Petrological Characteristics and Genesis of the Central Indian Ocean Basin Basalts

Page | 1

Author version: Acta Geol. Sin., vol.86(5); 2012; 1154-1170

Petrological characteristics and genesis of the Central Indian Ocean Basin basalts

PRANAB DAS1, SRIDHAR D. IYER1 AND SUGATA HAZRA2 1NATIONAL INSTITUTE OF OCEANOGRAPHY (COUNCIL OF SCIENTIFIC AND INDUSTRIAL RESEARCH) DONA PAULA, GOA – 403004 INDIA 2SCHOOL OF OCEANOGRAPHIC STUDIES, JADAVPUR UNIVERSITY KOLKATA – 700 032 INDIA

ABSTRACT

The Central Indian Ocean Basin (CIOB) basalts are plagioclase rich while olivine and pyroxene

T are very few. The analyses of forty five samples reveal high FeO (~10-18 wt%) and TiO2 (~1.4-2.7 wt%) indicating these a ferrobasaltic affinity. The basalts have typically high incompatible elements (Zr

63-228 ppm; Nb ~1-5 ppm; Ba ~15-78 ppm; La ~3-16 ppm), a similar U/Pb (0.02-0.4) ratio as the N-

MORB (0.16±0.07) but the Ba/Nb (12.5-53) ratio is much higher than that of the normal mid-oceanic ridge basalt (N-MORB) (~5.7) and Primitive Mantle (9.56). Interestingly almost all of the basalts have negative Eu anomaly (Eu/Eu* = 0.78-1.00) that may have resulted by the removal of feldspar and pyroxene during crystal fractionation. These compositional variations suggest that the basalts were derived through fractional crystallisation together with low partial melting of a shallow seated magma.

Key words: Basalts; CIOB; genesis

Page | 2

1. INTRODUCTION

The Indian Ocean is exemplified by the Central Indian Ridge (CIR), Southwest Indian Ridge (SWIR) and Southeast Indian Ridge (SEIR). The Central Indian Ocean Basin (CIOB) extends from the Ninetyeast Ridge in the east to the south of the Chagos–Lacaddive Ridge and the CIR in the west and is bounded in the south by the Rodriguez Triple Junction and the northern part of the SEIR and in the north by India and Sri Lanka.

The half spreading rate of the CIOB crust has been recorded to be 80 mm/yr (McKenzie and Sclater, 1971) and 80 to 20 mm/yr (Patriat and Segoufin, 1988). Kamesh Raju and Ramprasad (1989) documented that during A25 to A23 the average rate was 80 mm/yr but decreased to an average of 36 mm/yr during A23 to A21. In contrast, Mukhopadhyay et al. (1997) reported a rate between 80 and 36 mm/yr while Rajendran and Rao (2000) suggested an average rate of 78 mm/yr. Dyment (1993) more precisely calculated the spreading rate as 68 mm/yr to 92 mm/yr till A24 that decreased to 45 mm/yr after A24. Hence, it is evident that the CIOB has witnessed several episodes of spreading at variable rates.

The CIOB (avg. water depth 5100 m, 4500-5600 m) hosts several morpho-tectonic features like the trace of triple junction on the Indian Plate (TJT-In, Dyment, 1993), fracture zones (FZ), seamounts and lineations (Fig. 1) that have notably affected the stability and volcanic activities of the basin. Dyment (1993) suggested that propagating rifts may have influenced the evolution of the TJT-In during chrons A28 to A21 (~68 to 50 Ma), the approximate time when the CIOB was formed.

Bathymetry reveals abundant isolated seamounts and seamount chains sub-parallel to each other and to the major FZ (73° E, 79° E and 75°45′ E). A 200 seamounts occur either as isolated edifies or along eight sub-parallel chains, that trend almost N-S and probably formed from the ancient propagative fractures. A majority of these near-axis seamounts may be the products of the temporally widespread (Cretaceous ~65 Ma to late Eocene <49 Ma) collision between India and Eurasia. The mutual effect of the regional stress patterns retained the orientation of the chains while the local stress regime aided upwelling of magma and construction of the seamounts. Evidences indicate that the morphotectonic structures developed concurrently with the formation of the oceanic crust (Das et al., 2007).

Iyer and Karisiddaiah (1990) reported the characteristics of the basalts and later Iyer (1995), Mukhopadhyay et al. (1995) and Das and Iyer (2007) identified these as normal mid-oceanic basalt (N- MORB) while spilites occur sporadically (Karisiddaiah and Iyer, 1992). In the Indian Ocean, Page | 3 ferrobasalts have been recovered from Deep-Sea Drilling Project (DSDP) Sites 214, 216 and 254 on the Ninetyeast Ridge, Site 256 from the Wharton Basin (Thompson et al., 1978), Southeast Indian Ridge (SEIR) (Anderson et al., 1980), the Southwest Indian Ridge (SWIR) (le Roex et al., 1982) and the Australian–Antarctic Discordance (Klein et al., 1991). Iyer et al. (1999) reported ferrobasalts to occur in areas of morphological highs and enhanced magnetic amplitude in the CIOB.

Here we present the geochemical characteristics of the CIOB basalts and evaluate and interpret their distinctive features that may shed light on the mantle source and the conditions responsible for melt generation. We draw attention to the fact that this study pertains only to the basalts collected from the seafloor while those from the seamounts would be dealt later.

2. MATERIALS AND METHODS

A majority of the recovered samples are fragments and pillow basalts with slightly to highly altered glassy layer. For this study samples with no glass or with a very thin veneer were selected. To avoid the alteration effects we removed the glass and in the latter case selected the interior part of the samples. Forty-five samples were sliced using a diamond embedded saw and then abraded with sandpaper to remove the saw trace and remaining visible alteration. The fresh samples were cleaned using acetone and distill water and were coarsely crushed in a hydraulic piston crusher before powdering in a tungsten-carbide ring mill. To minimise contamination effects all the samples were powdered for 2 minutes. Major element oxides concentrations were determined by using a X-ray fluorescence (XRF) following the analytical procedure described by Rhodes (1996). The minor, trace and rare earth elements (REE), were determined through an inductively coupled plasma-mass spectrometry (ICP-MS, Perkin Elmer’s Elan DRCe and Perkin Elmer’s Elan DRC-II). The sample powder (0.05 gm) was digested with a mixture of HF, HNO3 and HClO4 to remove silica. Diluted HNO3 was added to dissolve the digested material and later distill water was added to make a 100 ml aliquot for use with the ICP-MS. A blank solution was run for each set of five samples and the reading was used to correct the results for contamination (if any) during solution preparation. Repeated measurements of BHVO-1, BRC-1 and JB- 2 were made for calibration and calculation of the accuracy of the analysis. The relative standard deviations for the major oxides was < ± 1% while for the trace elements and REE analysis it was ±0.6 to 3%.

The mineral analyses were performed with Cameca SX100 microprobe. For the analysis, the operating voltage was 15KeV a probe current of 20 nA was used, and the beam diameter was 5 μm. Page | 4

Calibration was carried out using natural mineral standards (BRGM, Orleans Cedex, France). Precision was better than 5% for each element and counting time was 10s for each element. The obtained data were ZAF corrected internally to eliminate all possible element interferences (after Philibert, 1963). Semiquantative energy-dispersive spectra analysis (EDS) of five plagioclasen grains were carried out with a JEOL SEM-EDS link system (JEOL, Tokyo Japan) at the National Institute of Oceanography, Goa, India and rock magnetic studies were conducted by using a Mole spin instruments at Indian Institute of Geomagnetism, Alibagh, India.

3. RESULT AND INTERPRETATION

3.1. Hand specimens

Several dredging operations have been performed in the CIOB and a variety of rocks have been recovered (Fig. 1; Table 1). A total of forty five samples were examined during the course of this work. The rocks are pillow lavas, with the outer rind made of glass (Fig. 2). Altered glass is seen in a few basaltic fragments. The small fragments of glass/ basalt chips were probably dislodged from larger outcrops. Most of the rocks are either sparsely phyric or aphyric basalts. In some samples plenty of vesicles of irregular to rounded shapes are noticeable. Not much variation is apparent in hand specimens among the samples.

3.2. Petrography

Mineralogically CIOB basalts are plagioclase phyric basalts. A majority of the basalts have plagioclase as a dominant mineral phase while olivine and pyroxene occur as subordinate minerals. Plagioclase occurs as phenocryst and acicular grains as a constituent of ground mass and some of the grains have jagged ends and corroded margins. Plagioclase phenocryst exhibits prominent lamellar and cross-hatched twinning (Fig. 3a) while sector zoned plagioclase is rare (Fig. 3b). A few vesicles occur and these are mainly rounded in shape sometimes have secondary minerals.

Plagioclases occurring as phenocrysts or microlites are quite fresh. Plagioclase phenocrysts showing fracturing and wavy extinction and rarely, fine tiny and spherical inclusions are observed within the plagioclase phenocrysts. At places two sets of lamellar twinning are present in a single crystal (Fig. 3c). In places, plagioclase phenocrysts with prominent crystal outline on three sides and corroded on one side are observed. Plagioclase phenocrysts with basaltic groundmass extending and into the twin lamellae are noticeable. Plagioclase phenocrysts showing reaction margin with the glassy groundmass Page | 5

are common. A few phenocrysts of plagioclase show both twinning and zoning (Fig. 3d) while, grains showing reaction relations with the glassy groundmass are also present (Fig. 3e).

Generally, olivine grains are altered to iddingsite (brown colour mineraloid) and fractured. In one sample unaltered euhedral olivine occurs in a glassy matrix (Fig. 3 f). Olivine grains exhibit inclusion of the plagioclase laths into them.

The groundmass consists of acicular plagioclase (may be due to the effect of quenching) exhibiting radiating fibrous structure. The other component of groundmass is glass, which depict devitrification, and spherulitic texture. Commonly, the basalts show porphyritic and partly hyalophitic texture (Fig. 3g). Thin plagioclase laths are oriented in one direction forming a flow texture (Fig. 3h), while intersertal texture is also present.

3.3. Mineral chemistry

Microprobe of a few plagioclases indicates two compositions that compositionally they belong to

two generations. The plagioclase variety-I is labradorite to bytownite in composition (An64-89 Ab11-36

Or0-0.28) while variety-II is more sodic (An36-61 Ab35-63 Or0-6.54) . The higher contents of Ab and Or is either an inherent characteristic of the magma or else an effect of alteration. This aspect is discussed

later. Analysis indicates that the phenocrysts are Ca-rich (plagioclase variety-I, An64-89) in contrast to the

groundmass plagioclases (plagioclase variety-II, An36-61). The Fe content in the phenocryst is higher in the more fractionated basalts and is positively correlated with the An contents. Bryan (1974) noted Fe to occupy the tetrahedral sites of plagioclase and the variation of FeO content or occupancy of Fe in the tetrahedral site may be mainly controlled by the fO2 (oxygen fugacity) and probably form NaFeSi3O8 and CaFe2Si2O8 in plagioclase. This happens when pyroxene is in a liquidus phase and the increase of Fe from phenocryst to groundmass plagioclase may be due to Fe/Mg fractionation by pyroxene (Kennedy et al., 1993; Hammer, 2006).

