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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
1
NATIONAL INSTITUTE OF OCEANOGRAPHY (COUNCIL OF SCIENTIFIC AND INDUSTRIAL
RESEARCH) DONA PAULA, GOA – 403004 INDIA
2
SCHOOL OF OCEANOGRAPHIC STUDIES, JADAVPUR UNIVERSITY
KOLKATA – 700 032 INDIA
ABSTRACT
The Central Indian Ocean Basin (CIOB) basalts are plagioclase rich while olivine and pyroxene
are very few. The analyses of forty five samples reveal high FeOT (~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 NMORB (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
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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 (NMORB) while spilites occur sporadically (Karisiddaiah and Iyer, 1992). In the Indian Ocean,
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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 JB2 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.
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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
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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.
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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 Timagnetite 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
140 10-8m3kg-1) 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).
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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)
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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
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accord with the evolved major element chemistry, the CIOB basalts are relatively enriched in Zr (48228 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; ~25166; ~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 NMORB yet, faint signatures of P-type MORB are noticeable in the mixing relation between N- and Ptypes 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).
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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
P a g e | 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.
In the CIOB basalts, FeOT and TiO2 steadily increase with decreasing MgO for the range of NMORB 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
P a g e | 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
P a g e | 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
P a g e | 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 longlived 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 longterm 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
P a g e | 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 highlevel 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 LREEenriched 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-
