, vanadium is present in the pentavalent state as vanadate oxyanions (HVO2− and H2 VO−4 ). In pelagic and hemi4 pelagic sediments, V is strongly coupled with the redox cycle of Mn (Tribovillard et al., 2006). Similar vertical distribution of V and Mn in the studied samples (Fig. 10(a)) with strong positive correlation (Fig. 10(b)) of V with Mn and Fe (r ¼ 0.98 and 0.8 respectively, po 0.001) suggests adsorption of vanadate onto Mn- and Fe-oxyhydroxides (Calvert and Piper, 1984). The V/Cr ratio is a redox indicator which reflects changes in the scavenging efficiency as a function of redox conditions (Gallego-Torres et al., 2010; Jones and Manning, 1994; Riquier et al., 2006). Under oxic conditions vandate adsorbs more strongly than chromate, whereas in reducing condition Cr (III) forms stronger surface complexes than VO2 þ . Jones and Manning (1994) suggested that V/Cr ratios o2 refers oxic conditions, 2–4.25 suboxic conditions, and 44.25 anoxic conditions. The average V/Cr ratio in the studied samples is low ( o2) which suggests an oxic environment of deposition (Fig. 3(a)). Unlike V which starts precipitating in suboxic condition, Mo only starts to precipitate when dissolved sulfide (H2S) is available (euxinic condition). Thus, when V/Mo ratio in sediments approaches the seawater ratio (o2.0 approximately), it indicates an anoxic condition of deposition (Gallego-Torres et al., 2010). This ratio between 2 and 10 indicates suboxic conditions, and between 10 and 60 it indicates normal oxygenation (Gallego-Torres et al., 2010). In the studied sediments, V/Mo ratio ranges from 19 to 48 (Fig. 3(d)) and lies within the range of oxic condition of deposition. 4.4.5. Authigenic redox proxies The bulk chemical composition of trace metals includes detrital input which is supplied by rivers to the marine system, biogenic input which is delivered by biological activity in the ocean and input from bottom ocean waters. It is this latter fraction of the trace metals (authigenic) which is used by some workers in detrital dominated system (Calvert and Pederson, 1993; Tribovillard et al., 2006 and references therein) to evaluate the redox condition of deposition. The redox ratios of the authigenic fractions of trace metals estimated from average shale (Wedepohl, 1971) composition reveals the oxic

Fig. 10. (a) Vertical distribution of Mn and V in the sediments. (b) Fe and Mn variation with reference to V in the sediments.

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condition of deposition. From the A–CN–K and Al–Ti–Zr plot (Figs. 5 and 6), it is clear that the studied sediments are deposited through the Netravati river (Section 4.1) from a source rock of granodiorite to tonalitic composition. Therefore, upper continental crust (Mclennan, 2001), a local granodioritic rock (Jayananda et al., 2000) and composition of saprolite after granodioritic gneisses (sample A1, Tripathi and Rajamani, 2007) were also used to compare the authigenic fraction of redox-sensitive trace metals. The above results (Table 5) show that the depositional environment remains oxic throughout the deposition of studied sediments. Although, V/Cr ratio for saprolite and Ni/Co ratio for local granodioritic rock shows sub-oxic nature (Table 5), they can not be used as a reference for suboxic condition of deposition, because of two reasons: 1) the source rock composition of studied samples varies from granodiorite to tonalite, which might affect this estimation from a single granodioritic source rock, and 2) according to Jones and Manning (1994) and Rimmer et al. (2004), the various trace metal ratios must be considered collectively to evaluate the redox condition of deposition instead of giving weightage to the individual ratios. Therefore, an oxic condition of deposition can be traced from the aforementioned redox proxies.