3.4. Major oxides

The samples have enhanced (Fe2O3 + FeO), TiO2, P2O5 and K2O and a low MgO contents except for a few that have high MgO (> 6.5 wt%; Table 2). The basalts are mainly hypersthene normative but a few have normative quartz component. The total alkali content points to a sub-alkalic nature of the basalts. Page | 6

SiO2 shows a narrow range of variation and is moderately correlated (R=0.38) with MgO and decrease for MgO < 4 wt% indicating removal of olivine and pyroxene (Fig. 4a). An overlap between

the Indian Ocean Ridge (IOR) and the CIOB basalts occurs for 13.5 to 15 wt% Al2O3 for MgO concentrations between 4 and 6 wt% (Fig. 4b). An increase in Al2O3 (16-17 wt%) with decreasing MgO (2.5-4.5 wt%) again signifies the removal of olivine and enhanced crystallisation of plagioclase during fractionation.

TiO2 increases with decreasing MgO and a distinct overlap between the IOR and the CIOB

basalts and a relative enrichment of TiO2 occurs and the samples with high FeO and TiO2 fall in the

Indian Ocean ferrobasalts field (Fig. 4c). The CIOB basalts show relatively low concentrations of TiO2 and FeO, with respect to MgO, than the Bouvet Triple junction, Spiess Ridge segment, Broken Ridge and Kerguelen Plateau (le Roex et al., 1982, 1983; Mahoney et al., 1995) and thus, may reflect the source characteristics.

Several of the CIOB basalts showsignificant Fe enrichment with decreasing MgO (Fig. 4d) and samples with MgO between 4.5 and 7 wt% and FeOT <12 wt% significantly overlap with the IOR basalts (Fig. 4d). The TiO2 and FeO contents of the CIOB basalts probably suggest the onset of Ti- magnetite crystallisation but petrographic study did not reveal identifiable Fe-Ti-minerals. The low oxygen fugacity (fO2) condition prevented the formation of distinctly visible Fe-Ti minerals, rather these phases were preserved in the glassy groundmass as minute grains. The size of the ferromagnetic grains ranges from 1 to 0.06 μm are of single domain as noted from the significantly high χlf values (100 to -8 3 -1 140 10 m kg ) and the BOCR (coercivity of remanence) value that varies from 25 to 58 mT.

CaO contents (~8 to 12 wt%; Table 2) decreases with decreasing MgO (Fig. 4e) indicating removal of Ca-pyroxene during crystallisation (cf. Byerly, 1980; Flower, 1980; Perfit and Fornari, 1983; Cox, 1993) and this may have to led to enhanced Fe in the CIOB basalts.

The Na2O concentrations of the CIOB basalts (~2.4-3.5 wt%) are quite consistent (Table 2) and relatively lower than the Spiess Ridge segment (~2.6-4.68 wt%), SWIR (>3.5 wt%) and Bouvet Triple

Junction (~3.21-4.68 wt%) (le Roex et al., 1982, 1992). An increase in Na2O with decreasing MgO (Fig. 4f) can be accounted by the accumulation of Na-plagioclase in the groundmass, as reveled through microprobe analysis. This is similar to that observed in the Efate Island Group (Raos and Crawford, 2004) and in the Broken Ridge and Kerguelen basalts (Mahoney et al., 1995). Page | 7

The K2O contents (0.24 to 1.6 wt%) in the CIOB basalts increase e with decreasing MgO (Fig.

4g; Table 2) and resemble the K2O, Zr/Nb and (La/Yb)N contents of the SWIR segment (Michard et al., 1986; Dosso et al., 1988; le Roex et al., 1982, 1992). This either points towards a K-enriched magma or may be due to fractionation and or assimilation processes during evolution.

The CIOB samples have elevated TiO2 (~1.36-2.7 wt%), and a moderately high P2O5 (0.09-0.25 wt%) and K2O (0.25-1.1 wt%) than N-MORB (Table 2). An increase in K2O contents accompanied by a

comparable increase in TiO2 and P2O5 (Fig. 4h; Table 2), may be due to the difference in bulk distribution coefficient of those elements during partial melting of the sub-oceanic mantle i.e., DK< DP < DTi (Sun and McDonough, 1989).

The crystallisation of Ca-rich pyroxene is evident from the plot of CaO/Al2O3 vs MgO (Fig. 5a), while the narrow range of CaO/Al2O3 ratio (~0.5 to 0.9) attests to the removal of Ca-rich plagioclase.

3.5. Trace Elements

To interpret the trace elements data and identify the source and differentiation of the magmas, we compare the behaviour of the hygromagmatophile elements (Allegre and Minister, 1978), as these highly incompatible elements are either imperceptibly or not at all fractionated during magmatic evolution.

The CIOB basalts show considerable compositional ranges in trace elements (e.g., Zr= 48-228

ppm; Nb= ~1-5 ppm; Ba= ~15-78 ppm; La= ~3-16 ppm; Table 2)that vary with SiO2. For instance a

steady decrease of Ce, Rb and Nb occurs with increasing SiO2 (Table 2). V shows a positive trend against FeOT (Fig. 5b) that indicates ferric components to increase with progressive crystallisation. This is typical of Fe-Ti rich basalts and may be accounted by fractionation of olivine ± clinopyroxene and is

supported by the CaO/Al2O3 ratio (0.50-0.90) of the samples. The Ce/Yb and Nb/Zr ratios (1.73-3.89 and 0.008-0.044, respectively) in the CIOB basalts show a narrow range of variation and indicate that the incompatible element distribution was mainly controlled by partial melting of the source rock.

There is a fairly constant U/Pb ratio (0.02-0.4) and a relative enrichment of Th and U in a few of the CIOB samples, however, the strong enrichment of Pb relative to the N-MORB is enigmatic. The U/Pb ratio of the CIOB samples is similar to N-MORB whereas Ce/Pb and Ce/U ratios (1.27 to 8.17 and 10.4 to 157.36, respectively) are low compared to N-MORB (~25, >150; Hofmann et al., 1986) and Primitive Mantle (PM 28, 97; McDonough and Sun, 1995; Hannigan et al., 2001). The Ba/Nb (12.5-53) Page | 8 ratio is much higher than N-MORB (~5.7) and Primitive Mantle (9.56) and decreases with increasing MgO indicating the role of partial melting during the formation of the CIOB basalt (Fig. 6). A comparison of the present data with the other oceanic basalts (Newsom et al., 1986; Watson et al., 1987; Halliday et al., 1990) indicates that U is more incompatible than Pb during mantle melting in the CIOB.

Generally, for ease of comparison and to understand the deviation from a primitive composition. The analysed data are normalised with respect to either on estimated primitive mantle composition, chondritic meteorites or the primitive MORB resulting a spidergram (Rollinson, 1993). The representative plots of the incompatible trace element concentrations were normalised to N-MORB (Fig. 7a) and PM (Fig. 7b). U, La and Pb contents are relatively higher and Th and Nb relatively lower than N-MORB, whereas the less incompatible elements (e.g., Nd, Sm and Yb) show a flat N-MORB normalised distribution (Fig. 7a). N-MORB normalised distribution of the incompatible elements of the CIOB basalts show concentration factor of ~2 to 10 higher than the typical N-MORB. The PM normalised incompatible element distribution of the CIOB basalts also shows an increase in U, La, Pb and depletion of Th and Nb (Fig. 7b) indicating a single parental source for the magma.

In general, the incompatible element ratios although are variable (Zr/Nb = 9-166, Y/Nb = 7-63, Sm/La = 0.4-1.7) but are relatively higher vis-á-vis N-MORB (Fig. 7a; Table 2). The ratios of the highly incompatible element e.g., Ba/La (~1.5-17) is greater by a factor of ~8 than N-MORB (Ba/La ~ 1.96), whereas moderately incompatible element ratio such as Sm/La (0.4 to 1.7) is very close to N-MORB (Sm/La ~ 1.04). All the samples have steady and low elemental enrichment relative to MORB from right to left in the spider diagram (Fig. 7a) and show the flat pattern in mid-right part of the diagram, as well as prominent negative Nb and positive Pb anomalies. The variation and trend are nearly similar for all the samples, irrespective of the location of the samples, suggesting a similar source and mode of evolution.

When compared with the IOR basalts Zr, Ce, Rb, Ba and Sr show a variable distribution against Nb (Fig. 8), Zr and Ce are positively correlated with increasing Nb in the CIOB basalts (Figs. 8a, b) while most of these elements are more concentrated in the IOR field and also plot outside the field for a given concentration of Nb. Rb and Ba show a scatter with Nb content (Figs. 8c, d), but interestingly most of Ba is concentrated within the IOR field whereas most of Rb largely lies outside the field. Also, for a given Nb content Rb and Ba concentrations in the CIOB basalts are relatively higher than those of IOR basalts. Sr shows a flat distribution with Nb and is restricted within the IOR field (Fig. 8e). In Page | 9

accord with the evolved major element chemistry, the CIOB basalts are relatively enriched in Zr (48- 228 ppm), and Y (31-86 ppm) than the average N-MORB (Zr ~ 51 ppm, Y ~ 25, Sun et al., 1979).

A plot of the trace elements against Zr together with the data of ferrobasalts, basalts from Site 215 and K-rich basalts from the SWIR (Fig. 9) shows Rb in the CIOB basalts to have a positive trend similar to the Site 215 and SWIR K-rich basalts whereas the ferrobasalts are enriched in Zr and Rb than the CIOB basalts (Fig. 9a). Y shows a good correlation and a positive trend with Zr (Fig. 9b) in the CIOB basalts and is relatively enriched than the ferrobasalts, basalts from the Site 215 and K-rich basalts from the SWIR for a given concentration. Sc is fairly positively correlated with Zr and is relatively enriched than the K-rich basalts from the SWIR and ferrobasalts (Fig. 9c). The co-variation of Zr and Nb (Fig. 8a) and a range of Zr/Nb (9 – 166) suggest variable melting or alternatively the basalts were derived from a heterogeneous source. The high Ti/Nb, Zr/Nb and Y/Nb ratios (~310-1999; ~25- 166; ~7-63, respectively) suggest a very low degree of alteration and therefore the data could be used for the petrogenetic models (as discussed later).

The CIOB basalts have high Zr/Nb (>25) similar to N-MORB (> 30; Wilson, 1989). The plot of Ce/Y vs Zr/Nb indicates a close association of the CIOB basalts with the SEIR whereas Kerguelen and Site 215 basalts are relatively enriched in Y while Ninety East Ridge is relatively depleted (Figs. 9d). The extents of distribution of these elemental ratios of the CIOB basalts suggest fractional crystallisation. The plot (La/Sm)N vs Zr/Nb (Fig. 9e) indicates that the CIOB basalts are typical N- MORB yet, faint signatures of P-type MORB are noticeable in the mixing relation between N- and P- types and this may be indicative of a low degree of partial melting of the source rock. The La/Yb and Ce/Y ratios (~0.7-2.7 and ~0.15-0.62, respectively) of the CIOB basalts are close to the chondrite values (La/Yb ≈ 1.39 and Ce/Y ≈ 0.39; Sun and McDonough, 1989) and were probably affected by olivine and pyroxene crystallisation.