CaO/Al2O3, K- P2O5, FeOT- FeOT/MgO,TiO2-Zr, FeOT-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
slope; 2), MgO-Al2O3, FeOT-MgO, Na2O-MgO, K2O-MgO, Ba/Nb, and (La/Sm)N-Zr/Nb
P a g e | 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 #???
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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.
P a g e | 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.
P a g e | 23
Figure 2 Representative photographs of the hand specimen.
P a g e | 24
P a g e | 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.
P a g e | 26
Figure 4
P a g e | 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.
(c)–(d) TiO2 and FeOT 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).
P a g e | 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.
P a g e | 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.
P a g e | 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.
P a g e | 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).
P a g e | 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
P a g e | 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.
P a g e | 34
Figure 11
Both (a) FeOT 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.
P a g e | 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.
P a g e | 36
Table 1: Sample location of basalts from the Central Indian Ocean Basin
Sample
Latitude (°S)
Longitude (°E)
Water
Depth (m)
13° 00’ 00”
11° 00’ 32”
12° 24’ 0”
13° 00’ 26”
11° 30’ 45”
12° 33’ 25”
13° 00’ 26”
12° 45’ 00”
12° 57’ 30”
12° 23’ 42”
11° 59’ 00”
12° 59’ 50”
11° 40’ 19”
14° 05’ 00”
11° 00’ 00”
13° 20’ 42”
11° 15’ 22”
13° 45’ 24”
13° 00’ 26”
12° 59’ 00”
13° 7’ 02”
12° 00’ 00”
12° 46’ 08”
12° 33’ 22”
11° 04’ 02”
13° 10’ 00”
13° 00’ 00”
75° 59’ 00”
73° 59’ 12”
76° 14’ 30”
76° 00’ 25”
76° 15’ 30”
78° 45’ 00”
76° 00’ 25”
74° 45’ 24”
75° 00’ 06”
76° 13’ 30”
76° 50’ 57”
76° 00’ 08”
77° 47’ 16”
76° 18’ 00”
77° 55’ 24”
77° 30’ 00”
73° 15’ 46”
75° 45’ 20”
76° 00’ 25”
76° 30’ 00”
77° 44’ 31”
77° 55’ 24”
78° 23’ 13”
77° 12’ 54”
77° 52’ 00”
75° 59’ 00”
76° 00’ 00”
5374
3600
4400
5374
4830
4550
4732
5348
5350
5402
5210
5190
5300
5050
4972
5320
5370
5400
5380
5440
5250
5210
5300
5330
5374
HRX21
11° 00’ 00”
11° 59’ 00”
12° 00’ 00”
77° 55’ 24”
76° 30’ 57”
76° 30’ 00”
5300
5402
5400
HRX28
HRX29
HRX30 (1)
HRX31
HRX34
HRX35
HRX36
HRX37
SS13-TS 90/91
AAS22 DR#19
SS11
11° 15’ 22”
12° 28’ 04”
14° 44’ 49”
13° 00’ 00”
12° 30’ 00”
9° 59’ 00”
12° 00’ 00”
12° 59’ 55”
13° 40’ 00”
14° 20’ 55”
12° 23’ 20”
73° 14’ 46”
76° 52’ 01”
77° 59’ 20”
75° 59’ 00”
78° 12’ 00”
79° 15’ 00”
76° 30’ 00”
76° 30’ 00”
79° 30’ 00”
76° 33’ 25”
76° 13’ 30”
4972
5083
5374
5120
5212
5400
5400
5040
5125
5325
PD1
PD2
PD3
PD5
PD6
PD7
PD8
PD10
PD11
PD14
PD16
PD18
PD19
PD20
PD21
PD22
PD25
PD27
PD28
PD30
PD31
PD32
