Sediments

C H A P T E R

F O U R

Sediments

Contents 117 117 124 135 137 139 146 149 151

1. Distribution and Source 1.1. Distribution 1.2. Source 2. Sedimentary Processes 2.1. Dissolution of carbonate 2.2. Bottom water mass and sedimentation 2.3. Depositional environment 2.4. Diagenesis 2.5. Sediment consolidation

Sediments are formed from disintegrated rocks as a result of physical and chemical weathering. The action of various agents, like ice, wind, water and variable temperature, helps fragment the rocks into smaller particles and leach the more soluble minerals. Weathering also results from reaction of seawater with basalt, erupting at the crests of the mid-oceanic ridges and at other submarine volcanic features, contributing in the process considerable amounts of materials to seawater. Another significant contributor has been extraterrestrial matter. Transported either in a dissolved or suspended state, all these materials ultimately form four types of sediment—biogenous and hydrogenous through precipitation, and lithogenous and cosmogenous in a clastic detrital state. Among the biogenous deposits are calcite, aragonite, opal, phosphorites and organics. The hydrogenous types are evaporites, zeolites, manganese nodules and polymetallic sulphides, while silica, feldspar and rock fragments constitute deposits of lithogenous type. The contributors to cosmogenous deposits have been the cosmic spherules, microtektites and minitektites (Chester, 1993; Lisitzin, 1996). In general, mechanical weathering dominates in temperate climates at high latitudes, where water in the form of ice is the chief weathering agent. Chemical weathering (i.e. leaching), on the other hand, is favoured by high rainfall, variable temperature and dominates in tropical areas. Hence, any sedimentary basin reflects the palaeo- and present environmental conditions of sedimentation including the processes and provenances. The distribution of various sediment types in the world’s oceans has been recorded with reliable accuracy (Fig. 4.1; Table 4.1). There have been considerable activities in recent years to understand the sediment distribution and sedimentary processes in the Central Indian Ocean Basin (CIOB), a basin which Handbook of Exploration and Environmental Geochemistry, Volume 10 ISSN 1874-2734, DOI: 10.1016/S1874-2734(07)10004-8

#

2008 Elsevier B.V. All rights reserved.

115

116

Mukhopadhyay, Ghosh and Iyer

160W

100

40W

0

40E

100E

60 N

30

0

30

60 S

Ice rafted

Carbonate

Siliceous

Red clay

Terrigenous

Siliceous/red clay

Figure 4.1 Sediment distribution in the world oceans (Kolla and Kidd, 1982; The Open University,1995).

Table 4.1

Distribution (%) of various sediment types in the world oceans

Major type

Minor type

Red clay Siliceous ooze

Diatom ooze Radiolarian ooze Calcareous ooze Foraminiferal ooze Pteropod ooze

Atlantic

Indian

Pacific

All oceans

26 07 – 65 02

25 20 0.5 54 0.1

49 10 05 36 –

38 12 03 47 0.5

Sources: Berzukov (1960) and Stowe (1996).

hosts the sediments of the two largest rivers of the world, the Ganges and the Bramhaputra, and also holds the second richest and second largest manganese nodule field in the world oceans. Hence, the sediments described throughout this chapter, if not specifically mentioned, are essentially those of the CIOB, which also includes the Indian Ocean Nodule Field (IONF).

117

Sediments

1. Distribution and Source 1.1. Distribution The nature and distribution of seafloor sediments in the Indian Ocean are principally controlled by five interrelated factors: (1) climatic and current pattern, (2) nutrient and organic production in surface waters, (3) relative solubility of calcite and silica, (4) submarine topography and (5) detrital input. Based on their interactions, four major types of sediments occur—terrigenous, calcareous, siliceous and pelagic (Fig. 4.2). Covering about 35% of the CIOB, calcareous sediments, with a sedimentation rate of 4–6 mm/103 year, are common along the equatorial highproductivity areas and at lesser depths, such as near the seismic and aseismic ridges (the Chagos Ridge, the Ninetyeast Ridge) and the shallow areas of seamount

60E

70

80

90 10 N

0

10 S

20

30

Calcareous Ooze

Siliceous Clay

Calcareous Clay

Brown Clay

Terrigenous Clay

Terri-Siliceous Clay

Mixture of Terrigenous, Calcareous, Siliceous Clay Southern limit of Indonesian Volcanic Tephra

Figure 4.2

Sediment distribution in the Indian Ocean (Udintsev,1975; Kolla and Kidd,1982).

118

Mukhopadhyay, Ghosh and Iyer

summits. Siliceous clay (and ooze) and red-brown clay, on the other hand, are the dominant sediment types in the deeper parts of the basins in the CIOB, and cover about 35 and 16% of the surface area, respectively. These sediments are found in areas where the rate of sedimentation is low (<2 mm/103 year) and terrigenous contribution is absent. The red-brown clay is eupelagic and contains <25% of the coarser fraction of either lithogenous or volcanic derivation. Terrigenous sediments, on the other hand, are hemipelagic and contain more than 25% terrigenous or coarse volcanic material. This sediment type, with the highest rate of sedimentation, is largely detrital brought about from the continents by the large rivers and covers
14% of the CIOB (Chester, 1993; Lisitzin, 1996). In the CIOB, the broad distribution of sedimentary facies is now available based purely on qualitative information from sediment samples obtained through box cores, Pettersson grabs and Van Veen grabs (Table 4.2). The enormous load of terrigenous materials (about 1.67 million tons/year) in the central and the eastern Indian Ocean, including the northernmost part of the IONF, is contributed mainly by the Ganges and the Bramhaputra rivers. This amounts to about 24% of total load carried by all world rivers (Milliman and Meade, 1983). These two rivers bring materials largely from the mighty Himalayas and from the Indo-Gangetic plains. The southern extent of terrigenous sediment distribution in the CIOB appears intriguing. Kolla and Biscaye (1973) identified terrigenous influence up to 10 S, based on the clay mineralogy and up to 8 S, based on chemical analyses of surface sediments (Nath et al., 1989). The major element ratio Al/(Al þ Fe þ Mn) was used to determine terrigenous extension. Four long-gravity cores of sediment, one each from the four sectors in the IONF were studied in detail to trace the limit of terrigenous influence—vertically and horizontally. Decreasing contents of detrital components such as Al, Ti, K and Fe and a decreasing ratio of Al/(Al þ Fe þ Mn) towards the southern sectors indicate diminishing influence of terrigenous influx. The ratio lowers the cut-off limit of 0.65 at around 10 S. The Al/(Al þ Fe þ Mn) ratio varies from 0.5 to 0.77 in these sediment cores and is lower (<0.65) near the top and base of the cores, with the ratio reaching a maximum value (0.77) at an intermediate depth corresponding to about 140–100 ka at 14 S (Mudholkar et al., 1993) and to the last interglacial maxima (Curray and Moore, 1971). All these studies in the CIOB approximately constrain the extension of terrigenous sediment in the northern part up to about 7 S, siliceous in the central part between 8 S and 15 S, calcareous sediments to the west of 74 E bordering mid-oceanic and Chagos ridges and brown clay in the southern part south of 15 300 S (Banerjee, 1998; Kolla and Kidd, 1982; Nath, 2001). The areas roughly between 7 S and 8 S and between 15 S and 15 300 S are composed of mixed terrigenous–siliceous sediments and siliceous–pelagic clays, respectively. Among other materials reported to occur in the CIOB are volcanic ash (tephra), pumice, aerosols and extraterrestrial materials. Much of the volcanic ash contributed by erupting volcanoes and carried by wind is finely dispersed, and sometimes forms several centimetre thick layers in deep-sea sediments. Some of these ash layers are correlatable over great distances, marking periods of major volcanic eruptions. A well-known example is the ash that the Toba supereruption spewed out in

Table 4.2

Location of sediment sampling stations in the IONF

Latitude ( S) 

0

Longitude ( E) 

10 15 –10 26

10 00

10 00

10 00

12 00

12 30

0



0



75 59.5 –76 15

75 15

75 37

76 00

76 07.50

75 52.50

0

Sampler type

Water depth (m)

BC

5400

BC

5400

BC

BC

BC

BC

5400

5400

–

–

Sample ID by author

T1

R1

A1

R2

T2

Type of studies made

▪ 12 nos. core in the region ▪ Physical properties ▪ Grain size þ clay mineralogy ▪ Geotechnical properties ▪ Grain size þ clay mineralogy ▪ Geotechnical properties ▪ Grain size þ clay mineralogy ▪ Geotechnical properties ▪ Grain size þ clay mineralogy ▪ Geotechnical properties ▪ Grain size þ clay mineralogy ▪ Geotechnical properties

Reference

Sector

a

C

b

B

a,c

B

b

B

a,c

B

b

B

a,c

B

b

C

a,c

C

b

C

a,c

C

(continued)

120

Table 4.2 (continued) Latitude ( S)

Longitude ( E)

Sampler type

Water depth (m)

Sample ID by author

Type of studies made

15 300

73 00

SC þ PG

4650

129

12 00

75 300

SC þ PG

–

128

11 00

75 450

SC þ PG

5039

287

12 010

73 00

SC þ PG

5000

124

11 300

81 280

SC þ PG

–

139

15 00

73 300

SC þ PG

4900

231

12 010

73 020

SC þ PG

5075

121

14 00

72 00

SC þ PG

5000

126

14 00

73 00

SC þ PG

4389

56

12 500

78 00

SC þ PG

4977

81

13 500

74 00

SC þ PG

5150

47

11 15.60

75 00.70

BC

–

AAS 4/1

12 00.10

75 29.90

BC

–

AAS 4/2

▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Biogenic silica concentration ▪ Tektites, minitektites ▪ Tektites, minitektites

