the nature and origin of authigenic chlorite and...

THE NATURE AND ORIGIN OF AUTHIGENIC CHLORITE AND RELATED CEMENTS IN OLIGO– MIOCENE RESERVOIR SANDSTONES, TAPTI GAS FIELDS, SURAT DEPRESSION, OFFSHORE WESTERN INDIA

J. M. Huggett1*, S.D. Burley2&3 now 4&3, F. J. Longstaffe5, S. Saha2&6,now 7 and M. J. Oates 2 now 8

Reservoir sandstones in the Mid- and South Tapti gas fields in the Surat Depression (Mumbai Offshore Basin, western India) have been investigated using a range of petrographic techniques, isotope geochemistry and basin modelling. Authigenic chlorite is abundant in the shallow-marine sandstones of the Miocene Mahim Formation, a major reservoir rock in the Mid- and South Tapti fields, which are described here in terms of their quality and diagenetic characteristics. The sandstones are currently at burial depths of between ~1500 and 2800m. The authigenic chlorite has had a significant impact on the resulting reservoir quality of the sandstones and is interpreted to have originated as odinite clay of the verdine facies that replaced faecal or pseudo-faecal pellets, together with volumetrically small but abundant grain coatings and grain rims, and formed at the site of major riverine iron influx onto the shallow-marine shelf during periods of relatively low sea level. Pellets have been variably compacted to form pseudomatrix. Reservoir sandstones from similar depositional settings on the west coast of India or other sub-tropical settings are likely to exhibit comparable diagenetic effects on reservoir quality.

Compositionally, the chlorite is the iron-rich form known as chamosite. The chemistry of all the chlorite morphologies is the same in all studied samples. Oxygen isotope analyses of carbonate cements in the Mahim Formation sandstones have provided an approximate temperature framework for diagenesis of the non-carbonate cements. Oxygen isotope results for the chlorite, however, suggest much higher temperatures than its position in the paragenetic sequence would warrant. These results suggest that the clay formed first as 1:1 layer clays, in this case odinite, which were then transformed to Fe-chlorite as burial depths and temperatures increased.

Reservoirs in the Mahim, Daman and Mahuva Formation sandstones are thus greatly influenced by the diagenesis of authigenic chlorite and locally by the precipitation of carbonate cements. Reservoir quality is good where thick, continuous chlorite rim cements are present and where chlorite pellets are sufficiently indurated for them not to be compacted. Chlorite rim cements have reduced the extent of quartz overgrowth cementation in the sandstones.

1 Petroclays, The Oast, Sandy Cross Lane, Heathfield, East Sussex.2 BG India, BG House, Hiranandani, Powai, Mumbai, India.3 Basin Dynamics Research Group, School of Earth Sciences, The University of Keele.4 Murphy Exploration and Production Ltd, Petronas Towers, Kuala Lumpur 50088, Malaysia.5 Department of Earth Sciences, The University of Western Ontario, London, N6A 5B7, Canada.

Key words: Oligo-Miocene, tidal deposits, chlorite diagenesis, reservoir, sandstones, Surat Depression, Mumbai Offshore Basin, India, Tapti gasfields.6 Department of Geology, IIT Bombay, Hiranandani, Powai, Mumbai, India.7 Shell International Inc., Houston, Texas, USA.8 BG Norge, Løkkeveien 111, 4007 Stavanger, Norway.

*Corresponding author, email [email protected]

383Journal of Petroleum Geology, Vol. 38(4), October 2015, pp 383-410

© 2015 The Authors. Journal of Petroleum Geology © 2015 Scientific Press Ltd

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INTRODUCTION

The Oligo-Miocene Mahuva, Daman and Mahim Formation sandstones are important reservoir rocks for natural gas in the Mid- and South Tapti fields in the Surat Depression, in the northernmost part of the Mumbai Offshore Basin, western India (Fig. 1). In these fields, the sandstones are preserved at present-day burial depths of ~1500 - 2800m TVD ss (True Vertical Depth sub-sea). The sandstones contain abundant authigenic chlorite that occurs as pellets, pore-filling matrix and grain-coating platelets, and minor cements that include authigenic calcite and siderite

The Surat Depression is located at the southern entrance to the Gulf of Khambhat (Cambay), which has alternated between an estuarine embayment and a coastal plain since the Oligocene through to the present day, with broadly similar depositional environments occurring throughout this time (Pandey, 1986; Gombos et al., 1995; Raju et al., 1999; Saha et al., 2015 in review). Verdine (odinite-rich green clay) is known to be widespread in sandstones in this part of the modern western continental shelf of India (Thamban and Rao, 2000), suggesting that the abundant chlorite in the Oligo-Miocene reservoir sandstones of the Tapti fields is a diagenetic burial product of verdine maturation. Understanding the origin and nature of authigenic chlorite is important because it exerts a significant influence on sandstone reservoir quality and may contribute to preservation of porosity in more deeply buried sandstones, a continuing exploration target in the Mumbai Offshore Basin (Lal et al., 2009; Pandey et al., 2013). In this study, subsurface samples of sandstones from the Mahuva, Daman and Mahim Formations were analysed using a range of techniques including XRD, isotope geochemistry and electron microscopy, in order to determine the precise nature of the influence of chlorite on reservoir quality in the Tapti fields.

Authigenic chlorite in reservoir sandstone has been the focus of numerous studies, largely due to its ability to inhibit the extent of quartz cementation in reservoirs when the chlorite precipitates as rims on sand grains. As a consequence, reservoirs with chlorite rims may have anomalously high porosities at burial depths and temperatures considerably in excess of those at which the majority of reservoirs are tightly quartz cemented (Ehrenberg, 1993; Anjos et al, 2003; Berger et al, 2009; Adjukiewicz and Larese, 2012; Dowey et al., 2012; Bahlis and Ros, 2013).

Although authigenic chlorite is present throughout the Mahuva, Daman and Mahim Formation sandstones in the Tapti fields, it is particularly abundant in the tidal bar deposits of the Miocene Mahim Formation, which is the focus of this paper. Reference is also made to the occurrence of authigenic chlorite in the other

formations to highlight differences in abundance and distribution. Siderite and calcite cements were analysed in terms of carbon and oxygen stable isotope ratios, and provided constraints on the timing and origin of chlorite diagenesis. Simple 1D burial models were constructed to constrain the temperature evolution of potential source rocks and Miocene reservoir sandstones in the Tapti fields, and to place diagenetic processes in a time-temperature framework.

Holocene sediments in the Chinchini Formation in the Gulf of Khambhat are of similar lithological facies to the Oligo-Miocene reservoir sandstones, but with negligible diagenetic alteration. Accordingly, sandstone samples from the Chinchini Formation were included in this study to investigate pellet mineralogy and its modification during early diagenesis.

GEOLOGICAL SETTING

The Mid- and South Tapti fields are located in the Surat Depression (also known as the Tapti-Daman sub-basin) of the Mumbai Offshore Basin (Zutshi et al., 1993; Wandrey 2004; formerly known as the Bombay Offshore Basin) (Fig. 1). Structures at these fields comprise simple anticlines with four-way closures (Fig. 2) formed as inversion folds, probably in response to compression related to the Himalayan orogeny (Pangtey, 1996). In the Surat Depression, basement consists of Deccan basaltic lavas (“Traps”’), which are unconformably overlain by a thick succession of Paleocene to Pliocene sedimentary rocks (Fig. 3) (Zutshi et al., 1993; Wandrey 2004). Oligo-Miocene sandstones form major reservoirs for oil and gas in fields such as North, Mid and South Tapti, Lakshmi, Gowri, and the ‘B’ and ‘C’ Series fields (Basu et al., 1982; Wandrey, 2004; Goswami et al., 2007; Sanyal et al., 2012). Source rocks are inferred to occur in the Paleocene to Eocene Panna Formation, but may also be present in the marine Oligocene mudstones of the Mahuva Formation (Goswami et al., 2007). The petroleum system of the Mumbai Offshore Basin was comprehensively summarised by Wandrey (2004).

In the Mumbai Offshore Basin, the Oligo-Miocene interval comprises the Mahuva, Daman and Mahim Formations (Fig. 3). The Mahim Formation encompasses a range of intertidal and estuarine deposits, while the Daman Formation comprises fluvial and tidal channels deposited on a broad, vegetated coastal plain. The Mahuva Formation is predominantly sub-tidal to shallow marine in the area around the Tapti fields but becomes more offshore immediately to the SW.

These formations were deposited in a marine to estuarine embayment, which was flanked by a coastal alluvial plain (Deb et al., 1997; Saha et al., 2007; 2014; Dabholkar et al., 2011). Sedimentation in this

384 Chlorite cements in Oligo-Miocene reservoir sandstones, Tapti fields, offshore W. India

Huggett.indd 384 11/09/2015 12:38:02

embayment and coastal plain continued in a similar sedimentary and climatic setting up to the present day. Palaeo-shorelines shifted laterally by hundreds of kilometres as a result of relative sea level changes (Pandey, 1986; Raju et al., 1999; see Fig. 3), due to a combination of eustasy modified by local tectonics. During relative sea-level falls, the palaeo-shoreline shifted westwards as the coastal plains expanded, and sub-tropical weathering cut deep into the flanking lateritic hinterlands. In particular, the Daman Formation records a major sea-level fall of some 80m, which resulted in progradation of the late Oligocene shoreline to the SW by some 400 km relative to its present-day location. The embayment was flooded in the early Miocene transgression, which resulted in the deposition of the Mahim Formation.

The sand-dominated tidal-bar sediments of the Mahim Formation are characterised by a high proportion of authigenic green clay that is now

mineralogically chlorite. At the present day, green clay replacement of faecal and biogenic pellets (verdine) occurs in shallow-water (~55m) continental shelf sediments in the Mumbai Offshore Basin away from the main influence of tidal current activity (see Unnikrishnan et al., 1999; Nayak and Shetye, 2003), and distal from the main development of present-day tidal ridges (Off, 1963; Saha et al., 2007, 2015 in review). Sedimentation rates in this setting are low, typically less than 2mm/yr (Borole, 1988), with fine-grained suspended material being transported across the continental shelf by shoreline-parallel currents from the Indus Delta immediately to the north of the Mumbai Offshore Basin (Rao and Rao, 1995). Verdine development has been documented to be most extensive where sedimentation rates are low (Thamban and Rao, 2000; Huggett et al, 2010). If similar pellets of verdine mineralogy were the precursor of the chlorite pellets in the Mahim Formation, it is likely that they

Fig. 1. Outline map of the Gulf of Khambat, western India, showing the Cambay Rift and the Surat Depression in the northernmost part of the Mumbai Offshore Basin with the main structural elements and location of the Tapti fields highlighted (based on Biswas, 1987; and Goswami et al., 2007). The Tapti fields licence concession area is outlined.

SURAT DEPRESSION Major Faults

Km0 4020

Oil FieldsGas Fields

Mid TaptiField

North Tapti

Laxmi

Kosamba

DabkaUmra

Nada

72°E

Tapti River

South TaptiField

Tapt

i Con

cess

ion

Area

Dahej

ElavKim

Matwan

Gandhar

Namad

a Rive

r

Pakhajan

INDIA

CAM

BAY

RIFT

73°E

22°N

21°NGowri

385J. M. Huggett et al.

Huggett.indd 385 11/09/2015 12:38:03

also formed in shelfal, low energy environments and were reworked into tidal sands during transgressions.

