Verdine and other associated authigenic (glaucony, phosphate) facies from the surficial sediments of the southwestern continental margin of India

Marine Geology, 111 (1993) 133-158

133

Elsevier Science Publishers B.V., Amsterdam

Verdine and other associated authigenic (glaucony, phosphate) facies from the surficial sediments of the southwestern continental margin of India V. P u r n a c h a n d r a Rao a, M. L a m b o y b and P.A. Dupeuble b aNational Institute of Oceanography, Dona Paula-403 004, Goa, India bDepartment of Geology, University of Rouen, BP 118, 76134 Mont Saint Aignan, France (Received June 15, 1992; revision accepted December 3, 1992)

ABSTRACT Purnachandra Rao, V., Lamboy, M. and Dupeuble, P. A., 1993. Verdine and other associated authigenic (glaucony, phosphate) facies from the surficial sediments of the southwestern continental margin of India. Mar. Geol., 111: 133-158. Green grains, pale green-brown infillings of foraminifer tests and brown friable aggregates occur in the coarse fraction of the surface sediments from the southwestern continental margin of India, between the Periyar river in the north and Quilon in the south. Study of the nature, distribution, mineralogy and internal structure of these particles resulted in (1) the discovery of a verdine facies and (2) determination of the relationships of phosphate with glaucony and verdine in these sediments. Two distinct verdine facies associated zones, a shelf zone at about 40 m and a slope zone between 100 and 280 m water depth, are distinguished. Verdine grains are abundant (up to 30%) where biogenic debris dominates in the sediments and are rare (< 5%) where terrigenous detritals dominate. On the continental shelf the verdine grains are dark green and predominantly occur as angular to subrounded grains, whereas on the slope they are represented by the coexistence of different types of pellets (dark green glossy and green to light green rugose pellets) and infillings. Scanning electron microscopy (SEM) studies indicate that the shelf grains are mostly homogeneous and the slope grains contain several detrital components. The most common nannostructure of the authigenic clays is the association of closely spaced contorted clay blades and globules. X-ray mineralogy suggests that these grains are a mixture of verdine dominated minerals. Phyllite C is the principal verdine mineral in the shelf zone. On the continental slope phyllite V dominates between 100 and 205 m water depth followed by phyllite C at about 280 m. Glaucony and phosphate predominantly occur as pale green to brown infillings of foraminifers and as phosphatic friable aggregates in the sediments of the terrace at 330 m water depth. Glauconitic smectite and carbonate fluorapatite are the respective mineral phases. The terrace sediments are sandy and often associated with non-glauconitized and non-phosphatized mollusc shells. Phosphatized verdine grains occur at 170 m water depth. Scanning electron microscopy studies indicate that phosphate is in the form of globular to rod-shaped bacteria-like structures or microbial filaments, or both. This type of phosphate occurs within the foraminifer tests and as cement in the friable aggregates. Glauconitized foraminifers are enclosed in phosphatic matrix and phosphate occurs within the glaucony matrix. The intergrowth of green clays and phosphatic globules is observed locally in phosphatized glaucony aggregates and also in phosphatized verdine grains. The formation of verdine, glaucony and phosphate in these sediments is relict. The age of their formation is difficult to estimate precisely, but appears to be older than mollusc shells dated at 5710 yrs B.P. Planktonic foraminifers from the terrace indicate a Quaternary age. Verdine formation on the shelf and on the slope is diachronous. The shelf verdine facies most probably formed during the Late Pleistocene high sea level stands or during the Early Holocene. The slope verdine may have formed contemporaneously with glaucony and phosphate on the terrace during low sea level stands. The terrace most probably acted as a sort of hardground facilitating glauconitization and phosphatization of the sediments. Upwelling, oxygen minimum zone and low terrigenous sediment accumulation provided adequate conditions for phosphatization which was mediated by microbial processes. The intergrowth of phosphate with the authigenic clays of verdine and glaucony may be due to the local confinement of the microenvironment, which was probably controlled by microorganisms.

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134

Introduction Authigenic marine clays occurring on the ocean floor are many and diverse. They are useful indicators of the sedimentary environment at the time of their formation. Among them, there are varieties of granular green clays which are similar in colour and morphology but which differ in their mineralogical composition. It is necessary to identify the specific green clay minerals as they reflect distinct environments of sedimentation (Odin, 1988). One such newly distinguished authigenic green clay facies is the verdine facies (Odin, 1985). One mineral in this facies has recently been designated as odinite (Bailey, 1988). It is a ferric magnesian dioctahedral-trioctahedral 1 : 1 clay mineral which was initially named 7 A phyllite V (Odin, 1985). Another authigenic clay coexisting with odinite is phyllite C, which has an intermediate structure between chlorite and smectite (Odin, Bailey et al., in Odin, 1988). It was shown that the verdine facies forms at less than 60 m water depth and its presence is now known from 13 locations in the present oceans (Odin, Debenay et al., in Odin, 1988; Odin, 1990). As this facies commonly grades into deeper water glaucony facies, studies of these facies are useful in interpreting the palaeogeography of continental margins. This is the main purpose of the first part of this paper. Glaucony minerals and phosphate minerals are important authigenic marine minerals formed in areas of low sedimentation under suboxic to anoxic conditions and mostly at shallow depths ( < 1000 m) on the sea floor. Glauconitization generally precedes phosphatization; however, synchronous formation and even phosphatization preceding glauconitization have been reported (Parker, 1975; Birch, 1980; Burnett, 1980; Marshall and Cook, 1980; Glenn, 1990; O'Brien et al., 1990). The formation of glauconite and carbonate fluorapatite (CFA) in sediments rich in organic matter and within the oxygen minimum zone occurs as follows (Glenn, 1990; O'Brien et al., 1990). Glauconite forms early in the sediments close to the sediment-water interface (suboxic environment) fed by the partial reduction of inherited ferric iron.

v PURNACHANDRA RAO

The same reduction process enables the release of phosphorus from the surface of ferric oxyhydroxides and supplies an important source of dissolved phosphorus to the pore waters, in addition to the phosphorus released from the decay of organic matter; this phosphorus subsequently accumulates as CFA. Does 'glauconite' grow contemporaneously with CFA? Several workers have used the term 'glauconite' in a broader manner to represent the glaucony facies rather than the mineral itself. Glauconitic smectite first forms de novo, then recrystallizes to glauconitic mica with an increasing potassium content in the interlayers (Odin and Matter, 1981). Glauconitic mica is the recommended term for glauconite (see Odin and Fullagar, in Odin, 1988). Therefore, the younger/present day forming glaucony-phosphate sediments should contain glauconitic smectite rather than glauconitic mica which needs more time (105-106 years) to form (Giresse et al., 1980; Odin and Letolle, 1980). The results presented here show the contemporaneous formation of glauconitic smectite with CFA. However, as other green clays corresponding to the verdine facies have been reported to occur in continental slope environments (Odin, 1988), not all the authigenic green clays associated with phosphate necessarily correspond to minerals from the glaucony facies. The relationships of phosphate with glaucony and verdine in upper slope and terrace sediments will be discussed in the second part of this paper.

Previous work and physiographic setting Verdine or glaucony minerals have not been reported from the sediments of the western continental margin of India. However, phosphorites consisting of minor amounts of phosphate (< 10% P205) in algal nodules, fossil corals and tubular structures (Baturin, 1982; Nair, 1985; Rao and Nair, 1988; Rao et al., 1990) and a high phosphate content (up to 2.2% P205) in surficial sediments (Rao et al., 1987) have been reported. The study region is located on the southwestern part of the western continental margin of India

135

AUTHIGENIC FACIES OF THE CONTINENTAL MARGIN OF INDIA

% ,d

~oe ~

Cochin

"

LAKE

k..\Oe' •

j

J"

ASHTAMUDI BACK WATERS

"~,

"~, qk

~uilon

k o, Km I

\ 76" r'r" Fig. 1. Sample location map. The samples shown by double circles are where phosphate is found together with glaucony or verdine. The sample locations without numbers (near to the coast) are inner shelf clays where there are no green grains.

(Fig. 1). A tropical climate prevails in this region and the annual rainfall is about 300 mm. The Periyar river drains through Precambrian gneissic rocks. The width of the continental shelf is 45 km. Sedimentary facies on the shelf are parallel to the coast. The inner continental shelf (up to 35 m depth) is smooth and covered by recent muds. The outer shelf has a rugged topography consisting of relict sandy sediments (Nair and Pylee, 1968). The shelf break occurs at about 80-100 m depth. The slope sediments are again clayey except on a 35 k m wide terrace occurring off Quilon at a depth of 330-420 m (Fig. 1) where they are sandy.

Part I. Verdine and glaucony facies Results

Composition of the sedimentary coarse fraction The coarse fraction ( > 63 pm) of the sediment samples across the outer continental shelf and slope was studied. Three size fractions ( > 1 mm, 0.5-1 m m and 0 . 1 2 5 - 0 . 5 m m ) were considered. Bored bivalve shells, gastropods and fragments of bryozoans are present in the > 1 m m fraction, although this fraction is found in few shelf samples.

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V. P U R N A C H A N D R A

The distribution of terrigenous detrital minerals (mostly quartz and feldspar), carbonate detrital minerals and green grains was investigated in the other two fractions. Some continental slope sediments do not contain the >0.5 mm fraction and in some other samples there are no green grains in this fraction. Wherever green grains are present in this fraction ( > 0 . 5 mm), their distribution is similar to that in the 0.125-0.5mm fraction of the sediment. The percentage distribution of these components in the 0.125-0.5 mm fraction is shown LEGEND

• verdine grains

• carbonate particles 100%

[ ] terrigenous particles

Line

a

5o

sample number depth

== I~

1452 100 m

1453 55 m

1454 45 m

RAO

in Fig. 2. Green grains are rare ( < 5 % ) where terrigenous detritals dominate ( > 65%) in the sediments; the percentage of green grains increases when the carbonate detrital grains become more important. Green grains predominantly occur in two zones which are parallel to the coast: one at about 40 m depth and the other extending from 100 to 330 m on the continental slope. The sediments between 40 and 100 m depth are predominantly quartz and feldspar sands. Here, the green facies mostly occurs as fillings in the fractures of quartz and feldspar grains. The highest concentrations of green grains in the two zones correspond to the sediments off Periyar and Vembanad lake (Fig. 2). The type of green grains varies. For example, on the shelf zone they are predominantly angular grains, whereas on the slope different forms, including pellets of different colours and morphology and infillings of foraminifer tests, coexist. On the terrace, they principally occur as infillings in planktonic foraminifers.