The CIOB basalt have enhanced high field strength elements (HFSE) e.g., Zr 48 to 228 ppm (Table 2). A narrow range of Zr/Ba ratio (1.5-5) indicates that these elements remained incompatible during magmatic evolution. Zr/Nb ratio (~9-166) (larger than primitive chondritic mantle of about 16, Sun et al., 1979) decreases with decreasing MgO and indicating clinopyroxene fractionation (cf. McCallum and Charette, 1978).

Page | 10

3.6. Rare earth elements

The chondrite normalised plot of REE of the CIOB basalts shows slightly enriched light rare earth element (LREE) than N-MORB and a flat heavy rare earth elements (HREE) patterns (Fig. 10a) similar to the N-MORB. Interestingly DSDP Site 215 basalts also enriched in LREE (Bougault, 1974; Thompson et al., 1974; Reddy et al., 1978; Frey et al., 1977). However, the CIOB samples have a negative Eu anomaly [Eu/Eu* = 0.78 – 1.00] (Fig. 10 a). Though the Eu anomaly is chiefly controlled by crystallisation of feldspar, particularly in felsic magma, because divalent Eu is accommodated in plagioclase in contrast to the other incompatible trivalent REE and plagioclase crystallisation would result in a negative Eu in the melt. Clinopyroxene fractionation from the early-formed basaltic liquid under low oxygen fugacity condition could also probably cause a negative Eu anomaly (McKay et al., 1986). A small negative Ce anomaly is present in the CIOB basalts. The negative Ce anomaly may have resulted from localized incipient weathering (Price et al., 1991), although no petrographic evidence for weathering has been detected. The CIOB basalts have a very gentle slope for REE distribution (with

[La/Yb]N = 0.62 – 1.6) which is akin to REE distribution in typical N-MORB (0.4-1.1; Humphris et al., 1985; Rolinson, 1993) and could be explained by the fractionation of pyroxene from an early formed melt. This view is attested by a relatively low CaO content in the CIOB basalts than in the N-MORB.

The CIOB basalts exhibit a fairly systematic LREE enrichment (Fig. 10a) as also shown by the increase in (Ce/Yb)N relative to Ce ppm (Fig. 10b). The minor variations in REE could be a consequence of crystal fractionation during emplacement of the basalts. Simple batch melting models suggest that the large range in (Ce/Yb)N ratio (0.48-1.6) may be due to variable extents of melting of a garnet lherzolite or because of fractional melting of a single source (Das, 2009).

4. DISCUSSION

The CIOB basalts are moderately alkali enriched with variable K2O/P2O5 (0.81 to 5.9). Although, Rb and Sr are sensitive to seawater alteration (Pearce, 1976; Verma, 1981) but K/Rb (300 to 800), Ba/Sr (0.2 to 0.5) and Ba/Rb (2 to 6) ratios suggest these basalts to be quite fresh. REE concentrations, which are believed to be immobile during seawater alteration (Kempe and Schilling, 1974) are also diagnostic to comprehend the nature of the source region and the magmatic processes.

The CIOB samples ferrobasaltic characteristics are equivalent to some of the more highly evolved ones from the global oceans (Clague and Bunch, 1976; Sinton and Hey, 1979; Byerly, 1980; Christie and Sinton, 1981; le Roex et al., 1982; Iyer et al., 1999). The CIOB basalts are similar in Page | 11

composition to Bouvet Island Hawaiites (SiO2 = 50.1 wt%, TiO2 = 3.5 wt%, Al2O3 = 15.4 wt%, FeO =

12.0 wt%, MgO = 4.3 wt%, CaO = 8.7 wt%, Na2O = 3.4 wt%, K2O = 1.3 wt%; Verwoerd et al., 1974)

and Spiess Ridge segment of SEIR (SiO2 = 50.52 wt%, TiO2 = 2.45 wt%, Al2O3 =14.39 wt%, FeO =

11.62 wt%, MgO = 5.51 wt%, CaO = 9.97 wt%, Na2O = 3.6 wt%, K2O=0.66 wt%; le Roex et al., 1982,

1992), except that the CIOB basalts show a relatively low TiO2 (~ 1.4 to 4 wt%) than the Bouvet and Spiess Ridge basalts. The Fe and Ti contents in the CIOB basalts show a good correlation with both MgO (Fig. 4 c, d) and Zr (Fig. 6a) and extend to compositions similar to some of the more Fe-rich differentiates from the Spiess Ridge (Fig. 4c, d).

The high FeO and TiO2 contents in the CIOB basalts indicate fractional crystallisation of olivine as also supported by the variation in the incompatible elements (Figs. 4, 6). Evidently, in terms of major elements the CIOB basalts, dredged from the 62-49 Ma old oceanic crust, are more evolved than the

Atlantic and Pacific basalts. In the CIOB basalts, SiO2 ranges between ~46 and 51% for MgO between

~2.8 and 7.1 wt% and Al2O3 and CaO vary from 15 to 18 wt% and from 8 to 12 wt%, respectively. The basalts with Mg# ranging from 27 to 57 indicates extensive fractionation, while the consistently high Fe and low Mg contents point towards a Fe-rich magma.

The fact that the CIOB basalts have quite consistent CaO/Al2O3, Na2O/TiO2 and Sr/TiO2 ratios but moderately high abundances of K2O and P2O5 (Table 2) strongly indicate that a single or similar magmatic source could account for the near constant composition of these basalts.

T In the CIOB basalts, FeO and TiO2 steadily increase with decreasing MgO for the range of N- MORB to Fe-Ti basalts (Fig. 4c, d). Fe-Ti basalts recovered from south transform intersection of Inca

Plate in Pacific Ocean typically have a narrow range of MgO (3.5-4.9 wt%) and SiO2 (50-52.2 wt%)

(Fornari et al., 1983) that are similar to the CIOB basalts. The range of FeO and TiO2 of the CIOB basalts at 2.8 to 5.21 wt% MgO has resulted in Fe-Ti enrichment which may be due to crystal fractionation. The Fe-Ti enrichment in the CIOB basalts is due to predominant removal of olivine and Ca-pyroxenes through fractionation as reflected by decreasing Ca and Mg with increasing Fe contents.

The bulk Kd of CaO, Al2O3 and MgO, ~1 for SiO2, and <1 for FeO, TiO2 and Na2O reflect that in the CIOB basalts FeO and TiO2 were concentrated in the liquid and eventually ilmenite and/or titanomagnetite cristallised from iron-rich basaltic liquids. Ilmenite crystallises only from melts developed below the oxygen partial pressures of the QFM (fugacity of oxygen that is fixed by the Page | 12

assemblage quartz-fayalite-magnetite at a given pressure and temperature; Carmichael and Ghiorso, 1986) buffer and is relatively scarce in N-MORB.

Here we used MgO as a proxy for the degree of differentiation since plagioclase is of limited importance in post-melting fractionation of MORB (Albarede, 1992). This is illustrated in a plot of FeOT vs FeOT/MgO (Fig. 11a) where apparent liquid lines of descent are shown. The apparent liquid line of descent for the CIOB basalts demonstrates the dominance of plagioclase over olivine and this in turn over clinopyroxene and by the variable Zr, Nb, Ce and Eu contents (Table 2).

Interestingly, K2O in the ferrobasalts at DSDP Site 216 (0.90 wt%), Spiess Ridge (0.5-1.11 wt%), Chain Ridge (0.77-0.83 wt %), Iceland (0.74 wt%), Broken Ridge (0.37-2.06 wt %) and Kerguelen Plateau (0.72-1.95 wt%) (le Roex et al., 1982; Dosso et al., 1988; Weis et al., 1993; Mahoney et al., 1995) all show high values as to the CIOB basalts. MOR tholeiitic trends are maintained by fractionation of olivine (Osborn, 1959), or olivine, clinopyroxene and plagioclase (Clague and Bunch,

1976; Bender et al., 1978; BVSP, 1981) in a low fO2 environment, resulting in a marked iron enrichment in the magma. Bender et al. (1978) showed that high-pressure fractionation of olivine and clinopyroxene is viable for fractionating MORB. Clague and Bunch (1976) calculated a ratio of 1:7.7:9.3 of crystal fractionation of olivine, clinopyroxene and plagioclase for differentiated ferrobasalts from East Pacific Rise (EPR), Galapagos Spreading Center (GSC) and Juan de Fuca Ridge.

The systematic variation in the ratios of highly incompatible elements in the CIOB basalts also supports the generation of single parent magma beneath the ancient SEIR system that formed the CIOB

crust (Fig. 9d). On a plot of (La/Sm)N versus TiO2 content (Fig. 11d), the scattered distribution could result from different extent of partial melting of a relatively heterogeneous source followed by fractional

crystallisation or differential melting of multiple sources that had similar REE and TiO2 contents.

The CIOB basalts are indistinguishable from N-MORB in terms of their trace element contents and their ratios (Figs. 8, 9) but are enriched in LREE relative to HREE and show a negative Eu anomaly (Fig. 10a). The variations of the ratios of alkali to less mobile elements e.g., Sm/Nd vs Rb/Sr plot (Fig. 12a) lacks any trend indicating that their distribution is affected by partial melting followed by the

fractionation of early formed plagioclase ± pyroxene . The positive trend of (Eu/Eu*)N vs (Sr/Nd)N points to an evolved nature of the basalts that have retained N-MORB signatures (Fig. 12b). The distribution of (Zr/Y)N against normalised Zr (Fig. 12c) indicates progressive partial melting, because the process would lead to enhanced Zr/Y ratio more effectively than a single stage partial melting. The Page | 13

higher Zr content and Zr/Y ratio help to distinguish basalts from fast spreading relative to slow spreading ridges and results from an open-system fractional crystallisation (Pearce and Norry, 1979).

Artificial neural network study of the CIOB also helped to geochemically delineate the CIOB basalts to be largely N-MORB but instances of enriched-MORB and ocean island basalt were noticed (Das and Iyer, 2009).

The K/Nd ratio (≈ 172 to 752) reflects that the magmatic source of the CIOB crust was enriched in alkalis relative to the REE. The U/Pb ratio (~0.02-0.4) close to the PM and MORB (~0.11 and ~0.16±0.07; Sun et al., 1979; Sun and McDonough, 1989; Halliday et al., 1995) whereas in the CIOB samples the Ba/Ce ratio (1.9-8.9) is higher and Ce/U ratio (10-157) is lower than the MORB (1.1±0.6, 180±79 respectively; Halliday et al., 1995) and the PM (3.9, 85 respectively; Halliday et al., 1995). This suggests a relative depletion of Ce and enrichment of Ba in N-MORB and thus rules out alteration of the CIOB basalts. Ba and U enrichment and Nb and Pb depletion in basalts from the Atlantic region reflect the recycling of the ancient crust (Halliday et al., 1995) but in case of the CIOB the depletion of Nb may be related to the partial melting of Nb depleted source rocks. The relation of Nb with La and other trace elements indicates that these N-MORB are enriched in Fe, Ti, LREE, LIL elements and depleted in Nb suggesting a small degree of partial melting of the source rock. In summary, the CIOB basalts are uniformly moderately enriched in LREE and alkali elements similar to ferrobasalts from other parts of the global oceans.