Reference

Sector

d

D

d

C

d

C

d

C

d

C

d

D

d

C

d

D

d

D

d

C

d

D

e

C

e

C

12 30.50

76 30.90

BC

–

AAS 4/5A

12 36.90

78 30.70

BC

–

AAS 4/6

12 370

78 300

BC

–

AAS 4/6

11 44.60

73 59.70

PG

4896

22

14 470 –15 59.80

73 440 –76 59.70

PG

4969–5161

7 sites

13 080

75 010

SC

5270

SK226

14 00

9 990

76 00

77 920

SC

SC

5050

5250

SS657

NR-1

▪ Tektites, minitektites ▪ Tektites, minitektites ▪ Tektites, minitektites ▪ Geochemistry, mineralogy ▪ Geochemistry, mineralogy ▪ General geochemistry ▪ Major elements, Toba tuff ▪ Glass shards, excess Al ▪ General geochemistry ▪ Major elements, Toba tuff ▪ Glass shards, excess Al ▪ REE concentration ▪ U, Th isotope, transition ▪ Major elements, Toba tuff

e

C

e

C

e

C

f

C

f

D

g

C

h i g

D

h i j j

A/B

h

121

(continued)

122

Table 4.2 (continued) Latitude ( S)

11 00

11 970

10 00–15 00

10 00–15 00

15 00–16 00

Longitude ( E)

78 490

78 490

72 00–74 00

75 00–80 00

72 00–82 00

Sampler type

SC

SC

PG

PG

PG

Water depth (m)

5325

5450

–

–

–

Sample ID by author

NR-21

NR-35

21, 22, 23, 33

1, 17, 19, 26, 27, 29, 32

2, 30, 31

Type of studies made

▪ Glass shards, excess Al ▪ U, Th isotope, transition ▪ Major elements, Toba tuff ▪ Glass shards, excess Al ▪ U, Th isotope, transition ▪ Major elements, Toba tuff ▪ Glass shards, excess Al ▪ XRD, infrared, DTA, clay ▪ Major elements ▪ REE analyses ▪ XRD, infrared, DTA, clay ▪ Major elements ▪ REE analyses ▪ XRD, infrared, DTA, clay ▪ Major elements ▪ REE analyses

Reference

Sector

i

C

j h i j

C

h i k

B/C/D

l m k

B/C/D

l m k

D

l m

11 00–15 00.50

72 01.60 –81 490

PG, BM

4300–5380

12 00

76 300

SC

5450

12 00

77 00

SC

5430

– F200B

F88B

14 00

74 00

SC

5240

SK176

12 00

76 400

SC

5300

GR1–120E

12 500

76 000

SC

5250

SS667

9 00–14 00

75 590 –76 010

SC

5158–5352

SPC14–19

9 00

77 00

BC

5400

–

13 030

75 440

BC

5099

GC02

▪ Radiolarian zonation ▪ U, Th, biogenic Si, chem. ▪ Radolarian zonation ▪ U, Th, biogenic Si, chem. ▪ Radolarian zonation ▪ U, Th, biogenic Si, chem. ▪ U, Th, biogenic Si, chem. ▪ U, Th, biogenic Si, chem. ▪ Geochemistry assessing biogenic and detrital influence ▪ Geotechnical properties

n

C

o,p

C

REE

q o,p

C

q o,p

D

o,p

C

o,p

C

r

A/B/ C/D

s

A

t

C

Sources: a, Khadge (2002); b, Valsangkar and Ambre (2000); c, Khadge (2000); d, Pattan et al. (1992); e, Prasad and Khedekar (2003); f, Banerjee (1998); g, Mudholkar et al. (1993); h, Pattan et al. (1999); i, Pattan and Shane (1999); j, Pattan and Banakar (1997); k, Rao and Nath (1988); l, Nath et al. (1989); m, Nath et al. (1992); n, Gupta (1996); o, Borole (1993a); p, Borole (1993b); q, Gupta (1988); r, Banakar et al. (1998); s, Khadge (1998); t, Pattan et al. (2005). Note: BC, box core; PG, Pettersson grab; SC, spade core; BM, boomerang core. Refer figure for sector reference. DTA, differential thermal analysis; IONF, Indian Ocean Nodule Field; REE, rare earth elements; XRD, X-ray diffraction.

123

124

Mukhopadhyay, Ghosh and Iyer

Sumatra (Indonesia) at 74 ka, which spread over a distance ranging between 1500 and 3500 km (Pattan et al., 2002). Aeolian contributions from the Australian, African and Asian landmasses are also considerable.

1.2. Source 1.2.1. Mineralogical indicators The distribution of various clay minerals in the sediments of the CIOB can be used to determine the source of the sediments. The clay minerals along with lithogenous and hydrogenous components form the dominating part of the four types of sediments available in the basin. Detailed study of clay minerals of different grain sizes in the surface as well as subsurface layers of sediment from 140 locations in the CIOB was made. The major clay minerals in the nodule field are smectite, illite, chlorite, kaolinite and mixed-layer derivatives. Other minerals are feldspar, pyroxene, quartz (lithogenous), zeolite, ferromanganese oxides and hydroxides (hydrogenous or authigenic). The hydroxides provide much of the amorphous iron-oxide minerals, which impart red to brown colour to the sediment. Grain size distribution and clay mineral analyses were carried out on 21 box core sediments collected broadly from two latitudes—15 cores from around 10 S in sector A and 6 cores from around 12 S in sector C (Tables 4.3 and 4.4). The lengths of the cores were at least 35 cm. Lithological and particle size investigations show that sediments are predominantly clayey silt, with mean grain size (M2F) of 7.0–8.6 for surface and 6.6–8.6 for subsurface sediment. The eastward increase of smectite and kaolinite and the southward decrease of illite have been noted (Valsangkar and Ambre, 2000). Variations in distribution and mineralogy of clay minerals in different size fractions of the IONF sediments are interesting (Table 4.4). The conventional treatment of the separation of clay minerals through settling velocity, followed by differential thermal analysis (DTA) suggests that the concentration of illite, kaolinite and chlorite in <2 mm fractions is higher in the northern part of the nodule field (sector A), while the concentration decreases in the southern part (sector D). The abundance of these three minerals in the terrigenous and siliceous sediments indicates influence of continental influx. The concentration of these minerals decreases towards the pelagic clay-dominated sedimentary regime in the south with a corresponding increase in smectite (derived from local young eruptions) and montmorillonite [resulting from weathering of ancient ridge volcanics; Banerjee (1998)]. The higher concentration of montmorillonite in the pelagic clay-rich domain in the southern areas of the IONF appears to have been caused by submarine weathering of basic volcanic rocks (halmyrolysis). Iron-rich montmorillonite can also be found in <1 and 1–2 mm fractions of the siliceous, and 1 mm fraction of the pelagic clays. It is possible that Fe-rich montmorillonite might have been formed by the interaction between iron hydroxides and biogenic silica (opal-CT) during early diagenesis in siliceous sediments having high contents of biogenic silica. In calcareous sediments, there is, however, an absence of iron-rich montmorillonite and predominance of diagenetically mobilised minerals following higher biogenic sedimentation. While montmorillonite has been abundant in <1 and 1–2 mm

Table 4.3

Grain size distribution in the IONF sediments

Sector and location ( S/ E)

Water depth (m)

Sediment layer (cm)

Graphic mean

Standard deviation

Graphic skewness

Graphic kurtosis

Sector B 10 000 /75 150 Sector B 10 000 /75 450 Sector B 10 000 /76 000 Sector C 12 000 /76 150 Sector C 12 22.50 /75 450

5150–5340

Surface 05–10 depth Surface 05–10 depth Surface 05–10 depth Surface 05–10 depth Surface 05–10 depth

7.61 7.62 7.70 7.84 7.59 7.53 7.29 7.13 7.72 7.25

2.43 2.45 2.58 2.37 2.55 2.80 2.52 2.94 2.41 2.27

–0.24 –0.22 –0.68 –0.17 –0.06 0.11 –0.06 0.24 –0.14 –0.15

0.71 0.80 0.80 0.81 0.39 0.79 0.64 0.77 0.59 0.66

5200–5400 5180–5400 5200–5380 5120–5320

Source: Valsangkar and Ambre (2000). Note: Sector B values are averages of five sediment cores from each location, while sector C values are averages of three sediment cores from each location. IONF: Indian Ocean Nodule Field.

Table 4.4 Distribution of clay minerals (%) in the IONF sediments Sector and location ( S/ E)

Sediment layer depth (cm)

Smectite average (range)

Illite average (range)

Kaolinite average (range)

Chlorite average (range)

Sector B 10 000 /75 150 Sector B 10 000 /75 450 Sector B 10 000 /76 000 Sector C 12 000 /76 150 Sector C 12 22.50 /75 450

Surface 05–10 depth Surface 05–10 depth Surface 05–10 depth Surface 05–10 depth Surface 05–10 depth

34.9 (32.4–37.4) 40.5 (–) 34.5 (21.1–53.0) 33.1 (23.1–43.1) 47.0 (36.5–55.2) 40.5 (23.4–57.8) 13.9 (5.66–23.7) 12.0 (10.2–15.2) 16.9 (13.8–21.4) 9.16 (5.17–13.3)

35.6 (35.2–36.0) 27.97 (–) 30.4 (21.4–40.3) 36.0 (27.3–53.3) 25.9 (14.5–35.2) 27.8 (17.1–36.8) 41.3 (37.7–49.2) 41.7 (35.7–49.2) 32.3 (25.0–37.5) 42.7 (34.5–53.3)

09.4 (6.10–12.6) 17.16 (–) 13.5 (8.55–17.8) 13.5 (11.7–15.4) 12.7 (6.30–20.6) 13.2 (9.26–20.6) 13.7 (7.89–23.6) 12.5 (9.42–14.8) 17.0 (15.5–17.8) 19.9 (12.5–25.3)

20.1 (18.9–21.4) 14.30 (–) 21.5 (17.1–30.0) 18.1 (10.3–28.1) 17.1 (10.6–23.1) 18.5 (9.26–29.6) 31.1 (28.8–33.0) 33.7 (29.5–40.2) 33.7 (29.3–35.7) 28.2 (20.8–38.4)

Source: Valsangkar and Ambre (2000). Note: Sector B values are averages of five sediment cores from each location, while sector C values are averages of three sediment cores from each location. IONF: Indian Ocean Nodule Field.

