1

DELTAIC SEDIME�TATIO� DURI�G CRETACEOUS PERIOD I� THE

�ORTHER� CAUVERY BASI�, SOUTH I�DIA: FACIES ARCHITECTURE,

DEPOSITIO�AL HISTORY A�D SEQUE�CE STRATIGRAPHY

1M.Ramkumar,

2V.Subramanian,

1D.Stüben,

1Institut für Mineralogie und Geochemie, Universität Karlsruhe, Germany. 2Department of Geology, National College, Tiruchirapalli – 620 001, India.

ABSTRACT

The Santonian-Campanian sequence of the Cauvery basin was documented with lesser detail owing to its lesser fossiliferous nature and relatively highly fossiliferous bounding strata. Micro-mesoscale lithofacies analysis coupled with documentation of sedimentary and tectonic structures, supplemented by bio and ichnofacies data of this sequence revealed that this sedimentary record represents the development of Gilbert type delta. Various stages of delta development were interpreted to have resulted during a third order glacio-eustatic sea level cycle. It is surmised that faulting at the dawn of Santonian that brought down topographic and structural highlands into lows permitted marine transgression and creation of steeply sloping river valley, augmenting intense continental erosion and influx of detrital sediment into the basin. With due course of time, smoothening of valley slope, submergence of river mouth by rising sea level coupled with cessation of detrital influx led to the demise of the deltaic deposition. The information that the bounding surfaces of this Santonian-Campanian sequences are recognised to be of sea level lowstands, that led to generation of good reservoir quality in the ensuing depositional products, when coupled with gas and oil pools in the sequence in offshore area of this basin necessitates intense exploration activities. This study has also indicated the presence of three types of variability of reservoir characteristics as defined by three systems tracts. Key words: Gilbert-type delta, Santonian, Campanian, Depositional environments,

Sequence analysis. I�TRODUCTIO�

Gilbert-type deltas are produced by progradation of alluvial or fluvial systems into

a standing body of water, either lacustrine (Winsemann and Asprion, 2001) or marine

(Sohn et al. 1997), usually in a basin with a steeply inclined margin. They show two

major types, viz., sand dominated and gravel dominated foreset deposits. The latter type

contains deposits with irregular stratification and abrupt changes in bed thickness,

grading pattern and grain fabric, which thwarted precise description of its facies

characteristics and evolutionary history. Predominance of inverse grading, lack of mud

matrix, common presence of well-sorted openwork gravel layers and lenses and

coarsening of clast size toward the downdip margin of a bed or the base of foreset slopes

in sedimentary successions were found to be helpful in recognizing Gilbert deltaic

sequences (Sohn et al. 1997).

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A more or less complete Barremian-Danian succession is exposed in the Ariyalur-

Pondicherry depression of the Cauvery basin. However, the onland exposures of the

Santonian-Campanian part have not been recorded systematically from this basin,

although equivalents of them are recorded in offshore regions (Govindan et al. 1996).

This is due to low topographic relief that makes distinguishing lithofacies successions in

the field difficult, thick soil cover that limits examination of fresh exposures and presence

of poor, patchy and weathered exposures making lucid description and stratigraphic

correlation of exposed strata difficult. The recent developments in infrastructure in the

area permitted access to exposures and the paper attempts to document the Santonian-

Campanian record in part of the Cauvery basin (Fig.1).

METHODS A�D MATERIALS

Systematic field mapping in the scale of 1:50,000 was conducted to collect data

on lithology, sedimentary structures, facies and faunal association (mega and ichno) and

tectonic structures from natural exposures, dug wells and mine sections. From these data,

lithostratigraphic setup of the study area has been compiled (after Ramkumar et al. 2002;

Table 1). Lateral and vertical variations of lithostratigraphic units in the Cauvery basin

are presented (Fig.2) and their geographic distribution is also shown in figure 3. The

stratigraphic information coupled with interpretations on depositional environments,

major geological events (such as faulting, sea level changes, depositional breaks, etc.)

with field checks and laboratory analyses, supplemented by available palaeontologic data

have enabled elucidating the depositional history and prevalent geologic processes. In

this paper, tectonic structure, gross lithology, stratigraphy and environmental

interpretation have been examined to cull out the depositional history of Santonian-

Campanian strata of the Cauvery basin.

TECTO�IC SETTI�G

According to Prabakar and Zutshi (1993), the Cauvery basin was formed during

Late Jurassic-Early Cretaceous and continued to evolve till the end of Tertiary through

rift, pull-apart, shelf sag and tilt phases. Seismic and deep well data revealed that the

basin consists of number of sub-basins (Fig.4), differentiated by many highs (Sastry et al.

1977; Kumar, 1983). Among them, the northernmost Pondicherry sub-basin contains

three mappable exposures. Based on the field structural and lithological criteria, contact

relationships and displacement features of strata, major tectonic movements of the

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Cauvery basin were interpreted following the standard procedures in terms of lithological

contact relationships, continuation/displacement of strata, change in attitude of beds and

other characteristics by Ramkumar et al (2002) viz., initial block faulting (F1 in Fig.3),

movement of fault blocks during Cenomanian (F2 in Fig. 3), reactivation of older fault

blocks and creation of new fault during Santonian (F3 in Fig.3) and reactivation of fault

blocks during post Danian-pre Quaternary (F4 in Fig. 3). Among these faulting events,

Santonian faulting event was considered to be intense after Late Jurassic-Early

Cretaceous faulting (Sundaram and Rao, 1986). It was also associated with beginning of

separation of Indian subcontinent from Madagascar and Antartica (Bandel, 2000). Fault

movement during Santonian brought the areas that remained topographic and structural

highlands since inception of this basin, due to which the Pondicherry sub basin became

wide and open shelf.