PD35
HRX2
HRX3
HRX5
HRX9
HRX18
HRX20
P a g e | 37
Table 2: Major, trace and REE analysis of CIOB basalts
PD1
PD2
PD3
PD5
PD6
PD7
PD8
PD10
PD11
PD14
PD16
PD18
Total
Y
Sc
Nb
Th
Sr
Rb
Ba
Co
Cu
V
Ga
Zn
Zr
U
Pb
49.26
2.04
15.31
15.02
2.25
11.49
0.2
4.67
9.53
2.93
0.89
0.15
1.25
101.25
39.38
48.48
1.65
0.19
131
13.28
50
44.2
97.77
325
18.13
80.84
125.85
0.406
2.360
50.7
1.9
14.68
14.26
2.14
10.91
0.24
4.9
9.59
2.46
0.9
0.19
1.3
101.12
73.43
50.93
3.219
0.254
143
18.63
77
51.08
158.33
396
17.91
200.58
175.90
0.718
5.617
50.37
2.5
15.18
14.53
2.18
11.11
0.29
4.33
8.82
2.94
0.88
0.16
1.15
101.15
50.3
47.83
1.96
0.17
141
22.87
78
71.52
142.93
255
19.59
83.63
136.99
0.826
3.455
51.08
1.59
14.86
14.14
2.12
10.82
0.17
3.63
9.98
2.99
1.44
0.11
1.6
101.59
36.41
47.89
1.002
0.113
156
13.86
81
42.38
194.85
264
15.72
144.48
86.58
0.78
5.709
49.67
2.22
14.63
14.84
2.23
11.35
0.21
4.64
9.97
2.84
0.83
0.15
1.2
101.2
49.15
52.76
2.282
0.199
142
13.82
71
46.94
156.92
357
16.84
164.85
130.08
0.687
6.067
49.24
2.17
15.27
15.32
2.3
11.72
0.2
4.51
9.26
2.88
0.97
0.18
1.15
101.15
46.4
48.32
1.86
0.28
144
15.82
235
42.83
85.25
317
22.94
95.64
114.67
0.471
3.088
49.84
1.61
14.06
12.13
1.82
9.28
0.18
7.09
11.69
2.91
0.33
0.15
1.2
101.19
36.65
40.18
3.64
0.27
141
7.13
26
40.78
90.91
298
14.62
64.06
111.1
0.467
3.107
49.32
1.48
15.06
12.32
1.85
9.42
0.15
6.08
11.92
2.91
0.53
0.25
1
101.02
38.9
39.17
1.7
0.11
159
7.65
15
33.82
142.84
296
14.65
71.96
84.25
0.483
6.484
49.14
1.98
14.69
14.75
2.21
11.28
0.18
5.21
10.5
2.78
0.61
0.17
1.12
101.13
44.4
48.61
1.91
0.21
121
12.57
37
41.3
79.67
331
16.62
81.21
102.7
0.580
1.355
49.38
1.35
16.68
11.85
1.78
9.06
0.16
5.46
11.58
2.94
0.47
0.14
0.88
100.89
31.94
38.83
0.67
0.03
148
6.59
15.21
45.07
129.09
267
14.97
101.55
73.35
0.442
3.119
50.32
2.18
14.5
13.61
2.04
10.41
0.18
5.3
10.24
2.84
0.64
0.19
1
101
49.68
51.07
1.72
0.13
139
14.79
45.62
42.8
79.11
299
17.56
85.11
110.9
0.802
1.923
50.98
1.44
14.7
11.2
1.68
8.57
0.16
6.21
11.82
2.98
0.43
0.09
1.1
101.11
32.55
46.91
1.17
0.17
151
9.73
43.30
45.19
126.51
274
14.06
121.76
93.23
0.277
4.585
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
5.27
11.29
2.26
11.78
4.12
1.44
5.31
1.02
6.73
1.49
3.93
0.62
4.16
0.64
8.984
16.75
2.865
19.63
5.95
1.995
9.39
1.77
10.40
2.25
7.42
1.22
6.74
1.01
9.54
16.02
3.19
14.18
5.22
1.74
6.75
1.23
8.13
1.8
4.75
0.74
4.81
0.75
7.056
7.52
1.646
11.04
3.36
1.34
5.51
1.01
5.63
1.23
3.79
0.62
3.33
0.50
6.721
12.01
2.125
14.59
4.54
1.61
6.88
1.34
7.70
1.65
5.23
0.88
4.67
0.74
4.54
12.45
2.34
11.12
4.44
1.58
5.83
1.08
7.27
1.64
4.38
0.7
4.63
0.73
6.45
13.44
2.51
11.08
3.78
1.32
4.94
0.88
5.84
1.3
3.4
0.53
3.45
0.54
9.31
10.05
2.83
13.82
3.86
1.33
5.04
0.88
5.73
1.29
3.38
0.52
3.44
0.54
4.14
11.07
2.15
10.24
4.06
1.42
5.43
1.02
6.78
1.52
4.04
0.65
4.19
0.68
7.42
7.34
2.16
10.49
3.32
1.22
4.26
0.78
5.17
1.16