Sediments

127

fractions, suggesting their crystallisation and segregation preference in these sizes, the percentage of illite, on the other hand, increases with increasing size. Illite shows low crystallinity in finer sediments caused by early diagenesis (Rao and Nath, 1988). 1.2.2. Geochemical indicators During the last few decades, valuable contributions have been made to improve our understanding of how the Indian Ocean works as a chemical system. The R-mode factor analysis of the geochemical data of sediment samples showed five important sources of supply for the various elements in the IONF sediment—detrital (loaded with Fe, Ti, Al, P and K), combined hydrogenous and diagenetic (Mn, Ni, Cu and Co), biogenic (Si), sea salt (Na and Mg) and dissolution residue (Ba). Geochemical and isotopic investigations [high Th/Ta ratio (12.8–21.1) and strontium-neodymium (Sr-Nd) data] suggest Himalayan rocks and materials of the Indo-Gangetic plains as the chief source area for CIOB sediments (Fagel et al., 1997). The major elements in the sediment samples from 140 locations covering the four sediment domains of the CIOB (including the IONF)—terrigenous, pelagic, calcareous and siliceous—were determined (Table 4.5; Fig. 4.3). The concentration of major elements, inter-element correlation, their possible source and mechanism of precipitation are briefly discussed here. Organic carbon concentration in the sediment is also given in Fig. 4.3. Aluminium, believed to be an index element for the land-derived alumino-silicates (terrigenous source) and also contributed by weathering of basalts and hydrothermal derivatives, varies from 8 to 13% in the IONF sediments. It is higher in sediments from the northern sectors (sectors A and B, 11–13%) than in the southern sector sediments (sectors C and D, 8–10%). This indicates a decrease in terrigenous input towards the south. Al shows good correlation with Ti (r ¼ 0.50), supporting its detrital origin (Mudholkar et al., 1993). Bulk and excess element concentration in the IONF sediments are shown in Table 4.6. Silica, which is normally transported into the marine environment from landderived alumino-silicates and from skeletons of diatoms, radiolarians, sponges and silico-flagellates, varies between 48 and 68%. Pattan et al. (1992) reported that
50% of silica in cores 226 (sector C) and 657 (sector D) is of biogenic origin, attributable to the higher surface productivity and better preservation. The higher content of SiO2 at 8–10-, 18–20- and 26–28-cm depths in core 226 is due to the probable presence of pumice at these levels. Silica, in general, does not show any correlation with Al and Ti, indicating dominance of a biogenic source. The average SiO2/ Al2O3 ratio and excess SiO2 values, which help in understanding the origin of the elements (Bischoff et al., 1979), are highest (6.7 and 27.97, respectively) in siliceous sediment, followed by intermixed siliceous and terrigenous areas (between 7 S and 8 S, 4.54 and 14.04) and pelagic and siliceous areas (between 15 S and 16 S, 4.25 and 12.27) areas. The major contributor for excess SiO2 has been biogenic silica, which was forthcoming due to the proximity of the northern IONF to the subequatorial region of high biological productivity (Nath et al., 1989). Titanium is an extremely immobile element in the marine environment, and shows input largely from the continents and sometimes from the weathering of oceanic basalts. The TiO2 content in the CIOB sediment varies from 0.28 to 0.44%

128 Table 4.5

Mukhopadhyay, Ghosh and Iyer

Concentration of major, trace and excess elements in the IONF sediments Calcareous ooze Element

Excess

Siliceous ooze Element

Major (%) and excess element (ppm) Si 12.75 – 27.4 Al 2.55 38 5.44 Fe 1.60 35 3.24 Mn 0.26 92 0.64 Ca 25.50 99 0.69 Na 0.96 – 1.67 K 0.46 29 1.62 P 0.07 77 0.10 Ti 0.12 – 0.23 Trace (ppm) and excess element (ppm) Ba 1494 87 3056 Co 24 85 64 Cu 141 92 305 Ni 110 88 220 Pb 20 82 44 Sr 958 95 175 Zn 83 78 160

Pelagic clay

Terrigenous clay

Excess

Element

Excess

Element

– 33 39 93 50 – 40 78 –

20.09 6.60 4.10 1.41 1.41 1.82 1.88 0.17 0.27

– 31 44 97 64 – 30 76 –

– 7.6 5.6 0.56 1.9 – – 0.08 0.47

90 87 93 89 89 50 79

2483 163 556 600 109 138 226

86 93 97 97 97 20 78

1662 46 235 – – 226 –

Sources: Nath et al. (1989), Borole (1993 a,b), Banakar and Jauhari (1994), Pattan et al. (1994), Banakar et al. (1998) and Jauhari and Pattan (2000). Note: IONF, Indian Ocean Nodule Field.

and decreases from the northern to the southern locations, indicating again a reduction in terrigenous input towards the south. It does not show much variation with depth except for some high values at intermediate depth, suggesting increased supply of terrigenous material. Ti generally shows a linear relationship with Al (r ¼ 0.91), Fe (r ¼ 0.72) and Mg (r ¼ 0.69), indicating its homogeneous distribution in the clay-hydrolysate fraction. Iron is normally found fractionated between pelagic clay, hydrogenous metals and metalliferous components. The average Fe2O3/Al2O3 ratio varies between 0.40 and 0.44, which is much lower than those in the average pelagic clays (0.58). A lack of excess Fe2O3 and a good correlation of Fe with Al (r ¼ 0.88) and Ti (r ¼ 0.72) suggest that iron is derived either from a hydrogenous source or from detrital clays. There are, however, large amounts of Fe-rich montmorillonite in the sediment (Rao and Nath, 1988). Potassium in the CIOB sediment (average content 1.88%, Mudholkar et al., 1993) is preferentially absorbed into clays. The presence of large quantities of illite and phillipsite and other products of diagenetic origin suggests that these components act as favourable sinks for potassium (Shankar et al., 1987). The abundance of illite is greater in siliceous-terrigenous area (north of sector A, 46–51%), compared to its characteristic lower abundance in siliceous clay (sectors B and C, 35–40%) and pelagic–siliceous clay (south of sector D, 33–36%). The average K2O content for

129

Sediments

2.0

20

Laxs (ppm)

Fexs (%)

2.5

1.5 2.0 0.5

15 10 5

1.5

Pxs (%)

Alxs (%)

2.0

2.0 0.5

Baxs (ppm)

Opal (%)

0.6 0.2

35 25 15 5

3500 2500 1500 500 80

Coxs (ppm)

1400

Srxs (ppm)

1

1000 600 200

60 40 20

400

Cuxs (ppm)

60 40 20

300 200 100

20

0.9

15

0.7

Mnxs (%)

C.F. (%)

CaCo3 (%)

80

10 5

0.5 0.3 0.1

4N 2

0

2

4

6

8

10 12 14S

4N 2

0

Latitude

2

4

6

8

10 12 14S

Latitude

Figure 4.3 Latitudinal variations in excess elements, opal, coarse fraction (CF) and carbonate in the Indian Ocean Nodule Field (IONF) sediments (Banakar et al.,1998). Table 4.6 Concentration of bulk and excess elements in the IONF sediments

Bulk Excess Excess: bulk

Al

Fe

Mn

P

Ba

La

Ce

Yb

5.9 1.7 29

3.6 1.6 44

0.73 0.68 93

0.13 0.09 69

2763 2521 91

29 10 34

80 32 40

3.5 1.9 54

Source: Banakar et al. (1998). Note: Al to P in %, Ba to Yb in ppm. IONF, Indian Ocean Nodule Field.

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Mukhopadhyay, Ghosh and Iyer

these three domains is 2.49, 2.16 and 2.16%, respectively, which roughly follows the trends of illite distribution. The excellent positive relationship between K and Al in most of the cases supports a homogeneous nature of the lithogenous material. The average sodium content of the sediments derived from continental weathering of rocks is moderate (Na2O, 1.97%). The presence of biogenically derived calcium in these sediments depends mainly on the position of the calcite compensation depth (CCD, also known as carbonate compensation depth, or carbonate line). Other sources of calcium include clay minerals and feldspars (Rao and Nath, 1988). Magnesium occupies the lattice of montmorillonite. The average MgO content in the IONF sediments is 2.26% showing a good correlation with Ti (r ¼ 0.69) and Fe (r ¼ 0.57). The MgO content decreases from north (2.36%) to south (2.05%) of the IONF, with a positive correlation with Al2O3 (r ¼ 0.73). Phosphate concentration increases along the sediment depths, as the mobilisation of iron oxyhydroxides (forming smectite) does facilitate an upwards diffusive flux of phosphate into the overlying water. A close relationship is found between the dissolved phosphate concentration and the organic carbon content along the sediment column (Nath and Mudholkar, 1989). In general, the presence of higher Al excess in both siliceous (where carbonate is absent) and calcareous (where opal is absent) sediments in the IONF and its strong positive correlation with biogenic matter (both opal and carbonate) suggest association of excess aluminium (Alex) with surface water productivity (Fig. 4.4). The average opal content in carbonate ooze and detrital clay to the north of sector A is about 5%, whereas in the siliceous ooze area (sectors B, C and part of sector D) it is
26%. Again, Al/Ti values in the subsurface sediment reach 48.5, an unusual increase, three times higher than average shale and potential crustal sources. The importance of a biogenic contribution for enhancing the Al/Ti ratio appears possible, which is also supported by sediment trap data showing that trap materials comprise 90–95% biogenic components. The upper sediment traps have average concentrations of 0.12% Al and 341 ppm Ti (Al/Ti ratio ¼ 3.6), while bottom sediment traps have average concentrations of 0.4% Al and 619 ppm Ti (Al/Ti ¼ 6.5, Pattan et al. (1992)). It is also seen that the highest Al/Ti and Alex values have a bearing on the sediments enriched in volcanic glass. These volcanic glass shards in ash layers were recovered from eight sediment cores at and around 14 S in the CIOB. The shards are fresh, colourless, with no signs of alteration, and constitute
60–70% of the coarse fraction. They are rich in SiO2 (76.8%) and total alkalis (8.4%), suggesting rhyolitic composition, with very low MnO concentration (0.05%). Electron microprobe analyses of 60 of glass shards show average concentrations of 6.74% Al and 0.04% Ti. The Al:Ti ratio of glass varies from 125 to 188 with a mean value of 175. The composition is consistent with that of the glass shards produced by the youngest Toba eruption in northern Sumatra (age 74 ka), suggesting such eruptions as a possible source of CIOB glass shards (see also Chapter 3). The volcanic ash (tephra) built aprons of several tens of metres, supporting bihemispheric dispersal of the ash cloud and dispersal of gas and aerosols into the CIOB (Pattan et al., 1999). Additionally, occurrence of metalliferous sediments and high concentration of volcanic spherules associated with sediments of 10 ka in the CIOB are suggested to have