SAMPLES AND METHODS

Core was available from the Mid- and South Tapti gas fi elds for the Oligo-Miocene sandstones of the Mahuva, Daman and Mahim Formations over a range of present-day burial depths between ~1500 and ~2800m. A total of 74 sandstone samples from four South Tapti and four Mid Tapti wells (locations shown on Fig. 2) were selected (Table 1), together with three samples of Holocene sandstone from a nearby shallow borehole cored for seabed geotechnical studies. The samples represent the range of depositional environments and burial depths available. Sampling was biased towards sandstones with a high proportion of chlorite and carbonate cements. Blue-dyed epoxy-impregnated thin sections, stained for carbonates and K feldspar, were prepared and described for each sample. Modal analysis was performed on all samples, with 300 points per slide counted. Core-derived porosity

and Klinkenberg-corrected permeability data were compiled and compared to porosity determined in thin sections.

X-ray diffraction (XRD) analysis was carried out on the <2µm fraction of selected sandstone samples using a Philips 1820 X-ray diffractometer with Ni-fi ltered CuKα radiation. For samples where XRD analysis indicates that chlorite-rich clay is the only clay present, the percentage of clay with a 0.7nm 001 spacing that is interlayered with the chlorite (1.4nm basal spacing) was calculated from the width at half height of the 001 diffraction (Hillier and Velde, 1992). The presence of 0.7nm layers is thought to indicate a 0.7nm clay precursor, whilst the abundance of these layers is believed to give an indication of the maturity of the chlorite.

Back-scattered scanning electron microscopy (BSEM) was carried out on carbon-coated polished thin sections using a JEOL 6310 SEM with an Oxford instruments energy dispersive X-ray analyser with INCA software. Analytical transmission electron microscopy (ATEM) was performed on two samples.

Fig. 2. Depth-structure map of the Mid- and South Tapti fi elds constructed on the top- Daman Formation horizon showing the general anticlinal structure of the fi elds, and the location of the study wells and the Tapti production platforms (based on Rana et al., 2006). Colour-coded for depth: purple-blue is deepest, red is shallowest. The Tapti concession boundary is shown on Fig. 1. Probable kitchen for hydrocarbon generation is the deep basinal areas to the south and SW of the South Tapti fi eld. Note the location of the South-West Tapti-1 well (SWT-1), used to model hydrocarbon generation from the source rock kitchen.

0 5 10 km

SWT-1

STE-1

STD-6 STC-52

STCSTB-1

STB

STA

MTA MTA-1MT-5

MT-2

MT-4Mid Tapti

Field

Tapti Concession Block Boundary

Depth colour bar

Mapped FaultsPlatformsStudy Wells

2320m

South Tapti Field 2200m

2000m

1800m

1600m

386 Chlorite cements in Oligo-Miocene reservoir sandstones, Tapti fi elds, offshore W. India

Huggett.indd 386 11/09/2015 12:38:03

Fig. 3. Summary stratigraphic chart for the Surat Depression, Mumbai Offshore Basin, illustrating general lithologies and the main sandstone reservoir intervals. Stratigraphic column from Deb et al. (1997). The palaeo-shoreline shifts through geological time (modified from Pandey, 1986) show the positions of shorelines in kilometres with respect to the current shoreline; positive values represent transgression whilst negative values represent regression.

Table 1. Summary of thin section point count data by formation (mean ± SD) indicating detrital and authigenic components for samples from the Mahim, Daman and Mahuva Formations.

Stratigraphy

AgeMa Series Formation Lithology

Lateral palaeo shoreline shifts

Pleistoceneto Recent

PlioceneChinchini

Tapti

Mahim

Daman

Mahuva

Diu

Belapur

Panna

Deccan VolcanicsCretaceous

10

20

30

40

50

60

Upp

erM

idM

idM

idU

pper

Upr

Lwr

Upr

Low

er

Low

er

Distance from current shoreline+100 0 -100 -200 -300km

0

65

Mio

cene

Olig

ocen

eEo

cene

Pale

ocen

e A

uthi

geni

c ch

lorit

e in

mic

rofo

ssils

Chl

orite

gra

in c

oatin

gs

Chl

orite

gra

in ri

ms

Non

-Fe-

calc

ite c

emen

t

Fe-

calc

ite c

emen

t

Fe-

calc

ite re

plac

emen

ts

Fe-

dolo

mite

Fib

rous

& fr

actu

re-fi

lling

calc

ite

Sid

erite

gra

in c

oatin

g

Mat

rix-re

plac

ing

side

rite

Sid

erite

repl

acin

g gr

ains

Por

e-fil

ling

side

rite

cem

ent

Qua

rtz o

verg

row

th c

emen

t

Aut

hige

nic

pyrit

e ce

men

t

Ti o

xide

s

Por

e-fil

ling

kaol

inite

Che

rt ce

men

t

Tot

al C

hlor

ite

Tot

al p

ore-

fillin

g ca

rbon

ates

Tot

al p

ore-

fillin

g ce

men

ts

Tot

al a

uthi

geni

c ce

men

ts &

cla

y

n = 42 47.7 4.1 0.4 0.5 0.2 0.2 0.8 12.0 6.0 4.4 0.1 0.4 1.5 3.1 1.7 0.1 0.0 0.0 1.7 0.4 0.9 3.1 1.5 0.8 0.0 0.0 0.4 24.4 9.5 14.2 30.6 4.0 3.6 7.6

Std.Dev. 5.3 3.3 0.7 0.6 0.3 0.4 1.4 12.7 5.4 4.4 0.3 0.9 2.2 6.3 6.0 0.4 0.1 0.0 3.9 2.5 2.3 6.9 2.2 0.8 0.0 0.0 2.3 4.3 12.6 10.8 8.8 4.3 2.3 6.6

n = 11 54.6 3.4 0.9 0.9 0.3 0.2 1.0 20.9 2.1 2.0 0.0 0.0 0.8 0.6 0.1 0.0 0.1 2.0 0.2 0.6 0.1 0.1 0.1 2.6 1.0 0.2 0.1 25.8 3.1 7.1 30.0 2.7 2.6 5.4

Std.Dev. 9.8 2.4 0.7 1.9 0.3 0.2 2.4 19.2 4.1 3.7 0.1 0.0 1.8 1.9 0.2 0.0 0.3 6.2 0.6 0.8 0.2 0.2 0.3 3.0 1.4 0.4 0.4 4.8 7.2 7.9 14.1 4.1 1.7 5.8

n = 21 42.6 2.9 1.9 0.2 0.2 0.1 3.6 9.9 1.3 5.3 0.2 0.1 0.9 0.9 3.6 0.6 0.9 0.0 1.6 9.7 0.0 1.7 3.0 1.5 0.2 0.5 0.2 17.8 8.7 15.0 30.1 4.6 1.3 5.9

Std.Dev. 19.3 2.6 4.9 0.3 0.3 0.2 8.9 16.6 1.6 9.8 1.0 0.3 1.5 3.9 10.1 1.3 2.4 0.1 4.2 14.0 0.0 4.7 4.0 2.3 0.6 1.1 0.7 5.1 12.6 12.3 13.6 7.4 1.5 9.0

Daman

Mahim

Mahuva

Detrital components Authigenic componentsSamples

Che

rt

Cha

lced

ony

Pla

gioc

lase

Fel

dspa

r

Porosity

STR

ATI

GR

APH

IC F

OR

MA

TIO

N

NU

MB

ER O

F SA

MPL

ES

Chlorite Carbonates Others Totals

Mon

ocry

stal

line

quar

tz

Pol

ycry

stal

line

quar

tz

Sec

onda

ry v

isib

le p

oros

ity

Tot

al v

isib

le th

in s

ectio

n po

rosi

ty

K-F

elds

par

Bio

clas

ts

Cla

y m

atrix

and

com

pact

ed c

hlor

ite p

elle

ts

Chl

orite

pel

lets

Slig

htly

com

pact

ed c

hlor

ite p

elle

ts

Prim

ary

visi

ble

poro

sity


Huggett.indd 387 11/09/2015 12:38:04

For this a carbon-coated copper grid was passed through diluted <2µm clay suspension and allowed to dry prior to examination using a JEOL 2000FX with an Oxford Instruments energy dispersive X-ray analyser with INCA software. Fracture surfaces of samples with chlorite rim cements were examined using a Philips Field Emission Scanning Electron Microscope (FESEM). For chemical analysis of the Fe2+/(Fe2+ + Fe3+) content, powdered samples of <2µm fraction clay were digested in HF/H2SO4 and the oxidisable species titrated against potassium permanganate. All chemical analyses and electron microscopy were carried out in the Natural History Museum, London.

Stable carbon and oxygen isotope analyses of carbonate cements were carried out at the Scottish Universities Environmental Research Centre and are reported in the standard δ-notation relative to VPDB (carbon) and VSMOW (oxygen). Samples were crushed and ground to a powder prior to analysis. In all cases the cement zoning occurs on too fine a scale to separate individual zones by micro drilling. Approximately 2g of carbonate were then reacted with 102% phosphoric acid to release CO2 for isotopic analysis. Reaction conditions were 100°C for 36 hours for pure siderite samples. Carbon dioxide was then manually extracted and purified prior to isotopic measurement using a dual-inlet, triple collector Micromass mass spectrometer. For samples containing both siderite and calcite, the selective acid extraction method of Al-Aasm et al. (1990) was used. The samples were first reacted at 25°C for 3 hours to extract CO2 from the calcite fraction, then at 100°C for 40 hours to obtain CO2 from the siderite. Fractionation factors of 1.01025 (calcite) and 1.00881 (siderite) were used to calculate the carbonate δ18O from the measured CO2 compositions (Rosenbaum and Sheppard 1986). Replicate measurements of the internal laboratory standard (Carrara Marble) yielded analytical precisions of ± 0.1‰ for both carbon and oxygen isotope measurements. The precision for mixed carbonate mineralogy and sequential extraction is sample-dependent, but is better than ± 0.4‰.

The chlorite oxygen isotope analyses were carried out in the Laboratory for Stable Isotope Science at the University of Western Ontario. The chlorite δ18O values are reported relative to VSMOW. Approximately 8mg of sample powder were placed into spring-loaded sample holders, which were then heated at 150°C for 12 hours under dynamic pumping. The samples were then top-loaded into nickel reaction vessels, where they were then heated at 150°C for a further 3 hours under dynamic pumping. Chlorine trifluoride was then added to the nickel reaction vessels, and the vessels sealed and reacted for 16 hours at 580°C to release the silicate-bound oxygen; the oxygen was then recovered and quantitatively converted to CO2 for oxygen isotope

analysis (Clayton and Mayeda, 1963; Borthwick and Harmon, 1982). The following average δ18O values were obtained for duplicate standard analyses during the course of these measurements (accepted value in parentheses): standard quartz, +11.7 ±0.1‰ (+11.5‰); kaolinite standard KGa-1, +21.4 ±0.1‰ (+21.5‰), and laboratory standard carbon dioxide, +10.30 ±0.03‰ (+10.30‰). Precision of ±0.2‰ (n = 5) was obtained for duplicate analyses of the chlorite samples.