100% Line

b

5o

1449 300 m

1451 85 m

1446 55 m

1,443 37 m

t ~9% "le C

1436 170 m

1~5 80 m

1453 50 m

1452 40 m

~

~100%

Line d

300m

1,427 205 m

1425 82 m

1422

0

40 m

~ 330 m

The clay mineralogy of the sediments shown in Fig. 1 has been reported by Nair et al. (1982), Rao et al. (1983) and Rao (1991). Smectite (55%) and kaolinite plus chlorite (39%) are the predominant clay minerals, followed by illite (6%) and trace amounts of gibbsite, in both inner shelf and continental slope clays.

Morphology of the green grains

50

1428

1403

Composition of the sedimentaryfine fraction (clays)

- 100% Line

e

50

-o

1402 330 m

1401 280 m

13~ 90 m

1391 85 m

Fig. 2. Distribution of various components in the coarse fraction (0.125-0.5mm fraction) of the sediments (see lines on Fig. 1). The green grain distribution in line e and at station 1428 (line d) represents both phosphatic and green clay components.

Two types of grains are found in the outer shelf (about 40 m depth: zone 1): angular to subrounded grains and various infilled tests and test fragments, the former being the most predominant type. Benthic foraminifers, gastropod tests and echinoderm test fragments acted as substrates for fillings of green clay and these infillings constitute less than 10% of the coarse fraction. The colour of a few of these substrates varies from red to ochre, indicating post genetic alteration. The grains are shiny and dark green in colour, showing some thin and narrow cracks on their surfaces (Fig. 3A and B). There are a few grains which are ochreous yellow,

AUTHIGENIC FACIES OF THE CONTINENTAL MARGIN OF INDIA

137

Fig. 3. Characteristic green grains. (A) and (B), angular to subrounded grains from the shelf-- (A) photomicrograph, (B) scanning electron micrograph showing narrow surface cracks. (C)-(F) pellets from the slope. (C) Green and light green rugose pellets (photomicrograph); (D) scanning electron micrograph showing rough surface texture of a pellet; (E) glossy dark green pellets (photomicrograph); (F) scanning electron micrograph showing smooth surface texture of a glossy pellet; and (G) green infillings of planktonic foraminifers from the terrace (photomicrograph). representing the subsequent oxidation of green grains. Owing to the smooth surfaces and fragmented nature of some of the grains and the altered nature of other grains, these may be considered as older and possibly reworked. The green grains from the continental slope (100-330 m water depth: zone 2) distinctly differ from the shelf zone. Several types of grains coexist. These are (1) smooth textured dark green glossy pellets (Fig. 3E and F), (2) green and light green pellets with a rough surface texture (Fig. 3C and D) and (3) infillings of planktonic foraminifers (Fig. 3G). Glossy dark green pellets are more numerous at stations opposite the mouths of the Vembanad lake and Periyar river. Green and light green rugose pellets are ellipsoidal and are morpho-

logically identical to green faecal pellets found elsewhere (Odin and Masse, in Odin, 1988). Biogenic debris can be seen on their surfaces (Fig. 3D). These occur down to a depth of 205 m. At the deeper depths ( > 200 m) green grains are mostly infilled planktonic foraminifers with some pellets and grains. On the terrace (330-420 m), infillings predominate, followed by more light green grains in the landward part of the terrace and more phosphatic friable aggregates in the seaward part of the terrace. The colour of the infillings of foraminifer tests also varies from pale green on the landward to brown on the seaward direction of the terrace. There are no dark green grains. Pale green infillings sometimes occur as moulds without surface cracks.

138

Internal structures of the green grains Microstructures Broken surfaces of the green grains were examined by SEM. The angular grains from the shelf are mostly homogeneous at low magnification (Fig. 4A). However, at high magnification the authigenic clays are characterized by small contorted blades (1 Ixm long) at some places and compact clays (Fig. 4B and C) at others. Some pellet-like grains from the shelf, however, contain few heterogeneous components such as silt-sized quartz and carbonate particles. In contrast, the light green pellets from the slope show the most heterogeneous character, containing several types of shell fragments and detrital grains (Fig. 4D and E). Green pellets from the slope are also heterogeneous (Fig. 4F). The internal structure is rather porous and authigenic clay does not complete fill the primary pores. Compact clays contain coccoliths; dissolution affects some of them and small blades of clay begin to grow on their surfaces (Fig. 4G and H). Within the cavities the authigenic clay blades sometimes form small spheres resembling the lepispheres described by Odin and Matter (1981) and aggregate around the cavity and on the surfaces of coccotiths (Fig. 4H). Glossy dark green pellets from the slope show a dual character. At the peripheral parts, they contain homogeneous compact clays like those in shelf angular grains and, towards the cores, they contain heterogeneous components like those in rugose pellets (Fig. 4I and J). Echinoderm test fragments are known to be a good material in which to study the processes of glauconitization and phosphatization (Odin, 1975; Lamboy, 1976, 1987; Odin and Lamboy, in Odin, 1988). Verdinized echinoderm fragments are sometimes abundant in the shelf samples. Broken surfaces of these fragments indicate that authigenic clays fill the primary pores (Fig. 5A). Dissolution affects the calcite of the stereome (Fig. 5B). Green clays partially fill the secondary pores. A compact to subparallelly arranged lamellar clay structure occurs in the primary pores (Fig. 5C), whereas in the secondary pores these clays look like globules (Fig. 5D). Some secondary pores contain pyrite framboids. When compared, these observations on

V.PURNACHANDRARAO verdinization correspond to the second stage of evolution in the glauconitization of the same substrate (Odin, 1975; Lamboy, 1976; Odin and Lamboy, in Odin, 1988). The pale green infillings from the terrace are glauconitic (see under X-ray mineralogy). Their microstructure is also heterogeneous (see Part II) and cannot be differentiated from that of the green pellets of the slope.

Nannostructures of the authigenic clays The most commonly found structure in different green grains from the continental shelf and slope is the association of long contorted clay blades (l~tm long) with ill-defined small globules (0.5-1 ~tm diameter) (Fig. 6A). At high magnification these irregular blades seem to have grown perpendicular to the planes (Fig. 6B). The globules are apparently made up of the filing of clay sheets (Fig. 6C). The porous clay structure shows widely spaced long (1-2 ixm) and rather straight blades (Fig. 6D) and the compact clay structures (Fig. 6E) are also common in some other green grains. The glauconite blades in pale green infillings from the terrace (see Part II) are similar to those reported in various glaucony grains (Odin, 1975; Lamboy, 1976). X-ray diffraction studies Green grains were separated from all the stations using a Frantz isodynamic separator with a longitudinal slope of 25 ° and a lateral slope of 17°, which have been used in previous studies (Odin, Bailey et al., in Odin, 1988). Free green grains and tests infilled with green material were attracted at a 0.6 A current, which is higher than suggested for both verdines and glauconies. Subsequently, pure grains were separated from the magnetic fraction by hand picking. Overall, the separation is good. However, there are a few apparently altered brownish grains, especially in shelf samples; it is impossible to separate them as they were dark green on one side and brown on the other side. Five representative samples at different depths (40,' 100, 205, 280 and 330 m) were selected for detailed X-ray diffraction studies. They were ground to a fine size and randomly oriented powders were

AUTHIGENICFACIESOF THE CONTINENTALMARGINOF INDIA

139

Fig. 4. Representative internal structures of the main kinds of green grains (scanning electron micrographs of the broken surfaces). (A-C) Angular green grain from the shelf; (A) at low magnification showing its frequent homogeneous nature, (B) detail of (A) showing closely spaced small contorted authigenic clay blades, (C) detail of (A) showing clays with compact nature. (D, E) Rugose light green pellet from the slope; (D) at low magnification showing several heterogeneous components, (E) detail of (D) showing different detrital particles embedded in clays. (F-H) Rugose green pellets from the slope; (F) at low magnification showing heterogenous nature, (G) detail of (F) showing compact clays with enclosed coccoliths, (H) another green pellet showing dissolution of skeletal components and the formation of lepispheres of clay blades in cavities and on the surfaces of coccoliths (coccoliths in the process of dissolution). (I-K) Aspects of dark green glossy pellet from the slope; (I) broken surface near to the edge of the grain showing compact homogeneous periphery and heterogeneous grains similar to other pellets towards the center, (J) detail of (I) showing compact clayey nature at the peripheral part, and (K) high magnification of the surface of the glossy pellet showing smooth surface texture.

140

V. PURNACHANDRARAO

Fig. 5. Growth of authigenic clays within the verdinized echinoderm test fragments (scanning electron micrographs). (A) Part of the echinoderm test showing the primary pores filled with clay; (B) partial dissolution of the calcitic stereome, completely infilled primary pores and partly filled secondary pores; and (C) and (D) details of (B) showing compact structure in the primary pores and globular aspect of the clay in the secondary pores, respectively.

Fig. 6. Nannostructures of authigenic clays from green grains of verdine facies (scanning electron micrographs - - all of them apparently coexist in one green grain). (A) Predominantly observed character of green clays in numerous grains showing the association of closely spaced contorted blades [white arrows, (B) high magnification] and clay globules [black arrow, (C) high magnification]; (D) porous structure showing widely spaced contorted clay blades; and (E) compact aspect of the clays.