131

Sediments

A

B

1.0 0.8 0.6

1.0

0.2

0.5 0.0 0.0

0.5

C 4000

1.0 1.5 Alxs

2.0

0

2.5

10

20 Opal

30

40

20 Opal

30

40

20

30

40

D 4000

n = 19 g = 0.77

n = 19 g = 0.77

3000 Baxs

3000 Baxs

1.5

0.4

0.0

2000 1000

2000 1000

0.0 E 4000

0.2

0.4 0.6 Mnxs

0.8

0

1.0 F 2.5

n = 18 g = 0.61

Alxs

2000 1000

10

n = 18 g = 0.88

2.0

3000 Baxs

n = 19 g = 0.77

2.0 Alxs

Mnxs

2.5

n = 19 g = 0.59

1.5 1.0 0.5 0.0

0

10

20 OpalCFB

30

40

0

10

OpalCFB

Figure 4.4 Inter-element relationship among opal and excesses of Al (wt%), Ba (ppm) and Mn (wt%). Note possible burial of Ba through Mn phase following a direct relation (see C), and a strong direct relation between Alexcess and productivity (biogenic silica) (see F). CFB ¼ Carbonate free basis (Banakar et al.,1998).

derived from mid-plate volcanism and hydrothermal activities at a local scale (Iyer et al., 1997a,b). The other material falling (and sinking) in the CIOB is pumice—a product of volcanic eruptions (local as well as distal). Being highly porous, it can float and move with ocean currents for long distances, even transporting attached organisms. In addition, the aerosol particles brought by wind are the other contributors.

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The aerosols flying into the northern portion of the Indian Ocean (size 0.1 mm to <0.001 mm) are delivered from central Asia, the Arabian Peninsula and the Indian Subcontinent and are characterised by high concentrations of biogenic components (Corg up to 20%, in certain cases up to 40%; Lisitzin, 1996). Among the extraterrestrial materials are tektites (size 1–1.5 mm), microtektites (size 200–900 mm) and minitektites (size 1–1.7 mm), which are found in the CIOB in various shapes—club, dumbbell, spherical, elongated, oval and so on. (Fig. 4.5; Table 4.7). The source materials of CIOB sediments are suggested to be mostly derived from the Himalayas, with minor contributions of illite-rich sediment from southern continental India, slumped sediment from the Indonesian arc and kaolinite-rich wind-blown particles from Australia (Wijayananda and Cronan, 1994). Thus, two major fractions (lithogenous detritus and calcium carbonate) and two minor fractions (authigenic Fe-Mn hydroxide colloids and biogenic detritus) constitute the sediment in the CIOB (Table 4.8). The detrital material derived from the Himalayas contributed to the bulk of sedimentation during
400 to
1100 ka (Pattan et al., 2005). Banakar and Jauhari (1994) divided the CIOB into six tentative geochemical zones from north to south (Fig. 4.6): (1) carbonate-dominated low Mn, low Fe sediment close to the equator; (2) carbonate-free high Al, high Ba siliceous sediment; (3) siliceous sediment with low Al, low Fe; (4) carbonate ooze with low Al, Ti and Fe (west of 74 E); (5) carbonate-dominated red clay with high Mn; and (6) terrigenous sediment with high Al, Ti and Fe to the north of the IONF. In addition, the interelement correlation (Figs. 4.3 and 4.4) suggests three strong positive mutual associations in the CIOB sediments, all having r > 0.7. These are (1) Ba and Mn, Ca and Sr; (2) Al, Mg, Fe, Ti and Sc and (3) Mn, Cu, Co and Ba. These associations may suggest well-defined carriers of the metals to the sediments (Banakar et al., 1998). The uranium and thorium contents in the IONF sediments are given in Table 4.8. The high Th/Ta ratio (12.8–21.1) in the CIOB sediments is indicative of orogenic calc-alkaline source rocks of Himalayan origin. The Sr-Nd isotope data are scattered between two-end members—the Himalayas and the seawater (Fagel et al., 1997; Nath, 2001). The Nd-isotope studies suggest supply of sediments from two major sources in the CIOB—the Indian shield (low E0 values between –14 and –17) and the Indonesian arc (E0 > 0; Fig. 4.7). From the barium distribution in this sediment (high barium suggesting higher productivity), Schmitz (1987) assumes that the sediment would record higher Ba concentration when passing through the high-productivity region and therefore will give an idea of plate movement. Accordingly, positions of the Indian Subcontinent at different geological times have been deduced from seafloor magnetic and biogenic barium data (Fig. 4.8). This deduction may also suggest that the Indian– Eurasian collision occurred during Late Eocene to Oligocene and uplift of the Himalayas reflected by a high-spreading rate, occurred during the Late Miocene. The source of volcanic ash layers in the CIOB sediments remains questionable. These ash layers have been reported from various depths in the sediment cores and may have different sources (Mascarenhas–Pereira et al., 2006). For example, the ash layers might be the product of secondary eruptions and have formed locally in situ in the basin, or may have been erupted through Indonesian arc volcanism (IVA) and later transported to the CIOB by seawater and air. There are many evidences in favour of in situ formation of ash layers in this magmatically active basin:

133

Sediments

A

c

b

a

d

e

g

f

B

a

b

c

d

h

Figure 4.5 Tektites in the Indian Ocean Nodule Field (IONF) sediment: (A) various shapes and dimensions of mini- and microtektites. (B) Different types of crater (pit, sandblasted) formed on the surface of tektites (Prasad et al., 2003).

Table 4.7

Composition of different tektites, and distribution of microtektites in the IONF

A. Compositional range (wt%)a

SiO2 Al2O3 FeO MgO CaO K2O Na2O TiO2

Microtektite

Tektite Australasian

IONF

CIOB

Minitektite CIOB

62.2–79.7 8.90–17.7 3.57–8.63 1.31–7.95 1.37–9.77 0.90–2.81 0.62–3.91 0.49–1.00

56.26–72.52 12.22–19.01 4.39–9.16 2.38–10.62 2.48–4.14 0.38–2.02 0.43–1.04 0.71–1.03

49.6–77.0 7.50–22.1 3.00–8.10 1.90–17.1 1.00–5.80 0.10–3.70 0.20–2.80 0.50–1.0

66.2–72.4 13.0–18.1 5.00–6.20 2.80–5.40 2.60–3.50 0.60–2.20 0.20–0.60 0.70–0.90

B. Distribution of microtektitesb Microtektite abundance

a

Sediment core

125^250 mm

>250 mm

Number of microtektites with impacts

AAS 4/1 AAS 4/2 AAS 4/5A AAS 4/6

659 379 662 217

445 277 401 364

73 19 28 30

Cassidy et al. (1969), Prasad and Sudhakar (1999) and Prasad et al. (2003). Prasad and Khedekar (2003). IONF: Indian Ocean Nodule Field. b

Size (mm) of impacted microtektite Range

Average

300–1400 280–1725 370–2000 295–2400

558 666 821 810

135

Sediments

Table 4.8 Uranium and thorium concentrations in representative the IONF sediments Sectors A þ B Sample 232 230 interval Th Th (cm) (dpm/g) (dpm/g)

238

232

230

238

232

00–10 10–20 20–30 30–40 40–50 50–60 60–70 70–80 80–90

1.28 1.67 1.79 0.75 0.64 0.78 0.78 0.78 0.77

2.81 3.06 1.98 2.99 2.56 2.35 2.42 2.43 2.61

47.39 12.31 1.79 1.30 1.85 0.93 0.93 0.81 0.73

0.88 1.07 1.10 1.51 – – 1.08 – 0.81

3.88 4.35 3.60 3.28 – – 3.20 – 1.80

1.97 1.45 1.33 – – – – – –

28.16 16.38 9.16 – – – – – –

Sector C

Sector D

230 U Th Th U Th Th (dpm/g) (dpm/g) (dpm/g) (dpm/g) (dpm/g) (dpm/g)

1.37 1.32 1.08 1.10 – – 0.97 – 0.63

Sources: Borole (1993a) and Banakar et al. (1998). Note: One core from each sector (A þ B, C and D)—NR-1, F88B and SK176, respectively—is furnished here. IONF: Indian Ocean Nodule Field.

(1) the association of ash layers with 10-ka-old radiolarian assemblage (Gupta, 1988), (2) occurrence of a large field of ‘in situ formed’ pumice associated with ash layers (Iyer and Sudhakar, 1993a,b), (3) confirmation of several episodes of local eruptions of fractionated mid-ocean ridge basalt (MORB) melt (see secondary eruptions in Chapter 2; Mukhopadhyay et al., 1995, 1997), (4) discovery of 10-ka-old high concentrations of volcanogenic hydrothermal material including spherules and rhyolitic glass shards (Iyer et al., 1997a,b) and (5) geochemical signatures and analyses of shards suggest an intra-plate volcanic event (Mascarenhas–Pereira et al., 2006). However, the fibrous and cuspate shape morphology and geochemistry [Ti/Al ratio, enrichment of Zr and light rare earth element (LREE), chondrite normalised pattern] of some of these tephra layers suggest their formation from Sumatra/Java IVA (Martin-Barajas and Lallier-Verges, 1993) with some from the Youngest Toba Tuff (YTT at 74 ka; Pattan et al., 2002). Also the ‘fresh’ pumice which was associated with these ash layers in the IONF is suggested to be chemically more akin to the IVA (Mudholkar and Fujii, 1995). Nevertheless, the enigma that an episode having an age of 74 ka is associated with 10-ka-old ash layers or with 200-ka-old sediments keeps the ‘IVA-transport theory’ under question (Nath, 2001). On the other hand, it is more likely that both the ‘in situ eruptions’ as well as the ‘transported eruptions from IVA’ contributed to the addition of secondary volcanics in the CIOB.