A simple 1D burial history model was constructed in BasinModTM for the Mid- and South Tapti fields based on the stratigraphic sequence encountered in the STB-1 well using Burnham (1989) kinetics and calibrated to present-day Horner-corrected borehole temperatures. Additionally, a burial model was created for the kitchen area to the SW of the Tapti fields by using the SWT-1 well (see Fig. 2 for location). In the model calculations, mechanical compaction was linked to 1D fluid flow using the default BasinModTM algorithms and a modified Kozeny-Carman permeability calculation. The geothermal calculations were based on steady-state heat flow with a maximum temperature gradient in the Paleocene and a present-day sediment-water interface temperature of 22°C, and rounded to the nearest degree. Surface temperatures were increased to 24°C for the early Miocene using the predictive charts of Wygrala (1989).

RESULTS

Depositional EnvironmentsThe Holocene samples of the Chinchini Formation are representative of the shallow continental shelf sediments offshore in the Gulf of Khambhat (Cambay) away from the main influence of tidal current activity (Nayak and Shetye, 2003), and distal from the main development of present-day tidal ridges.

In the Tapti wells studied, the Mahim Formation comprises stacked, cleaning- and coarsening-upward sandstone packages enclosed in marine shales. In seismic time slices, the sandstones form elongate, lozenge-shaped bodies up to 2km wide and 10km long (Fig. 4), consistent with tidal bars comparable to those described for the present-day Gulf of Khambhat (Off, 1963; Saha et al., 2007). Core descriptions show that the sand bars are 10-15m thick, with a rapid increase in sand content from a rippled and bioturbated base. Sedimentary structures in the lower parts of the sand bars are dominated by lenticular and rippled fine-grained sandstones and banded heteroliths, commonly bioturbated with Planolites sp. and Thalassinoides-type burrows. The upper parts of the sand bars are typically cross-stratified and rippled, locally massive, and medium grained. The tops of the bars include pebbly sandstones with scour and erosional surfaces indicating extensive tidal reworking.


Huggett.indd 388 11/09/2015 12:38:04

The Mahuva and Daman Formation sandstones comprise interbedded fluvial and tidal channel units characterised by sharp, erosive bases and fining-upward grain-size profiles associated with smaller tidal bars interpreted to have been deposited in an estuarine setting or at the mouths of fluvial channels. Core studies show a range of different channels that are distinguished by extent and type of bioturbation. End-member fluvial channels are characterised by an absence of bioturbation; tidally-influenced channels near to estuary mouths are characterised by bioturbated top foresets, whilst sub-tidal channels are extensively bioturbated throughout.

Burial HistoryA 1D burial history plot (Fig. 5) was constructed for a representative on-structure South Tapti field well, based on well STB-1 (location in Fig. 2). Throughout the Oligocene and for the early part of the Miocene, periods of relatively slow deposition alternated with intervals of more rapid deposition. When deposition was slow, sediment packages remained within 100m of the coeval geomorphic or seabed surface for long

periods of time, and sediment palaeo-temperatures were in general less than 40°C. However, significant subsidence in the Burdigalian associated with the development of accommodation and marine transgression, caused rapid burial of Late Oligocene and Early Miocene sediments to depths of ~500m, resulting in temperature increases of 30°C over a relatively short period of time. The Middle through Late Miocene was characterised by continuous accommodation development and sedimentation with approximately 1km of burial to the start of the Pliocene. Very rapid deposition was initiated in the last 5 Ma with a further 1km of burial during the Pliocene. The burial history plot shows that palaeo temperatures in the Mahim Formation reached 60°C in the Middle Miocene because of this mid- to Late Miocene – Pliocene rapid burial, whilst the Mahuva Formation reached 80°C (Fig. 5). Present-day (downhole) temperatures exceed 100°C. Minor on-structure recent uplift is associated with the inversion responsible for the formation of the anticlines at the Tapti fields.

Source rocks for the Tapti reservoirs are inferred to occur in the interbedded marine shales of the Paleocene

Fig. 4. Seismic amplitude extraction map of the Tapti concession area (located in Fig. 1) from the stratigraphic horizon representing the main transgressive event in the Mahim Formation (modified from Saha et al., 2007; and Parks et al., 2006). Image shown is acoustic amplitude, with dark blue representing shale and pale blue-green-yellow colours representing sand. The distribution of sand is interpreted to represent tidal sand bars, revealing their size, geometry and orientation.

0 5 10 km

Acoustic Amplitude

0

50

100150200250

300350400450

500

Tapti Concession Block Boundary


Huggett.indd 389 11/09/2015 12:38:05

Panna Formation and Oligocene Mahuva Formation. The source rocks reached gas maturity in the Late Miocene and Pliocene in the modelled well in the main kitchen areas to the SW (SWT-1, see Fig. 2). On-structure in the South Tapti fi eld, Mahuva Formation potential source rocks are modelled to have entered the oil generation window in the mid-Miocene but have only recently entered the peak oil generation.

An estimate of porosity loss and compaction rate is an outcome from the 1D basin model. Fig. 6 shows that most porosity loss due to compaction in the sandstones of the Mahuva, Daman and Mahim Formations took place during the Early Miocene very soon after deposition, after which the sandstone framework stabilised and subsequent porosity loss was small.

Petrography and MineralogyModal analysis data for detrital and authigenic components of the 74 study sandstone samples from the Mahuva, Daman and Mahim Formations are summarised in Table 1, whilst details of authigenic chlorite habit and abundance are presented in Table 2. Compositional data for chlorite are given in Tables 3 and 4, and for carbonates in Table 5 along with their

stable isotope compositions. The oxygen isotope compositions of pure chlorite separates (<2 and <1µm size-fractions) are summarized in Table 6.

A summary diagenetic sequence for the Mahim Formation sandstones is shown in Fig. 7. This diagram indicates the main diagenetic processes and authigenic minerals observed in the sandstones and places them in a time framework based on relative textural relationships and at approximate temperatures by inference with the burial history.

Chinchini FormationThe Holocene sandstones from the Chinchini Formation sampled from the shallow borehole are friable, fi ne- or medium-grained sublitharenites (using the Folk 1968 classifi cation) dominated by quartz with minor feldspar and opaque minerals, and rare green clay pellets, which appear to be detrital. The sandstones are loosely cemented by minor grain-rimming, low-Mg calcite. The clays interbedded with the sandstones are dominated by either illite with rare illite-smectite, kaolinite and chlorite, or illite-smectite with rare illite, chlorite and kaolinite. These clays are all detrital in origin. Green clay pellets from these

Fig. 5. 1D burial history model for a crestal on-structure representative well (STB-1) from the South Tapti fi eld (location in Fig. 2) showing the stratigraphy above the Mahuva Formation in the well and modelled maturation zones. Burial model colour-coded for lithology: yellow is sandstone and brown is shale, with a colour overlay for organic maturation zones.

Early Mature (oil)0.5 to 0.7 (%Ro)

Mid Mature (oil)0.7 to 1.2 (%Ro)

Temperature Isotherms

Maturation Zones

Sediment Lithologies

Sandstone

Shale

Rupelian Chattian

500

Dep

th S

ubse

a (m

) 1000

1500

2000

2500

Aq LauBurdigalianLower Middle Upper

Formation

t=0

1Ma

3.2Ma

11.3Ma

14.5Ma

15.1Ma

16.5Ma17.7Ma17.7Ma

24.8Ma

30.5Ma

34Ma01020

Age (Ma)30

Mahim Sands

120°C

110°C

100°C

90°C

80°C

70°C

60°C

50°C

40°C30°C

Daman Sands

Tapti Sands

Chinchini Sand

Chinchini ShaleChinchini Formation

Tapti Formation

Mahim Formation

Daman Formation

Mahuva Formation

MIOCENEOGLIOCENEPLIOCENE

Serravallian Tortonian MesPle

Holocene


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sandstones are micro-porous and were difficult to analyse quantitatively with BSEM EDS. However, replicate measurements indicate that the Holocene green clay pellets have higher Mg and Si and lower Al and Fe than the Oligo-Miocene chlorite pellets (Table 3). XRD analysis of the Chinchini green clay pellets was not possible due to there being too few of them to separate out and analyse.

Detrital components in the Oligo-Miocene Sandstones Regardless of depositional environment, the Oligo-Miocene sandstones of the Mahuva, Daman and Mahim Formations are arenites of variable texture, being poorly to well sorted, and very fine to very coarse grained. Minor metamorphic polycrystalline quartz, chert, chalcedony and plagioclase feldspar, plus trace

Fig. 6. Predicted porosity reduction rate (expressed as a fraction of total porosity lost per million year) plotted against time for the basal Mahim Formation Sandstone from the 1D basin history model in Fig. 5. Compaction rate is most pronounced early during burial.

Fig. 7. Paragenetic summary chart based on petrographic data and linked to the burial history model, showing diagenetic processes and authigenic minerals that have affected the Mahim Sandstones in the Tapti fields. Bar width indicates the extent of the process or the abundance of the diagenetic mineral. Background colour is indicative of shallow- (blue), intermediate- (yellow) and deep-burial (red) processes.

0

24

22

20

18

16

14

12

10

8

6

4

2

0

10

Tim

e M

a

20 30

Porosity reduction rate (fraction/my)

Lower M

ioceneM

iddle M

ioceneU

pper M

ioceneP

lioceneP

leHeat Flow 62 Mw m -2

20 30 40 50 60 70 80

?

? ? ? ?

recrystallisation

EstimatedTemperature

Burt

al P

roce

sses

Sedi

men

t Wat

er In

terf

ace

Pellet compaction

Chlorite

Quartz overgrowth

Fe Calcite

Fe dolomite

Burrowing

Boring of bioclasts

Odinite pellets

Grain coatings

Verdinisation

Time Lower Miocene Present Day

?

?

CaMg siderite

Pyrite

Pellet dissolution

Chlorite rims


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Table 2. Chlorite abundance and textural habit from XRD and point count modal analysis for representative samples from all study wells.

Table 3. Compilation of chemical compositional data for studied chlorites from the shallow borehole, Mid Tapti MTA-1 well and other Tapti wells. The upper three rows contain data from BSEM EDS analyses; the bottom row contains data from ATEM analyses. All data are normalised to 100% (anhydrous).

Table 4. Normalized ATEM EDS compositional analyses of single authigenic chlorite crystals (numbered 1-9) with calculated cation distribution based on O20(OH)16 stoichiometry. The Fe2+/(Fe2+ plus Fe3+) ratio is assumed to be 0.65.