3.02
0.46
3.1
0.47
4.85
13.16
2.47
13.17
4.61
1.62
6.04
1.14
7.63
1.73
4.65
0.72
4.89
0.73
3.63
8.53
1.42
9.86
3.12
1.18
4.82
0.898
5.11
1.093
3.46
0.594
3.08
0.484
SiO2
TiO2
Al2O3
T
Fe2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P 2O 5
LOI
P a g e | 38
PD19
PD20
PD21
PD22
PD25
PD27
PD28
PD30
PD31
PD32
PD35
HRX2
SiO2
TiO2
Al2O3
T
Fe2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P 2O 5
LOI
Total
Y
Sc
Nb
Th
Sr
Rb
Ba
Co
Cu
V
Ga
Zn
Zr
U
Pb
48.08
2.98
13.79
18.2
2.73
13.92
0.24
3.86
9.03
2.59
0.97
0.25
1.3
101.29
73.43
50.93
3.219
0.25
143
18.63
77.60
51.08
158.33
396
17.91
200.59
175.89
0.718
5.617
51.13
2.34
13.94
13.73
2.06
10.5
0.2
5.06
9.95
2.68
0.78
0.19
1.1
101.1
52.13
53.84
2.38
0.24
133
21.10
70.53
40.82
120.60
288
17.29
164.08
133.91
1.404
9.618
47.98
4
13.15
19.39
2.91
14.83
0.23
3.42
7.44
2.88
0.98
0.31
1.5
101.28
84.68
51.29
4.13
0.33
151
22.97
57.95
42.28
172.60
391
17.71
246.46
225.57
0.882
6.935
49.1
1.97
14.61
15.07
2.26
11.53
0.19
5.17
10.07
2.83
0.84
0.15
1.15
101.15
39.32
42.55
1.30
0.10
116
10.01
41.23
36.91
64.36
294
14.22
70.90
90.71
0.526
4.470
50.57
1.95
13.67
12.8
1.92
9.79
0.21
6.22
11.33
2.65
0.45
0.14
1
100.99
46.76
45.76
2.52
0.24
136
10.73
67.48
50.39
102.65
323
15.71
136.86
124.44
1.178
5.981
49.91
1.44
14.84
12.02
1.8
9.2
0.17
5.89
11.94
3.12
0.53
0.15
1
101.01
34.03
45.74
0.75
0.04
147
10.98
17.54
45.26
127.52
274
15.24
53.05
88.38
0.542
1.587
50.49
1.55
18.25
13.05
1.96
9.98
0.16
2.81
9.03
3.08
0.97
0.19
1.3
100.88
35.27
38.73
1.41
0.23
174
14.71
62.38
44.59
221.38
204
14.54
127.33
97.24
0.821
3.816
49.28
2.27
15.09
15.46
2.32
11.82
0.19
4.47
9.22
2.91
0.95
0.16
1.25
101.25
48.87
50.51
0.82
0.35
141
14.77
44.19
38.65
137.79
303
17.06
101.24
136.37
0.771
2.901
50.87
1.65
14.28
12.65
1.9
9.67
0.22
6.52
10.55
2.78
0.38
0.12
1
101.02
33.33
43.32
1.87
0.21
134
7.71
24.09
50.21
151.02
327
13.98
147.06
106.39
0.244
2.599
49.79
1.59
14.49
12.47
1.87
9.54
0.19
6.39
11.52
2.85
0.53
0.16
1
100.98
40.57
42.59
1.36
0.28
136
10.24
21.78
43.46
130.98
290
15.63
69.15
95.86
0.339
2.954
50.59
2.07
13.82
12.87
1.93
9.84
0.19
5.92
11.17
2.54
0.63
0.21
1.5
101.51
49.92
50.28
2.5
0.206
134
13.65
58.19
40.16
132.32
357
15.58
147.28
113.35
0.542
4.702
48.57
1.66
17.23
11.12
1.87
9.54
0.12
5.68
10.79
3.03
0.46
nd
1.06
101.07
57.38
48.25
3.48
0.8
154
22.41
61.26
31.93
147.18
396
17.86
195.51
145.14
0.635
6.046
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
8.98
16.75
2.87
19.63
5.95
1.99
9.39
1.77
10.40
2.248
7.42
1.216
6.74
1.013
6.56
12.84
2.28
15.34
4.65
1.64
7.31
1.393
7.95
1.708
5.57
0.934
4.95
0.76
17.32
21.32
4.16
27.24
7.75
2.47
11.91
2.222
12.54
2.763
8.82
1.483
7.80
1.2
3.54
9.97
1.95
9.23
3.75
1.3
5.04
0.92
6.21
1.41
3.79
0.58
3.92
0.61
5.05
12.56
2.04
13.71
4.19
1.5
6.54
1.252
7.07
1.512
4.84
0.812
4.39
0.692
4.39
8.19
1.93
9.87
3.28
1.24
4.36
0.8
5.37