N-MORB normalised multi-element patterns (Fig. 7a) for the CIOB samples show an enrichment of U, La and Pb and depletion of Nb, Th, Ce and Pr. Most of the elements show a relative enrichment than the N-MORB. The HFSE abundances in the CIOB basalts are always significantly above the N-MORB levels, although the late phase fractionation, suggest derivation of the magma either from a less depleted, refractory peridotitic source or due to a low degree of partial melting. Notably, Nb/Y values (~ 0.01 to 0.13) in the CIOB are less than the MORB levels (~ 0.06 to 0.48), despite the enrichment of HFSE. Extensive partial melting of the source rocks can be ruled out during the evolution of the CIOB basalts because increased melting would result in decreased LREE (Nicholls and Harris, 1980), as attested by the low MgO contents (2.8 to 7.14 wt%) and high contents of incompatible elements (TiO2 = 1.4-2.7 wt%, La = 3-16 ppm, Rb > 6 ppm) that characterise low partial melting. Moreover, the observed high content of HFSE is not due to re-melting of a depleted mantle source (cf. Sun and Nesbitt, 1977) but Page | 14

may typical of the source. Therefore, it is evident that the CIOB basalts originated by a process of low partial melting coupled with fractional crystallisation.

The variations in Zr/Y and Ti/Y in the CIOB are probably related to the occurrence of a long- lived heterogeneous magma chamber, as also demonstrated by Zr/Nb and Y/Nb ratios (Table 2), that may be sustained because of high heat flow and magma flux (Michael and Cornell, 1998). A consensus of these studies is that olivine (± spinel) and plagioclase are the dominant phases that control the evolution of MORB. Interestingly, the CIOB basalts show a similar variation irrespective of the spreading regime i.e., the basalts recovered for this study encompass a range of half spreading rate (90 to 55 mm/yr, Dyment 1993; Das et al., 2007).

Attempts have been made to explain the formation of ferrobasalts by considering various magmatic processes. The along-axis petrologic variations of fast spreading ridges, such as the EPR (Langmuir et al., 1986) and the GSC (Clague and Bunch, 1976) and that of the moderate to fast spreading ancient SEIR (Klein et al., 1991) are explained by variable degree of shallow crystallisation or by difference in the depths of magma generation (Scheidegger, 1973). According to these studies, the commonly evolved magmatic compositions remain essentially constant over millions of years and represent small dilutions of a chamber maintained at a near steady state by repeated increments of mixing and fractionation. However, the basic questions involved are how and under what conditions magma is retained at a shallow depth for long durations and then undergoes subsequent differentiation. In this context was proposed the mechanism of the neutral buoyancy-zonation structure of magma reservoirs together with the role of fractional crystallisation in altering melt buoyancy (Ryan, 1994). The horizon of neutral buoyancy (HNB), defined “as that depth interval within which the melt magma density and the aggregate country rock density is equal” occurs sub-lithospherically and has a narrow vertical and a wide lateral extent. Beneath the HNB region, magma ascends due to of positive buoyancy and is stabilised at shallow depths (2-4 km), while above this region the magma descends by negative buoyancy. The neutral buoyancy of tholeiitic melts thus provides favourable conditions for their long- term stability over millions of years and the existence of magma reservoirs (Ryan, 1994).

Assuming ridge eruptions to be infrequent along an intermediate spreading ridge, as in our study area, considerable amounts of each new magma batch might remain in shallow along-axis magma chambers which could further fractionate and be subsequently joined by the next batch of ascending magma. Depending on the relative width of the zone(s) of intrusion, the frequency of magma Page | 15 replenishment would be low and hence magma residing in the chamber could undergo extremes in crystal fractionation (cf. Clague and Bunch, 1976). Fractional crystallisation is a theme that has often been used to explain the formation of ferrobasalts at the DSDP Sites 214 and 216 (Thompson et al., 1978), GSC (Byerly, 1980), the Conrad FZ (le Roex and Dick, 1981) and the Spiess Ridge (le Roex et al., 1982). and this may also be applicable to the CIOB samples. According to Ryan (1994), upward migrating off-axis melts may continue to be trapped in oceanic crust as old as 25-30 Ma and those high- level sills may occur at shallow depth. On this basis, the conjugate crust of the CIOB and the ancient SEIR may have been underlain at relatively shallow depths by rather small magma reservoirs (Mukhopadhyay et al., 1995; Das et al., 2005; Das et al., 2007).

Spreading rate by itself cannot bring about the extreme differentiation required for the formation of the ferrobasalts, implying the need to consider factors, such as rate of magma supply, low degree of partial melting and presence of relatively large, stable magma chamber.

5. CONCLUSION

The petrological investigations of the CIOB basalts permit to conclude that:

1. The basalts are nearly homogeneous and show well-defined chemical criteria that

distinguish them from other oceanic basalts, particularly by their high K2O, TiO2 and

P2O5 and low MgO contents. These characteristics indicate a ferrobasaltic affinity for the CIOB basalts.

2. The Fe-Ti rich magma was probably derived from varying extent of melting of LREE- enriched mantle sources.

3. A consistent CaO/Al2O3, Na2O/TiO2 and Sr/TiO2 ratios but moderately high abundances of K2O and P2O5 as well large range in (Ce/Yb)N and the systematic variation in the ratios of highly incompatible elements indicate a single or similar magmatic source .

4. Three geochemical co-variation diagrams, including : 1), SiO2-MgO, MgO-CaO, MgO- T T T CaO/Al2O3, K- P2O5, FeO - FeO /MgO,TiO2-Zr, FeO -Zr, FeO-V, (Zr/Y-Zr)N,

(Eu/Eu*)N-(Sm/Nd)N, (Ce/Yb)N-Ce, Y-Zr/Nb, Zr-Nb, Rb-Nb, and Ce-Nb show a positive T slope; 2), MgO-Al2O3, FeO -MgO, Na2O-MgO, K2O-MgO, Ba/Nb, and (La/Sm)N-Zr/Nb Page | 16

display a negative slope ; and 3), Ce/Y-Zr/Nb, Sr-Nb, and Sm/Nd-Rb/Sr exhibit a near zero slope (level, i.e. nearly parallel to abscissa), together indicate a progressive evolution :progressive partial melting and progressive fractional crystallization.

Acknowledgements: The samples were collected under the project “Surveys for Polymetallic Nodules” project funded by Ministry of Earth Sciences, (previously Department of Ocean Development), New Delhi. PD acknowledges the Council of Scientific and Industrial Research, New Delhi, for financial assistance in the form of a Research Fellowship. We acknowledge our colleagues at the NIO and Director for permission to publish this work. We are greatful to the Wadia Institute of Himalayan Geology (Dehradun), National Geophysical Research Institute (Hyderabad), Allahabad University (Allahabad) and Indian Institute of Geomagnetism (Alibagh) for analytical facilities. We thank Prof. F. A. Frey for suggestions on an earlier version of this manuscript and thank Dr. M. Shyam Prasad for his pep talks. We thankful to Zhou Guoqing and other anomalous reviewer for their critical review and Dr. F. Hongcai for considering the publication in the journal. This is NIO’s contribution #???

REFERENCES Albarede, F., 1992. How deep do the common basaltic magmas form and differentiate? Journal of Geophysical Research 97, 10997-11009. Allegre, C. J., Minster, J. F., 1978. Quantitative models of trace element behaviour in magmatic processes. Earth and Planetary Science Letters 38, 1-25. Anderson, R. N., Sparisou, D. J., Weissel, J. K., Hayes, D. E., 1980. The interrelation between variations in magnetic anomaly amplitudes and basalt magnetization and chemistry along the Southeast Indian Ridge. Journal of Geophysical ResearchJournal of Geophysical Research 85, 3883-3898. Basaltic volcanism study project (1981). Pergamon Press, New York, 1286 p. Bender, J. F., Hodges, F. N., Bence, A. E., 1978. Petrogenesis of basalts from the project FAMOUS area: experimental study from 0 to 15 kbars. Earth and Planetary Science Letters 41, 277-302. Bougault, H., 1974. Distribution of first series transition elements in rocks recovered during DSDP Leg 22 in the north eastern Indian Ocean. In: Initial Reports of the Deep Sea Drilling Project, Bougault H., Cande S. C. et al. (Ed.) Washington (U.S. Govt. printing office) 22, 449-457. Bryan, W. B., 1974. Fe, Mg relationships in sector-zoned submarine basalt plagioclase. Earth and Planetary Science Letters 24, 157-165. Page | 17

Byerly, G. R., 1980. The nature of differentiation trend in some volcanic rocks from the Galapagos Spreading Center. Journal of Geophysical Research 85, 3797-3810. Carmichael, I. S. E., Ghiorso, M. S., 1986. Oxidation-reduction relations in basic magma: a case for homogeneous equilibrium. Earth and Planetary Science Letters 78, 200-210. Christie, D.M., Sinton, J. M., 1981. Evolution of abyssal lavas along propagating segments of the Galapagos spreading centre. Earth and Planetary Science Letters 56, 321–335. Clague, D. A., Bunch, T. E., 1976. Formation of ferrobasalts at East Pacific mid-ocean spreading centers. Journal of Geophysical Research 81, 4247-4256. Das, P., 2009. Morphotectonic Evolution and Petrochemistry of Central Indian Ocean Floor. Ph.D thesis submitted to Jadavpur University, India. Das, P., Iyer, S. D., 2007. An investigation of basalts from the Central Indian Ocean Basin. 17th Annu. VM Goldschmidt Conference, Goldschmidt 2007 - "atoms to planets" August 19 - 24, Cologne, Germany, A202. Das, P., Iyer, S. D., 2009. Geochemical characterization of oceanic basalts using Artificial Neural Network. Geochemical Transactions, 10:13 doi: 10.1186/1467-4866-10-13. Das, P., Iyer, S. D., Kodagali, V. N., 2007. Morphological characteristics and emplacement mechanism of the seamounts in the Central Indian Ocean Basin. Tectonophysics 443, 1-18. Das, P., Iyer, S. D., Kodagali, V. N., Krishna, K. S., 2005. A new insight into the distribution and origin of seamounts in the Central Indian Ocean Basin. Marine Geodesy 28, 259-269. Donnellya, K. E., Goldsteina, S. L., Charles, H., Langmuir, C. H., and Marc, S., 2004. Origin of enriched ocean ridge basalts and implications for mantle dynamics. Earth and Planetary Science Letters 226, 347– 366 Dosso, L., Bougault, P., Beuzart, J. Y., Calvez, J. Y., Joron, J. L., 1988. The geochemical structure of the South-East Indian Ridge. Earth and Planetary Science Letters 88, 47-59. Dyment, J., 1993. Evolution of Indian Ocean Triple Junction between 65 and 49 Ma (anomaly 28 to 21). Journal of Geophysical Research 98, 13863-13878. Flower, M. F. J., 1980. Accumulation of calcic plagioclase in oceanic tholeiite: an indication of spreading rate. Nature 287, 530-532. Fornari, D. J., Perfit, M. R., Malahoff, A., Embly, R., 1983. Geochemical studies of abyssal lavas recovered by DSRV Alvin from Eastern Galapagos Rift, Inca Transform and Ecuador Rift. 1: major element variations in natural glasses and spatial distribution of lava. Journal of Geophysical Research 88, 10519-10529. Frey, F. A., Haskin, M. A., Poetz, J. A., Haskin, L. A., 1968. Rare earth abundances in some basic rocks. Journal of Geophysical Research 73, 6085-6098. Page | 18

Frey, F. A., Jones, W. B., Davies, H., Weis, D., 1991. Geochemical and petrological data for basalts from Sites 756, 757 and 758: implications for the origin and evolution of Ninetyeast Ridge. Proc. ODP Sci. Results, College Station, TX (Ocean Drilling Program) 121, 611-659. Halliday, A. N., Davidson, J. P., Holden, P., DeWofl, C., Lee, D. C., Fitton, J. G., 1990. Trace-element fractionation plumes and the origin of HIMU mantle beneath the Cameroon line. Nature 347, 523-528. Halliday, A. N., Lee, D. C., Tommasini, S., Davies, R. G., Paslick, C. R., Fitton, J. G., James, E. D., 1995. Incompatible trace elements in OIB and MORB and source enrichment in the sub-oceanic mantle. Earth and Planetary Science Letters 133, 379-395.