Table 1 XRD analyses of the < 2 µm fraction of selected samples

MT-1 1553.00 - 20 - - 100 1 11

MT-2 1634.20 - 15 - - 24 1 10

MT-2 1635.00 - 4 - - 100 0 9

MT-2 1637.80 4 8 4 2 100 0 6

MT-4 1669.00 4 22 7 3 100 1 14

MT-5 2000.50 - 4 2 2 100 1 21

MT-4 2023.00 - 18 3 6 99 1 13

MT-5 2039.00 - 36 - - 94 1 14

MT-4 2065.00 - 64 - - 64 1 13

MT-4 2094.00 - 24 - - 93 1 13

MTA-1 2105.8 - 15 4 - 100 1 10

MT-4 2110.00 - 21 - - 82 1 18

MTA-1 2132.0 - 37 1 - 100 1 11

MTA-1 2176.0 - 34 1 - 100 1 10

MT-4 2204.00 - 1 1 2 40 1 20

MTA-1 2539.2 - 41 - - 100 1 11

MTA-1 2556.0 - 26 - - 100 1 13

STB-1 1779.00 - 2 0 5 64 1 18

STE-1 1877.00 1 3 - - 98 1 12

STC-5z 2089.70 - 1 - - 93 1 12

STC-5z 2095.50 - 14 - - 57 1 15

STC-5z 2100.90 - 2 1 4 63 1 21

STD-6 2128.50 1 14 1 23 97 1 18

SWT-1 2182.10 - 1 - - 66 1 22

SWT-1 2190.75 - 1 - - 96 1 11

SWT-1 2291.65 - 1 4 0 100 0 8

Mid Tapti Field

South Tapti Field

Ratio of pellets to rimsWell Sample

depth (m) Chlorite rims % % 7nm clay in chlorite

Chlorite 001 peak °2θ at half height width

% chlorite from XRD

Chlorite pellets & compacted pellets %

Chlorite coatings %Field

Table 2 Oxide weight Percentages and cation compositions for EDS data obtained by ATEM

Component 1 2 3 4 5 6 7 8 9

SiO2 34.27 30.84 32.35 35.08 32.08 30.66 32.34 33.49 32.64

Al2O3 22.14 23.71 20.74 19.72 22.68 22.51 21.02 22.24 21.85

Fe2O3 7.26 7.77 7.75 7.40 7.40 7.90 7.83 7.52 7.60

FeO 29.03 31.08 30.98 29.62 29.62 31.60 31.31 30.07 30.41

MgO 7.30 6.59 8.17 7.94 8.22 7.33 7.50 6.68 7.47

Si 6.08 5.60 5.87 6.26 5.75 5.60 5.88 6.00 5.88

Al 1.92 2.40 2.13 1.74 2.25 2.40 2.12 2.00 2.12

Al 2.70 2.67 2.30 2.44 2.53 2.44 2.38 2.69 2.52

Fe3+ 2.17 2.37 2.36 2.22 2.23 2.43 2.39 2.26 2.30

Fe2+ 4.02 4.41 4.39 4.13 4.15 4.51 4.44 4.21 4.28

Mg 1.93 1.78 2.21 2.11 2.20 2.00 2.03 1.78 2.01

10.82 11.23 11.27 10.90 11.11 11.37 11.25 10.94 11.11

Number of cations on the basis of O20(OH)16

∑ octahedral

∑ tetrahedral 8.00 8.00 8.00 8.00 8.00 8.00 8.00 8.008.00

Sample Age Number of analyses MgO Al2O3 SiO2 Fe2O3/FeO

PJ Pilot Hole (TD = 39m) Holocene n = 6 15.2 14.3 42 24.5

MTA-1 Oligo-Miocene n = 26 6.9 23.9 34.2 35

Other Tapti wells Oligo-Miocene n = 56 6.6 18.4 27.8 47.2

ATEM analyses Oligo-Miocene n = 9 7.5 21.9 32.6 31

Chemical composition


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rutile and zircon, are also present (Table 1). Occasional shelly bioclasts have been replaced by ferroan carbonates. Small amounts of detrital clay are present in many samples, mostly as discontinuous laminae and compacted burrow linings. X-ray diffraction indicates the presence of illite, illite-smectite and kaolinite. The detrital clay is considered to consist predominantly of these components.

Authigenic components in theMiocene Mahim Formation Sandstones

Chlorite Chlorite is the most widespread and abundant authigenic mineral in the sandstones of all three formations (Fig. 8, Table 1), particularly in the tidal bar sandstones (Table 7). Authigenic chlorite occurs

Table 5. Cation proportions from EDS (mole fraction) and stable isotope data (‰) for siderite- and calcite-cemented samples. N/D = not determined. See interpretation of carbonate stable isotope results for discussion of mineral precipitation calculations.

Mg Ca Mn Fe δ13C ‰ VPDB

δ18O ‰ VPDB

δ18O ‰ VSMOW

T °C seawater δ18O

= –3‰


= –2‰


= –1.2‰


= 0‰

MT-2 1636.5Microcrystalline siderite cement in chlorite pellet-rich argillaceous tidal sand

0.09 0.08 0.05 0.79 –14.4 –5.4 +25.4 38 44 49 58

MTA-1 2001.2 Pore-filling microspar siderite in tidal ridge sand N/D N/D N/D N/D –16.3 –6.6 +24.1 46 53 58 67

MTA-1 2105.4 Microspar siderite coating chlorite rims in tidal bar sand 0.06 0.08 0.04 0.82 –12.4 –9.1 +21.5 64 71 78 88

MTA-1 2105.8 Microcrystalline siderite in tidal bar sand N/D N/D N/D N/D –22.8 –5.9 +24.8 42 48 53 62

MTA-1 2120.5 Microspar siderite-cemented clay in tidal ridge sand N/D N/D N/D N/D –16.9 –8.8 +21.8 61 69 75 86

MTA-1 2128.6 Microspar siderite-cemented clay in tidal channel sand N/D N/D N/D N/D –17.7 –9.5 +21.1 66 74 81 92

MTA-1 2151.2Microcrystalline siderite rims overgrown by ferroan calcite in tidal channel sand with rare chlorite

0.16 0.09 0.04 0.71 –21.4 –5.4 +25.3 39 45 50 58

MTA-1 2160.3Microspar siderite in tidal ridge sand with chlorite pellets and abundant compacted pellets

0.11 0.10 0.04 0.76 –28.1 –3.7 +27.1 29 34 39 47

MTA-1 2173.9Microcrystalline siderite replacement or infilling of clay pellets in tidal channel sand

N/D N/D N/D N/D –25.7 –6.2 +24.5 44 50 55 64

MTA-1 2645.5 Spaerheroidal grain-coating siderite in tidal sheet sand 0.07 0.13 0.02 0.78 –12.5 –11.2 +19.3 81 90 98 110

MTA-1 2647.5 Pore-filling microspar siderite in bioclastic sand N/D N/D N/D N/D –6.9 –8.7 +22.0 60 67 74 84

MTA-1 2658.5 Zoned siderite microspar in claystone 0.10 0.08 0.01 0.82 –7.4 –4.7 +26.1 34 40 45 53

MTA-1 2719.2 Microcrystalline siderite-cemented claystone N/D N/D N/D N/D –8.2 –3.5 +27.3 28 33 38 45

STB-1 1777Concretion of microcrystalline, intergrown non-ferroan calcite and siderite

0.13 0.15 0.04 0.67 –18.7 –3.3 +27.5 27 32 37 44

STD-6 2177.5Early microcrystalline siderite overgrown by ferroan calcite at top of fluvial sand with tidal influence; ~1% chlorite pellets, most partially pyritized

N/D N/D N/D N/D –25.3 –3.3 +27.5 27 32 37 44

STE-1 1906Microspar ferroan calcite and microcrystalline Mn-siderite cement in Intertidal laminite

0.05 0.15 0.24 0.57 –17.1 –10.7 +19.9 76 84 92 104

MT-4 1672 Ferroan calcite-cemented, chlorite pellet-rich sand N/D N/D N/D N/D –15.0 –10.4 +20.2 55 61 67 75

MT-4 2106 Ferroan calcite-cemented bioclastic sand N/D N/D N/D N/D –7.4 –10.5 +20.1 55 62 67 76

MT-4 2025 Fibrous calcite-cemented sand 0 0.93 0.03 0.04 –15.3 –10.9 +19.6 58 65 71 80

MT-5 2000.2Minor ferroan calcite in sand with compacted chlorite-pellet matrix and low porosity

N/D N/D N/D N/D –15.4 –12.2 +18.4 66 74 80 90

MTA-1 2105.4Ferroan calcite and Mn calcite cement in tidal bar sand with minor chlorite pellets and rare bioclasts

0.09 0.01

0.57 0.91

0.26 0.03

0.08 0.05

–11.2 –9.9 +20.7 51 58 63 72

MTA-1 2105.8Ferroan calcite microspar cement in tidal bar sand with minor chlorite pellets and rare bioclasts

N/D N/D N/D N/D –13.0 –3.2 +27.6 17 21 25 31

MTA-1 2120.5Ferroan calcite cement in tidal ridge sand with bioclasts and abundant chlorite pellets

N/D N/D N/D N/D –15.9 –9.9 +20.7 51 58 63 72

MTA-1 2128.6Ferroan calcite cement in tidal channel sand with abundant chlorite pellets and rare bioclasts

0.01 0.92 0.04 0.03 –16.2 –10.4 +20.2 55 61 67 75

MTA-1 2151.2Ferroan calcite that has overgrown siderite rims and rare chlorite in tidal channel sand

N/D N/D N/D N/D –11.5 –10.5 +20.1 55 62 67 76

MTA-1 2173.9Ferroan calcite that has overgrown siderite-cemented, compacted clay pellets in tidal channel sand

N/D N/D N/D N/D –16.6 –10.7 +19.9 56 63 69 78

MTA-1 2658.5 Ferroan calcite cement in rootlet cavities and fractures 0 0.95 0.02 0.03 –15.0 –11.3 +19.3 60 67 73 82

MTA-1 2645.5Poikilotopic ferroan calcite that has overgrown spaheroidal siderite and chlorite rims in tidal sheet sand

0 0.94 0.04 0.02 –14.2 –12.3 +18.2 68 75 81 91

MTA-1 2647.5Ferroan calcite that has cemented/replaced bioclasts in transgressive, bioclastic sand

N/D N/D N/D N/D –8.0 –9.5 +21.1 49 55 61 69

MTA-1 2719.2Ferroan calcite cementing microfractures in siderite cemented claystone

N/D N/D N/D N/D –14.3 –9.4 +21.2 48 55 60 68

STB-1 1777Concretion of microcrystalline non-ferroan calcite and siderite with rare fragments of chlorite pellets

0.04 0.86 0.02 0.07 –18.5 –3.5 +27.3 18 22 26 33

STB-1 1779Ferroan calcite cement in sand; chlorite pellets compacted to matrix where ferroan calcite absent

N/D N/D N/D N/D –18.0 –10.2 +20.4 53 60 65 74

STD-6 2177.5Ferroan calcite microspar that has overgrown earlier siderite; at top of tidally influenced fluvial sand

N/D N/D N/D N/D –23.4 –2.5 +28.3 14 18 22 27

STE-1 1887.9Calcite nodule in estuarine mudstone with brown-green, uncompacted chlorite pellets and bioclasts

0.03 0.87 0.02 0.08 –14.3 –10.6 +20.0 56 62 68 77

STE-1 1906 Microspar ferroan calcite cement and microcrystalline Mn-rich siderite cement in intertidal laminite

0.1 0.88 0.05 0.06 –13.3 –11.8 +18.8 64 71 77 86

Calcite

Siderite

Calculated mineral precipitation temperature

Well Sample depth (m) Sample description

Isotopic composition


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mostly as pellets together with compacted pellets that form pseudomatrix, while rims (clay platelets radial to the grain they coat) and grain coatings (clay platelets tangential to the grain) are a widespread but quantitatively minor component (Table 2). Pellets and compacted pellets (Fig. 8a) are volumetrically an order of magnitude more abundant than either rims or coatings. Pellets in the Mahim Formation sandstones have a similar appearance to those in the Holocene Chinchini Formation (Fig. 9a). Some dissolution of pellets has occurred after carbonate cementation (Fig. 8b). Some pellets, possibly those that are less mature, have been leached after formation of chlorite matrix (Fig. 9b). Distinguishing between detrital matrix and thoroughly compacted chlorite pellets was not always possible in thin section; hence the two were counted together in the modal analysis (Table 1). Green clay matrix is interpreted to be composed of compacted clay pellets, whilst brown and frequently birefringent

clay matrix was assumed to be detrital. Unequivocal chlorite matrix (i.e. clearly compacted chlorite pellets) was most abundant in the tidal ridge/bar and sub-tidal channel sandstones (Table 7). Discrete chlorite pellets occur almost exclusively in the tidal ridge/bar and tidal channel sandstones and are rare in the sub-tidal channels. Pellets are uncompacted where poikilotopic carbonate cement is present. Evidence of leached pellets in the form of voids is most apparent in carbonate-cemented sandstones (Fig. 8b). This leaching can amount to as much as 10% of the sample volume.