AUTHIGENICFACIESOF THE CONTINENTALMARGIN OF INDIA

analysed on a Philips X-ray diffractometer using iron-filtered Co K~ radiation scanned at 1o 20/min. and operated at 40 kV and 25 mA.

Shelf green grains These grains (0.25-1 mm size) are from 40 m water depth. The diffractograms of the untreated samples show distinct reflections culminating at 14.3, 7.29, 4.6 and 3.57 A. The peaks at about 7 A and 14 A are broad (Fig. 7Aa) and the peak heights above the baseline are approximately equal. After ethylene glycol treatment, the peak at 14 ]k shifted to about 16 A with no change in 7 A reflection, indicating the presence of swelling type minerals associated with these grains (Fig. 7Ab). The 0.1 N HC1 treatment (at 70°C for l h) was applied (Fig. 7Ac). The intensity of the peak at 7 A was considerably reduced and the peak at 14 A changed into a broad hump. Heat treatment was also applied (Fig. 7B). With increasing temperature the peak at 7 A reduced and finally disappeared after T

i,o

I~,"

1,4 i~ A"

~"

T

~

" 2e

j,4 j,6 A"

tl Normal

390 °

- lh

Fig. 7. X-Ray diffractograms from the dark green angular grains (0.25-1 m m fraction) from the shelf at 40 m water depth (sample No. 1432); EG = after ethylene glycol treatment.

141

heating to 600°C for 4 h. However, the height of the 14 A reflection also apparently initially reduces until 390°C. After heating to 490°C/2 h and to 600°C/4 h (Fig. 7Bcd), the development of a dome at about 14 A and a portion of this dome shifting to low 'd' spacings at about 10 A are seen. Quartz, feldspar and carbonate minerals are seen as impurities in all the X-ray diffraction patterns. The green grains in the finer size fraction (0.125-0.25 mm) of these samples showed a similar X-ray diffraction behaviour during different treatments. Authigenic green clays (verdine and glaucony facies minerals), K-smectite and kaolinite and chlorite all have characteristic X-ray diffraction peaks at 7 and 14 A. As we have analysed magnetic green grains from the marine environment, these reflections may predominantly correspond to authigenic green clays. The equal size at 7 and 14 A has been documented in grains formed with the minerals (phyllite V and phyllite C) from verdine facies. However, phyllite V does not expand on glycolation and, during moderate heat treatment, the 7 A peak decreases and the 14 A peak increases. Moreover, the 14 A peak develops on HC1 treatment for phyllite V. As the peaks in our diffractograms (Fig. 7A and B) do not satisfy these characteristic features suggested for phyllite V (Odin, Baily et al., in Odin, 1988), we suggest that the 14 A peak of the authigenic clay mineral is probably dominated by phyllite C. The initial lowering of the 14A peaks at low temperatures (below 390°C) during heat treatment observed in our X-ray diffractograms (Fig. 7B) supports the presence of phyllite C. The subsequent development of the 14 A peak into a broad dome (at higher temperatures) is a property of both phyllite C and phyllite V. The shift from 14 to 16 A during glycolation is also diagnostic of phyllite C. However, the 7 A peak is supposed to remain during HC1 treatment for phyllite C (Odin, Bailey et al., in Odin, 1988), but this peak disappeared in our X-ray diffractograms (Fig. 7A); this may be due to some admixture of kaolinite. The presence of kaolinite in our X-ray diffractograms is further supported by a large 7 A peak (compare with the diffraction peaks of pure phyllite C of the Senegalese shelf, Odin and Masse, in Odin, 1988, where the 7 A peak is ten times smaller than the 14 A peak). The behavi-

142

V. PURNACHANDRA RAO

our of K-smectite is similar to that of phyllite C on glycolation. As K-smectite and kaolinite are the dominant clay minerals in the surrounding sediments (Rao, 1991), we suggest that the minerals in these grains are a mixture of dominant phyllite C with some other detrital clay minerals. The broad nature of the peaks may be due to alteration and poor evolution.

Slope green grains Pellets from the upper slope: The dark green glossy pellets (0.25-0.5 mm size) from 100 m water depth and green and light green pellets (0.25-0.5 mm size) from 205 m water depth were analysed. The untreated X-ray diffraction patterns of these pellets (Fig. 8A and B) differ from the dark green shelf grains (Fig. 7A) by showing a large peak at 7 A and a small, broad dome centred at about 14 A. 7 t

T

I0 i

,,o

14 i

16 A = =

~ ,~a"

~" ÷" "2o Fig. 8. (A) X-Ray diffractograms of the dark green glossy pellets (0.25-0.5 mm fraction) from the slope at 100 m water depth (sample No. 1452); (B) X-ray diffractogramsof green and light green rugose pellets from the slope at 205 m water depth (sample No. 1427).

Although the colour of the pellets differs, the diffractograms are essentially similar (even their behaviour after different treatments), except that the intensity of the authigenic mineral peaks is slightly lower in green and light green pellets. There is a small peak at 10 A. A slight peak (14 A) shift is observed in glycolated samples (Fig. 8A and B). When the samples were heated to 390°C/2h and 490°C/2h, the 7 A peak was destroyed and a small dome developed at about 14 A. The relative heights of quartz and calcite peaks in green and light pellets are larger than in dark green glossy pellets, confirming the SEM observations. The X-ray diffraction patterns obtained from the 0.125-0.25 mm fraction of the dark green pellets at different treatments are similar to those described from the 0.25-0.5 mm fraction. The large 7 A reflection and the very small 14 A reflection of these clays represent the basic character of young phyllite V (Odin, Bailey et al., in Odin, 1988) which has been named as odinite (Bailey, 1988). The lowering of the 7 A peak during heat treatment also supports phyllite V. The slight expansion of the 14 A reflection during glycolation and the partial collapse of the 14 A peak during heat treatment suggest the possibility of a certain proportion of phyllite C or detrital clay minerals, or both, mixed with odinite.

Infilled tests and pellets from 280m water depth: Green infilled tests and pellets in the size fraction 0.25-0.5 mm were separately purified. To remove the carbonate material, these infillings were treated with acetic acid (1 N) for l h at room temperature before subjected to X-ray analysis. The pellets were used directly without any pretreatment. The diffraction peaks obtained from them are similar except that the 14 A peak is larger in infillings (Fig. 9A and B); this change might be due to pretreatment with acetic acid (Odin, Bailey et al., in Odin, 1988) but may also be due to the different stage of evolution of the authigenic green clay. The diffractograms are characterized by a large 14 A peak and a small 7 A peak. A slight shift is observed in the 14 A reflection and in the 7 A reflection during ethylene glycol treatment. The heat treatment at 250°C/1 h did not modify the 14 A peak from the normal peak for infillings

AUTHIGENICFACIESOFTHECONTINENTALMARGINOFINDIA

T lnfilllngs

tt0k 5

,?

,4 ,pA" ~ .

2~0o:,.,¢?,

2e PIletsel

I0

t4 ~A"

e I~'°

÷° "ze

Fig. 9. (A) X-Ray diffractograms of the green infillings from the slope at 280 m water depth (sample No. 1401); and (B) X-ray diffractograms of the green pellets from the same sample.

(Fig. 9A). However, this peak already collapsed to lower d spacings at this temperature for the pellets (Fig. 9B). The peak at 7 A disappeared and the 14/k peak modified into a broad hump between 10 and 13 A after heating to 490°C/2 h in both separates. The large 14 A peak and the very small 7/k peak (Fig. 7A and B) are the characteristic features of both phyllite C and glauconitic smectite. However, the different treatments used here suggest that the dominant mineral is phyllite C. As in the shelf samples, the characteristics of phyllite C are not well clefined. For example, the 14 A peak did not decrease by a factor of two during moderate

143

heat treatment (Fig. 9A) at 250°C/1 h, which is a characteristic of phyllite C (Odin, Bailey et al., in Odin, 1988). Nevertheless, there is a slight shift in the 7 and 14 A reflections during glycolation and a shift in the 14 A reflection towards high angles at 490°/2 h (Fig. 9A and B) similar to phyllite C. On the other hand, the 14 A peak partially collapsed to lower d spacings even at low temperatures (250°C) in the case of pellets (Fig. 9B); this behaviour is similar to that of smectite (compare Fig. 9B with Fig. 10, glauconitic smectite, at this temperature), indicating more unstability of the authigenic clay in pellets than in infillings. This collapse is not completed at high temperature at 490°C/2 h to confirm glauconitic smectite/smectite. It appears that the authigenic clay mineral tends more towards distinct phyllite C in the infillings and more towards a mixture with smectitic minerals in the pellets. The nature of the substrate (for example, tests are good substrates for growth as well as for the preservation of authigenic clays, whereas alteration may be feasible in pellets) might have contributed to this difference. We therefore suggest that the authigenic clay mineral is phyllite C, with typical intermediate characters between chlorite and smectite.

Infillings of planktonic foraminifers from the terrace (330 m water depth): The X-ray diffractograms of the untreated samples (pale green infillings) do not show peaks at 7 A nor at 14/k; instead, they display a broad hump between 11 and 13 A (Fig. 10) and well developed peaks at 4.5 and 3.3 ~. However, in the acetic acid treated samples, this broad hump shifted to low angles and centred at about 14.6/k. In glycolated samples, this hump marginally shifted to low angles. The 14 A reflection was clearly shifted to high angles when heated to 250°C/1 h and was represented by a large well defined peak at about 10A when heated to 490°C/2 h. The absence of the principle reflections at 7 and 14/k in the normal diffractograms from untreated grains (Fig. 10) indicates that the authigenic mineral is not related to the verdine facies. Acid treated samples show a broad peak at 14.6/k. Although this peak has not completely expanded to 17/k during glycolation, its complete collapse into 10 A

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Fig. 10. X-Raydiffractogramsof the pale green infillingsfrom the terrace at 330 m water depth (sample No. 1403). layers at 490°C/2 h clearly resembles the characters of glauconitic smectite. Odin (1988) reported a similar behaviour for the green grains of the glaucony facies of the Congolese and Ivory coast and suggested that it is difficult to observe a shift after ethylene glycol treatment when there is interlayer potassium. Therefore, the mineral is a glauconitic smectite which represents an early stage of formation of glaucony minerals.