2. Sedimentary Processes The inter- and intra-oceanic variables that generally control the sedimentation pattern in deep-sea environments are: (1) sediment type and grain size; (2) eustatic and local sea-level changes; (3) tectonics; (4) rates of sediment supply and accumulation; (5) geometry and size of receiving basin; and (6) ocean current circulation patterns,

136

Mukhopadhyay, Ghosh and Iyer

A

B 80E

20N

100

80E

100

Fe (%)

Al (%)

20N

0

0

20S

20S > 6.5 6.5 - 5.5 5.5 - 4.5 4.5 - 3.5

> 5.5 5.5 - 4.5 4.5 - 3.5 3.5 - 2.5

3.5 - 2.5 2.5 - 1.5 1.5 - 0.5 < 0.5

C

2.5 - 1.5 1.5 - 0.5 < 0.5

D

20N

Smectite (%)

>70 50 - 70 30 - 50 10 - 30 <10

MYANMAR

INDIA

2

1

20N

3 4

4 Sri Lanka

3

0

0

00

40

3

00

40

5

3 5

20S

20S 80E

100

80E

100

Figure 4.6 Geochemical zones in the Indian Ocean Nodule Field (IONF) sediments showing distribution of (A) Al, (B) Fe and (C) smectite. The overall sediment distribution process (Wijayananda and Cronan, 1994) is sketched in (D): 1 ¼ Mg, Al, Ti, V, Cr, Fe and smectite-rich subcontinental sediments, 2 ¼ Himalayan range sediments, 3 ¼ sediments from subduction zone, 4 ¼ terrigenous sediments, 5 ¼ Mn, Co, Cu and Ni-rich hydrogenous sediments (after Nath, 2001).

governed in part by the Coriolis effect (cf. Blatt et al., 1980). These six variables are further governed by even larger-scale processes such as the relative rates of generation and destruction of oceanic crust, the disposition of the continents, global climatic changes and possibly Milankovitch cyclicity. The factors that may control the sedimentation in the CIOB are discussed later.

137

Sediments

30N INDIA 20 −14

AFRICA

−14.7 Old continental area particles

−7.9

−7.2

−7.2 −1.9 0 −2.6 −7.3 −2.8 −3.1 −7.0 10

20 −7.7

−5.8 −5.9

−10.8

40 00

30

40

+2.9 −4.9

Antarctic waters

20

10

−7.2

−12.2

−7.4 +0.06

−2 −3

−7.8 −7.7

−8.4

−12 Atlantic waters

10E

−17.4

−7.2

−9

30

40

50

−6.8

50

60S

−18.9

60

70

80

90

100 110E

Figure 4.7 Inferred sources of sediments in the Indian Ocean Nodule Field (IONF) based on Neodymium (Nd) isotopes (expressed as e0 units; Dia et al.,1992).

2.1. Dissolution of carbonate The primary factors influencing the nature of deep-sea sediments have been productivity and preservation, that is, the production and supply of planktonic organisms, along with the depth of carbonate dissolution. Planktonic organisms, in fact, are the major supplier of materials for the formation of deep-sea biogenous sediments (siliceous and calcareous clays). About half of the world’s deep-ocean floor is covered by biogenous ooze composed of coccoliths (5–30 mm), foraminifer tests (50–500 mm), diatom frustules (5–50 mm) and radiolarian frustules (40–150 mm). Some contribution also comes from seawater by chemical processes, in particular, by the precipitation of dissolved material. These sediments are called chemogeneoushydrogenous (Lisitzin, 1996), are subject to favourable physico-chemical conditions and facilitate the formation of ferromanganese oxide deposits in the world oceans including the IONF. From the surface, calcareous skeletal parts descend rapidly down the water column and are ultimately deposited as carbonate sediment (calcareous ooze). However, in deeper areas in excess of 4000 m, in general, the seawater becomes sufficiently acidic (with the increase in concentration of CO2) and becomes undersaturated with respect to CaCO3. As a result, calcareous material begins to dissolve.

138

Mukhopadhyay, Ghosh and Iyer

62 Ma

53 Ma

10N

IN IN

0

DI

17 Ma

6 Ma

217 DI

A

217

217 IN

DI

6 Ma

A

216

217

A

216

216 215

10

216

217

217

20

215

216

215 214

216 30

40 Ma

215 214

214

213

215

213

214

213

214 213 213

215 214

40S

9.9 cm/yr

5.6 cm/yr

2.3 cm/yr

1.4 cm/yr

11.3 cm/yr

213 −217 = DSDP cores

Figure 4.8 Paleo-position of the Indian Subcontinent deduced from biogenic barium and seafloor magnetic data (Schmitz, 1987). Sediment used for this experiment came from 5 DSDP cores 213 to 217.

Dissolution of shells is enhanced by high hydrostatic pressure and low water temperature. The depth at which calcium carbonate shows an accelerated dissolution is called the lysocline and the depth at which the proportion of carbonate falls below 20% is known as CCD. The siliceous ooze are found commonly in deep abyssal plains at water depths of about 5000 m. This indicates that the distribution and composition of deep-sea sediments are controlled by the productivity and preservation of planktonic organisms. The production and supply of such organisms are guaranteed close to the zone of high biological productivity (best in equatorial region), while preservation of skeletons is influenced by the CCD. As the solubility of skeletal remains is deep and temperature dependent, the CCD is depressed close to the equator, along with the elevated mid-oceanic ridge regimes and other shallow areas in the ocean, where biological productivity is generally greater than that in the open ocean. The average level of CCD is an indicator of the rate of removal of atmospheric CO2 to the deep sea, which has been variable over the geological past (Fig. 4.9). There are two principal ways by which CO2 gets into the ocean: (1) directly, through solution of CO2 from the atmosphere; and (2) indirectly, through transportation of weathered land products by rainwater as carbonic acid. If the first process is dominant, the ocean becomes more acidic and both lysocline and the CCD rise; if the indirect process of CO2 intake is more active, the opposite happens. Fluctuation in CCD through geological time and its effect on sedimentation pattern and distribution in the Indian Ocean are very pronounced, and this helps determine paleoclimatic conditions fairly accurately (Pickering et al., 1989).

139

Sediments

Paleo - CCD (Peterson and Backman, 1990) ~ 26S, 58E

Paleo - depth (km)

3

Paleo - depth curve 4

~22S 62E

5

~16S 75E

~12.5S 78E

~20S 69E

~18S 71E

Oxyhydroxide drought

2

1

6 0 PS

10 PL LM

20 MM

EM

30 LO EO LE

40

50 ME

EE

60 Ma P

Epoch

Figure 4.9 Position of the ancient calcite compensation depth (CCD) in the Indian Ocean vis-a-vis paleo-depth of the crust (Banakar and Hein, 2000). P ¼ Palaeocene, EE ¼ Early Eocene, ME ¼ Middle Eocene, LE ¼ Late Eocene, EO ¼ Early Oligocene, LO ¼ Late Oligocene, EM ¼ Early Miocene, MM ¼ Middle Miocene, LM ¼ Late Miocene, PL ¼ Pliocene, PS ¼ Pleistocene, 1 ¼ seamounts formation began, 2 ¼ crustal accretion commenced.

Despite the fluctuations in the global sea level since Cambrian (Fig. 4.10), the global average of CCD (Fig. 4.10) appears to be deeper now than at any time in the past 250 million years. In the Atlantic, the CCD has lowered from
3500 m during Jurassic– Cretaceous time. In the Indian Ocean, the change has also been stark, mean CCD deepened from about 3900 m at 135 Ma to 4000 m at 35 Ma and further to 5100 m today (Kolla and Kidd, 1982). A much lower CCD level (
5700 m) is recorded from a 3-m calcareous clay sediment core recovered from a location adjacent to 73 E fracture zone (Vishnu FZ, Mukhopadhyay et al., 1994). The Pacific Ocean, however, does not show much variation, with the CCD lowered by only 200 m since Jurassic period. Paleo-levels of CCD could be ascertained from long sediment cores and from the age– depth relationship of oceanic crust (as oceanic crust is believed to be subsiding with time because of cooling and contraction, The Open University, 1995).

2.2. Bottom water mass and sedimentation The sedimentary process in the CIOB is highlighted by interaction of various alien water masses influencing the precipitation, erosion and non-deposition of sediments over the geological past. For example, the mixing of two bottom water masses, namely, nutrient-rich Antarctic bottom water (AABW) and the nutrient-poor North Atlantic deep water (NADW), appears to have influenced circulation and distribution of nutrients, both in vertical and in horizontal directions (Fig. 4.11). There is evidence that as Antarctica moved to its present polar position during the Eocene, there was a general deterioration of climate, ultimately resulting in glaciation near the end of the Eocene or beginning of the Oligocene. This resulted in the production of copious amounts of cold, aggressive bottom water that must

140

Mukhopadhyay, Ghosh and Iyer

0

2

4

6

570 Ma Cambrian 500 Ordovician 430 Silurian 395 Devonian 345 Carboniferous 280 Permian 225 Triassic 190 Jurassic 135 Cretaceous Paleoc. and Eocene Oligocene Miocene

0

2

4

6

65 38 25 5 Ma

Plio - pleistocene

100's of m above present sea - level

Figure 4.10 Estimated global eustatic sea level since 570 Ma (Hallam,1984).

have moved northward into the Indian Ocean through numerous fractures in the Southeast Indian Ridge (Kennet et al., 1975). This bottom water slowly grew into a vigorous circulating force, the AABW and consequently inhibited sedimentation and caused erosion at many places. The strength of the circulating water probably decreased by the Oligocene, drawing an end to the hiatus and permitting sedimentation to resume in most places. The AABW, which entered the CIOB through the saddles along the Ninetyeast Ridge at 5 S, sank into the basin and made a westward movement, possibly playing a double role—acting as an erosive agent to remove the top thin veneer of sediment and enriching the sediment and the nodules with additional metal input (Johnson and Nigrini, 1982; Warren, 1982). The NADW, on the other hand, entered through the major fracture zones across the central Indian Ridge on the west leaving a trail of influence on the water circulation pattern in the CIOB. The turbulence in the deep-water movement