Chlorite coatings occur on all types of grains, including chlorite pellets (Fig. 8c) and carbonate clasts; chlorite also lines borings in shell fragments. Authigenic chlorite also lines intragranular pores resulting from pellet dissolution (Fig. 8c), and more rarely secondary pores after dissolved detrital feldspar. Chlorite grain rims tend to be thickest on clay-rich

Table 6. Summary of oxygen isotope compositions of individual chlorite separates. Results of duplicate analyses are shown in bold.

Table 7. Summary of point count data (mean ± SD) for porosity-reducing components, calculated thin section minus cement porosity, and core-derived helium (air-dried) porosity and Klinkenberg-corrected permeability data. The clay matrix is dominated by chlorite, and in many instances it is only chlorite. See text for discussion.

Table 8. Semi-quantitative X-ray diffraction data for sandstone samples with both clay pellets and clay matrix. The fact that the clay is entirely chlorite implies that the matrix is compacted chlorite pellets.

Av. SD Av. SD Av. SD Av. SD Av. SD Av. SD Av. SD Av. SD

Tidal ridge/bar (n = 16) 14 12 6 5 6 4 11 9 25 14 8 11 25 6 232.6 612.5

Tidal channel (n = 19) 31 13 0.2 0.2 1 0.8 8 8 39 7 8 10 21 6 46 128.2

Sub-tidal channel (n = 5) 23 10 2 3 3 3 17 8 40 14 9 9 10 5 1.4 4.5

Measured permeability (mD)

Minus cement porosity (%)

Measured porosity (%)

Clay matrix + pore-filling cements (%)

Compacted chlorite matrix (%)

2.4

Depositional EnvironmentChlorite pellets (%) Pore-filling cements

(%)

3

Total clay matrix (%)

Table 6. Oxygen isotope compositions of chlorite separates

δ18O ‰ Yield δ18O ‰ YieldVSMOW µmol O2/mg VSMOW µmol O2/mg

<2µm <2µm <1µm <1µm

MTA-1 2105.8 +8.92* 12.47 +9.57 13.00MTA-1 2132 +8.73 11.1 +8.79 12.67MTA-1 2176 10.52 12.76 10.54 13.15MTA-1 2539.2 +9.70 12.44 +9.93 12.63MTA-1 2556 +9.21 12.58 10.39 12.75

*Results of duplicate analyses shown in bold

Sample depth (m)

Well

WellSample depth

(m)Pellets Matrix Chlorite Kaolinite Illite swelling

clay

MTA-1 2105.8 9% 9% 100% 0% 0% 0%

MTA-1 2132.0 4% 33% 100% 0% 0% 0%

MTA-1 2176.0 1% 33% 100% 0% 0% 0%

MTA-1 2539.2 tr 41% 100% 0% 0% 0%

MTA-1 2556.0 1% 25% 100% 0% 0% 0%

MTA-2 1635.0 24% 35% 100% 0% 0% 0%

MTA-2 1637.8 13% 17% 100% 0% 0% 0%

Chlorite type Clay mineralogy from XRD


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Fig. 8. Representative thin section images of Mahim Formation sandstones taken in plane polarised light. Blue is porosity. (A) Pseudomatrix-supported chloritic sandstone. Rounded quartz grains are colourless. Extreme compaction of chlorite pellets has produced pseudo-matrix and greatly reduced porosity. Pyrite (black) has cemented intergranular porosity prior to compaction. Scale bar = 100µm. (B) Grain-rimming siderite (short arrows) has overgrown chlorite rims (long arrow) and is in turn enclosed by ferroan calcite (red arrow). The ferroan calcite cementation pre-dates compaction of chlorite pellets. Some chlorite pellets (LCP) have been leached after formation of carbonate cement. Scale bar = 100µm. (C) Chlorite pellets (LCP) have been leached after formation of chlorite matrix. Scale bar = 100µm. (D) Sandstone with low clay content, discontinuous chlorite rims (arrows) are present on quartz grains, and are not extensive enough to have prevented a small amount of quartz overgrowth cement (e.g. arrow on the left). Scale bar = 100µm.

Fig. 9. Representative SEM images. (A) Verdine granule (V) from the Holocene Chinchini Formation. Scale bar = 100µm. (B) BSEM image of a chlorite coating around partly leached chlorite grain, Mahim Formation. The coating has been overgrown by siderite (arrow) and both have been overgrown by poikilotopic calcite (Ca). Scale bar = 100µm. (C) SEM image showing detail of grain-rimming chlorite, Mahim Formation. Scale bar = 5µm. (D) Sphaerosiderite (arrow) has precipitated on quartz grains prior to the grain-rimming chlorite, Mahim Formation. Scale bar = 50µm.


Huggett.indd 395 11/09/2015 12:38:13

detrital grains and chlorite pellets, and comprise densely intergrown equant platelets (Fig. 9c). Modal analysis data indicate that there is a weak negative correlation between the proportions of chlorite rims and quartz overgrowth cement (Fig. 10a), and between total chlorite and the proportion of quartz overgrowths (Fig. 10b).

Although authigenic chlorite is present in all the shallow-marine and estuarine sandstones, the absence of authigenic chlorite from the fluvial sandstones of the Daman and Mahuva Formations, and the lower overall abundance of chlorite of all modes of occurrence in these formations, is evidence of facies control on the distribution of chlorite rims.

XRD analyses show strong even number 00l reflections, consistent with the chlorite being Fe-rich. The proportion of 0.7nm thick unit-cell layers in chlorite (1.4nm thick unit cell) is suggested to be an indication of the temperature of chlorite formation (Hillier and Velde, 1992). For samples where XRD analysis indicates that chlorite is the only clay present, calculated values for the proportion of 0.7nm clay interlayered with the 1.4nm clay layers were determined following the method of Hillier and Velde (1992). All except four of the values calculated by this method indicate <20% 0.7nm layers in the chlorite. There is no clear relationship between depth and the percentage of 0.7nm layers, or between the ratio of rims to pellets and the percentage of 0.7nm layers (Table 2).

X-ray energy dispersive spectra (EDS) obtained from polished thin sections and ATEM data (Table 4) indicate that all of the chlorite samples are the Fe-rich chlorite, chamosite, as defined by Hayes (1970). Most SEM-EDS analyses have low totals due to the analysis volume being larger than the particle or particles being analysed, and to the inclusion of resin in the analyses. More accurate analyses were obtained by ATEM (Table 4). A further inaccuracy in the chemical quantification results from slight oxidation of chlorite in the older

cores. Post-coring oxidation is evident from the bright margins (due to iron oxides) of chlorite rims in BSEM images of the cores cut the greatest time ago, and is inferred from the overly high iron oxide weight percentages measured for chlorite in the older South and Mid Tapti cores compared with the most recent (fresh, unweathered) cores from well MTA-1 (Table 3). Iron oxides as well as siderite can be avoided in ATEM analyses, which therefore provide the most reliable data of all compositional analyses (Table 3).

The Fe2+/(Fe2+ + Fe3+) ratio values determined for the <2µm fraction of three samples (for which XRD indicates chlorite is the only clay) are 0.6-0.7. All of these samples include minor authigenic siderite, which occurs as crystals that are >>2µm in size. Hence, any contamination of the <2µm chlorite is considered to be negligible and the Fe2+/(Fe2+ + Fe3+) values are assumed to reflect chlorite compositions correctly.

The δ18O values of five, high purity chlorite pellet separates range from +8.7 to +10.5‰ for the <2µm size-fraction and from +8.8 to +10.5‰ for the <1µm size fraction of the same samples (Table 6). On average, the smaller size fraction is enriched in 18O by ~0.4‰ relative to the larger size fraction of the same sample (Table 6).

Quartz cementOvergrowths of quartz are only present in significant amounts in the matrix poor, “clean” Mahim Formation sandstones in which detrital clay is largely absent. Quartz overgrowths are widespread in sandstones with negligible pore-filling chlorite and typically comprise up to 10% of the sandstone volume (Fig. 10). The average abundance of quartz cement is highest in the sub-tidal channel sandstones, although the high standard deviations indicate that little significance can be attributed to the averages (Table 7). Quartz cement has overgrown siderite rims that have been subsequently enclosed by ferroan calcite.

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Page 38 of 48

For review only

Journal of Petroleum Geology

123456789101112131415161718192021222324252627282930313233343536373839404142434445464748495051525354555657585960

Fig. 10. Point-counted quartz overgrowth abundance versus (A) chlorite rims, and (B) total chlorite amount visible in thin section. Data are from all formations, but mainly the Mahim Formation.

A B


Huggett.indd 396 11/09/2015 12:38:14

Carbonate cementsSiderite and calcite (ferroan and non-ferroan) cements are present in all three formations (Table 1) and all depositional facies studied. Most of the siderite cement is chemically impure with a typical cation fraction of 0.1-0.2 Ca, Mg and Mn substitution for Fe (Table 5). This siderite is hereafter referred to as CaMg-siderite. Microcrystalline CaMg-siderite forms concretions and has cemented clay-rich matrix and chlorite pellets. In some of the clay-poor sandstones, this CaMg-siderite occurs as coatings on grains, including chlorite pellets, and on grain-coating chlorite rims (Fig. 8b). CaMg- siderite occurs in all of the tidal and sub-tidal facies and has been overgrown by ferroan calcite (Fig. 8b) and quartz. Microspar and macrospar CaMg-siderite, sometimes with sweeping extinction, have cemented sandstones locally, particularly in the Mahuva Formation, and have partly replaced chlorite rims. The textural relationship of the microspar and macrospar with the micritic siderite is not clear.

Rarely, grain-coating siderite has a spheroidal form. This siderite is in some instances Mn-rich (Mn ~ 0.2 cation fraction), is overgrown by chlorite rims (Fig. 9d), and very rarely has been overgrown by dolomite as well as the more abundant Fe-calcite. Spheroidal siderite also occurs as micro-nodules around former rootlets and as grain coatings in the Mahuva Formation sub-tidal sandstone. The textural relationships of the various CaMg-siderite habits suggest that siderite formed at several different times during the early diagenetic evolution of these sandstones.

Fe-calcite is weakly to strongly ferroan (0.02-0.08 Fe cation fraction), although it is less variable in composition and morphology than the siderite. The ferroan calcite occurs as macrospar or poikilotopic cement that has overgrown siderite and engulfs chlorite pellets and chlorite rims. Macrospar Fe calcite has been overgrown by quartz cement. Ferroan dolomite has replaced bioclasts and clastic grains; it has not cemented primary porosity. Its textural relationship with other carbonates is not clear.

Siderite follows a weak trend from low δ13C (minimum –28.1‰) and high δ18O (maximum +27.5‰), to relatively high δ13C (maximum –6.9‰) and low δ18O (minimum +19.3‰). No systematic trends were observed between siderite or calcite chemistry and isotopic compositions, or between carbonate chemistry and isotopic compositions with depth (Table 5). Calcite cements fall into two groups in terms of stable isotope compositions, one with δ18O values of +27.3 to +28.3‰, and another with δ18O values of +18.2 to +21.2‰. These groups do not correspond to any particular calcite morphology observed in thin section. The δ13C values range from –23.4 to –13.0‰ for the higher 18O group, and –16.6 to –7.4‰ for the lower 18O group. The highest δ13C

values are associated with bioclastic sandstones containing bivalve shell fragments. The cluster of three samples with lower carbon and higher oxygen isotope compositions represent both early micritic calcite and microspar ferroan calcite.