Synthesis about the mineral phases Although we have analysed few representative samples at different depths compared with the large number of samples present in the study area, the following conclusions can be drawn. The shelf grains are dominated by phyllite C, whereas on the slope phyllite V dominates down to a depth of 205 m, followed by phyllite C down to a depth of 280 m, glauconitic smectite is abundant on the terrace at 330m water depth. When compared with the verdine mineral reflections of New Caledonia, the Senegalese shelf and the French Guiana shelf (Odin and Masse, in Odin, 1988), the X-ray reflections in our verdine samples are broad, complex in their behaviour and never represent a single authigenic mineral. The complex nature and crystallization state of the grains may be due to three reasons. (a) The green grains under study certainly

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represent a mixture of dominant authigenic minerals and other detrital clay minerals. The high current (0.6 A) used for separating the grains under the magnetic separator support this fact. The presence of a small 10 A peak in some diffractograms confirms the mixture. It is possible that there could be some initial clay inherited from the substrate along with authigenic clay. Scanning electron microscopy studies revealed that most of the green grains contain detrital minerals and clays with different structures. The abundance of detrital remains may also be responsible for the poor evolution of authigenic clays. (b) There may be two successive stages of neogenic products, for example, phyllite C + phyllite V or phyllite V + glaucony minerals sheltered in the same grain. (c) The alteration of neogenic minerals (represented by a brown colour in a few grains) is probably responsible for the unclear diffraction properties and behaviour with regard to treatment. Discussion

Origin and evolution of verdine grains and distinction between verdine and glaucony grains The verdine facies develops in a shallow marine oxidizing environment at depths between 5 and 60 m (Odin and Sen Gupta, in Odin, 1988). However, the authigenic clay minerals in this facies never form in contact with the oxidizing waters, but form in sheltered microenvironments within the granular substrates such as bioclastic debris, porous mineral debris and faecal pellets. The shelf green grains are free irregular grains ranging from angular to pellet-like morphologies. Many grains are homogeneous. The initial substrate for these grains may not be bioclastic debris, because within the infilled echinoderm test fragments the green clay exhibits only the second stage of evolution (Fig. 5); it would be contradictory to imagine, in the same sample, a carbonate substrate which was completely dissolved and gave rise to free green grains and undissolved carbonate substrate (even internally). The green pigment occurring within the fractures of mineral debris has been found between 40 and 100 m water depth; however, this debris is not a major substrate at about 40 m

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depth where dark green particles are concentrated. Therefore, the initial substrate may be larger clayrich faecal pellets inside which the authigenic clay was grown. Some pellets were subsequently broken down into different types of grains, probably due to reworking and to some extent by mineral evolution. The surfaces on these verdine grains with thin cracks correspond to the mineral evolution. In other words, these grains represent a long depositional history and have had enough time to evolve. The altered nature of some green grains represents a post-genetic phase of the history of the shelf facies. In contrast with the shelf grains, the initial substrate of all the green particles from the continental slope verdine facies is preserved and thus can be identified as either a faecal pellet or a carbonate test. It is difficult to explain the differences such as the shiny smooth aspect of the dark green glossy pellets and the rugose aspect of green and light green pellets. Scanning electron microscopy studies show more heterogeneous material in green and light green pellets, suggesting their less evolved nature compared with dark green pellets. The glossy and smooth texture of the glossy pellets may be due to the absence of heterogeneous components and the presence of small-sized clay blades and compact clay structures at the peripheral parts. The internal composition of the initial faecal pellet before the growth of authigenic clays may also be responsible for the smooth surface texture. The colour of the verdine grains has so far been used as an index of maturity (Odin, Bailey et al., in Odin, 1988; Bailey, 1988). However, our studies indicate that in spite of the identical colour (dark green) of the shelf grains (Fig. 3A) and the glossy pellets of the slope (Fig. 3C), (a) the authigenic mineral phases in them are different (Figs. 7A and 8A), and (b) the shelf grains show the dominance of altered phyllite C and the slope pellets show young phyllite V, indicating different stages of maturity. On the other hand, although the colour and morphology of the pellets (dark green glossy and green to light green rugose pellets, Fig. 3E and C) on the slope region vary, the authigenic clay in them does not show different stages of maturity (Fig. 8A and B). It is therefore suggested

that, in the case of the verdine facies, the colour misleads and cannot always be taken into account as a criterion to characterize the mineral and its evolution stage. The colour may roughly reflect (a) the type of material present in them, i.e. more clays in dark green grains and more carbonates and detrital minerals in green and light green pellets, and (b) the compactness of the clay structure (the dark coloured grains are more compact). It is difficult to distinguish the grains of the verdine from those of the glaucony facies. Excluding X-ray mineralogy, the colour, microstructure and nannostructure of the authigenic clay forming verdine facies are similar to those observed in the glaucony facies (Odin, 1975; Lamboy, 1976; Odin and Matter, 1981). In evolved verdine grains, the deep surface cracks are absent, unlike those in evolved glaucony grains. Wherever narrow cracks are present, the fresh looking substrate fragments are intimately mixed with the authigenic clay. Under SEM, the dark green parts of the verdine grains still contain some heterogeneous substrate fragments, whereas the dark green glaucony grains always contain pure evolved glauconitic minerals (Odin, 1975; Lamboy, 1976). The dark green infillings corresponding to verdine and their close association with the non-cracked foraminifer tests have to be linked to the mineralogical evolution. In dark green glaucony infillings, the alteration of the inherited clay, crystallization of glauconitic smectite and its recrystallization into glauconitic mica are associated with morphological grain transformations. It appears that the mineralogical evolution is less important in verdine grains and not associated with major morphological modifications, unlike in the glaucony grains.

Age of the verdinefacies Is verdinization a present day process? Before this study, the verdine facies was known from 13 locations from the present day oceans; it covers a long range of geographical areas in tropical latitudes. The age of the verdine facies so far reported from the continental margins is younger than 20,000 yrs B.P. (Odin and Sen Gupta, in Odin, 1988). The only exception to this is the Miocene verdine facies off the Niger delta (Por-

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renga, 1966). In several locations (Odin, Debenay et al., in Odin, 1988) it is found that the deeper verdine grains are more evolved than those at shallower depths; this has been logically explained by the fact that the sea stayed longer (between 20,000 yrs and the present) at deeper water depths and therefore more grains evolved there. The situation on the western continental margin of India differs from the above. Firstly, there are two discontinuous and distinct verdine facies associated zones, one on the outer shelf (about 40 m depth) and the other on the continental slope (100-280 m depth), wherein the nature of the green grains distinctly differs. Moreover, the shelf grains are more altered and the slope grains represent a young stage of verdine; this is contrary to the earlier reported observations (see Odin, Debenay, et al., in Odin, 1988). Secondly, in the shelf region there are a few bioclastic substrates which shelter the green clays and some of these substrates are altered. Numerous fresh looking bioclasts associated with the green grains do not show any sign of verdinization. This suggests that verdine formation may be a relict process. Thirdly, within the slope facies, young phyllite V occurs between 100 and 205 m water depth, followed by phyllite C down to 280 m depth. Finally a few verdine grains on the slope are coated by phosphate at 170 m depth (Part II). Moreover, the verdine at 280 m and glaucony/phosphate at 330 m are associated with non-glauconitized and non-phosphatized relict molluscs (5710 yrs B.P.). These suggest that the green facies and the phosphate facies on the slope are also relict and most probably formed simultaneously before 5710 yrs B.P. (see Part II). In view of these points, we believe that the formation of verdine facies at shelf and slope zones is relict and diachronous.

Hypothesis L related to Pleistocene formation The shelf grains may have formed during one of the Pleistocene sea-level highstands on the shelf and were reworked during subsequent regressive stages. The alteration of the foraminiferal tests and yellowish ochre colour coated green grains on the shelf may support the regression of sea level during which they developed these characters. On the continental slope, between 100 and 280 m

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water depth, pellets of different types and infillings are mixed. Although the colour and morphology of the pellets vary, they have not shown significant differences to indicate the stage of evolution. Odin and Fullagar (in Odin, 1988) reported that within a single deposit the grains may be heterogeneous and a mixture of different stages of evolution can be found within a single sediment. In view of this, these are considered to have formed in a single event. Compared with the shelf verdine grains, the slope verdine grains mainly show young phyllite V and are not altered; it is therefore assumed that they formed during the last glacial maximum when the sea level was at about 110-120 m below the present sea level, or during early transgression, or both.