Hammer, J. E., 2006. Influence of fO2 and cooling rate on the kinetics and energetics of Fe-rich basalt crystallisation. Earth and Planetary Science Letters 248, 618-637. Hannigan, R. E., Basu, A. R., Teichmann, F., 2001. Mantle reservoir geochemistry from statistical analysis of ICP-MS trace element data of equatorial mid-Atlantic MORB glasses. Chemical Geology 175, 397–428. Hofmann, A. W., Jochum, K. P., Seufert, M., White, W.M., 1986. Nb and Pb in oceanic basalts: new constraints on mantle evolution. Earth and Planetary Science Letters 79, 33-45. Humphris, S. E., Thompson, G., Schilling, J. G., Kingsley, R. H., 1985. Petrological and geochemical variations along the Mid-Atlantic Ridge between 46° S and 32° S: influence of the Tristan da Cunha mantle plume. Geochimica et Cosmochimica Acta 49, 1445-1464. Iyer, S. D., 1995. A study of the volcanics of the Central Indian Ocean Basin and their relationship to the ferromanganese deposits. Unpubl. Ph.D. thesis, Jadavpur Univ., India 222 p. Iyer, S. D., 1999. Alteration of basaltic glasses from the Central Indian Ocean. Journal Geological Society of India 54, 609-620. Iyer, S. D., Karisiddaiah, S. M., 1990. Petrology of ocean floor rocks from Central Indian Ocean Basin. Indian Journal of Marine Science 19, 13-16. Iyer, S. D., Mukhopadhyay, R., Popko, D. C., 1999. Ferrobasalts from the Central Indian Ocean Basin. Geo-Marine Letters 18, 297-304. Karisiddaiah S. M. and Iyer S. D., 1992. A note on incipient spilitisation of Central Indian Basin basalts. Jour. Geol. Soc. India, 39, 518-523. Kamesh Raju K. A. and Ramprasad T., 1989. Magnetic lineations in the Central Indian Ocean Basin for the period of A24-A21: a relative study in relation to the Indian Ocean Triple Junction trace. Earth Planet. Sci. Lett., 95, 396-402. Kempe, D., Schilling, J.-G., 1974. Discovery Tablemount basalt: petrology and geochemistry. Contribution Mineralogy Petrology 44, 101-115. Page | 19

Kennedy, A. K., Lofgren, G. E., Wasserburg, G. J., 1993. An experimental study of trace element partitioning between olivine, orthopyroxene and melt in chondrules; equilibrium values and kinetic effects. Earth and Planetary Science Letters 115, 177-195. Klein, E., Langmuir, C. H., Staudigel, H., 1991. Geochemistry of basalts from the Southeast Indian Ridge, 115°-138° E. Journal of Geophysical Research 96, 2089-2108. Langmuir C. H., Bender J. F. and Batiza R., 1986. Petrologic and tectonic segmentation of the East Pacific Rise, 5°30’ N-14°30’ N. Nature, 322, 422-429. le Roex, A. P., Dick, H., 1981. Petrology and geochemistry of basaltic rocks from the Conrad fracture zone on the America-Antarctic Ridge. Earth and Planetary Science Letters 54, 117-138. le Roex, A. P., Dick, H. J. B., Watkins, R. T., 1992. Petrogenesis of anomalous MORB from the Southwest Indian ridge: 11° 53’ E to 14° 38’ E. Contribution Mineralogy Petrology 110, 253- 268. le Roex, A. P., Dick, H. J. B., Reid, A. M., and Erlank, A. J., 1982. Ferrobasalts from the Spiess ridge segment of the Southwest Indian ridge. Earth and Planetary Science Letters 60, 437-451. le Roex, A. P., Dick, H. J. B., Erlank, A. J., Reid, A. M., Frey, F. A., Hart, S.R., 1983. Geochemistry, mineralogy and petrogenesis of lavas erupted along the Southwest Indian Ridge between the Bouvet Triple Junction and 11 degrees East. Journal of Petrology24, 267-318. McDonough, W. F., Sun, S. -S., 1995. Composition of the Earth. Chemical Geology120, 223-253. Mahoney, J. J., Jones, W. B., Frey, F. A., Salters, V. J. M., Pyle, D. G., Davies, H. L., 1995. Geochemical characteristics of lavas from Broken Ridge, the Naturaliste Plateau and southernmost Kerguelen Plateau: Cretaceous plateau volcanism in the southeast Indian Ocean. Chemical Geology 120, 315-345. McKenzie D. P. and Sclater J. G., 1971. The evolution of the Indian Ocean since the Late Cretaceous. Geophys. J.R. Astron. Soc., 25, 437-528. McCallum, I. S., Charette, 1978. Zr and Nb partition coefficients: Implications for the genesis of mare basalts, KREEP, and sea floor basalts. Geochimica eta Cosmochimica Acta 42, 859-869. McKay, G., Wagstaff, J., Yang, S. R., 1986. Zirconium, hafnium and rare earth element partition coefficient for ilmenite and other in high-Ti lunar mare basalts: An experimental study. Proc. Lunar. Planet. Sci. Conf. 16th. Journal of Geophysical Research 91, D229-D237. Michael, P. J., Cornell, W. C., 1998. Influence of spreading rate and magma supply on crystallization and assimilation beneath mid ocean ridge: evidence from chlorine and major element chemistry of mid-ocean ridge basalts. Journal of Geophysical Research 103, 18325-18356. Michard, A., Montigny, R., Schlich, R., 1986. Geochemistry of the mantle beneath the Rodriguez Triple Junction and the Southeast Indian Ridge. Earth and Planetary Science Letters 78, 104-114. Mislankar, P. G., Iyer, S. D., 2001. Petrographical indicators of petrogenesis: examples from Central Indian Ocean Basin basalts. Indian Journal of Marine Science 30, 1-8. Page | 20

Mukhopadhyay, R., Batiza, R., Iyer, S.D., 1995. Petrology of seamounts in the Central Indian Ocean Basin: evidence of near axis origin. Geo-Marine Letters 65, 343-352. Mukhopadhyay R., George P. and Ranade G.R., 1997. Spreading rate dependent seafloor deformation in response to India-Eurasia collision: result of a hydrosweep survey in the Central Indian Ocean Basin. Mar. Geol., 140, 219-229. Newsom, H. E., White, W. M., Jochum, K. P., Hofmann, A. W., 1986. Siderophile and chalcophile element abundance in oceanic basalts, Pb isotope evolution and growth of the earth’s core. Earth and Planetary Science Letters 79, 33-45. Nicholls, G. D., Nalwalk, D., Hayes, E. E., 1964. The nature and composition of rock samples dredged from the Mid-Atlantic Ridge between 22° N and 52° N. Marine Geology. 1, 333-340. Osborn, E. F., 1959. Role of oxygen pressure in the crystallization and differentiation of basaltic magma. American Journal of Science 257, 609-647. Patriat P. and Segoufin J., 1988, Reconstruction of the Central Indian Ocean, Tectonophys., 155, 211- 234. Pearce, J. A., 1976. Statistical analysis of major element patters in basalts. Journal of Petrology 17, 15- 43. Price, R. C., Gray, C. M., Wilson, R. E., Frey, F. A. & Taylor, S. R. (1991). The effects of weathering on rare-earth element, Y and Ba abundances in Tertiary basalts from southeastern Australia. Chemical Geology 93, 245–265. Rajendran S. and Prakasa Rao T. K. S., 2000. Analysis of a north-south magnetic profile over the Central Indian Ocean. Curr. Sci., 78, 1378-1381. Raos, A. S., Crawford, A. J., 2004. Basalts from the Efate Island group, central section of the Vanuatu arc, SW Pacific: geochemistry and petrogenesis. Journal of Volcanology and Geothermal Research 134, 35-66. Reddy, V. V., Subbarao, K. V., Reddy, G. R., Matsuda, J., Hekinian, R., 1978. Geochemistry of volcanics from the Ninetyeast Ridge and its vicinity in the Indian Ocean. Marine Geology 26, 99-117. Rhodes, J. M., 1996. Geochemical stratigraphy of lava flows sampled by the Hawaii Scientific Drilling Project. Journal of Geophysical Research 101, 11729–11746. Robinson, S.G., 1986. The late Pleistocene paleoclimatic record of North Atlantic deep-sea sediments revealed by mineral magnetic measurements. Physics of the Earth and Planetary Interiors . 42, 22– 47. Ryan, M. P., 1994. Neutral-buoyancy controlled magma transport and storage in Mid-Ocean Ridge magma reservoirs and their sheeted-dike complex: a summary of basic relationship. In: Magmatic systems, Ryan M. P. (Ed.), Academic Press San Diego, 97-135. Page | 21