DISCUSSION

Formation of iron-rich claysGreen-coloured clay minerals (verdine) occur in modern low-latitude near-shore settings associated with major riverine input (Odin and Sen Gupta, 1988). Verdine pellets comprise an assemblage of clays dominated by either odinite or odinite-rich mixed-layer clay associated with detrital minerals and ferric chlorite, and occur only in sediments that have not been buried more than a few cm beneath the sediment-water interface (e.g. Odin and Sen Gupta, 1988; Rao et al., 1995). Odinite is a ferric, iron-rich clay with a 0.7nm basal spacing (Bailey, 1988a).

At the present day, verdine forms only in low latitude, oxic sand-rich sediments, preferentially in water depths of 20 - 60m, where the water temperature is >20°C and where major river systems deposit iron-rich sediment (Odin et al., 1988). An example of this environmental setting is the western continental margin of India, where recent verdine has been described by Thamban and Rao (2000). Here, the source of iron-rich lateritic sediment is derived from the weathering of the basaltic Deccan Traps. The coastal hills known as the Western Ghats represent the tilted rifted margin of the Deccan lavas (Radhakrishna, 1993) and provided laterite-rich sediment to the basins offshore to the west throughout the Tertiary (Widdowson, 1997).

The distribution of verdine is of relevance to reservoir geology because verdine is potentially a precursor mineral for chlorite rims, the distribution of which can be important in preserving reservoir quality in sandstones (McIlroy et al., 2003; Dowey et al., 2012). Chlorite pellets in ancient shallow-marine sandstones deposited at low latitudes have a similar distribution to odinite (the precursor of chlorite in faecal pellets and ooids) in Recent sediments (e.g. Odin and Sen Gupta, 1988; Thamban and Rao, 2000). The relationship between these two minerals has not been widely studied or reported due to a lack of available sedimentary successions that include both Recent and more deeply-buried diagenetically modified clays. Previous studies have found chlorite rim cement to occur in a range of near-shore sandstones, not all of which are associated with riverine input, and no specific depositional environment is favoured overall (e.g. Ehrenberg, 1993; Hillier et al., 1996; Aargard et al., 2000; Dowey et al., 2012).

If the chlorite in sandstones in the Mahuva, Daman and Mahim Formations originated as clay pellets that


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have a similar composition to the those in the Holocene verdine in the Chinchini Formation, analyses of the latter will provide an estimate of the composition of the Oligo-Miocene chlorite prior to diagenetic modification. The data in Tables 3 and 4 indicate that diagenesis has resulted in a reduction in Mg and Si, and an increase in Al and Fe. The mineralogical transformation that this reflects is not yet well understood, but is thought to involve transformation of ferric iron-rich serpentine clay (ferric-rich 0.7nm clay, odinite; see Bailey, 1988b) and smectite to ferrous iron-rich clays (berthierine, plus swelling chlorite and other mixed-layer clays), and ultimately to ferrous iron-rich chlorite (Aargard et al., 2000; Ehrenberg, 1993; Ryan and Reynolds, 1996; McCarty et al., 2004; Huggett et al., 2006).

Whilst the thermal stability of odinite (ferric iron-rich) is not known, its absence as a discrete clay mineral from pre-Quaternary sediments (Ryan and Hillier, 2002) suggests that it is replaced by berthierine or chlorite (both rich in ferrous iron) at a low burial temperature. In addition to iron reduction, chlorite formation involves replacement of 0.7nm by 1.4nm thick crystal layers (Hillier and Velde, 1992; Huggett et al., 2006). The 0.7nm layers are commonly present in chlorite, which suggests that there is a 0.7nm clay precursor to the 1.4nm chlorite (Hillier and Velde, 1992). Ryan and Hillier (2002) and Hillier (1994) demonstrated that the replacement of 0.7nm by 1.4nm layers is temperature-controlled. However, the measured percentage of 0.7nm layers in the chlorite in the analysed sandstones does not show a decrease with increasing burial depth (Table 2). In fact the shallow samples from the Mid Tapti field have overall lower 0.7nm contents than the deeper samples from the South Tapti field. Nor is there a correlation between the ratio of chlorite pellets to chlorite rims and the percentage of 0.7nm layers in the <2µm fraction.

Compaction has affected a high proportion of chlorite pellets in the Mahim, Daman and Mahuva Formation sandstones and is the principal cause of reservoir quality reduction in the Mid- and South Tapti fields. However, the precise effects of compaction are difficult to assess because much of the clay described as matrix in the modal analysis may in fact be pellets that have undergone intense compaction. The XRD data (Table 8) suggest that this matrix is pure chlorite in the high porosity sandstones, consistent with it being compacted pellets. The intensity of compaction is inferred to have been greater for softer (i.e. less mineralogically mature) pellets. Some chlorite pellets enclosed in compacted chlorite pellet matrix have a blue-green colour compared with the matrix (olive-green). This colour difference is consistent with the matrix being less diagenetically altered, as well as less indurated.

EDS X-ray analyses indicate a similar chemistry for chlorite pellets, matrix and rims. However, a change in mineralogy to a more stable clay assemblage does not necessitate a change in chemistry. It is quite probable that invertebrate faecal pellets vary from species to species and genera to genera in terms of mineralogical composition, amount of binding mucus (i.e. reactive organic matter: Anderson et al., 1958; Satterberg et al., 2003) and density. Certain invertebrates may also modify the clay in their guts, thereby pre-disposing the clay to alteration or replacement by other clay minerals (McIlroy et al., 2003; Needham et al., 2006). All of these factors affect the diagenetic evolution of the pellets during burial. The mineralogical composition and abundance of mucus may affect the rate of verdinisation, while the density will affect the susceptibility to compaction.

Sediment-feeding invertebrate species, and hence their pseudo-faeces, are typically confined to a narrow range of environments (e.g. Fenchel, 1996). It follows that the degree of induration and the mineralogy of the clay pellets will also reflect the sedimentary environment. The modal analysis data (Table 1) indicate that the proportion of compacted pellets is highest in the Mahim Formation, and lowest in the Mahuva Formation. Modal analysis data presented by depositional environment show that each environment has different abundances of pellets, slightly compacted pellets, and matrix that may once have been clay pellets (Table 7). This is consistent with the pellets having been more or less prone to recrystallisation (which makes the pellets harder and more resistant to compaction) to varying degrees in different sedimentary environment, according to the species of invertebrate(s) that made the pellets.

Interpretation of carbonate stable isotope results Inferred carbonate formation temperatures using the oxygen isotope data are dependent upon the δ18O value assumed for the water from which they precipitated or re-equilibrated. Given the shallow-marine to estuarine depositional environment of the Mahim Formation sandstones in the Tapti fields, the maximum probable oxygen isotope composition is that of seawater (0‰ at present, versus –1.2‰ for an ice-free ocean prior to the Early Miocene: Shackleton and Kennett, 1975). Significant freshening (and hence lowering of δ18O) of this seawater due to dilution with meteoric water is very likely. For example, Von Rad et al. (2006) found that marine sediments associated with the Holocene monsoon in the Indian Ocean record δ18O values of ~ –2‰. Development of a monsoonal climate, linked to the uplift of the Himalayas, began in the Miocene (Valdiya, 1999; Zheng et al., 2004; Clift et al., 2002) and possibly as early as the Eocene (An Zhisheng et al., 2001; Najman et al., 2010),


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and resulted in local marine δ18O values as low as ~ –3‰. Considering the depositional environment of the Oligo-Miocene sandstones in the Tapti fi elds, the infl uence of freshwater mixing from monsoonal run-off and potential diagenetic reactions, it is reasonable to assume a water δ18O range between its present value of 0‰ and a minimum value of –3‰.

Substantial early formation of secondary authigenic minerals in a closed system has also been shown to produce up to 3‰ lowering of marine porewater during low-temperature, closed-system interaction with oceanic sediments rich in volcanic ash or through direct interaction with basalt (see reviews by Savin and Yeh, 1981; Longstaffe, 1987). The sediments of the Mahim Formation, however, do not have a detrital or diagenetic history indicating such neoformation in a closed marine porewater system on such a massive scale. Sediment-water interaction during deeper burial, by comparison, is known to drive trapped porewater δ18O to higher values (see Longstaffe, 1987 and references therein). Hence it is reasonable to consider that later stages of diagenesis of the analysed sandstones involved porewaters enriched in 18O relative to its starting composition.

The following mineral-water oxygen-isotope geothermometers were used to calculate carbonate crystallization or isotopic re-equilibration temperatures (temperature, T is in Kelvin in all equations):

(1) Calcite (0-500°C calibration; O’Neil et al., 1969 as revised by Friedman and O’Neil, 1977); 1000lnαcalcite-water = [(2.78 x 106)/T2] - 2.89; and(2) Siderite (45-75°C calibration: Zhang et al., 2001): 1000lnαsiderite-water = [(2.56 x 106)/T2] +1.69.

where α is the temperature dependent fractionation factor.

Zhang et al. (2001) reported that oxygen isotope fractionation between siderite and water over the temperature range of their study (45-75°C) was indistinguishable for abiotic and microbially mediated siderite precipitation. Mortimer and Coleman (1997) described an oxygen isotope geothermometer for microbially mediated precipitation of Mg-Ca siderite at low temperatures. Calculated temperatures using that relationship are, on average, 15-35°C lower than those from Equation (2). However Equation (2) is used in this study, because the resulting data are more consistent with paragenesis interpreted from other fractionation equations and independent textural and temperature data concerning siderite-water oxygen isotopic fractionation (e.g. Carothers et al., 1988).

For the three calcite cement samples from the Mahim Formation with high δ18O values (Fig. 11), the calculated crystallization temperatures range from ~15°C (δ18Owater = –3‰) to ~30°C (δ18Owater = 0‰) (Table 5). For the larger grouping of calcite cements with lower δ18O values, calculated temperatures range from ~60 ± 10°C (δ18Owater = –3‰) to ~80 ± 10°C (δ18Owater = 0‰) (Table 5). Both sample populations include early concretionary calcite and later pore-fi lling ferroan calcite. This suggests that some isotopic re-setting occurred during burial, and/or that later cement was added to the early concretionary calcite.

For CaMg-siderite samples from the Mahim Formation sandstones, the wide range of δ18O values is consistent with the variety of siderite morphologies and textures present. Most of the 18O-rich CaMg-siderite (δ18O = +24.1 to +27.5‰; ~35 ± 10°C for δ18Owater

Fig. 11. Carbon versus oxygen isotopic compositions of calcite and siderite cements from the Mahuva, Daman and Mahim Formations. The grey arrow indicates the general trend in isotopic compositions expected for carbonate cements produced during a progressive shift from diagenesis dominated by early and shallow microbial sulphate reduction to microbial fermentation and deeper abiotic processes. Blue-fi lled diamonds = ferroan calcite; green-fi lled squares = micritic and microspar calcite; black-fi lled triangles = siderite.