Hypothesis II: related to Holocene formation It is known that verdine is very susceptible for alteration and easy to destroy (Odin, Bailey et al., in Odin, 1988). Its survival during sea-level regression is doubtful (Odin , 1988). If shelf verdine formed during the Pleistocene as just suggested, it would be expected to have been destroyed during subsequent emergent conditions of the shelf during the last glacial maximum. Phyllite V grows faster (1000 years as an order of magnitude) than glauconitic minerals, which need 10 to 100 times longer (Giresse et al., 1980; Odin and Matter, 1981; Odin, 1988). Considering the susceptible nature and fast growth of verdine and also its occurrence as a facies of formation in recent times (between 15,000 and 3000 yrs B.P.), we suggest an alternative explanation and propose that the shelf verdine might have formed during the early Holocene. The warm shallow water shelf environments and proximity of the river and lagoonal systems, as suppliers of metals, were probably favourable to a faster evolution of verdine. However, to explain their relict and altered nature, it is necessary to assume a change in the sedimentary environment; this could be a sea-level regression of short duration during the middle Holocene. Although there is no eustatic sea-level change known at this time in the world ocean, a sea-level regression has been reported in the northwestern continental shelf of India and attributed to tectonic movements. This low sea level is indicated in the form of vadose

AUTHIGENIC FACIES OF THE CONTINENTAL MARGIN OF INDIA

diagenetic textures in the limestones collected at 80-90 m depth on the shelf (Rao and Nair, 1992). Furthermore, the limestones collected at 80 m depth on the continental shelf (about 200 km north of the study area) show that the corals (dated at 8300 yrs B.P.) were encrusted by coralline algae (dated at 8000 yrs B.P.) suggesting a change in sedimentary environment, maybe a sealevel regression, at about this time. Such a sealevel regression is possible in the study area, but has not yet been reported. It is not known up to what depth sea-level regression has taken place due to tectonic disturbances and how long it remained. As young verdine is associated with the slope, it is suggested that this facies formed during this regression. Dating of the verdine may be useful in confirming the regression.

Deeper water glaucony facies (330 m water depth) The infillings and pellets at 330 m contain glauconitic smectite; this indicates that the deeper water glaucony facies succeeds the verdine facies. A similar sequence of authigenic phases has been observed on the continental margin of Gabon and Congo (Giresse and Odin, 1973). Glauconitization occurs close to the sediment-sea water interface and represents a low rate of sedimentation at the time of its evolution. Its association with abundant carbonate sediments and phosphate on the terrace confirms a low terrigenous input. The glaucony most probably formed (as discussed in Part II) contemporaneously with slope verdine facies occurring towards the land and phosphate facies towards the sea.

Controlling factors for the verdineformation The abundance of the verdine grains coincides with the abundance of carbonate detrital minerals (Fig. 2) in zones 1 and 2. However, the carbonate detrital minerals in zone 1 appear fresh compared with the associated verdine grains and we propose a pelletal substrate for these grains. It therefore seems that their relationship in this zone is incidental. In contrast, there appears to be a true relationship between carbonate substrates and verdine grains found on the continental slope and terrace.

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The low content of verdine in areas of terrigenous detrital minerals indicates that these may not be a preferable substrate for verdine growth. As verdine is found in pellets or in tests, some sort of substrate appears to be necessary for their growth. This agrees with the suggestion of Odin and Sen Gupta (in Odin, 1988) that the availability of initial substrate plays a major part in verdine formation, as in the formation of glaucony facies (Odin, 1975; Lamboy, 1976; Odin and Matter, 1981). Verdine grains are predominant in the sediments off Vembanad lake and the Periyar river mouth; this reflects the fact that these systems provided a large flux of metals, especially iron, and favoured formation verdine. As the study area is in tropical latitudes, these physiographic settings are in agreement with the previous observations on verdine facies (Odin and Sen Gupta, in Odin, 1988). Regarding the depth of verdine formation, shelf verdine is within the depth range (< 60 m) proposed by Odin and Sen Gupta (in Odin, 1988). However, slope verdine occurs between 100 and 280 m depth. If we assume that the slope verdine formed during the last glacial maximum when the sea level was at about -120 m (hypothesis I), it seems that verdine was forming within the depth range of 160 m (280-120 m = 160 m). Reworking of sediments during subsequent sea-level transgression can sometimes transport verdine laterally from shallower to deeper slope depths; however, it is not the case in the verdine seen here because the verdine (phyllite V) occurring at 100 205 m depth distinctly differs from the verdine (phyllite C) at 280 m depth and there is no mixing of these phases; this implies that verdine may have formed at more than 60 m water depth. A similar explanation was proposed for the slope verdine on the Senegalese continental margin (Odin and Masse, in Odin, 1988). If, however, we consider slope verdine formation during the Holocene regression (hypothesis II), there are relict molluscs (5710 yrs B.P.), whose normal living depths range from circalittoral to the upper bathyal zone, which are associated with verdine at 280 m depth. Consequently, the verdine probably stayed between 80 (upper limit of the circalittoral zone) and 280 m (present day depth) at about 5700 yrs B.P. To determine the depth of verdine formation, we need

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to deduct from these depth values the sea-level change, which is unknown. If the relict molluscs lived at 80 m water depth, verdine possibly formed at less than 60 m water depth; in this instance, tectonic movements should be considered to deepen the sea floor from 80 m (depth at about 5700 yrs B.P.) to 280 m (present day depth). If the relict molluscs lived at 280 m water depth, verdine probably formed at more than 60 m.

Verdine facies from the Indian Ocean The verdine deposit studied here is relict and covers an area of about 6000 km 2. This is the second deposit reported from the Indian Ocean. The first known verdine facies is located in the coral reef lagoon of Mayotte (Comoro Islands, near Madagascar) where the present day formation of verdine takes place at depths between 3 and 24 m (Odin, Debenay et al., in Odin, 1988). Although there are Holocene reefs on the central west coast of India and relatively large rivers, bringing much sediment to the continental shelf, both on the east and west coast of India, the verdine facies has not yet been reported in these regions. There are only some reports of green grains attributed to glauconite in the shelf sediments of the east coast of India. Detailed mineralogical studies are required on these grains and in other areas to identify new verdine facies in coastal regions of the Indian margins.

Part II: Phosphate association with glaucony and verdine Results

Phosphate-glaucony association in the terrace sediments (330 m water depth) The terrace sediments are sandy and consist of several types of skeletal components and friable phosphate aggregates.

Skeletal components The coarse fraction (> 1 mm) of the terrace sediments contains abundant non-glauconitized and non-phosphatized mollusc shells together with

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a few fish elements (teeth and otoliths). There are two types of associations: in the first type, bivalves (maximum 1 cm diameter) are predominant, followed by a few gastropods; within the bivalves, Veneridae (Timoclea) are dominant and Arcidae (Bathyarca sp.) are common. This type occurs at two stations (1401 and 1403) and the depths are 280 and 330 m. In the second type of association, bivalves, gastropods and scaphopods are found together at station 1428 (420m water depth); bivalves mainly belong to Malletiidae (Malletia) and Nuculidae (Nuculana), gastropods to Rissoidae (Benthonellania sp.) and scaphopods to Gadilidae (Cadulus sp.). It seems that the first type of association normally occurs from the circalittoral to upper bathyal zone, whereas the second association represents a typical upper bathyal region (P. Lozouet, pers. commun.) and their normal living depths agree with the depth of the samples. The 14C date on the shells from station 1401 gives an age of 5710___ 125 yrs B.P. Although the depths correspond to the living depths of these molluscs, living specimens are not observed at these stations. This could be due to the coincidence of terrace depths with the oxygen minimum zone (150-1500 m) reported in this region (Sen Gupta and Naqvi, 1984), whereas at about 5700 yrs B.P. oxidizing bottom conditions might have existed there. The 0.5-1 mm size fraction contains mostly white planktonic foraminifer tests, whereas in the 0.125-0.5mm fraction there are pale green to brownish and white planktonic foraminifers of the same species. The colour of the foraminifer tests is due to the infilling material. Tests without infilling material are white. Tests infilled with pale green material are distributed more landwards and tests infilled with brown material increase seawards. However, some foraminifer tests show both chambers infilled with green material adjacent to chambers infilled with brown material.

Distribution of foraminifer species The distribution of various planktonic foraminifer species in the 0.125-0.5 mm fraction is shown in Table 1. Two points should be observed from this table. Species such as Globigerinoides ruber, Globigerina bulloides and Globigerinita glutinata, which are susceptible to highly susceptible to dissolution,

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TABLE 1 Distribution of planktonic foraminifer species in the coarse fraction of the sediments of the 330-420 m terrace

Planktonic foraminifera

Distribution A

F

Globigerina bulloides Globigerinoides ruber G. trilobus t

F = frequent;

R

X

X

T

X

G. sacculifer G. fistulosus G. conglobatus Globorotalia menardii G. menardiiflexuosa G. tumida G. crassaformis Globoquadrina dutertrei G. conglomerata Pulleniatina obliquiloculata Globigerinella aequilateralis Orbulina universa Sphaeroidinella dehiscens A = abundant;

C

X

X X

X X X

X X X

C = common;

R = rare; and

T =

trace amounts.

are sparsely distributed (rare to trace amounts) or absent in these sediments. On the other hand, species moderately susceptible to dissolution and highly dissolution resistant species such as Globor-

otalia menardii, G. menardii flexuosa, G. tumida, Globoquadrina dutertrei, G. conglomerata, Pulleniatina obliquiloculata, Globigerinella aequilateralis and Globigerinoides trilobus are common to abundant in these sediments. The benthic foraminifers (both infilled and not infilled) are dominated by Buliminidae (Bolivina, Bulimina, Uvigerina). This association indicates a low oxygen content and high organic carbon and phosphorus content at the sediment-water interface (Boersma, 1990; Boersma and Mikkelsen, 1990). In one sample (1403) Miliolids and Ammonia occur in rare to trace amounts.