141

Sediments

created due to its low stability in this basin (Levitus, 1982) might be responsible for the intense erosion of younger sediments. To understand the nature of sedimentation in the CIOB over the last 300 ka, the isotope decay rates were measured using @ 13C and 230Th. Variations in 13C data (Geochemical Ocean Sections Study, GEOSECS project) in benthic foraminifera from different oceans suggest that the southern ocean is an ideal region to monitor fluctuations in the global influence of the NADW. The nutrient properties of circumpolar deep water reflect the mixing of deeper water masses of the world. The bottom water mass causes sedimentation as well as erosion. To quantify the erosional capability of the bottom water mass, the distribution of 230Th and the index radiolarian species in the sediment cores in the IONF were studied. For this, three sediment cores, NR-1 (sector A, 9.99 S, 77.92 E), NR-21 (sector B, 11.00 S, 78.49 E) and NR-35 (sector C, 11.97 S, 78.49 E) were retrieved using a spade corer (dimensions 20 cm  30 cm  45 cm). The lengths of the recovered sediment columns were 28 cm (NR-1), 32 cm (NR-21) and 23.5 cm (NR-35) from water depths of 5250, 5325 and 5450 m, respectively. The bottom sediment at these localities is siliceous ooze, with <3% carbonate (Nath et al., 1989). Some chemistry of sediments down these three cores and that from a nearby core is given in Fig. 4.12. The exponential decay of 230Th is clearly evident in cores NR-1 and NR-35, which help in determining their average accumulation rates (NR-21 shows A

Warm surface current

Figure 4.11

Deep cold current

(Continued)

142

Mukhopadhyay, Ghosh and Iyer

Present day

B

Productivity

Depth (km)

0

728 3

752

OMZ

721

754

757 722 731

6 Late Miocene−Early Pliocene Productivity

Depth (km)

0

721 728 3

752

OMZ

754

757

722 731

6 0

4000 Distance (km)

8000

Figure 4.11 (A) Distribution and exchange of warm surface and colder deep water along the conveyor belt in the world’s oceans (Einsele, 2000). (B) Fluctuation in oxygen minimum zone (OMZ): present day and at Late Miocene (Dickens and Owen,1994).

a non-uniform distribution). Again, the ratio of the 230Th flux in the sediments (Fa) to its production rate (Fp) in the overlying water column provides an evidence for sediment erosion on the seafloor (Mangini et al., 1982). The production rate (Fp) of 230Th and the flux of the 230Th to the sediments (Fa) is 2.4 dpm/cm2 ka for core NR-1 and 3.1 dpm/cm2 ka for core NR-35. The respective Fa/Fp ratios (0.18 and 0.22) are far less than that expected under ideal conditions (Fa/Fp ¼ 1) of nonerosion and non-deposition. These extremely low Fa/Fp ratios indicate removal of the younger sediments by some dynamic agent. The expected surface activities of 230Th under ideal conditions of non-erosion and non-deposition are 210 and 161 dpm/g for NR-1 and NR-35, respectively, assuming a constant flux of 230Th. By extrapolating the best-fit decay curves for these values, the total thickness of the sediment column removed should be around 30 and 37 cm, respectively, for the NR-1 (sector A) and NR-35 (sector C) cores. With an average accumulation rate of 2 mm/ka, the effective chronological record thus eroded can be estimated to be around 175 ka (Banakar et al., 1991). Additional evidences of erosion were obtained by studying the distribution of index Neogene radiolarian zone in the NR-1 core. Biostratigraphically important radiolarian species such as Collosphaera invaginata (first appearance datum, FAD, 0.15–0.2 Ma), Collosphaera tuberosa (FAD, 0.4–0.5 Ma), Collosphaera orthoconus

143

Sediments

(FAD, 0.65 Ma), Stylatractus universus (last appearance datum, LAD, 0.42 Ma) and Lamprocystis nigriniae (FAD, 0.9 Ma) were encountered in subsamples of NR-1 core. C. invaginata, the index species of the very first and topmost zone of the Neogene radiolarian biostratigraphy, is absent not only in the upper subsamples but also in the entire length of core NR-1 (Gupta, 1988, 2000; Johnson et al., 1989). C. orthoconus and C. tuberosa normally co-occur in the 0–25-cm depth interval. However, the disappearance of the top zone (Neogene radiolarian zone 2), characterised by C. tuberosa until the 25-cm depth and the absence of S. universus (LAD, 0.42 Ma) throughout core NR-1, indicate that the top sediment layers were eroded. The sediment in the core spans a period between
200 and
400 ka. The age derived for the top 25-cm depth in this core with an accumulation rate of 1.6 mm/ka is around 155 ka. Considering the
175 ka erosion of the sedimentary record based on the Fa/Fp ratio, the time span thus derived for this core, using the radiochemical method, is between
175 and
350 ka, which agrees well with the biostratigraphic age derivation (Banakar et al., 1991).

0 10

Al/Ti (g/g) 10 30

0

Alex (%) 20 40

Biogenic opal (%) Terrg. matter (wt%) 20 30 40 20 25 30

Volcanic glass Abundant

Pass

A

20

Rare

30

Absent

40 NR -1

50 10

30

50 20

40

60

20

30

40

20

30

40

10 Pass

20

Rare

Abundant

30 NR -21 40 0

10

20

30

20

40

60

20

30

40

20

30

40 Rare

10 20 30

Pass

Core depth (cm)

0

Abundant NR -35

Figure 4.12

(Continued )

144

Mukhopadhyay, Ghosh and Iyer

B

20

Fe/Ti

16 12 8 4 0

Mn/Ti

3 2 1

Ce-anomaly

0 0.10 0.05 0.00

−0.05 −0.10 0

100

200

300

400

500

Core depth (cm)

Figure 4.12 (A) Distribution and concentration of terrigenous material Al, biogenic opal and volcanic glass from coarse fraction in three sediment cores NR-1, NR-21 and NR-35 [PAAS has Al/Ti ratio as 16.7] (Pattan and Shane., 1999). (B) Distribution of cerium (Ce) anomalies and Mn/ Ti and Fe/Ti ratio in the Indian Ocean Nodule Field (IONF) sediment (Pattan et al., 2005).

Mottled zones were encountered between the depths of 6 and 14 cm in the NR-1 and NR-35 cores. Nearly horizontal 230Th decay patterns in these mottled zones can be attributed to sediment mixing due to bioturbation. Similar mottled bioturbated zones between the depths of 10 and 20 cm have been noticed in core NR-21 also (Nath and Mudholkar, 1989). A sharp decrease in nitrate values and high levels of organic carbon in these not-so-infrequent mottled zones in the IONF suggest anoxic conditions. Thus, it can be said that the AABW-induced sediment erosion as recorded in NR-1 and NR-35 and intense bioturbation as shown by all three cores might have provided progressively favourable conditions for maintaining the ferromanganese nodules in the IONF at the sediment–water interface. However, it is difficult to ascertain whether there was a single erosional event lasting nearly 175 ka, or the erosion was accomplished in several intermittent short-lived events.

145

Sediments

The sedimentation rate (in cores NR-1, NR-35, SK226 and SS657) decreases from north to south with corresponding increase in the accumulation of biogenic silica (Table 4.9). The increase in biogenic silica and its accumulation rate towards the south reflect decrease in terrigenous input and increase in siliceous microfossil concentration. For example, intensified monsoon in Asia and the Indian Subcontinent between 125 and 75 ka (Prell and Kutzbach, 1987) appears to have been responsible for discharge of considerable amounts of sediment by rivers like the Ganges and the Bramhaputra resulting in least concentration of biogenic silica to the north of the IONF (Pattan et al., 1992). The effect of bottom water mass on depositional condition can also be shown by the distribution of radiolarians. As the majority of radiolarians inhabit in the upper 100 m of the oceanic water column, their surface distribution in the CIOB indicates three well-defined water masses in the basin (Gupta, 1996). The water masses are (1) high salinity (>34.5/psu), cooler (<27 C) highly productive (>0.2 mgC/m3/h) water mass in the south-western part during the southwest monsoon, represented by high Pyloniids assemblage; (2) comparatively low saline, warmer, low-productive water mass in the north-western part during the southwest monsoon, characterised by high Euchitoniids assemblage; and (3) a transitional one in the remaining areas. Moreover, down-core fluctuation of the ratio of Pyloniids to Spongodiscids could be a function of monsoon. Because of the CIOB’s proximity to the equator, these oceanographic variations can be strongly related to long-range transfer function of monsoon intensities (Prell and Kutzbach, 1987). Sedimentation in this basin is further influenced by a pronounced front of the bottom water at 10 S latitude, with the characteristic hydro-chemical structure separating the reversing monsoon and the subtropical gyres. Seasonal reversal of winds and currents during the monsoon is also noteworthy due to its influence on the distribution of nutrients. Spectacular monsoon reversal in the northwest Indian Ocean involves northeastward flow of the Somali Current between April and September, which supports strong upwelling and high biological productivity.

Table 4.9 Concentration and accumulation rate of biogenic silica, rate of sedimentation and sediment density in different sectors of the IONF Biogenic silica

Sector

A B C D

Water depth (m)

Core length (cm)

Concentration (%)

Accumulation rate (g/cm2/ year)

Sedimentation rate (mm/ka)

Density (g/cm3)

5250 5325 5450 5270 5050

41 35 28 32 36

24.71 33.54 30.41 26.76 22.59

1.14  105 – 1.69  105 4.05  105 1.45  105

1.6 4.6 2.2 2.0 2.0

0.287 0.287 0.254 0.757 0.321

Sources: Banakar et al. (1991), Pattan et al. (1992) and references therein. IONF: Indian Ocean Nodule Field.

146

Mukhopadhyay, Ghosh and Iyer

In contrast, during October–March, water current flows towards the southwest, without any upwelling activities. The record of biogenous sedimentation suggests that the monsoon has been active since the Mid Miocene (10–12 Ma). Upwelling and consequently monsoon was stronger during glacial and weaker during interglacial periods, such as at present. Also, the monsoon and related atmospheric circulation in the CIOB depend on the uplift of the Himalayas (more rapid during the last 3 Ma), as this huge mountain chain acts as a barrier between high atmospheric pressure during winter and low atmospheric pressure during summer in the Indian Subcontinent ( Johnson and Nigrini, 1982; Kolla and Kidd, 1982).