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= –3‰; ~55 ±10°C for δ18Owater = 0‰; Table 5) is micritic and occurs as a replacement of clay matrix or rimming grains in sandstone. This CaMg-siderite is inferred to have formed during slow burial in the Early Miocene over a period of several million years during which the temperature is unlikely to have exceeded ~50°C. The temperatures calculated for the earliest grain-rimming siderite are probably the most accurate, as later siderite may have precipitated from isotopically evolved pore water, for which the δ18O values are unknown. Most occurrences of macrospar and microspar CaMg-siderite (δ18O = +19.3 to +22.0 ‰) yield higher temperatures: ~70 ±10°C for δ18Owater = –3‰ and ~95 ±10°C for δ18Owater = 0‰. This range includes the single observed instance of CaMg-siderite that has overgrown earlier chlorite rims. It is therefore inferred from the combined textural and isotopic results that chlorite rims formed after early micritic siderite (average >35-55°C) but before later microspar and macrospar siderite (average ~70-95°C).

The equation of Mortimer and Coleman (1997) was initially used to estimate the formation temperature of one Mn-siderite sample (STE 1906m), assumed on textural and chemical grounds to be of very early diagenetic origin:

(3) Microbial Mn-siderite (18-40°C calibration; Mortimer and Coleman, 1997):

1000lnαsiderite-water = [(-0.98 x 106)/T2] + 33.4.

The calculated temperatures using this equation ranged from 30°C for δ18Owater = –3‰ to <0°C for δ18Owater = 0‰. While such temperatures are consistent with very early diagenesis, there is considerable uncertainty concerning the fractionation between microbially-mediated Mn-siderite and water at low temperatures. This sample also coexists with later ferroan calcite cement that has very similar carbon and oxygen isotopic compositions (Table 5). This siderite probably underwent post-crystallization isotopic exchange during burial diagenesis. Accordingly, isotopic temperatures are reported based on the equation of Zhang et al. (2001) in Table 5, assuming that such re-equilibration occurred. A similar situation exists for another texturally early siderite cement (MTA-1 2645.5m), which has isotopic compositions very similar to coexisting but paragenetically later ferroan calcite (Table 5).

The calcite δ13C values are consistent with a carbon source derived from seawater bicarbonate (~0‰) and carbon dioxide produced during microbial sulphate reduction which is typically –25‰ (Irwin et al., 1977). Methanogenesis was probably a source of some carbon; carbon dioxide resulting from this process is generally enriched in 13C while showing a wide range in isotopic composition, depending upon

the precise reaction (Whiticar et al., 1986). The most 18O-rich calcite cements tend to have lower δ13C values (Fig. 11), suggesting a stronger microbial role during their formation. Calcite cements with δ13C values > -11‰ occur in bioclastic sandstones, which suggests a greater contribution of carbon from dissolution of marine bioclasts.

Siderite is the principal carbonate cement in clay-rich intervals in sandstones of all three formations. Siderite is favoured where there is excess Fe2+ relative to sulphide (derived from sulphate in seawater), which is the case in intertidal sediments with a high riverine influx of iron (as is the case for the sediments described here). Consequently, the siderite may contain carbon derived from microbial sulphate reduction, methanogenesis, and deeper abiotic processes; the proportions acquired from the latter processes may be expected to increase with time and burial temperature. The broad pattern from 18O-rich/13C-poor to 18O-poor/13C-rich siderite in Fig. 11 is consistent with this model. Siderite from the Mahuva Formation has higher δ13C values than siderite from the Mahim Formation in well MTA-1, although for these samples higher δ13C values correlate with higher δ18O values. Assuming that higher δ18O values indicate lower crystallization temperature, this pattern may tentatively be linked to an earlier onset of methanogenesis in the Mahuva Formation.

Macro-quartz diagenesisQuartz overgrowths are volumetrically subordinate cements in the reservoir sandstones in the Tapti fields. Petrographic and SEM data indicate that the paucity of such quartz results from the inhibiting effect of the pervasive pore-filling and grain-rimming authigenic chlorite cements, an observation made for many reservoir sandstones (e.g. Ehrenberg, 1993; Adjukiewicz and Larese, 2012; Dowey et al., 2012; Bahlis and Ros, 2013).

According to the 1D basin model for well STB-1 (Fig. 5), the present-day downhole temperatures in the Mahuva, Daman and Mahim Formations indicate that temperature has not been a limiting factor in quartz precipitation; present-day reservoir temperatures are >100°C, higher than the threshold (80°C) for quantitatively significant quartz cementation (Giles et al., 2000; Worden and Burley, 2003). Fig. 5 shows that the reservoir sandstones in the Mahim Formation reached 80°C sufficiently long ago to have been cemented if quartz cement and nucleation sites were available (Mahuva Formation ~17 Ma, Mahim Formation ~7Ma). Quartz cementation in the few clean sandstones is ~10% of sandstone volume. Complete occlusion of porosity by quartz may have been limited by available silica sources. Dissolution of feldspar is frequently a major source of dissolved silica for quartz


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cementation during burial diagenesis (Bjorlykke and Egeberg, 1993) In the reservoir sandstones at the Tapti fields, however, there is <1% feldspar (as a result of the laterite source: Widdowson, 1997) and there is no evidence that substantial amounts of feldspar had been dissolved previously. What little feldspar present is partly leached, but it would not have provided sufficient silica to account for the observed volumes of quartz cement. An additional but also quantitatively minor potential source of silica may have been provided by the maturation of 0.7nm iron-rich clay to chlorite.

Time-temperature mineral relationshipsThe timing of chlorite rim cement formation relative to quartz cementation is believed to be critical to preservation of porosity in reservoir sandstones containing these cements (e.g. Ehrenberg, 1993; Ramm and Ryseth, 1996; Anjos et al, 2003; Bahlis and de Ros, 2013; Dowey et al., 2012). There is general agreement that chlorite rim cements originate as detrital clay coatings or iron-rich gels, and are found in similar facies to chlorite pellets (e.g. Ramm and Ryseth, 1996; Grigsby 2001; Bahlis and de Ros, 2013). In the Oligo-Miocene sandstones at the Tapti fields, chlorite rims in general pre-date chlorite pellet dissolution, which casts doubt on the likelihood of the rims having been derived from the pellets.

There is less agreement about the timing of chlorite rim formation. Aagaard et al. (2000) found chlorite rims only at depths below ~3000m in Haltenbanken, offshore Norway, from which they concluded that the clay must have precipitated at ~90°C. However, other factors such as facies may have controlled the depth distribution in Haltenbanken. Using oxygen-isotope data, Longstaffe (1986) and Ayalon and Longstaffe (1988) reported a wide range of temperatures for paragenetically early, grain-coating and pore-lining chlorite from Upper Cretaceous sandstones from Alberta. They noted that diagenetic chlorite or its precursors (i.e. odinite, berthierine) formed over a continuum from very early, shallow diagenesis at low temperatures to higher temperatures (75 - 110°C) associated with deeper burial. Also using stable isotope data, Grigsby (2001) determined precipitation temperature of 20-40°C for chlorite rims from the Oligocene Lower Vicksburg Formation, Texas. These studies used the empirically determined chlorite-water oxygen isotope geothermometer of Wenner and Taylor (1971), derived for hydrothermal systems. When extrapolated to lower temperatures, this geothermometer produces calculated temperatures that are on average ~50°C lower than the chlorite-water and berthierine-water geothermometers which have since become available for diagenetic systems (Cole et al., 1987; Savin and Lee, 1988; Hornibrook and Longstaffe, 1996). Ehrenberg (1993) also suggested

that chlorite rims form over an extended period of time, with continued growth being fed by dissolution of matrix and detrital grain coatings. Billault et al. (2003) demonstrated through SEM imaging of sections through chlorite rims that the chlorite particles increase in size from the first formed to the last on the outer face of the rim. They also interpreted this feature to indicate formation over an extended time period, thus accounting for the wide range of estimated precipitation temperatures.

Fig. 12 illustrates the temperature-water δ18O variation possible for the average <1µm Tapti chlorite (+9.8±0.7‰) according to the following four oxygen isotope geothermometers (T in Kelvin in all equations):

(4) chlorite-water (Savin and Lee, 1988; their composition c):103lnαchlorite-water = [(3.72 × 103)/T] + [(2.50 × 106)/T2] + [(–0.312 × 109)/T3] + [(0.028 × 1012)/T4] – 12.62

(5) chlorite-water (Savin and Lee, 1988; their composition e):103lnαchlorite-water = [(2.56 × 103)/T] + [(3.39 × 106)/T2] + [(–0.623 × 109)/T3] + [(0.056 × 1012)/T4] – 11.86

(6) chlorite-water (Cole et al., 1987):103lnαchlorite-water = [(3.47 × 106)/T2] + [(–5.79 × 103)/T]

(7) berthierine/odinite-water (Hornibrook and Longstaffe, 1996):103lnαberthierine-water = [(5.174 × 103)/T] + [(2.483 × 106)/T2] + [(–0.430 × 109)/T3] + [(0.039 × 1012)/T4] – 13.59

Assuming porewater δ18O values of –1.2 to 0‰ (i.e. unevolved seawater), the geothermometers suggest chlorite-porewater equilibration at ~85 ±10°C to ~110 ±10°C; somewhat higher temperatures would be implied (~110 ±10°C to ~120 ±10°C) if the berthierine-water geothermometer was employed. All calculated temperatures would be ~5°C higher if the average chlorite δ18O (+9.4 ±0.7‰) of the <2µm fraction was used in the calculation. Such temperatures are consistent with those reported for chlorite crystallization in such systems but are considerably higher than expected for chlorite formed early in diagenesis. In the case of the Tapti field sandstones, the temperatures are similar to or higher than those calculated for CaMg-siderite that coats the chlorite pellets and rims, assuming that the porewater δ18O value (–1.2 to 0‰) was the same for both minerals.

Uncertainties always accompany isotopic temperature calculations of this type. These uncertainties include whether the isotopic geothermometers represent the mineral-water system accurately, and whether isotopic equilibrium was achieved between the solid and liquid phases. That there is close


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agreement between the experimental (Cole et al., 1987) and empirical bond type (Savin and Lee, 1988) calculations provide some confi dence in the chlorite-water geothermometers used here (Fig. 12). Whether isotopic equilibrium between chlorite and porewater was established is more diffi cult to assess. The <1µm and <2µm chlorite size-fractions show the same range in δ18O values and their average compositions are not statistically different from each other; hence there is no obvious grain-size effect on the extent of (re)-equilibration. Nonetheless, inheritance of oxygen from the putative 1:1 layer odinite precursor requires evaluation. The oxygen isotope geothermometer of Hornibrook and Longstaffe (1996) for a mineral intermediate in composition between berthierine and odinite suggests that 18O is preferred in Fe-rich 1:1 layer clays relative to chlorite, other factors (temperature, water composition) being equal. Furthermore, oxygen inheritance would be most likely from the tetrahedral sheet, which is enriched in 18O relative to the octahedral sheet in clay minerals (e.g. Girard and Savin, 1996). In short, inheritance of oxygen by chlorite from a 1:1 layer precursor would increase its δ18O value, causing the calculated chlorite-water “temperature of formation” to be lower than the true temperature for complete isotopic exchange.

Assuming that the geothermometers are accurate and that isotopic equilibrium was established in the system, two main hypotheses emerge to explain the calculated chlorite formation temperatures:

(1) the early diagenetic porefl uids had δ18O lower than 0 to –1.2‰, as hypothesized earlier for the carbonate cements, and/or

(2) the chlorite oxygen-isotope geothermometer does not retain a low temperature signature of odinite formation, but rather records temperature-porewater conditions following the odinite-chlorite transformation.