Composition and nannostructures of the foraminifer tests infillings: The X-ray diffraction studies indicate that glauconitic smectite is the authigenic clay in pale green infillings of planktonic foraminifers (Fig. 10). Calcite and CFA are present in the

brown infillings. The presence of calcite is due to the test material. Investigations on freshly broken surfaces of the foraminiferal tests infilled with green material under SEM indicate that the infilled materials are mostly compact authigenic clays (Fig. 11). Detrital fragments and fragments of biogenic tests, sometimes smaller foraminifers, are enclosed with the glaucony in the chamber (Fig. l lA). The clay mouldings of chambers and test pores can be seen at the peripheral parts (Fig. I1B). Under high magnification the presence of globular to rodshaped apatite structures can be seen at more than one place within the chamber. For example, these structures (Fig. 11C and D) are concentrated in one chamber of the enclosed smaller foraminifer test and partly fill the other chambers. The rodshaped apatite particles have been referred as phosphatized bacteria by several workers. The investigations on brown foraminifer infillings indicate that some of the test chambers are occupied by phosphatized microbial remains (Fig. 12). The inner wall separating the two chambers is being dissolved (Fig. 12A and B). The inframicron-sized apatite globules are associated with numerous filaments (Fig. 12B and C). Elsewhere, authigenic clays and apatitic globules (Fig. 12D) appear in the same chamber. In another brown coloured foraminifer (Fig. 13A), irregular hollow structures composed of inframicron-sized apatite globules constitute the majority of the chamber (Fig. 13B and C). The origin of the hollow structures is not known; they most probably indicate that the organic matter within the chamber could have been mineralized into apatite by microbial processes. The inframicron apatite particles enveloped the hollow structures; this may further indicate that apatite formation is an early diagenetic process before the complete decay of organic matter which acts both as source of phosphorus and as a substrate for phosphate mineralization as suggested by Southgate (1986). Aggregates of inframicron-sized apatite particles are associated with phosphatized filamentous structures (Fig. 13D).

Friable phosphate aggregates Friable phosphate aggregates are brown to dark brown, irregular and granular textured particles in

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Fig. I I. Scanning electron micrographs from freshly broken surface of a glaucony infilled foraminiferaltest. (A) Low magnification shows the chamber is filled with green clays and biogenic particles; (B) detail of an area similar to those indicated on (A) by white arrow (after rotation) showing that authigenic clay is moulding the test (white dots) and the test pores (white arrows); (C) detail of (A) showing that a small foraminiferis enclosed within authigenic clay and some chambers of the foraminiferare filled with apatite particles; and (D) detail of (C) showing globular and rod-shaped apatite particles, some of them moulding the inner wall.

the sand size (0.5-1 mm) fraction. They are particularly abundant at station 1403. They correspond to poorly cemented rock fragments and contain several types of grains. Scanning electron microscopy studies on freshly broken surfaces of these aggregates indicate that they consist of micritic groundmass enclosing a complete or part of several calcareous skeletal fragments (Fig. 14A). Micropellets (0.1 mm) are common in these aggregates. The groundmass contains 1-1.5 lam globular and rod-shaped apatite particles (Fig. 14B and D) with filaments locally attached to them. Some of the rod-shaped particles are flattened to the surface (Fig. 14C, D and E) and others converge resembling rosettes (Fig. 14D). Furthermore, rod-shaped

particles are found within the pores of the foraminifer tests (Fig. 14C and E) and penetrate into the chambers. Within the chamber they accumulate more near to the pores and protrude into the interior of the chamber. Coccoliths are abundant in some aggregates and some of them are being moulded by globular apatite particles (Fig. 14F). The morphology of these apatite particles are similar to those in phosphatized foraminifers (Fig. 11) and phosphatized bacteria reported in both ancient and recent phosphorites (O'Brien et al., 1981; Riggs, 1982; Mullins and Rasch, 1985; Soudry and Lewy, 1990; Lamboy, 1990a,b, 1993; Rao and Burnett, 1990; Garrison and Kastner, 1990; Brbh+ret, 1991). The filament-like structures

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151

Fig. 12. Scanning electron micrographs from freshly broken surface of a phosphate infdled foraminiferal test. (A) General view of one chamber; (B) detail of (A) showing phosphatized microbial filaments in contact with chamber wall (probably partly dissolved); (C) detail of (A) showing coalesced phosphatized microbial filaments associated with apatite globules on their surfaces; and (D) detail of (A) showing authigenic green clay (white arrow) close to inframicron-sized apatite particles.

Fig. 13. Scanning electron micrographs from freshly broken surface of a phosphate infilled foraminiferal test. (A) Surface of a wall (white dot) and infilled part of a chamber at low magnification; (B) detail of (A) showing irregular phosphatized (organic?) structures; (C) detail of (B) showing that these structures are hollow and have inframicron-sized apatite particles on their surfaces; and (D) detail of (A) showing phosphatized filaments and aggregates of inframicron-sized apatite particles.

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Fig. 14. Scanning electron micrographs from broken surface of a phosphatic friable aggregate. (A) At low magnification showing that different types of particles are cemented; (B) detail of (A) showing surface of a calcitic particle (white arrow) and interstitial phosphate cement; (C) detail of (A) showing a section of foraminiferal test with large pores; (D) enlarged portion of (B) showing ovoid and rod-shaped particles constituting the phosphatic cement - - note the combination of two or three particles forming rosettes (white arrows) on the substrate; (E) enlarged portion of (C) showing the test pores occupied by rod-shaped apatite particles, some of them being flattened against the test; and (F) in another part of the aggregate showing coccoliths (white arrows) being moulded by globular apatite particles.

(Fig. 12), the presence of cell-like structures (with filaments adhered to them) within the test pores (Fig. 14C and E) and their arrangement within the chamber can be explained by the invasion of the test chamber by microbes which subsequently mineralized into apatite; this may not be possible to explain by pure chemical precipitation. Furthermore, these structures are also close to the morphologies of the apatite particles synthesized in laboratory experiments by microbial mediation (Lucas and Pr6v6t, 1984; Hirschler et al., 1990). The microorganisms are thus assumed to have played a major part in releasing

phosphorus and precipitating CFA during early diagenesis. In some other aggregates, glauconitized foraminifers are enclosed in the fine-grained phosphatic matrix (Fig. 15A). Certain chambers of foraminifers are glauconitized (Fig. 15B and C) and adjacent chambers are filled with globular apatite (Fig. 15C). Some chambers contain different detrital particles (Fig. 15B) along with authigenic glaucony minerals. Figure 16 shows the intergrowth of glaucony blades and apatite globules in glauconitized foraminifers embedded in a phosphatic matrix.

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Fig. 15. Scanning electron micrographs from broken surface of a friable aggregate. (A) A complete foraminifer is enclosed in the aggregate; (B) detail of (A) showing a chamber filled with green clays along with detrital particles; and (C) detail of the central part of (A) showing that the first small chambers are filled with apatite globules and the large adjacent chambers are filled with green clays which mould the test and are clearly visible at the location of the test pores (white arrows). globules are enriched with calcium (CaO, 60.48%) and phosphorus (PzOs, 27.17%).

Discussion

Possible sources of phosphorus and environmental conditions for phosphatization

Fig. 16. Intergrowth of glaucony blades and apatite globules in the pores of a foraminiferal test (one pore is shown between two white dots) enclosed in a phosphate aggregate.

Phosphate-verdine association in the slope sediments Some green grains from the continental slope at 170 m water depth show a brown peripheral coating (Fig. 17A). These grains are associated with clayey sediments rich in organic matter ( > 4%; Paropkari et al., 1992). X-ray mineralogy showed that odinite (a mineral of the verdine facies) is the dominant authigenic clay in the pellets. Scanning electron microscopy studies at the transitional zone showed the intergrowth of authigenic clay and apatitic globules (Fig. 17B and C) similar to the phosphatized bacteria discussed earlier. Microprobe analysis indicates that these

Apatite occurs as an infilling material in foraminiferal tests and as cement in aggregates, indicating a diagenetic origin. The distribution of planktonic foraminiferal species (both phosphatized and non-phosphatized) indicates the presence of a very high number of dissolution resistant species and the absence of susceptible to highly susceptible species in the terrace sediments. Such an association generally corresponds to depths near the foraminiferal lysocline (FL), which is located at 3300 m in the Arabian Sea (Cullen and Prell, 1984). As the sediments of the terrace are shallower than the FL, the dominance of the dissolution resistant species may be due to the influence of seasonal coastal upwelling, which is well documented in this region (Wyrtki, 1973). It is probable that the upwelling mechanism played a major part in the accumulation of high contents of organic matter in the terrace sediments at the time of phosphatization. The phosphorus released to the interstitial pores

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v. PURNACHANDRARAG

Fig. 17. Phosphatized green grain from the verdine faciesof the slope area. (A) Broken surface showing coating at the lower portion; (B) detail of (A); and (C) detail of (B) showing intimatelyjuxtaposed apatite globules and authigenic clay blades.

probably stems from various processes. These are anoxic sulphate reduction processes, suboxic organic degradation processes at the sediment - water interface and/or the release of adsorbed phosphorus from ferric oxyhydroxide cycling processes and the metabolic activity of sulphur oxidizing microbial mats (Froelich et al., 1988; F611mi, 1989; Berner, 1990; O'Brien et al., 1990). It is difficult to determine their specific contribution as the phosphatic sediments under study are relict and sandy (discussed later). However, no microbial mats are encountered in the sediments. The oxygen minimum zone ( < 0.5 ml/1 dissolved oxygen) with intense denitrification occurs along the western continental margin of India at water depths between 150 and 1500 m (Sen Gupta and Naqvi, 1984). Therefore, at the time of phosphatization, the bottom sediments of the terrace were probably in contact with low oxygenated waters; this allowed the sulphate reduction process in the bottom sediments and favoured the release of phosphorus from the decomposition of organic matter. As glauconitic smectite and CFA are found together in the infillings and in aggregates, the process related to the reduction of ferric iron and the release of adsorbed phosphorus from oxyhydroxide surfaces could have contributed some phosphorus to the interstitial waters. The dissolution of fish elements might also have provided some phosphorus to the interstices of the sediments; this has been identified as an important phosphorus

source in other regions (Suess, 1981; Froelich et al., 1988). Carbonate fluorapatite within the test chambers implies that the organic matter in the test itself might have acted as an additional source of phosphorus. The components within the friable aggregates are predominantly carbonate fragments, suggesting low terrigenous sediment accumulation or dilution of terrigenous material by high productivity on this terrace. The inferred conditions such as upwelling, oxygen minimum zone and low terrigenous sediment accumulation are consistent with the general conditions for phosphorite formation suggested by several workers (Burnett, 1980; Arthur and Jenkyns, 1981; Burnett et al., 1983).