5. Discussion The redox condition of marine sediments is the cumulative response of organic carbon and bottom water oxygen concentration. High organic carbon flux to the sediment and low bottom water oxygen concentrations favor the development of anoxic conditions close to the sediment-water interface (McKay et al., 2007). This redox condition can be traced by the trace metal distribution in sediments. However, in the present study trace metal redox proxies do not reflect the bottom water condition (Fig. 1(b)). This can be the result of the following scenarios: 1) the present day bottom water condition was different from that of the past or in other words the presently observed reducing condition is a recent phenomenon; 2) the original trace metal responses are modified by diagenetic signature after burial of the sediments; or 3) the trace metal redox proxies (Section 4.4) are not consistent with all depositional environments. The OMZ in Arabian Sea is the cumulative response of three factors: 1) very slow movement of water masses within the OMZ, which leads to depletion in oxygen concentration by organic matter decomposition, 2) very large scale oxygen consumption rate, which results from enormous primary production due to the strong upwelling in Arabian Sea and 3) low oxygen concentration of the water masses entering to the intermediate depths of Arabian Sea from Indian ocean (Swallow, 1984; Olson et al., 1993). The intensity of this OMZ has varied on orbital and sub-orbital time scales (Altabet et al., 1995; Reichart et al., 1998; Ten Kate et al., 1994). By using the proxies for surface water productivity, water column denitrification, winter mixing and aragonite compensation depth, Reichart et al. (1998) reported that this variability is mainly driven by the changes in surface water productivity and deep winter mixing. They reported that except in the short span of Heinrich events, the OMZ remains mostly stable in Arabian Sea. The same was also documented for a sedimentary core within OMZ of northern Arabian Sea by Schulte et al. (1999), who used organic geochemical proxy (C35/C31-n-alkane ratio) to study the redox conditions in the sediments. These studies suggest that there is no drastic change in OMZ condition that could have happened during the Holocene period. Agnihotri et al. (2003) observed low Mn/Al ratio throughout the deposition (from  10 ka BP to present) in 3268G5 core (water depth 600 m, 12.5°N, 74.2°E) near Mangalore offshore, that suggests that Mn was being mobilized out of the sedimentary column, a phenomenon commonly

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observed in reducing environment. This suggests that the bottom water condition of studied location remain static during this period of deposition. Recently, by using the enrichment of Mo and Cr, Naik et al. (2014) reported prevalence of suboxic depositional environment from late-Holocene to present in the sediments from Goa offshore (Core AAS9/19; water depth of 367 m,14°30.115′N; 73°08.515′E). The linear sedimentation rate in the present studied location is around 17 cm/kyr was calculated to be around 17 cm/ kyr from δ14C of core 3268G2 (12°31.1′N, 74°09.4′E) from a water depth of 370 m (Somayajulu et al., 1999), which suggests that the studied sedimentary column covers a time span of  18 kyr (approximately) and has overlain by nearly the same reducing bottom water condition as present. Therefore, the first assumption of variable bottom water condition can be ruled out. Postdepositional diagenesis is an important sedimentary process wherein Mn, Fe, U and sulfate are reduced at the expense of organic carbon destruction (Somayajulu et al., 1999 and references therein) which in turn modifies the original extent of the metal enrichments (Tribovillard et al., 2006). Under some conditions, oxidizing agents can be replenished from above after the development of reducing conditions in the sediments by abiogenic or bacterial organic matter oxidation. These post-depositional processes may modify the original depositional anoxic signatures of trace metals by remobilization of redox-sensitive elements such as U, Mo, V and Cd (Tribovillard et al., 2006, and reference therein). The impact of diagenetic overprint on the studied responses of trace metals can be ruled out based on the strong positive correlation of detrital elements Al and Ti with various redox-sensitive metals (Mn, V, Cu, Zn, Mo, Ni, Th and U) (rZ 0.88, p o0.001) which reflects lack of significant remobilization of elements after burial of sediments. The lack of significant diagenesis may be the result of dominant terrestrial organic matter input to the studied site (see Section 4.3). The recalcitrant terrestrial organic matters are less labile and less susceptible to the microbial activity and hence cause minor changes in the redox condition than the marine organic matter (Aller, 1998 and references therein). Therefore, the third assumption seems correct in the present context i.e. the trace metal redox proxies are not suitable at all depositional environments (shelf/deep Ocean). The response of trace metals in a particular depositional environment depends on the stability of the water column anoxia and the sedimentation rate which defines the exposure time of trace elements to that environment (Tribovillard et al., 2006). The sedimentation rate in this part of the Arabian shelf is relatively high (17 cm/kyr, value for 14C calibrated core 3268G2; Somayajulu et al., 1999), which may reduce the exposure time of trace elements to the suboxic water column and hence an oxic signature may be reflected in sediments. Similar observation was also made in continental margin sediments of NE pacific (core 01, McKay et al., 2007), where despite reducing conditions, redox-sensitive trace metals did not show any enrichment throughout the core. The current studied location is present at the upper edge of Arabian Sea OMZ (150–1500 m) which is also dominated by the strong seasonal phenomenon: the monsoons. The surface productivity and mid-depth water circulation respond to this seasonal phenomenon in Arabian Sea (Naqvi, 1987), both of which controls the O2 concentration in sub-surface water. This seasonal variation in O2 concentration may result in the instability of the water column anoxia in the studied location. To study the seasonal forcing on OMZ variation, dissolved oxygen concentration in gridded (1°  1°) World Ocean Atlas 2013 (WOA13, Garcia et al., 2014) database was studied along a traverse in the studied location. WOA13 provides the most updated and largest dataset of dissolved oxygen and nutrients ranging from years 1955–2012. We have used the analyzed fields of WOA13 dataset which are the arithmetic mean of interpolated values of observed depth levels to

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Fig. 11. Seasonal variation of dissolved O2 concentration in seawater along a traverse near Mangalore. The dissolved O2 concentrations are taken from World Ocean Atlas 2013 for four seasons: winter: January–March; spring: April–June; summer: July–September; fall: October–December. The data are plotted using the Ocean Data View software (ODV 4.6.2, Schlitzer, 2015).