As discussed above for the carbonate cements, lower porewater δ18O values could have arisen because meteoric waters had been intermixed with the seawater. The intertidal-estuarine depositional environment interpreted for the Mahim Formation allows for a meteoric water infl uence, especially in the context of monsoonal run-off. Thus pore waters with δ18O values as low as –3‰ may have been present during burial diagenesis. Such compositions would indicate chlorite crystallization at ~65 - 80 ±10°C, depending on the choice of chlorite-water geothermometer (Fig. 12). Higher temperatures nonetheless remain likely, because rock-water interactions during burial diagenesis would have caused enrichment in 18O of the

Fig. 12. Temperature versus average δ18O values of water for average <1µm chlorite (+9.8 ±0.7‰) using the geothermometers of Savin and Lee (1986), Cole et al. (1987) and Hornibrook and Longstaffe (1996) (the latter for berthierine). Red-coloured area illustrates the temperature range for isotopic equilibration of the <1µm chlorite with porewater having a δ18O of –1.2 to 0‰ (seawater), whereas the green-coloured area shows the same for a porewater δ18O of –3±0.5‰ (brackish/estuarine/evolved water). Using the average δ18O value of the <2µm chlorite size fraction (+9.4 ±0.7‰) would increase the calculated temperature by ~5°C for any given water δ18O value. Data are from the Mahim Formation.

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porewaters relative to its original composition when first trapped.

While the assumption of a δ18O value lower than that of normal seawater reduces the calculated temperature of chlorite pellet formation, this temperature remains much too high for crystallization very early in the paragenetic sequence. Instead it is proposed that the oxygen isotopic temperatures for chlorite carry the signature of prograde burial diagenesis, beginning at, and following, the thermal threshold at which Fe-rich 1:1 clay (e.g. odinite, berthierine) transforms to 2:1:1 Fe-rich chlorite. Hornibrook and Longstaffe (1996) showed that rims of berthierine-odinite intermediate minerals in the Clearwater Formation oil sands in Alberta had been preserved because temperatures during burial diagenesis did not exceed 70°C, and that transformation to Fe-chlorite appeared to begin at ~60°C. They noted more generally that sandstones containing Fe-chlorite rims had almost always experienced burial temperatures exceeding 70°C. Temperatures of 70-200°C have been proposed for the transformation to Fe-chlorite, with the lower end of this range most commonly cited (e.g. Velde et al., 1974; Iijima and Matsumoto, 1982; Jahren and Aagaard, 1989; Walker and Thompson, 1990; Longstaffe, 1993; Hornibrook and Longstaffe, 1996).

Fig. 13 places the diagenetic events and authigenic cements in the Mahim, Daman and Mahuva Formation sandstones in a time-temperature framework combining the results of petrographic observations and isotope geothermometry with the burial model.

Odinite clay formed syn-depositionally and during very early diagenesis by a process of verdinisation, replacing faecal or pseudo-faecal pellets and as grain coating rims, and was associated with the precipitation of early siderite cements. As initial burial rapidly compacted the sediment, pellets were squeezed to form pseudomatrix. Although textural relationships suggest that grain rims and coatings formed after the pellets, replacement of the precursor pelletal clay by chlorite may have occurred at the same time as the odinite rims precipitated. Most mechanical compaction took place in the first few hundred metres of burial.

The temperature of ferroan calcite spar crystallization is estimated to range from ~60 ±10°C (δ18Owater = –3‰) to ~80 ±10°C (δ18Owater = 0‰), which is inferred to equate to depths of ~1000m to 1500m, respectively (Fig. 13). For the Mahim Formation, this would have taken place during the Middle to Late Miocene. The inverse relationship between the abundances of ferroan calcite and compacted chlorite pellets may be interpreted in at least three ways:

(1) The pellets were compacted after the carbonate cement formed; if the calcite crystallization temperatures calculated using the stable isotope data are correct, this yields a rather high temperature (and

therefore a significant burial depth) for clay pellet compaction.

(2) Pellet compaction occurred in areas not cemented by carbonates, after the carbonate cement formed; the δ18O values of the calcite cements have then been variably re-set.

(3) The ferroan calcite and siderite macrospar preferentially cemented areas of high porosity and permeability where clay pellets were indurated and had not been compacted. Chlorite rims around mouldic pores in carbonate-cemented horizons indicate that the rims pre-date pellet dissolution as well as ferroan calcite and macrospar siderite cementation, while pellet dissolution post-dates Fe-calcite cementation.

Whichever of these scenarios is correct, clay rim recrystallisation from Fe-rich 1:1 layer odinite to 2:1:1 layer chlorite occurred prior to ferroan calcite cementation, at <80°C, and probably <60°C. Assuming that <60°C is the more realistic temperature, this is estimated to indicate that the clay rims precipitated in the Mahim Formation at a burial depth of less than 1000m during the Early Miocene (Fig. 13a). From this it can also be inferred that the presence of pore-filling ferroan calcite cement did not provide a barrier to subsequent transformation of this clay to chlorite. Ferroan calcite has also overgrown small quartz overgrowths, indicating that quartz cementation began at < ~60-80°C, but that widespread quartz cementation had not occurred at that time. Despite subsequent burial, a few samples of very early diagenetic calcite and siderite have retained oxygen isotopic compositions characteristic of crystallization close to the sediment-water interface at temperatures <35°C. Quartz cementation is suggested to have continued to the present day, but was inhibited by the presence of chlorite rim cements and the late gas charge (Fig. 13b). Gas generation took place in the surrounding kitchen areas in the late Pliocene and was available for migration over the last 3Myrs.

Chlorite and Reservoir QualityThe porosity and permeability data (Fig. 14) for the principal environments of deposition show that tidal ridge and bar sandstones have the best reservoir quality, while the sub-tidal sandstones (restricted to the Mahuva Formation) have the poorest reservoir quality among the sands analysed. The scatter in the porosity-permeability data results from variations in the proportions of authigenic clay pseudomatrix and total pore-filling cements (hence the high standard deviations in Table 7). The tidal ridge and bar sandstones have the lowest average contents of clay matrix and the highest porosity and permeability values. However, the lowest porosity and permeability values are not associated with the highest clay contents, but with the highest abundances of pore-filling cements (carbonates and quartz). From


Huggett.indd 403 11/09/2015 12:38:16

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Huggett.indd 404 11/09/2015 12:38:16

this it can be inferred that that micro-porosity in the clay is significant, both in terms of volume and connectivity, and that diagenetic modification is at least as significant a control of reservoir quality as is depositional environment.

Overall, permeability increases with increasing percentage of chlorite rims regardless of whether

the entire data set is considered (Fig. 15a) or only samples with <10% pore-filling carbonate (Fig. 15b). The influence of chlorite rims on reservoir quality may be partly masked by the fact that the clay rims are not found where compacted pellets are present. However, it is reasonable to infer that the chlorite rims in low matrix sandstone are insufficiently thick both

Fig. 14. Core-measured Klinkenberg-corrected permeability versus effective porosity for the principal depositional environments encountered in the Tapti fields.

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Page 43 of 48

For review only

Journal of Petroleum Geology

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Fig.15. (A) Permeability versus chlorite rim abundance (all data); (B) Permeability versus chlorite rim abundance (only samples with <10% carbonate cement). Both plots show that the highest permeability samples correspond to samples containing >3% chlorite rims.

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to block pore throats and cause a significant reduction in permeability.

The fact that in Recent sediments odinite-rich, verdine deposits are found nearshore and seaward of deltas and estuaries suggests higher rates of sedimentation than in the outer shelf, where glauconite is most commonly formed. Consequently, there is less chance of verdine pellets becoming sufficiently mineralised to resist compaction than there is for glauconitised pellets. Glauconitised faecal pellets may more frequently be at an advanced stage of maturation, and hence induration, when buried than are pellets replaced by odinite. This means that the presence of glauconite is less likely to result in poor reservoir quality than is odinite.

CONCLUSIONS

The Tapti fields Oligo-Miocene reservoirs are extensively cemented with grain rimming and pore-filling chlorite cements that formed as odinite cements and replacements of faecal or pseudo-faecal pellets. The reservoir quality of the sandstones varies from moderate, where detrital clay is sparse and chlorite pellets have been leached, and grain-rimming chlorite has inhibited quartz cementation; to poor where there is abundant clay matrix, compacted chlorite pellets and/or ferroan calcite cement. In this type of estuarine to shallow marine depositional setting, the presence of abundant iron-rich fine grained detrital material is likely to result in the authigenic development of chlorite grain coats and pellets, and exert a significant influence on reservoir quality by inhibiting the precipitation of quartz overgrowth cements. Preservation of porosity by early precursor chlorite clays may extend the porosity basement of Oligo-Miocence sandstones deeper into the Surat Depression.

Compaction of chlorite-rich pellets began early in the burial history of the sediments, and is the main factor in porosity loss. The other influences on reservoir quality in these sandstones are the distribution of authigenic chlorite and the proportions of detrital clay, compacted iron-rich clay pellets or pore-filling ferroan calcite. The firmness of clay pellets is interpreted to be a key control of reservoir quality in the Tapti fields. Firmness appears to depend upon the degree of iron-rich clay mineralization, which may be dependent upon either rate of sediment deposition or the texture and composition of the faecal pellet (i.e. the specific organism producing the pellets).

The authigenic chlorite is compositionally and mineralogically chamosite and is interpreted to have originated as pre-cursor odinite clay that replaced faecal or pseudo-faecal pellets. This chlorite formed at the site of major riverine iron influx, derived from the lateritic weathering of exposed Deccan basalts in

the source area hinterlands of the Western Ghats, and transported onto the shallow-marine shelf, possibly initially introduced during periods of relatively low sea level. The reaction path from precursor green clay to chlorite takes place by transformation of 0.7nm layer clay rich in magnesium and ferric iron, probably via one or more intermediate stages, to ferrous iron-rich 1.4nm layer chlorite, over a temperature range of 60-70 ±10°C.

Oxygen isotope data for the carbonate cements provide an approximate temperature framework for diagenesis of the non-carbonate cements, although post-depositional oxygen isotopic exchange has undoubtedly affected many samples. The oxygen isotope results for Fe-chlorite indicate equilibration temperatures of at least 55°C and most likely significantly higher. Such temperatures do not reflect original Fe-clay formation but recrystallisation reactions. The chlorite likely first precipitated as 1:1 layer Fe-clays such as odinite or berthierine, the oxygen-isotope signal of which was reset to higher temperatures during burial diagenetic transformation to Fe-chlorite.

ACKNOWLEDGEMENTS

BG Group is thanked for permission to publish. In particular, the support of Jim Brown, the former BG India Subsurface Director, in encouraging this technical work is gratefully acknowledged. We also thank Martin Gill for carrying out the XRD analyses, Tony Fallick for facilitating the carbonate oxygen and carbon isotope analyses, Lisa LeClair for assisting with the chlorite oxygen isotope analyses, and Alison Burley of Geology Graphics for expert drafting of figures. Financial support for the chlorite oxygen isotope analyses was provided by the Natural Sciences and Engineering Research Council of Canada and also made possible in part through the Canada Research Chairs programme. This is Laboratory for Stable Isotope Science Contribution #315. The authors extend their thanks to journal reviewers Pratul K. Saraswati, Peter Krois and Sadoon Morad whose constructive comments and suggestions much improved the paper.

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