2.3. Depositional environment The distribution of rare earth elements (REE) in sediment reflects the environment of deposition, and varies as a function of lithology, sediment type and bottom water condition (Fig. 4.13). As an example, the REE concentration in a sediment core from the IONF is furnished in Table 4.10. The total REE concentration in the carbonate sediment in the western part of the IONF is minimum (average 80 ppm), increases to a maximum in detrital clay in the northern part (average 183 ppm) and shows moderate concentration in siliceous sediment (average 169 ppm). The REE were carried to the IONF sediments in two phases—Ti phase (detrital indicator) and Mn phase (oxide indicator). The REE concentration pattern in the Ti and Mn phases has been contrasting—in the detrital phase REE concentration decreases from La towards Lu, while in the oxide phase concentration records a reverse trend (Nath, 2001; Pattan and Banakar, 1997). Distribution of REE, and in particular cerium (Ce) anomalies, in marine sediments are considered as indicators of depositional environment, such as, the limit of the spreading of marine anoxia (Liu et al., 1988), spread of the AABW (Glasby et al., 1987) and variations in surface productivity. In the IONF, a southward evolution of a Ce anomaly in sediment from negative to positive is evident. There is also a consistent hump across the middle rare earth elements (MREE—Sm, Eu, Gd and Tb) and a noticeable Eu anomaly (
1.2). The observed Eu anomaly suggests considerable aeolian input of detrital component to the deep-sea sediment of the IONF (Banakar et al., 1998). In the calcareous sediment, the Laxs, Cexs and Ybxs proportions are 35, 13 and 44% of their respective bulk component and in siliceous sediment Laxs remains unchanged, while Ybxs and redox-sensitive Cexs increase to 54 and 40%, respectively. A positive association among opal, Mnxs and Ybxs might suggest a major role of Mn-oxide coating on biogenic-settling particles in fractionating heavy REE (HREE) and Ce. The detrital clay zone has no excess La, Ce or Yb, as expected. Laxs does not yield any systematic latitudinal trend. In siliceous sediment, the Ce anomaly becomes prominent (up to 1.5) and the enrichment of HREE over LREE (Lan/Ybn ¼ 0.6) becomes more significant than in the calcareous zone (0.8). LREE is more prone to the adsorptive removal from the water and should lead to LREE-enriched patterns in the sediments (Elderfield, 1988). On the contrary, an HREE-enriched distribution might indicate a mixed mechanism. The regeneration associated with the release of carbonyl ions to the deep water might provide large scope for the deep-water complexation of

147

Sediments

2

Terrigenous sediments (type-I)

1

2533 2532, 2535 2501 2483, 2513 2537

0.4 2

Siliceous clays (nod. free reg.) (type-II)

1

101, 150 F156, F157 153

0.5 2

Siliceous clays (overlain by nods) (type-III)

1

NR120D NR124, 2528 NR129

0.4

Sample/shale

1

Calcareous sed. (type IV)

CaC

O : 3 47 %

CaC

O : 3 62%

CaCO

3

0.1

: 87%

S121 S124 2494 S127

0.05 6

Pelagic red/clays (type V) 210 164, 144 132 148 160 136 150, 231

6 0.6 1

MORB

0.07 La Ce Nd Sm Eu Gd Dy

Yb Lu

reg = region, nod = nodules, sed = sediment, MORB = Mid Ocean Ridge Basalt

Figure 4.13 Shale-normalised characteristic pattern of rare earth elements (REE) in various sediment types in the Indian Ocean Nodule Field (IONF). Flat pattern holds no signature of fractionation between light rare earth element (LREE) and heavy rare earth elements (HREE). Absence of Ce anomalies may indicate a mixed continental source (Nath, 2001).

regenerated REE (Murray et al., 1991; Piper, 1974). The cumulative effect of this, in addition to precipitation of colloidal ferromanganese oxide particles, might have left HREE-enriched patterns in the sediments along with the more pronounced Ce anomaly. Remineralisation of oxide coating at depth has been understood to balance

Table 4.10

REE concentration (ppm) in a sediment core from the IONF

Core depth (cm)

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Total REE

000–004 020–022 032–034 048–050 060–062 076–078 096–098 108–110 124–126 136–138 142–144 148–150 165–170 180–185 195–200 210–215 230–235 255–260 285–290 320–325 335–340 350–355 365–370 400–405 425–430 450–455 475–480 490–495

31.75 26.19 29.81 31.11 35.91 38.86 37.40 37.52 36.61 44.42 44.64 40.20 38.62 47.49 44.33 42.15 46.40 43.51 38.10 46.95 44.28 43.47 44.01 49.94 53.82 63.71 58.02 60.33

82.45 72.50 75.70 81.53 93.97 93.64 93.24 88.63 86.15 102.37 103.41 92.62 86.24 103.98 102.55 97.99 108.45 108.08 95.69 101.77 101.18 103.98 105.39 110.66 108.65 113.46 112.65 112.47

8.07 6.64 7.38 7.76 9.23 9.97 9.91 9.88 9.42 11.05 11.20 10.21 9.82 11.89 10.98 10.63 11.66 11.11 9.58 12.26 11.09 10.93 11.16 12.79 13.36 15.6 13.93 14.54

35.44 28.86 32.23 34.47 41.55 44.39 44.50 43.55 42.32 49.09 49.90 45.87 44.17 53.11 49.06 48.14 51.29 49.11 41.79 55.30 49.87 48.16 49.98 57.97 60.20 69.24 62.92 64.54

9.28 7.69 8.39 9.47 11.28 12.33 11.92 11.61 11.79 13.11 13.36 12.34 11.60 14.27 13.58 12.89 13.16 12.61 11.02 13.99 13.51 12.64 13.02 15.83 15.94 17.79 16.12 16.67

2.38 1.96 2.05 2.14 2.80 3.06 3.16 2.99 2.81 3.16 3.30 3.15 2.86 3.50 3.31 3.15 3.35 3.21 2.77 3.37 2.85 2.99 2.97 3.32 3.47 3.99 3.63 3.53

8.49 6.77 7.36 8.19 9.65 10.45 10.46 10.21 9.86 11.68 11.56 10.79 10.36 12.53 11.52 11.11 11.90 11.11 9.63 12.19 11.54 10.90 11.32 13.02 13.43 15.54 14.25 13.92

1.37 1.13 1.24 1.32 1.67 1.76 1.82 1.74 1.68 1.94 1.93 1.76 1.75 2.09 1.90 1.86 1.93 1.82 1.58 2.12 1.88 1.83 1.91 2.26 2.35 2.72 2.39 2.49

8.19 6.87 7.27 8.47 9.93 10.77 10.72 9.97 10.58 11.7 11.82 11.17 10.95 13.01 12.09 11.62 11.54 10.88 9.57 12.02 11.73 10.94 11.77 14.12 14.55 16.51 14.82 15.23

1.74 1.36 1.55 1.68 1.96 2.13 2.09 2.09 2.09 2.41 2.48 2.17 2.16 2.60 2.40 2.33 2.37 2.22 1.98 2.53 2.38 2.19 2.29 2.75 2.99 3.52 3.07 3.11

4.30 3.59 3.91 4.40 4.97 5.29 5.18 5.08 5.10 6.21 6.08 5.86 5.62 6.73 6.20 6.06 6.10 5.55 4.88 6.45 6.21 5.57 5.77 7.07 7.73 9.00 7.87 8.01

0.57 0.45 0.54 0.59 0.63 0.67 0.72 0.66 0.67 0.82 0.81 0.71 0.71 0.86 0.78 0.77 0.81 0.72 0.62 0.87 0.78 0.70 0.75 0.90 1.01 1.20 1.02 1.03

3.76 3.14 3.56 3.79 4.29 4.62 4.54 4.44 4.58 5.60 5.42 5.04 4.98 5.72 550 5.20 5.32 4.76 4.36 5.55 5.42 5.06 4.96 6.18 6.62 7.93 7.00 6.72

0.54 0.44 0.52 0.54 0.62 0.66 0.66 0.67 0.63 0.81 0.79 0.71 0.69 0.82 0.76 0.71 0.77 0.71 0.60 0.81 0.72 0.67 0.68 0.83 0.95 1.16 0.99 1.00

198 167 182 196 229 239 236 229 224 264 267 243 251 279 265 255 275 265 232 276 266 263 266 297 305 341 319 324

Source: Pattan et al. (2005). Note: IONF, Indian Ocean Nodule Field; REE, rare earth elements.

Sediments

149

the preferential removal of LREE in the upper water column (Sholkovitz et al., 1994). The very strong association of Ce with Mn in both bulk and excess element correlation matrices and the southward positive evolution of the Ce anomaly are suggestive of burial of Ce via the Mn-oxide phase (Banakar et al., 1998). REE concentration, which is generally higher in the upper oxic zones, suggests its upward diffusion through the sediment column and subsequent incorporation in the oxyhydroxide phase (Pattan and Banakar, 1997). REE fractionations are found to be characteristic for each sediment type—for example, flat shale-normalised patterns for terrigenous sediment, positive Ce anomaly in siliceous sediment, negative Ce anomaly in calcareous sediments and LREE-depleted patterns in pelagic red clay. Based on this, Nath et al. (1992) suggested that REE signatures and fractionations in the IONF are indicative of depositional setting, lithological variations and surficial diagenetic processes. The concentration and accumulation rates of 10Be in different sediments of the IONF also suggest extensive removal of the isotope along the continental margins towards regulating its distribution in deep sea sediments (Nath et al., 2007).