Time of phosphatization/glauconitization Two environmental conditions are evident from the terrace sediments. The presence of abundant coccoliths in the friable phosphate aggregates suggests that diagenesis took place in fine-grained sediments corresponding to low energy conditions. In contrast, the present day sediments on the terrace are sandy. The environments reflected by the friable phosphate aggregates are thus incompatible with the admixed sediments. It is therefore suggested that the diagenetic phosphate formation on the terrace is a relict process. The planktonic foraminifers present on the ter-

AUTHIGENIC FACIES OF THE CONTINENTAL MARGIN OF INDIA

race are relatively wide ranging species occurring from Late Pliocene to Recent age. However, the common presence of Globigerinoides trilobus, G. dutertrei and G. obliquiloeulata suggests a probable Pleistocene to Recent age. The quasi absence of the Quarternary index fossil Globorotalia truncatulinoides does not invalidate this point of view because many workers (Cullen and Droxler, 1984; B6 and Tolderlund, 1971) indicated that this species is rare in this region. The associated fossil molluscs are neither glauconitized nor phosphatized; this implies that glauconitization/phosphatization is older than 5700 yrs B.P. The actual time of phosphatization/glauconitization of the terrace sediments is presumed to have taken place during Late Pleistocene or Early Holocene low sea levels for the following reasons. (a) The terrace sediments contain glauconitic smectite which is an early stage of formation of glaucony; its presence indicates that the time for the formation of glaucony was short and therefore evolution has not reached the micaceous potassium end member of the glaucony family. (b) Verdine occurring at shallower slope (100-280 m) also shows its young stage (see Part I). (c) The presence of two early stage authigenic minerals of different but related facies at successive depths suggests that they formed at the same time over a short duration. Since the Late Pleistocene/Early Holocene was suggested for slope verdine (see Part 1), the same may hold good for glaucony. It is possible that this terrace acted as a sort of hardground for a short period at about this time, resulting in glauconitization and formation of an early diagenetic friable phosphate crust; during subsequent transgression reworking has taken place, fine-grained material from this terrace has been winnowed to deeper water and the early diagenetic crust disrupted into friable aggregates.

Relationships of phosphate with glaueony and verdine Verdine is the shallowest facies found at water depths between 40 and 280 m. The terrace sediments (at 330 m depth) are represented by an increase in glauconitic components towards the

155

land and an increase in phosphatic components towards the sea; this general distribution (the sequence of mineral phases from landward to seaward) respects the earlier contentions that these facies are neighbouring stages in the evolution of a sedimentary basin (Odin and Letolle, 1980). Subsequent processes complicated this distribution. For example, phosphate coated verdine grains occur at about 170 m depth (Fig. 17A). Similarly, the friable phosphate aggregates enclosed glauconitized foraminifers. The glauconitized test chambers are in contrast with the adjacent phosphate groundmass (Fig. 15A), indicating that phosphatization generally post-dates glauconitization and verdinization. However, foraminifers showing glaucony in one chamber and apatite globules in the adjacent chamber (Fig. 15C), and the presence of globular apatite within the foraminiferal test chamber surrounded by a glaucony matrix (Fig. l lC) are arguments in favour of the simultaneous formation of glaucony and apatite or their sequential formation close in time. The intergrowth of authigenic clay with apatite in phosphatized glaucony aggregates (Fig. 16) and in phosphatized verdine grains (Fig. 17B and C) shows their synchronous formation locally, which has not previously been reported in phosphorites. The formation of green clay (verdine/glaucony) and phosphatization are two different processes. Lithophile elements are required for verdine/glaucony formation, whereas organic matter and microbes (to concentrate phosphorus) are probably involved in phosphate formation. All these facies form in microenvironments and occur at or near the sediment-water interface. With respect to the time required to crystallize these minerals, Froelich et al. (1988) reported that phosphate pellets of 0.125-0.5 mm size form within 3 12 years. Odin and Fullagar (in Odin, 1988) suggested that glauconitic smectite (an early stage glaucony mineral) and verdine minerals crystallize within a few thousand years. On the other hand, Glenn (1990) and O'Brien et al. (1990) reported paragenetic sequences of these minerals suggesting that glauconite forms early, followed by the formation of apatite. The observed relationships of phosphate with glaucony and verdine may be explained by the

156

partial overlap of microenvironmental conditions and the formation time of minerals. Glaucony and verdine minerals form first in semi-confined environments. Their growth decreases when the microenvironment becomes more confined; this confined environment is probably suitable for apatite formation. Therefore, apatite growth begins before authigenic clay formation stops; this leads locally to the intergrowth of apatite and authigenic clay. Similar microenvironmental differences were assumed by Odin and Letolle (1980) for the growth of glaucony and phosphate. As apatite formation is microbial and concomitant with authigenic clay formation, microbial communities were probably responsible for changing the confinement of the microenvironment and controlling the contemporary formation of these minerals. Thompson (1987) suggested that bacterial respiration controls the regional microenvironments, which may hold the key to explaining the authigenesis of glauconite and phosphate at low oxygen conditions.

Conclusions The surficial sediments of the southwestern continental margin of India contain three types of relict authigenic mineral facies: a verdine facies at water depths between 40 and 280 m and an associated glaucony and phosphate facies on the terrace at about 330-420 m water depth. The verdine facies covers a wide area on both the continental shelf and slope and predominantly occurs in two zones: the shelf zone at about 40 m depth and the slope zone between 100 and 280 m water depth. Verdine grains from the shelf zone differ distinctly from those on the slope in colour, morphology, initial substrate, internal constituents and authigenic mineral phases. The shelf verdine facies is probably older than the slope facies. The actual age of the verdine is difficult to estimate precisely. It is proposed that the shelf verdine probably formed during the Pleistocene high sea levels or early Holocene and slope verdine during sea level lowstands. This verdine facies here is the fourteenth known location in the present day oceans. This facies is in agreement with those reported earlier with respect to the physiographic settings (tropical

V. PURNACHANDRA RAO

region and off the mouths of rivers and lagoons). As far as the depth of formation is concerned, our studies indicate that slope verdine may form at more than 60 m. However, the palaeodepth and its change after the formation of slope verdine are difficult to estimate. The terrace sediments consist of glaucony and phosphate facies; glauconitic smectite and carbonate fluorapatite are the respective mineral phases. Globular to rod-shaped apatitic bacteria-like structures or microbial filaments, or both, occur within the foraminifer test chambers and as cements in friable phosphate aggregates. The local simultaneous growth of apatite with authigenic clay within the phosphatized glaucony aggregates and phosphatized verdine grains was probably controlled by the microenvironment. It is suggested that, during the Late Pleistocene/ Early Holocene low sea levels, the terrace acted as a kind of hardground favouring the glauconitization/phosphatization of the pre-existing sediments. During subsequent transgression, reworking of terrace sediments has taken place and coarse grained glaucony/phosphate sediments remain as relict sediments.

Acknowledgements The authors thank Dr. G.S. Odin (P & M University, Paris) for looking at our samples and correcting our earlier interpretations of X-ray diffractograms. We thank Professor J.Ch. Fontes (University Paris-Sud) for radiocarbon dating and P. Lozouet (National Museum of Natural History, Paris) for identifying and providing some information on molluscs. Dr. G. Coquerel (University of Rouen) and M.N. Le Coustumer (CNRS, Caen) helped with the X-ray diffraction studies. Part of this paper was presented at the International Symposium on Phosphorites (February 22-March 1, 1992) held at Assiut (Egypt). Partial financial assistance to Rao was provided by project leaders IGCP 325 to attend the conference, tYNESCOprovided partial financial support. Rao thanks the Department of Science and Technology (India) and the Commission of the European Communities (Brussels) for offering a post-doctoral fellowship under which this work was carried out. He also

A U T H I G E N I C FACIES OF THE CONTINENTAL M A R G I N OF INDIA

thanks the Director and R.R. Nair of the National Institute of Oceanography, Goa, India, for their encouragement. The authors thank Dr. K. F611mi and an anonymous reviewer for their constructive comments. This paper is a contribution to IGCP project 325: "Correlation of palaeogeography with phosphorites and other authigenic minerals". References Arthur, M.A. and Jenkyns, H.C., 1981. Phosphorites and palaeoceanography. Proc. 26th Int. Geol. Cong. (Paris, France.) Ocean. Acta, Spec. Publ., pp. 83-96. Bailey, S.W., 1988. Odinite: A new dioctahedral-trioctahedral Fe rich 1:1 clay mineral. Clay Miner., 23: 237-247. Baturin, G.N., 1982. Phosphorites on the Sea Floor. (Developments in Sedimentology, 33.) Elsevier, Amsterdam, 343 pp. Br, A.W.H. and Tolderlund, D.S., 1971. Distribution and ecology of living planktonic foraminifera in surface waters of the Atlantic and Indian Oceans. In: B.M. Funnel and W.R. Riedel (Editors), Micropaleontology of Oceans. Cambridge Univ. Press, London, pp. 105 149. Berner, R.A., 1990. Diagenesis of phosphorous in sediments from non-upwelling areas. In: W.C. Burnett and S.R. Riggs (Editors), Phosphate Deposits of the World. Cambridge Univ. Press, pp. 27-31. Birch, G.F., 1980. A model of penecontemporaneous phosphatisation by diagenetic and authigenic mechanisms from the western margin of Southern Africa. In: Y.K. Bentor (Editor), Marine Phosphorites. SEPM Spec. Publ. 29: 79-100. Boersma, A., 1990. Late Oligocene to late Pliocene benthic foraminifers from depth traverses in the Central Indian Ocean. Proc. ODP, Sci. Results, 115: 315-380. Boersma, A. and Mikkelsen, N., 1990. Miocene-age primary productivity episodes and oxygen minima in the central equatorial Indian Ocean. Proc. ODP, Sci. Results, 115: 589-609. Br6hrret, J.G., 1991. Phosphatic concretions in black facies of the Aptian-Albian marnes bleues formation of the Vocontian basin (SE France) and at site DSDP 369: evidence of benthic microbial activity. Cretaceous Res., 121: 411-435. Burnett, W.C., 1980. Apatite glauconite associations off Peru and Chile: paleo-oceanographic implications. J. Geol. Soc. London, 137:757 764. Burnett, W.C., Roe, K.K. and Piper, D.Z., 1983. Upwelling and phosphorite formation in the ocean. In: E. Suess and J. Thiede (Editors), Coastal Upwelling, Part A. Plenum, New York, pp. 377-397. Cullen, J.L. and Droxler, A.W., 1990. Late Quaternary variations in planktonic foraminifer faunas and pteropod preservation in the Equatorial Indian Ocean. Proc. ODP, Sci. Results, 115: 597-588. Cullen, J.L. and Prell, W.L., 1984. Planktonic foraminifera of the northern Indian Ocean: distribution and preservation in surface sediments. Mar. Micropalentol., 9: 1-52. F611mi, K.B., 1989. Evolution of Mid-Cretaceous Triad. Lecture Notes in Earth sciences. Springer Verlag, Berlin, 153 pp. Froelich, P.N., Arthur, M.A., Burnett, W.C., Deakin, M.,