1.18
3.09
0.48
3.18
0.48
13.34
8.55
2.70
16.45
4.18
1.46
6.04
1.059
5.82
1.19
3.64
0.591
3.08
0.463
7.28
12.21
2.72
13.95
4.75
1.65
6.17
1.16
7.74
1.73
4.58
0.71
4.81
0.71
5.08
10.32
1.70
11.17
3.35
1.29
5.23
0.97
5.36
1.14
3.64
0.611
3.28
0.5
8.24
10.45
2.52
12.28
3.83
1.38
5.1
0.93
6.15
1.39
3.71
0.58
3.8
0.57
9.87
11.5
2.73
17.36
4.75
1.65
7.38
1.37
7.71
1.65
5.13
0.86
4.63
0.70
10.95
18.33
3.30
20
5.69
1.89
8.37
1.58
8.55
1.84
5.79
0.99
5.32
0.80
P a g e | 39
HRX3
HRX5
HRX8
HRX9
HRX13
HRX18
HRX20
HRX21
HRX22
HRX28
HRX29
SiO2
TiO2
Al2O3
T
Fe2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P 2O 5
LOI
Total
Y
Sc
Nb
Th
Sr
Rb
Ba
Co
Cu
V
Ga
Zn
Zr
U
Pb
48.26
2.62
15.59
12.28
1.84
9.39
0.18
7.14
9.96
3.62
0.35
nd
1.1
101.1
57.24
46.63
3.28
0.519
138
21.24
55.50
51.37
247.09
397
17.01
183.04
156.55
0.512
4.984
49.93
2.13
15.93
11.36
1.7
8.69
0.17
6.19
9.64
4.26
0.38
nd
nd
99.99
27.24
41.19
2.28
0.849
153
14.47
38.91
47.31
172.88
266
13.90
125.99
112.33
0.448
6.147
48.34
1.63
16.79
12.21
1.83
9.34
1.41
6.05
10.39
2.93
0.24
nd
1.3
101.29
29.76
36.48
1.30
0.201
108
0.88
16.67
49.41
100.51
244
11.95
94.59
81.02
0.082
2.797
48.88
1.69
15.3
12.14
1.82
9.29
0.18
5.31
10.88
3.06
0.41
0.22
2.2
100.27
4.20
38.69
2.05
1.321
41
74.42
95.90
63.48
284.28
45
10.86
148.95
83.81
0.765
5.354
50.22
1.46
17.65
11.08
1.66
8.48
0.1
5.47
10.18
3.46
0.37
nd
0.98
100.97
22.62
42.69
3.87
2.35
132
3.15
51.72
107.72
133.06
247
11.74
73.38
36.46
0.285
10.846
48.74
2.34
17.36
12.08
1.81
9.24
0.18
4.49
10.8
3.49
0.51
nd
nd
99.99
86.30
47.37
4.69
0.90
150
20.94
58.77
58.08
186.47
470
19.03
248.33
228.78
0.707
4.329
48.64
1.75
16.14
12.29
1.84
9.4
0.2
6.52
10.8
3.15
0.52
nd
1
101.01
42.08
41.47
1.84
0.76
121
9.55
32.54
64.27
206.69
297
14.68
111.97
77.17
0.42
3.357
46.7
2.31
15.49
17.75
2.66
13.58
0.35
4.5
8.38
2.75
0.93
0.17
1.8
101.13
45.39
48.78
2.62
0.51
151
20.36
107.21
59.67
150.36
309
16.02
159.87
129.12
0.639
7.422
49.05
1.72
16.41
11.12
1.67
8.5
0.18
6.63
9.99
3.5
0.38
0.23
1.5
100.71
18.02
30.20
1.41
1.02
111
1.76
32.21
43.05
99.63
199
11.27
61.80
48.21
0.124
2.948
48.94
1.7
16.44
13.48
2.02
10.31
0.18
4.51
10.87
3.5
0.38
nd
0.85
100.85
46.82
46.23
2.95
0.58
141
13.37
44.05
41.08
115.87
342
15.55
144.19
128.35
1.258
3.226
48.48
2.21
18.11
11.96
1.79
9.15
0.15
6.34
9.25
3.24
0.25
nd
0.88
100.87
34.40
43.92
2.50
1.11
150
14.44
44.99
67.50
203.24
285
14.05
114.06
110.54
0.355
5.259
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
11.54
16.97
3.07
19.29
5.46
1.81
8.19
1.51
8.57
1.83
5.83
0.97
5.25
0.78
5.78
14.30
1.68
10.77
2.99
1.15
4.45
0.81
4.41
0.93
2.95
0.50
2.68
0.41
2.32
8.11
1.18
8.41
2.72
1.03
4.27
0.80
4.58
0.98
3.11
0.54
2.87
0.44
4.84
17.88
0.97
5.05
1.11
0.33
1.44
0.23
1.12
0.22
0.66
0.12
0.73
0.11
7.15
44.85
1.9
10.92
2.77
1.02
4.41
0.72
3.99
0.83
2.57
0.42
2.31
0.36
16.31
27.80
4.51
29.21
8.25
2.57