2.4. Diagenesis The sediments in the IONF are overlain in most areas by ferromanganese nodules and crusts. Although the seawater largely contributes metals for the formation of these manganese oxides through chemical precipitation (hydrogenous method), a significant contribution of metals comes through diagenetic remobilisation. Pore water plays an important role in such diagenetic reactions. Consequently, detailed pore-water studies were made to determine nutrients, estimate diffusive fluxes and geochemical balances, understand the nature and sequence of oxidation and oxidants and the role of these in diagenesis of the organic matter. Some of the organic matters, CaCO3 and opaline silica, which form in the surface water due to biological productivity, settle to the bottom where they undergo decomposition and dissolution. Organic matter decomposition follows a sequence of processes depending upon the availability of the oxidant, involving successive utilisation of O2, NO3, MnO2, Fe(OH)3 and SO4 (Berner, 1982). The initial increase in nitrate at the surface sediment reflects oxygen respiration coupled with nitrification. At least two sediment cores (both from sector C) show mottling/bioturbation in the intermediate layer between 10 and 25 cm. This intermediate layer records a sharp fall in the nitrate values and along with increased levels of organic carbon in the solid phase indicates the effect of bioturbation and consumption of oxygen. This suggests that oxic and anoxic processes are active at different intervals. The subsurface layer between 10 and 18 cm in the studied cores shows oxic diagenesis, while the layers beyond these depths show the existence of anoxic conditions and hence favour denitrification. Thus, it appears that the process is being controlled by the reaction rates and bioturbation of sediments (Nath and Mudholkar, 1989). The relation between early diagenetic process and the nutrient levels in the pore water are examined from three sediment cores spread over different sedimentary regimes (Fig. 4.14). Interesting results with no diffusion gradient in nitrate

150

Mukhopadhyay, Ghosh and Iyer

Nitrate 20 25 30 35 70E

0

Organic carbon 2 4 6 8

80 INDIA Chennai

N

SRI LANKA

Nitrate 20 30 40 50

247

0

Organic carbon 2 4 6 8

0

Organic carbon 2 4 6 8

m

CIOB 246

5000

226

0

Nitrate 10 20 30 40

Core 226

10S

C EN TRAL I NDIAN

RID G

E

0

E C HA G OS LA CC ADI VE RID G

185

10N

Core 247

200 km

Core 246

0

Goa

Figure 4.14 Pore-water chemistry: concentration of nitrate and organic carbon in three sediment cores (Nath, 2001). Location of samples: 247 ¼ IAPB region, 246 ¼ sector A and 226 ¼ sector C.

concentration are found in terrigenous sediment regime (sample 247 in India-Australia Plate Boundary (IAPB) zone), whereas sediments from the southern siliceous area (sample 246, sector A) and pelagic clay area (sample 226, sector C) do show such gradient in diffusion of nitrate. This probably indicates a strong relationship between

Sediments

151

nature of pore-water chemistry and pulses of turbiditic sedimentation (Nath, 2001). The fact that early diagenetic processes control the pore-water nutrient levels has also been demonstrated in a study of another two cores from the IONF. The cores—one each from sectors B (NR-21) and C (NR-35)—show nitrification in the lithologically more uniform upper layers. At intermediate depths, the bioturbated zones contain organic carbon, which was probably not fully decomposed and showed signs of denitrification. Core NR-21 showed reworking and thorough mixing of the sediment throughout its length, as is evident from the nearly horizontal decay profile and the highly variable distribution of 230Th (Table 4.8). Core NR-35, on the other hand, provides typical transition metal profiles. The minimum solid-phase Mn is encountered at intermediate depths without any spikes in the transition metals and an increase of Mn, Ni and Cu towards the surface indicates anoxic diagenesis (Dymond et al., 1984), which facilitates the diagenetic growth of nodules. In sector A also, the oxidation of remobilised Mn2þ appears to have taken place down core (NR-1). The formation of micronodules below the surface of the sediments and the down-core oxidation of remobilised Mn2þ probably act favourably for transition metal movement to the surface to form ferromanganese nodules (Banakar et al., 1991; Mudholkar et al., 1993). Additional evidence for sediment mixing comes from recrystallised biogenic ˚ in X-ray diffraction (XRD) of the bulk opal, identified by a peak of 3.31–3.34 A sediments. The intensity of the peaks was normalised to one for each core. In core NR-21, the normalised intensity varies within 16%, while in the other two cores it varies within about 25%. Low variation in the intensity of the biogenic opal in core NR-21 (in sector B) might therefore be due to mixing of the older and younger sediments, leading to the nearly uniform distribution of biogenic opal with depth, unlike the random distribution pattern in the other two cores. It appears from the solid-phase Mn distribution in NR-21 that the diagenetic reactions and remobilisation of Mn were established subsequent to the redistribution of the sediment, which otherwise would have resulted in uniform distribution of solid-phase Mn within the core. However, later diagenetic reactions in this core have not influenced the 230Th distribution, which preserved the evidence for mixing and support the validity of the 230Th method of dating.

2.5. Sediment consolidation The physical processes associated with sedimentation and its consolidation is expected to have a profound bearing on the engineering characteristics of the sediment. Any information about those physical processes helps to reveal the degree of consolidation, including nature of compaction and plasticity of sediment. Knowledge of these geotechnical properties is also considered essential in order to understand the behaviour of sediments under dynamic loading, which is required for designing any seafloor mineral mining system. To ascertain the post-precipitation characteristics of sediments in the IONF, several sediment cores were examined (Khadge, 2000, 2002). The cores were recovered from three sectors in the IONF—1 gravity box core (at 9 S, 77 E) and 15 spade cores (between 10 S and 10 100 S, 75 E, and 76 E) from sector A, and 6 spade cores from sector C (between 11 550 S and 12 150 S, 75 450 E, and 76 E).

152 Table 4.11

Mukhopadhyay, Ghosh and Iyer

Physical properties of sediments from the central region of the IONF Maximum

Minimum

Average

Sector B Specific gravity Porosity Wet bulk density Void ratio Water content Shear strength Liquid limit Plastic limit

– (%) (g/cm3) – (%) (kPa) (%) (%)

2.86 94.4 1.21 16.9 515 4.79 280 156

2.15 84.8 1.10 5.6 410 1.63 160 48

2.32 91.1 1.15 10.6 460 3.12 228 108

Sector C Specific gravity Porosity Wet bulk density Water content Shear strength Liquid limit Plastic limit

– (%) (g/cm3) (%) (kPa) (%) (%)

2.60 94 1.27 553 10.1 280 166

1.88 88 1.12 380 1.6 206 103

2.16 90.5 1.13 441 3.77 233 127

Source: Khadge (2000, 2002). Note: IONF, Indian Ocean Nodule Field.

Incidentally, all these cores are bottomed by siliceous clay/ooze, at depths well below the CCD. Table 4.11 shows variation in physical properties of seafloor sediments. The sediments at the northern part of sector A contains abundant radiolarian tests and four clay minerals—montmorillonite (in predominate quantity), illite, kaolinite and chlorite. Water content, which varies with the textural change and clay mineralogy, decreases with depth (Fig. 4.15). To a depth of 470 cm, water content is consistent up to a value of 370% and then a decrease occurs down core. This water content is calculated on dry-weight basis that corresponds to about 81% on wet-weight basis. Overburden pressure of the sediments might have squeezed the water out. In contrast, sediments in the southern parts of sector A contain either smectite or illite as the dominant mineral phase. The water content calculated in dry-weight basis ranges from 369 to 577% in the north and from 380 to 553% in the south. The high water content is probably due to the abundant presence of smectite (including montmorillonite). Both these minerals have an expandable lattice structure and consequently have the highest water-holding capacity. This capacity was further augmented by the presence of about 60% microfossil/radiolarian tests, which are hollow and could therefore be filled with water. The Attenberg limits show consistency in depth-wise variation with water content. The liquid limit (LL) and plastic limit (PL) of the sediment, which reflect the capacity to withstand load, show higher values in the southern parts of sectors A and C (LL ¼ 223–239% and PL ¼ 105–136%) than in the northern areas of

PL, LL, WC (%) 0

100 200 300 400

Sp. Gr.

Wet Porosity (%) Density (gm/cc)

2.0 2.5 70

80

90

1.1 1.3 1.5

Shear strength (kPa) SD/ST/CL (%) 2

6

10

14

0

0

50 100

CL

SD

*

M/I/K + C (%)

50 100

3.1 3.9 4.7

PL

M

LL

Illite

K+C

WC Silt

Depth (m)

2.3

5.5 6.3 7.1

PL = Plastic Limit, LL = Liquid Limit, WC = Water Content, M = Montmorillonite, I = Illites, K = Kaolinite, C = Chlorite, SD = Sand, ST = Silt, CL = Clay, * = Disturbed sediment, not studied

Figure 4.15 Variation in physical properties down the sediment column. Note drastic change at 470-cm depth corresponding to Pliocene^ Pleistocene boundary (Khadge,1998).

154

Mukhopadhyay, Ghosh and Iyer

sector A (average LL ¼ 20%, average PL ¼ 109%; Table 4.11). The shear strength of the IONF sediment shows a depth-wise increase from the surface to the bottom and from the north to the south (2.2–13 kPa at
9 S, 1.6–7.2 kPa at sector A, and 1.6–5.4 kPa at sector C). The surface and bottom sediments are highly bioturbated and show medium to high plasticity. The value of shear strength of the IONF sediments is, however, lower than those reported from the nodule-rich areas of the Pacific Ocean (10–19 kPa; Hirst and Richards, 1975). The sediment column in the IONF shows an unconformity at a depth of about 470 cm (Fig. 4.15). The sediment at this depth is characterised by high-specific gravity, high density and low porosity compared wih lower values for sediments above and below this level. For example, the sediments of the sectors A and C are in general characterised by these average values: specific gravity 2.18, wet bulk density 1.14 g/cm3 and porosity 90.2%; these abruptly changes to 2.5, 1.54 g/cm3 and 71%, respectively, at 470-cm depth. Such an abrupt change may have been caused by variation in mineralogy (such as increase in montmorillonite from 40 to 67% at this depth) and increased degree of bioturbation in the sediment (Khadge, 2000, 2002). Again, the shear strength data, which reflect on the bearing capacity and settlement characteristics of sediment, spiked to 12.8 from 10 kPa on the top and lower layers. The unconformity at 470 cm, described earlier, spans about 16 cm in thickness and can be calibrated to the depositional history at the Pliocene–Pleistocene boundary. The clay mineralogy and physical properties show drastic changes at this layer, indicating the change in sedimentation characteristics. The drastic increase in montmorillonite content, spike in shear strength values and the presence of pumice at this depth imply change in mineral composition caused by increased input from ridge-weathered rock-sediment during the Early Pliocene. The total absence of radiolarian fossils at this depth also indicates either climatic change or volcanic activity that precluded their preservation. In summary, the IONF encompasses three major sediment types—siliceous (
80% area, having positive Ce anomaly), pelagic/red clay (
10%, LREE depleted) and terrigenous (
10%, flat shale-normalised REE pattern). Mineralogically, while illite, chlorite and kaolinite show increase in abundance from south to north (sectors D to A), smectite and montmorillonite show an opposite trend. About half of the silica is contributed to the CIOB biogenically. Secondary volcanic material may have been sourced from both local basin eruptions and IVA. The process of sedimentation in the IONF essentially depends on dissolution of carbonate, circulation of bottom water mass and depositional mechanism (diagenesis and consolidation).