Scheidegger A. 1973. On the prediction of the reach and velocity of catastrophic landslides. Rock Mechanics 5: 231-236. Sinton, H. R., Hey, R., 1979. Oceanic ferrobasalts, off-ridge magmas and propagating rifts along the Galapagos Spreading Center. EOS Transaction American Geophysical Union 60, 46, 971. Sun, S.S., Nesbitt, R.W., 1977. Chemical heterogeneity of the Archaean mantle, composition of the earth and mantle evolution. Earth and Planetary Science Letters 35, 429-448. Sun, S. S., McDonough, W. F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Magmatism in the ocean basins, Saunders A. D., and Norry M. (Ed.), Geological Society of London Special Publication, London 42, pp. 313-345. Sun, S. S., Nesbitt, R.W., Sharaskin, A. Y., 1979. Geochemical characteristics of mid-ocean ridge basalts. Earth and Planetary Science Letters 44, 119-138. Thompson, G., Bryan, W. B., Frey, F. A., Sung, C. M., 1974. Petrology and geochemistry of basalts and related rocks from Sites 214, 215, 216, DSDP Leg 22, Indian Ocean. In: Initial Reports of the DSDP, von der Borch C. C., Sclater J. G. et al. (Ed.), Washington, (U. S. Govt. printing office) 22, 459-468. Thompson, G., Bryan, W. B., Frey, F. A., Dickey, Jr. J. S., 1978. Basalts and related rocks from deep- sea drilling sites in the central and eastern Indian Ocean. Marine Geology 26, 119-138. Thy, P., 1989. Phase equilibrium constrains on the evolution of transitional and mildly alkalic Fe-Ti basalts in the rift zones of island. In: Evolution of Mid Ocean Ridges, Sinton J. M. (Ed.), American Geophysical Union Washington, D.C. 57, pp. 39-51. Verma, S. P., 1981. Seawater alteration effects on 87Sr/86Sr, K, Rb, Cs, Ba, and Sr in oceanic igneous rocks. Chemical Geology 34, 81-89. Verwoerd, W. J., Erlank, A. J., Kable, E. J. D., 1974. Geology and geochemistry of Bouvet Island. Proceeding Symposium. On Andean and Antarctic volcanology problems, Santiago, Chile, September. Watson, E. B., Othman, Ben, D., Luck, J. M., Hofmann, A. W., 1987. Partitioning of U, Pb, Cs, Y, Hf, Re and Os between chromian diopsidic pyroxene and haplobasaltic liquid. Chemical Geology 62, 191-208. Wilson M. (1989). Igneous petrogenesis. Chapman and Hall, Netherlands, 466 p. Weis, D., Frey, F. A., Leyrit, H., Gautier, I., 1993. Kerguelen Archipelago revisited: geochemical and isotopic study of the SE province lavas. Earth and Planetary Science Letters 118, 101-119. Weis, D., Frey, F. A., 1996. Role of Kerguelen plume in generating the eastern Indian Ocean seafloor. Journal of Geophysical Research 101, 13831-13849. Page | 22

Figure 1 (a) Generalised map of Indian Ocean showing the sampled sites from the CIOB (map prepared using GeoMap). (b) bathymetric map of the study area showing morphological features and sample locations.

Page | 23

Figure 2 Representative photographs of the hand specimen.

Page | 24

Page | 25

Figure 3 a) Plagioclase phenocrysts with cross-hatched and lamellar twinnings. The intergranular space is occupied by devitrified glass. Plagioclase phenocrysts are subhedral in nature, fractured and have inclusions. The rims of plagioclase grain show reaction relation. b) Subhedral plagioclase grain with corrugated grain boundary showing reaction relation with the groundmass. Sector zoning and lamellar twining are also present in plagioclase grains. The devitrified glassy material occupies the intergranular space. c) Large fractured plagioclase phenocrysts surrounded by small plagioclase grains and microlites. Lamellar twinning shown by both large and small plagioclase grains. Uneven grain boundaries of the plagioclase indicate reaction relation with the groundmass. Intergranular space occupied by devitrified and partly altered glass. d) Subhedral plagioclase grains show zoning and lamellar twinning by the plagioclase phenocrysts. Inclusions of plagioclase present in the plagioclase phenocryst. e) Plagioclase grains form a porphyritic texture, lamellar twinning displayed by the individual plagioclase grains and the devitrified glassy material occupies the intergranular space. f) Subophitic texture displayed by the plagioclase and olivine grains. Both olivine and plagioclase grains are fractured and embedded in the glassy matrix. Secondary minerals and glassy groundmass have occupied the fractures. Subhedral plagioclase grains show lamellar twinning. g) Euhedral fresh olivine grain embedded in the glassy matrix. The olivine grain is fractured and occupied by iddingsite. Microlites of plagioclase showing flow texture and the intergranular space is covered by the glassy groundmass.

Page | 26

Figure 4

Page | 27

(a) SiO2 varies between 44% and 51% for MgO content between 2% and 6% suggesting fractional crystallisation. The CIOB basalts are more evolved than those from the Indian Ocean Ridge (IOR; data from PETDB http://petdb.ldeo.columbia.edu).

(b) Al2O3 increases with decreasing MgO indicating fractionation of olivine and pyroxene and accumulation of plagioclase. Ferrobasalts from Spiess and Chain ridges show a good comparison with the CIOB basalts. The IOR basalts are relatively depleted in Al2O3 for a given concentration of the MgO.

T (c)–(d) TiO2 and FeO increases with decreasing MgO content indicating fractionation of olivine and pyroxene and accumulation of plagioclase. The CIOB basalts show a relatively enrichment in FeO content and a close association with the ferrobasalts from Spiess and Chain ridges (le Roex et al., 1982; Dosso et al., 1988; Weis et al., 1993; Mahoney et al., 1995).

(e) CaO varies considerably (8-12%) for a narrow range of MgO (2-6%) indicative of removal of olivine, pyroxene and Ca-plagioclase.

(f) Na2O increases with decreasing MgO indicating accumulation of Na-plagioclase during evolution of the CIOB basalts.

(g) K2O shows a sharp increase with decreasing MgO and a close association with the ferrobasalts from other oceans.

(h) P2O5 increases with increasing K2O. TheCIOB basalts are relatively depleted in P2O5 than the K-rich basalts from SWIR (le Roex et al., 1992) and DSDP Site 215 (Mahoney et al., 1995; Weis and Frey, 1996).

Page | 28

Figure 5

(a) Removal of olivine, pyroxene and Ca-plagioclase during magmatic evolution as substantiated by MgO vs CaO/Al2O3 ratio.

(b) V shows an increasing trend with increasing FeOT.

Page | 29

Figure 6

Ba/Nb ratio of the CIOB basalts shows a increasing trend with decreasing MgO indicate that the distribution of these incompatible elements were controlled by the fractional crystallisation of the magma.

Page | 30

Figure 7

Normalised spider diagrams for the trace elements of the CIOB basalts.

(a) N-MORB (Sun et al., 1979) normalised spider diagram of the CIOB basalts shows a relatively higher concentration than the N-MORB. U, La and Pb show a positive anomaly whereas Th, Nb, Ce and Pr show a negative anomaly relative to other elements.

(b) Primitive mantle (Sun and McDonough, 1989) normalised spider diagram of the incompatible elements of the CIOB basalts shows a relative enrichment for the highly incompatible elements. Again, U, La and Pb show a positive anomaly and Th, Nb, Ce and Pr show a negative anomaly relative to the other elements.

The positive Pb and U anomaly of the CIOB indicates that both U and Pb acted as incompatible elements and were relatively enhanced in the source rock. Page | 31

Figure 8

The incompatible elements Zr, Ce, Rb, Ba and Sr show a variable distribution with Nb of the CIOB basalts. The hatched area represents IOR basalts (data from PETDB http://petdb.ldeo. columbia.edu).

Page | 32

Figure 9

(a-c) The plots of Rb, Y and Sc against Zr of the CIOB basalts show positive trends. Ferrobasalts (le Roex et al., 1982), basalts from DSDP Site 215 (Weis and Frey, 1996) and K-rich basalts from SWIR (le Roex et al., 1992) are plotted for comparison. Rb content of the CIOB basalts shows a close association with the K-rich basalts and basalts from the DSDP Site 215 where as ferrobasalts are relatively enriched in Zr. The concentrations of Y and Sc in the CIOB basalts are relatively enriched than the DSDP Site 215 ferrobasalts and basalts.

(d) The variation in the Ce/Y vs Zr/Nb ratios of the CIOB basalts mostly falls in the SEIR domain and indicates a genetic relation. In comparison, the basalts from Ninety East Ridge and DSDP Site 215 Page | 33

(Weis and Frey, 1996) do not show any relation with the CIOB basalts.

(e) (La/Sm)N-Zr/Nb binary mixing diagram indicates that the CIOB basalts are mainly N-MORB but some component of P-MORB is also present. (For comparison DSDP Site 215 and Broken Ridge data of Weis and Frey, 1996 and Mahoney et al., 1995 have been plotted).

Figure 10

(a) Chondrite normalised (Sun and McDonough, 1989) REE distribution shows the CIOB basalts to be relatively enriched in LREE and to have a flat HREE. Generally, Ce and Eu show a negative anomaly except for sample HRX27 that shows a positive Eu anomaly and a decreasing HREE trend.

(b) (Ce/Yb)N-Ce diagram indicates that Ce variation is initially controlled by the melting of the source rock and later by fractional crystallisation.

Page | 34

Figure 11

T Both (a) FeO and (b) TiO2 show a strong positive correlation with Zr indicating their enrichment in the CIOB basalts during the partial melting of the source rocks.

(c) FeOT versus FeOT/MgO plot to demonstrate the dominance of mineral phases during fractional crystallisation. The data are subdivided into three groups on the basis of MgO content: open circles indicate MgO 2-4 wt%, light gray MgO 4-6 wt% and dark gray represent MgO > 6 wt%. Fractionation over this range of MgO is evidently dominated by olivine and clinopyroxene and even the most evolved samples show a limited evidence of plagioclase removal. The removal vectors are after Albarède (1992).

(d) TiO2 and chondrite normalised La/Sm ratio (Sun and McDonough, 1989) of the CIOB basalts show a scatter suggesting an heterogeneous source.

Page | 35

Figure 12

(a) Sm/Nd versus Rb/Sr ratios indicate their enrichment in the CIOB basalts.

* (b) Chondrite normalised (Sun and McDonough, 1989) (Eu/Eu )N versus (Sr/Nd)N ratio shows the enriched N-MORB nature of the CIOB basalts.