12.13
2.27
12.85
2.75
8.75
1.47
7.72
1.17
7.99
9.03
2.09
13.10
3.69
1.27
5.53
1.05
5.84
1.27
4.08
0.68
3.59
0.54
9.23
16.51
2.47
15.74
4.58
1.58
6.85
1.302
7.36
1.55
4.90
0.82
4.41
0.68
1.96
6.08
0.75
5.00
1.62
0.64
2.40
0.47
2.66
0.60
1.90
0.33
1.69
0.26
6.27
14.62
2.25
14.95
4.36
1.55
6.75
1.25
7.06
1.55
4.87
0.82
4.32
0.67
6.76
19.28
2.06
12.95
3.53
1.28
5.36
0.94
5.30
1.12
3.54
0.58
3.17
0.47
P a g e | 40
HRX30
HRX31
HRX34
HRX35
HRX36
HRX37
SS13TS
90/91
AAS22
DR#19
SS11
SiO2
TiO2
Al2O3
T
Fe2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K 2O
P2O5
LOI
Total
Y
Sc
Nb
Th
Sr
Rb
Ba
Co
Cu
V
Ga
Zn
Zr
U
Pb
47.96
1.53
16.56
11.13
1.67
8.51
0.18
6.94
11.74
3.5
0.46
nd
0.68
100.68
41.83
42.53
2.97
0.68
160
13.13
40.33
36.67
133.10
307
14.80
133.14
120.90
0.609
3.645
49.47
2.53
17.1
10.82
1.62
8.28
0.1
6.17
10.76
2.63
0.4
nd
1.2
101.18
31.47
36.95
1.32
0.24
143
7.51
23.67
58.59
179.03
246
12.85
133.99
85.38
0.24
4.886
49.5
2.02
18.88
10.92
1.64
8.35
0.65
3.64
9.58
3.39
0.95
nd
1.1
100.63
21.33
36.45
4.34
3.10
139
13.69
80.57
115.55
319.27
225
16.04
113.15
98.95
0.707
9.659
49.54
1.64
16.27
11.15
1.67
8.53
0.18
6.45
10.89
3.51
0.36
nd
0.65
100.64
40.43
40.94
5.32
1.56
149
14.33
71.38
74.15
249.00
336
15.87
159.09
134.86
0.444
7.347
49.01
2.15
17.5
12.59
1.89
9.63
0.15
5.12
10.12
2.95
0.41
nd
1
101
53.26
46.70
3.09
0.81
152
18.51
76.70
60.11
221.81
357
17.43
167.09
139.18
0.685
5.228
48.78
2.12
17.23
12.7
1.91
9.71
0.12
5.64
9.57
3.46
0.37
nd
1.2
101.19
52.65
48.97
2.80
0.76
137
22.08
64.12
73.63
143.56
310
17.39
164.78
134.23
1.009
3.607
50.35
1.91
16.07
13.77
2.07
10.53
0.2
4.14
9.47
2.9
0.9
0.18
1.15
101.04
47.86
49.39
3.04
0.40
145
20.23
71.43
47.04
112.95
225
15.75
147.01
111.32
0.656
5.072
50.68
2.17
13.62
12.34
1.85
9.44
0.29
6.06
11.24
2.72
0.73
0.17
1.6
101.62
50.18
47.03
2.66
0.44
121
12.71
69.73
47.41
122.41
358
16.14
150.64
134.92
0.623
11.069
50.99
1.27
13.54
11.67
1.75
8.93
0.17
6.65
12.41
2.87
0.34
0.09
0.99
100.99
48.23
51.11
1.22
0.12
121
9.00
70.11
44.99
120.22
270
15.77
167.25
99.59
0.666
8.125
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
8.28
15.28
2.47
15.47
4.29
1.49
6.34
1.16
6.42
1.36
4.32
0.71
3.81
0.58
6.54
10.23
1.87
11.58
3.14
1.15
4.64
0.84
4.79
1.03
3.21
0.53
2.86
0.44
12.40
29.96
2.84
15.81
3.68
1.21
5.0
0.81
4.11
0.82
2.50
0.40
2.20
0.33
8.87
25.12
2.58
15.63
4.32
1.47
6.29
1.14
6.26
1.33
4.25
0.70
3.76
0.59
10.61
15.36
3.08
19.34
5.51
1.85
8.02
1.50
8.34
1.76
5.63
0.93
4.86
0.72
9.10
18.15
2.83
17.93
5.16
1.73
7.71
1.42
7.99
1.73
5.54
0.93
4.85
0.74
9.15
12.63
2.17
14.02
4.05
1.45
6.32
1.22
6.82
1.51
4.84
0.79
4.32
0.66
8.48
14.07
2.66
16.97
4.91
1.69
7.53
1.41
7.73
1.68
5.30
0.89
4.74
0.73
7.31
12.02
1.82
12.56
3.27
1.21
6.22
1.32
6.98
1.09
4.29
0.71
3.93
0.67
nd - not determined
all Fe analysed as Fe2O3, FeO has been calculated considering 85% as FeO