102 standard depth levels (Garcia et al., 2014). The WOA13 dataset provides better resolution at shallow depths as compared to the earlier released databases. In the following description, we adopted the O2 concentration of 0.5 ml/l as the threshold value for OMZ from the study of Kamykowski and Zentara (1990) which has been used in several works of OMZ characterization (e.g. Helly and Levin, 2004; Fuenzalida et al., 2009; Paulmier and Ruiz-Pino, 2009). The oxygen distribution in the upper 400 m water depth is given in Fig. 11 for the four seasons along a traverse near Mangalore. During the spring (April–June) to summer (July–September) transition, there is a clear thickening of OMZ upper layer (Fig. 11; 0.5 ml/l contour) appears in the studied traverse as a result of which the studied location seems to be overlain by oxygen rich water column during this time period. Such thickening of OMZ was earlier reported by Paulmier and Ruiz-Pino (2009) who quantified this thickening of seasonal transition using World Ocean Atlas 2005 database. According to their estimation, during spring-summer transition there is a thickening of 20% OMZ thickness in Arabian Sea. This is associated with a shoaling of average upper OMZ CORE limit from 280 to 220 m and 430% decrease in mean O2 concentration in OMZ CORE, where OMZ CORE was defined by the water mass having dissolved oxygen concentration of o20 μM. During fall (October–December) to winter (January–March) transition, Fig. 11 suggests a clear shoaling of OMZ upper boundary layer. Therefore, the seasonal variation (Fig. 11) of O2 distribution suggests that the water column anoxia above the studied location is not stagnant throughout the year. This will put the similar impact as high sedimentation rate on trace metal response to water column anoxia i.e. reduction of exposure time to the anoxic water column. Therefore, the use of these trace metals proxies should be considered cautiously in continental shelf environments which are characterized by high sedimentation rate and significant seasonal variation of O2 concentration in the water column. 6. Conclusions The results of the present study allow the following conclusions to be drawn with the help of sediment geochemistry of SK-291/

GC-8 gravity core retrieved from the water depth of 180 cm of NE Arabian Sea, which lies within the present OMZ (150–1500 m). 1. The studied sediments are the result of low to moderate source area weathering of granodioritic to tonalitic composition. 2. Similar vertical distribution of major and trace elements along with their strong positive correlation with Aluminum and Titanium reflects the dominance of terrigenous input and lack diagenetic alteration throughout the core. 3. The TOC:TS and molar TOC:TN ratios in the studied samples suggest that the organic matters are mainly of terrestrial origin. 4. Various redox proxies (viz. enrichment of redox-sensitive elements, authigenic U, and ratios of trace metals such as U/Th, V/ Cr, V/Mo, Ni/Co and (Cu þMo)/Zn) indicates an uniform oxic depositional environment, despite the fact that the present location lies within the OMZ. 5. The estimated redox proxies from bulk composition and marine fraction of trace metals do not reveal any significant difference in the depositional environment. 6. The present oxic signature of trace metal redox proxies in an anoxic environment is the cumulative response of high sedimentation rate and significant seasonal forcing on dissolved O2 concentration in the water column. Therefore, the use of trace metal proxies as indicator of the redox state of deposition of sediments in continental shelf environment seems to be more tricky an affair than what is thought so far.

Acknowledgments We acknowledge support from the Ministry of Earth Sciences, Government of India in the form of a research project (Sanction number MoES/16/49/09/RDEAS). SSA acknowledges financial support from the Ministry of HRD, Government of India in the form a research scholarship available through the host Institute. The Chemical Engineering Department of the host Institute is acknowledged for instrumental support for bulk chemical analysis. We are greatly thankful to the National Centre for Antarctic and

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Oceanic Research for allowing us (MKP and SSA) to be a part of the cruise SK 291. We thank Nicolas Tribovillard and an anonymous reviewer of the Journal for their critical review that helped us to better our understanding and also the quality of our presentation. We are indebted to Tim Jickells, Associate Editor, for his supportive role in handling our manuscript.

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