(c) Chondrite normalised plot of (Zr/Y)N ratio and (Zr)N of the CIOB basalts shows a good correlation brought about by the progressive partial melting of the source rocks. Page | 36

Table 1: Sample location of basalts from the Central Indian Ocean Basin

Sample Latitude (°S) Longitude (°E) Water Depth (m)

PD1 13° 00’ 00” 75° 59’ 00” 5374 PD2 11° 00’ 32” 73° 59’ 12” 3600 PD3 12° 24’ 0” 76° 14’ 30” 4400 PD5 13° 00’ 26” 76° 00’ 25” 5374 PD6 11° 30’ 45” 76° 15’ 30” - PD7 12° 33’ 25” 78° 45’ 00” 4830 PD8 13° 00’ 26” 76° 00’ 25” 4550 PD10 12° 45’ 00” 74° 45’ 24” 4732 PD11 12° 57’ 30” 75° 00’ 06” 5348 PD14 12° 23’ 42” 76° 13’ 30” 5350 PD16 11° 59’ 00” 76° 50’ 57” 5402 PD18 12° 59’ 50” 76° 00’ 08” - PD19 11° 40’ 19” 77° 47’ 16” 5210 PD20 14° 05’ 00” 76° 18’ 00” 5190 PD21 11° 00’ 00” 77° 55’ 24” 5300 PD22 13° 20’ 42” 77° 30’ 00” 5050 PD25 11° 15’ 22” 73° 15’ 46” 4972 PD27 13° 45’ 24” 75° 45’ 20” 5320 PD28 13° 00’ 26” 76° 00’ 25” 5370 PD30 12° 59’ 00” 76° 30’ 00” 5400 PD31 13° 7’ 02” 77° 44’ 31” 5380 PD32 12° 00’ 00” 77° 55’ 24” 5440 PD35 12° 46’ 08” 78° 23’ 13” 5250 HRX2 12° 33’ 22” 77° 12’ 54” 5210 HRX3 11° 04’ 02” 77° 52’ 00” 5300 HRX5 13° 10’ 00” 75° 59’ 00” 5330 HRX9 13° 00’ 00” 76° 00’ 00” 5374 HRX18 11° 00’ 00” 77° 55’ 24” 5300 HRX20 11° 59’ 00” 76° 30’ 57” 5402 HRX21 12° 00’ 00” 76° 30’ 00” 5400 HRX28 11° 15’ 22” 73° 14’ 46” 4972 HRX29 12° 28’ 04” 76° 52’ 01” - HRX30 (1) 14° 44’ 49” 77° 59’ 20” 5083 HRX31 13° 00’ 00” 75° 59’ 00” 5374 HRX34 12° 30’ 00” 78° 12’ 00” 5120 HRX35 9° 59’ 00” 79° 15’ 00” 5212 HRX36 12° 00’ 00” 76° 30’ 00” 5400 HRX37 12° 59’ 55” 76° 30’ 00” 5400 SS13-TS 90/91 13° 40’ 00” 79° 30’ 00” 5040 AAS22 DR#19 14° 20’ 55” 76° 33’ 25” 5125 SS11 12° 23’ 20” 76° 13’ 30” 5325

Page | 37

Table 2: Major, trace and REE analysis of CIOB basalts

PD1 PD2 PD3 PD5 PD6 PD7 PD8 PD10 PD11 PD14 PD16 PD18

SiO2 49.26 50.7 50.37 51.08 49.67 49.24 49.84 49.32 49.14 49.38 50.32 50.98

TiO2 2.04 1.9 2.5 1.59 2.22 2.17 1.61 1.48 1.98 1.35 2.18 1.44

Al2O3 15.31 14.68 15.18 14.86 14.63 15.27 14.06 15.06 14.69 16.68 14.5 14.7 T Fe2O3 15.02 14.26 14.53 14.14 14.84 15.32 12.13 12.32 14.75 11.85 13.61 11.2

Fe2O3 2.25 2.14 2.18 2.12 2.23 2.3 1.82 1.85 2.21 1.78 2.04 1.68 FeO 11.49 10.91 11.11 10.82 11.35 11.72 9.28 9.42 11.28 9.06 10.41 8.57 MnO 0.2 0.24 0.29 0.17 0.21 0.2 0.18 0.15 0.18 0.16 0.18 0.16 MgO 4.67 4.9 4.33 3.63 4.64 4.51 7.09 6.08 5.21 5.46 5.3 6.21 CaO 9.53 9.59 8.82 9.98 9.97 9.26 11.69 11.92 10.5 11.58 10.24 11.82

Na2O 2.93 2.46 2.94 2.99 2.84 2.88 2.91 2.91 2.78 2.94 2.84 2.98

K2O 0.89 0.9 0.88 1.44 0.83 0.97 0.33 0.53 0.61 0.47 0.64 0.43

P2O5 0.15 0.19 0.16 0.11 0.15 0.18 0.15 0.25 0.17 0.14 0.19 0.09 LOI 1.25 1.3 1.15 1.6 1.2 1.15 1.2 1 1.12 0.88 1 1.1 Total 101.25 101.12 101.15 101.59 101.2 101.15 101.19 101.02 101.13 100.89 101 101.11 Y 39.38 73.43 50.3 36.41 49.15 46.4 36.65 38.9 44.4 31.94 49.68 32.55 Sc 48.48 50.93 47.83 47.89 52.76 48.32 40.18 39.17 48.61 38.83 51.07 46.91 Nb 1.65 3.219 1.96 1.002 2.282 1.86 3.64 1.7 1.91 0.67 1.72 1.17 Th 0.19 0.254 0.17 0.113 0.199 0.28 0.27 0.11 0.21 0.03 0.13 0.17 Sr 131 143 141 156 142 144 141 159 121 148 139 151 Rb 13.28 18.63 22.87 13.86 13.82 15.82 7.13 7.65 12.57 6.59 14.79 9.73 Ba 50 77 78 81 71 235 26 15 37 15.21 45.62 43.30 Co 44.2 51.08 71.52 42.38 46.94 42.83 40.78 33.82 41.3 45.07 42.8 45.19 Cu 97.77 158.33 142.93 194.85 156.92 85.25 90.91 142.84 79.67 129.09 79.11 126.51 V 325 396 255 264 357 317 298 296 331 267 299 274 Ga 18.13 17.91 19.59 15.72 16.84 22.94 14.62 14.65 16.62 14.97 17.56 14.06 Zn 80.84 200.58 83.63 144.48 164.85 95.64 64.06 71.96 81.21 101.55 85.11 121.76 Zr 125.85 175.90 136.99 86.58 130.08 114.67 111.1 84.25 102.7 73.35 110.9 93.23 U 0.406 0.718 0.826 0.78 0.687 0.471 0.467 0.483 0.580 0.442 0.802 0.277 Pb 2.360 5.617 3.455 5.709 6.067 3.088 3.107 6.484 1.355 3.119 1.923 4.585

La 5.27 8.984 9.54 7.056 6.721 4.54 6.45 9.31 4.14 7.42 4.85 3.63 Ce 11.29 16.75 16.02 7.52 12.01 12.45 13.44 10.05 11.07 7.34 13.16 8.53 Pr 2.26 2.865 3.19 1.646 2.125 2.34 2.51 2.83 2.15 2.16 2.47 1.42 Nd 11.78 19.63 14.18 11.04 14.59 11.12 11.08 13.82 10.24 10.49 13.17 9.86 Sm 4.12 5.95 5.22 3.36 4.54 4.44 3.78 3.86 4.06 3.32 4.61 3.12 Eu 1.44 1.995 1.74 1.34 1.61 1.58 1.32 1.33 1.42 1.22 1.62 1.18 Gd 5.31 9.39 6.75 5.51 6.88 5.83 4.94 5.04 5.43 4.26 6.04 4.82 Tb 1.02 1.77 1.23 1.01 1.34 1.08 0.88 0.88 1.02 0.78 1.14 0.898 Dy 6.73 10.40 8.13 5.63 7.70 7.27 5.84 5.73 6.78 5.17 7.63 5.11 Ho 1.49 2.25 1.8 1.23 1.65 1.64 1.3 1.29 1.52 1.16 1.73 1.093 Er 3.93 7.42 4.75 3.79 5.23 4.38 3.4 3.38 4.04 3.02 4.65 3.46 Tm 0.62 1.22 0.74 0.62 0.88 0.7 0.53 0.52 0.65 0.46 0.72 0.594 Yb 4.16 6.74 4.81 3.33 4.67 4.63 3.45 3.44 4.19 3.1 4.89 3.08 Lu 0.64 1.01 0.75 0.50 0.74 0.73 0.54 0.54 0.68 0.47 0.73 0.484

Page | 38

PD19 PD20 PD21 PD22 PD25 PD27 PD28 PD30 PD31 PD32 PD35 HRX2

SiO2 48.08 51.13 47.98 49.1 50.57 49.91 50.49 49.28 50.87 49.79 50.59 48.57

TiO2 2.98 2.34 4 1.97 1.95 1.44 1.55 2.27 1.65 1.59 2.07 1.66

Al2O3 13.79 13.94 13.15 14.61 13.67 14.84 18.25 15.09 14.28 14.49 13.82 17.23 T Fe2O3 18.2 13.73 19.39 15.07 12.8 12.02 13.05 15.46 12.65 12.47 12.87 11.12

Fe2O3 2.73 2.06 2.91 2.26 1.92 1.8 1.96 2.32 1.9 1.87 1.93 1.87 FeO 13.92 10.5 14.83 11.53 9.79 9.2 9.98 11.82 9.67 9.54 9.84 9.54 MnO 0.24 0.2 0.23 0.19 0.21 0.17 0.16 0.19 0.22 0.19 0.19 0.12 MgO 3.86 5.06 3.42 5.17 6.22 5.89 2.81 4.47 6.52 6.39 5.92 5.68 CaO 9.03 9.95 7.44 10.07 11.33 11.94 9.03 9.22 10.55 11.52 11.17 10.79

Na2O 2.59 2.68 2.88 2.83 2.65 3.12 3.08 2.91 2.78 2.85 2.54 3.03

K2O 0.97 0.78 0.98 0.84 0.45 0.53 0.97 0.95 0.38 0.53 0.63 0.46

P2O5 0.25 0.19 0.31 0.15 0.14 0.15 0.19 0.16 0.12 0.16 0.21 nd LOI 1.3 1.1 1.5 1.15 1 1 1.3 1.25 1 1 1.5 1.06 Total 101.29 101.1 101.28 101.15 100.99 101.01 100.88 101.25 101.02 100.98 101.51 101.07 Y 73.43 52.13 84.68 39.32 46.76 34.03 35.27 48.87 33.33 40.57 49.92 57.38 Sc 50.93 53.84 51.29 42.55 45.76 45.74 38.73 50.51 43.32 42.59 50.28 48.25 Nb 3.219 2.38 4.13 1.30 2.52 0.75 1.41 0.82 1.87 1.36 2.5 3.48 Th 0.25 0.24 0.33 0.10 0.24 0.04 0.23 0.35 0.21 0.28 0.206 0.8 Sr 143 133 151 116 136 147 174 141 134 136 134 154 Rb 18.63 21.10 22.97 10.01 10.73 10.98 14.71 14.77 7.71 10.24 13.65 22.41 Ba 77.60 70.53 57.95 41.23 67.48 17.54 62.38 44.19 24.09 21.78 58.19 61.26 Co 51.08 40.82 42.28 36.91 50.39 45.26 44.59 38.65 50.21 43.46 40.16 31.93 Cu 158.33 120.60 172.60 64.36 102.65 127.52 221.38 137.79 151.02 130.98 132.32 147.18 V 396 288 391 294 323 274 204 303 327 290 357 396 Ga 17.91 17.29 17.71 14.22 15.71 15.24 14.54 17.06 13.98 15.63 15.58 17.86 Zn 200.59 164.08 246.46 70.90 136.86 53.05 127.33 101.24 147.06 69.15 147.28 195.51 Zr 175.89 133.91 225.57 90.71 124.44 88.38 97.24 136.37 106.39 95.86 113.35 145.14 U 0.718 1.404 0.882 0.526 1.178 0.542 0.821 0.771 0.244 0.339 0.542 0.635 Pb 5.617 9.618 6.935 4.470 5.981 1.587 3.816 2.901 2.599 2.954 4.702 6.046