[ 57 Hensley, V., Jahnke, R., Kaul, L., Kim, K.H., Roe, K., Soutar, A. and Vathakanon, C., 1988. Early diagenesis of organic matter in Peru continental margin sediments: phosphorite precipitation. Mar. Geol., 80: 309-343. Garrison R.E. and Kastner, M., 1990. Phosphatic sediments and rocks recovered from the Peru margin during ODP Leg 112. Proc. ODP, Sci. Results, 112: 111-134. Giresse, P. and Odin, G.S., 1973. Nature minrralogique et origine des glauconies du plateau continental du Gabon et du Congo. Sedimentology, 20: 457-488. Giresse, P., Lamboy, M. and Odin, G.S., 1980. Evolution grom&rique des supports de glauconitisation: application ~i la reconstitution du pal~oenvironnement. Oceanol. Acta, 3: 251-260. Glenn, C.R., 1990. Pore water, petrologic and stable isotope data bearing on the origin of Modern Peru margin phosphorites and associated authigenic phases. In: W.C. Burnett and S.R. Riggs (Editors), Phosphate Deposits of the World, Vol. 3. Cambridge Univ. Press, Cambridge, pp. 46-61. Hirschler, A., Lucas, J. and Hubert, J.C., 1990. Bacterial involvement in apatite genesis. Microbial Ecol., 73, 211-220. Lamboy, M., 1976. Grologie marine du plateau continental au Nord de l'Espagne. Thesis, Univ. Rouen, 283 pp. Lamboy, M., 1987. Gen&e de grains de phosphate ~ partir de drbris de squelette d'&hinodermes: les processus et leur signification. Bull. Soc. Grol. Fr., 8, 4: 759-768. Lamboy, M., 1990a. Microstructure of a phosphatic crust from the Peruvian continental margin: phosphatised bacteria and associated phenomena. Oceanol. Acta, 13: 439-451. Lamboy, M., 1990b. Microbial mediation in phosphatogenesis: new data from the cretaceous phosphatic chalks from northern France. In: A.J.G. Notholt, and I. Jarvis (Editors), Phosphorite Research and Development. Spec. Pub1. Geol. Soc. London, 52: 157-167. Lamboy, M., 1993. Phosphatisation of calcium carbonate in phosphorites: ultrastructure and importance. Sedimentology 40, in press. Lucas, J. and Pr~v6t, L., 1984. Synth~se de l'apatite par voie bact+rienne fi partir de mati+re organique phosphat& et de divers carbonates de calcium dans des eaux douce et marine naturelles. Chem. Geol., 42:101-118. Marshall, J.F. and Cook, P.J., 1980. Petrology of iron and phosphorous-rich nodules from the East Australian continental shelf. J. Geol. Soc. London, 137: 765-771. Mullins, H.T. and Rasch, R.F., 1985. Sea floor phosphorites along the central California continental margin. Econ. Geol. 80: 696-715. Nair, R.R., 1985. Holocene phosphorites of the western continental margin of India, Mahasagar. Bull. Natl. Inst. Oceanogr. India, 18:273 279. Nair, R.R. and Pylee, A., 1968. Size distribution and carbonate content of the sediments of the western shelf of India. Bull. Natl. Inst. Sci. India, 38:411 420. Nair, R.R., Hashimi, N.H. and Rao, V.P., 1982. Distribution and dispersal of clay minerals on the western continental shelf of India. Mar. Geol, 50: MI M9. O'Brien, G.W., Harris, J.R., Milnes, A.R. and Veeh, H.H., 1981. Bacterial origin of East Australian continental margin phosphorites. Nature, 294: 442-444. O'Brien, G.W., Milnes, A.R., Veeh, H.H., Heggie, D.T., Riggs,

158 S.R., Cullen, D.J., Marshall, J.F. and Cook, P.J., 1990. Sedimentation dynamics and redox iron-cycling: controlling factors for the apatite-glauconite association on the East Australian continental margin. In: A.J.G. Notholt, and I. Jarvis (Editors), Phosphorite Research and Development, Spec. Publ. Geol. Soc. London, 52: 61-86. Odin, G.S., 1975. De glauconiarum, constitutione, origine, aetateque. Thesis, Univ. P. and M. Curie, Paris, 250 pp. Odin, G.S., 1985. La verdine, faci6s granulaire vert, marin et c6tier, distinct de la glauconie: distribution actuelle et composition. C.R. Acad. Sci. Paris, 301: 105-108. Odin, G.S., 1988. Green Marine Clays. (Developments in Sedimentology, 45.) Elsevier, Amsterdam, 445 pp. Odin, G.S., 1990. Clay mineral formation at the continent ocean boundary: the verdine facies. Clay Miner., 25: 477-483. Odin, G.S. and Letolle, R., 1980. Glauconitization and phosphatization environments: a tentative comparison. In: Y.K. Bentor (Editor), Marine Phosphorites, SEPM Spec. Publ., 29:227 237. Odin, G.S. and Matter, A., 1981. De glauconiarum origine. Sedimentology, 28: 611-641. Parker, R.J., 1975. The petrology and origin of some glauconitic and glauco-conglomeratic phosphorites from the south African continental margin. J. Sediment. Petrol., 45:230 242. Paropkari, A.L., Babu, C.P. and Mascarenhas, A., 1992. A critical evaluation of depositional parameters controlling the variability of organic carbon in Arabian Sea sediments. Mar. Geol., 107: 213-226. Porrenga, D.H., 1966. Clay minerals in recent sediments of the Niger delta. Proc. 14th Natl. Conf. Clays and Clay Minerals, Pergamon, Oxford, pp. 221-233. Rao, Ch.M., Paropkari, A.L., Mascarenhas, A. and Murty, P.S.N., 1987. Distribution of phosphorous and phosphatisation along the western continental margin of India. J. Geol. Soc. India, 30:423 438. Rao, V.P., 199l. Clay mineral distribution in the continental shelf and slope off Saurashtra, west coast of India. Indian J. Mar. Sci., 20:1 6. Rao, V.P. and Burnett, W.C., 1990. Phosphatic rocks and manganese crusts from seamounts in the EEZ of Kiribati and Tuvalu, Central Pacific Ocean. In: B.H. Keating, and

V. P U R N A C H A N D R A

RAO

B.R. Bolton (Editors), Geology and Offshore Mineral Resources of the Central Pacific Basin, Houston, Texas. Circum-Pacific Council for Energy and Mineral Resources, Earth Science Series, Vol. 15, pp. 285-296. Rao, V.P. and Nair, R.R., 1988. Microbial origin of phosphorites of the western continental shelf of India. Mar. Geol., 84: 105-110. Rao, V.P. and Nair, R.R., 1992. A re-evaluation of palaeoclimatic conditions during the Pleistocene and Holocene on the western continental shelf of India: evidence from the petrology of the limestones. In: B.N. Desai (Editor), Oceanography of the Indian Ocean. Oxford & IBH, New Delhi, pp. 423-438. Rao, V.P., Nair, R.R. and Hashimi, N.H., 1983. Clay mineral distribution on the Kerala continental shelf and slope. J. Geol. Soc. India, 24: 540-546. Rao, V.P., Natarajan, R., Parthiban, G. and Mascarenhas, A., 1990. Phosphatised limestones and associated sediments from the western continental shelf of India. Mar. Geol., 95: 17-29. Riggs, S.R., 1982. Phosphatic bacteria in the Neogene phosphorites of the Atlantic coastal plain continental shelf system. IGCP Project 156, Phosphorites Newsl., 11: pp. 34. Sen Gupta, R. and Naqvi, S.W.A., 1984. Chemical oceanography of the Indian Ocean, north of the Equator. Deep-Sea Res., 31: 671-707. Soudry, D. and Lewy, Z., 1990. Omission surface incipient phosphate crusts on early diagenetic calcareous concretions and their possible origin, upper Campanian, Southern Israel. Sediment. Geol., 66: 151-163. Southgate, P.N., 1986. Cambrian phoscrete profile, coated grains and microbial processes in phosphogenesis: Georgina Basin, Australia. J. Sediment. Petrol., 56: 429-441. Suess, E., 1981. Phosphate regeneration from sediments of the Peru continental margin by dissolution of fish debris. Geochim. Cosmochim. Acta, 45: 577-588. Thompson, J.B., 1987. A model for bacterially controlled deepwater anaerobic mineralization of glauconite and phosphorite. EOS, Am. Geophys. Union, 68: 1711. Wyrtki, K., 1973. Physical oceanography of the Indian Ocean. In: B. Zeitzschel (Editor), The Biology of the Indian Ocean. Springer, New York, pp. 18 36.