La 8.98 6.56 17.32 3.54 5.05 4.39 13.34 7.28 5.08 8.24 9.87 10.95 Ce 16.75 12.84 21.32 9.97 12.56 8.19 8.55 12.21 10.32 10.45 11.5 18.33 Pr 2.87 2.28 4.16 1.95 2.04 1.93 2.70 2.72 1.70 2.52 2.73 3.30 Nd 19.63 15.34 27.24 9.23 13.71 9.87 16.45 13.95 11.17 12.28 17.36 20 Sm 5.95 4.65 7.75 3.75 4.19 3.28 4.18 4.75 3.35 3.83 4.75 5.69 Eu 1.99 1.64 2.47 1.3 1.5 1.24 1.46 1.65 1.29 1.38 1.65 1.89 Gd 9.39 7.31 11.91 5.04 6.54 4.36 6.04 6.17 5.23 5.1 7.38 8.37 Tb 1.77 1.393 2.222 0.92 1.252 0.8 1.059 1.16 0.97 0.93 1.37 1.58 Dy 10.40 7.95 12.54 6.21 7.07 5.37 5.82 7.74 5.36 6.15 7.71 8.55 Ho 2.248 1.708 2.763 1.41 1.512 1.18 1.19 1.73 1.14 1.39 1.65 1.84 Er 7.42 5.57 8.82 3.79 4.84 3.09 3.64 4.58 3.64 3.71 5.13 5.79 Tm 1.216 0.934 1.483 0.58 0.812 0.48 0.591 0.71 0.611 0.58 0.86 0.99 Yb 6.74 4.95 7.80 3.92 4.39 3.18 3.08 4.81 3.28 3.8 4.63 5.32 Lu 1.013 0.76 1.2 0.61 0.692 0.48 0.463 0.71 0.5 0.57 0.70 0.80

Page | 39

HRX3 HRX5 HRX8 HRX9 HRX13 HRX18 HRX20 HRX21 HRX22 HRX28 HRX29

SiO2 48.26 49.93 48.34 48.88 50.22 48.74 48.64 46.7 49.05 48.94 48.48

TiO2 2.62 2.13 1.63 1.69 1.46 2.34 1.75 2.31 1.72 1.7 2.21

Al2O3 15.59 15.93 16.79 15.3 17.65 17.36 16.14 15.49 16.41 16.44 18.11 T Fe2O3 12.28 11.36 12.21 12.14 11.08 12.08 12.29 17.75 11.12 13.48 11.96

Fe2O3 1.84 1.7 1.83 1.82 1.66 1.81 1.84 2.66 1.67 2.02 1.79 FeO 9.39 8.69 9.34 9.29 8.48 9.24 9.4 13.58 8.5 10.31 9.15 MnO 0.18 0.17 1.41 0.18 0.1 0.18 0.2 0.35 0.18 0.18 0.15 MgO 7.14 6.19 6.05 5.31 5.47 4.49 6.52 4.5 6.63 4.51 6.34 CaO 9.96 9.64 10.39 10.88 10.18 10.8 10.8 8.38 9.99 10.87 9.25

Na2O 3.62 4.26 2.93 3.06 3.46 3.49 3.15 2.75 3.5 3.5 3.24

K2O 0.35 0.38 0.24 0.41 0.37 0.51 0.52 0.93 0.38 0.38 0.25

P2O5 nd nd nd 0.22 nd nd nd 0.17 0.23 nd nd LOI 1.1 nd 1.3 2.2 0.98 nd 1 1.8 1.5 0.85 0.88 Total 101.1 99.99 101.29 100.27 100.97 99.99 101.01 101.13 100.71 100.85 100.87 Y 57.24 27.24 29.76 4.20 22.62 86.30 42.08 45.39 18.02 46.82 34.40 Sc 46.63 41.19 36.48 38.69 42.69 47.37 41.47 48.78 30.20 46.23 43.92 Nb 3.28 2.28 1.30 2.05 3.87 4.69 1.84 2.62 1.41 2.95 2.50 Th 0.519 0.849 0.201 1.321 2.35 0.90 0.76 0.51 1.02 0.58 1.11 Sr 138 153 108 41 132 150 121 151 111 141 150 Rb 21.24 14.47 0.88 74.42 3.15 20.94 9.55 20.36 1.76 13.37 14.44 Ba 55.50 38.91 16.67 95.90 51.72 58.77 32.54 107.21 32.21 44.05 44.99 Co 51.37 47.31 49.41 63.48 107.72 58.08 64.27 59.67 43.05 41.08 67.50 Cu 247.09 172.88 100.51 284.28 133.06 186.47 206.69 150.36 99.63 115.87 203.24 V 397 266 244 45 247 470 297 309 199 342 285 Ga 17.01 13.90 11.95 10.86 11.74 19.03 14.68 16.02 11.27 15.55 14.05 Zn 183.04 125.99 94.59 148.95 73.38 248.33 111.97 159.87 61.80 144.19 114.06 Zr 156.55 112.33 81.02 83.81 36.46 228.78 77.17 129.12 48.21 128.35 110.54 U 0.512 0.448 0.082 0.765 0.285 0.707 0.42 0.639 0.124 1.258 0.355 Pb 4.984 6.147 2.797 5.354 10.846 4.329 3.357 7.422 2.948 3.226 5.259

La 11.54 5.78 2.32 4.84 7.15 16.31 7.99 9.23 1.96 6.27 6.76 Ce 16.97 14.30 8.11 17.88 44.85 27.80 9.03 16.51 6.08 14.62 19.28 Pr 3.07 1.68 1.18 0.97 1.9 4.51 2.09 2.47 0.75 2.25 2.06 Nd 19.29 10.77 8.41 5.05 10.92 29.21 13.10 15.74 5.00 14.95 12.95 Sm 5.46 2.99 2.72 1.11 2.77 8.25 3.69 4.58 1.62 4.36 3.53 Eu 1.81 1.15 1.03 0.33 1.02 2.57 1.27 1.58 0.64 1.55 1.28 Gd 8.19 4.45 4.27 1.44 4.41 12.13 5.53 6.85 2.40 6.75 5.36 Tb 1.51 0.81 0.80 0.23 0.72 2.27 1.05 1.302 0.47 1.25 0.94 Dy 8.57 4.41 4.58 1.12 3.99 12.85 5.84 7.36 2.66 7.06 5.30 Ho 1.83 0.93 0.98 0.22 0.83 2.75 1.27 1.55 0.60 1.55 1.12 Er 5.83 2.95 3.11 0.66 2.57 8.75 4.08 4.90 1.90 4.87 3.54 Tm 0.97 0.50 0.54 0.12 0.42 1.47 0.68 0.82 0.33 0.82 0.58 Yb 5.25 2.68 2.87 0.73 2.31 7.72 3.59 4.41 1.69 4.32 3.17 Lu 0.78 0.41 0.44 0.11 0.36 1.17 0.54 0.68 0.26 0.67 0.47

Page | 40

HRX30 HRX31 HRX34 HRX35 HRX36 HRX37 SS13- AAS22 SS11 TS DR#19 90/91

SiO2 47.96 49.47 49.5 49.54 49.01 48.78 50.35 50.68 50.99

TiO2 1.53 2.53 2.02 1.64 2.15 2.12 1.91 2.17 1.27

Al2O3 16.56 17.1 18.88 16.27 17.5 17.23 16.07 13.62 13.54 T Fe2O3 11.13 10.82 10.92 11.15 12.59 12.7 13.77 12.34 11.67

Fe2O3 1.67 1.62 1.64 1.67 1.89 1.91 2.07 1.85 1.75 FeO 8.51 8.28 8.35 8.53 9.63 9.71 10.53 9.44 8.93 MnO 0.18 0.1 0.65 0.18 0.15 0.12 0.2 0.29 0.17 MgO 6.94 6.17 3.64 6.45 5.12 5.64 4.14 6.06 6.65 CaO 11.74 10.76 9.58 10.89 10.12 9.57 9.47 11.24 12.41

Na2O 3.5 2.63 3.39 3.51 2.95 3.46 2.9 2.72 2.87

K2O 0.46 0.4 0.95 0.36 0.41 0.37 0.9 0.73 0.34

P2O5 nd nd nd nd nd nd 0.18 0.17 0.09 LOI 0.68 1.2 1.1 0.65 1 1.2 1.15 1.6 0.99 Total 100.68 101.18 100.63 100.64 101 101.19 101.04 101.62 100.99 Y 41.83 31.47 21.33 40.43 53.26 52.65 47.86 50.18 48.23 Sc 42.53 36.95 36.45 40.94 46.70 48.97 49.39 47.03 51.11 Nb 2.97 1.32 4.34 5.32 3.09 2.80 3.04 2.66 1.22 Th 0.68 0.24 3.10 1.56 0.81 0.76 0.40 0.44 0.12 Sr 160 143 139 149 152 137 145 121 121 Rb 13.13 7.51 13.69 14.33 18.51 22.08 20.23 12.71 9.00 Ba 40.33 23.67 80.57 71.38 76.70 64.12 71.43 69.73 70.11 Co 36.67 58.59 115.55 74.15 60.11 73.63 47.04 47.41 44.99 Cu 133.10 179.03 319.27 249.00 221.81 143.56 112.95 122.41 120.22 V 307 246 225 336 357 310 225 358 270 Ga 14.80 12.85 16.04 15.87 17.43 17.39 15.75 16.14 15.77 Zn 133.14 133.99 113.15 159.09 167.09 164.78 147.01 150.64 167.25 Zr 120.90 85.38 98.95 134.86 139.18 134.23 111.32 134.92 99.59 U 0.609 0.24 0.707 0.444 0.685 1.009 0.656 0.623 0.666 Pb 3.645 4.886 9.659 7.347 5.228 3.607 5.072 11.069 8.125

La 8.28 6.54 12.40 8.87 10.61 9.10 9.15 8.48 7.31 Ce 15.28 10.23 29.96 25.12 15.36 18.15 12.63 14.07 12.02 Pr 2.47 1.87 2.84 2.58 3.08 2.83 2.17 2.66 1.82 Nd 15.47 11.58 15.81 15.63 19.34 17.93 14.02 16.97 12.56 Sm 4.29 3.14 3.68 4.32 5.51 5.16 4.05 4.91 3.27 Eu 1.49 1.15 1.21 1.47 1.85 1.73 1.45 1.69 1.21 Gd 6.34 4.64 5.0 6.29 8.02 7.71 6.32 7.53 6.22 Tb 1.16 0.84 0.81 1.14 1.50 1.42 1.22 1.41 1.32 Dy 6.42 4.79 4.11 6.26 8.34 7.99 6.82 7.73 6.98 Ho 1.36 1.03 0.82 1.33 1.76 1.73 1.51 1.68 1.09 Er 4.32 3.21 2.50 4.25 5.63 5.54 4.84 5.30 4.29 Tm 0.71 0.53 0.40 0.70 0.93 0.93 0.79 0.89 0.71 Yb 3.81 2.86 2.20 3.76 4.86 4.85 4.32 4.74 3.93 Lu 0.58 0.44 0.33 0.59 0.72 0.74 0.66 0.73 0.67 nd - not determined all Fe analysed as Fe2O3, FeO has been calculated considering 85% as FeO