GROSS LITHOLOGY A�D STRATIGRAPHY

The Sillakkudi Formation that represents Santonian-Campanian age, consists of

three members namely, Varakuppai lithoclastic conglomerate member, Sadurbagam

pebbly sandstone member (both deposited during Santonian) and Varanavasi sandstone

member (deposited during Campanian). The Santonian deposits unconformably overlain

Anaipadi sandstone member of the Garudamangalam Formation and also offlap much

older Karai Formation.

The Varakuppai lithoclastic conglomerate member was termed as ‘lower beds’ of

Upper sandstone member (Sundaram and Rao, 1986) and ‘unnamed fluvial silty

sandstone’ (sensu Tewari et al. 1996). These beds are typically exposed in the regions

south of Alundalipur-northwest and north of Melarasur-south and west of Sadurbagam

particularly along stream and road sections. In a stream section, these are well preserved

and directly overlie Karai Formation with distinct erosional, angular unconformity. These

deposits show upward coarsening and cyclic large scale cross bedding. The erosional

intensity/quantum was so high and hence at places, they overlain directly the much older

Karai Formation clays. Three major palaeo channels were recognised with the help of

remotely sensed data and field mapping (Fig.3). Occurrence of large scale cross bedding,

mud drapes, fresh feldspar and sandstone clasts, made Tewari et al. (1996) to interpret

this unit as fluvial channel fill with some marine tidal influence during periods of

seasonal change in discharge. Similar to the facies sequence of Helvetifjellet Formation

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SE Korea, as reported by Gjelberg and Steel (1995), the lower portion of Varakuppai

member of the Cauvery basin is also devoid of marine influence and shows progressive

increase of marine influence towards top [Indicated by sedimentary structures and also by

an ichnofauna, thalassinoid (Photograph 1), typical of shallow coastal regions. Marine

influence during deposition has been inferred by George (2000) based on bioturbation in

Namurian strata of south Wales] signifying presence of a sequence boundary at the base

of this member. This member contains no insitu fossils except few thalassinoid burrows.

Subsurface sections of this member contain Marginotruncana coronata - Dicarinella

asymetrica zone (Govindan et al. 1996) of Santonian age. At places, in sub-surface it was

found to be unfossiliferous (Chidambaram, 2000).

The Sadurbagam pebbly sandstone member consists of coarse siliciclastics with

abundant marine fauna, shell fragments and varying proportions of calcareous matrix.

The beds are massive, thick-medium and parallel bedded. At places, pockets of shell rich

carbonate lenses with abundant siliciclastic admixture are found to occur that made the

earlier workers (Sundaram and Rao, 1986; Tewari et al. 1996) to interpret this member as

lower limestone member and Kilpaluvur grainstone member respectively. At the base of

this member, an erosional surface followed by distinct cobble-pebble quartzite

conglomerate is observed. Chandrasekaran et al. (1996) have recorded load casts, slump

folds, pillow structures and synerasis cracks in this member. The sandstone beds

frequently show normal grading and low angle cross bedding. Occasional development of

algal mounds, intimately associated with fault scarps/cliffs of older sedimentaries is also

found to signify this member. The sedimentary structural and lithological information

suggests a subtidal - intertidal clastic depositional environment for this member.

The Varanavasi sandstone member consists of planar sheet like sandstones,

constituted by massive, thick bedded, coarse-medium grained sandstones. These are the

most extensive lithologies in this part of the basin, developed mainly due to the constant

supply of sediment and stable shelf conditions that lasted for a long duration. They rest

over the pebbly sandstone member with non-depositional surface, may be due to flooding

of marine waters. Ayyasamy (1990) recognised a hiatus between Santonian (Sadurbagam

member) and this member based on ammonite zonation. These beds contain intermittent

fine pebble-coarse sand laminae, probably representing periods of higher energy and

sediment influx and show normal grading to re-establish deposition of thick, massive,

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medium sandstone deposits (Photograph 2). The sandstones are loosely packed and

cemented by calcareous material. Towards top, serpulid colonies are frequently observed

that thrived in a slowly waning sea (Ramkumar, 1997). Important fauna of this member

include Inoceramidae, serpulids, Turritella, Globotruncana arca (Cushman),

Globigerinelloides Cushman, Marginotruncana Marginata (Reuess), Whiteinella baltica

Douglas and Rankin and Archaeoglobigerina indicative of Campanian age. Govindan et

al. (1996) recorded typical Campanian benthonic foraminifera Bolivinoides culverensis

Barr and B. decoratus (Jones) in this member. This member also belongs to Karapadites

Karapadense ammonite zone (Sastry et al. 1968, 1972) and Globotruncana elevata,

ventricosa planktonic foraminiferal zone (Govindan et al. 1996). Upper surface of this

member represents erosional surface associated with a period of regression and was

recognised to represent a type 1 sequence boundary (Ramkumar, 1999).

FACIES SUCCESSIO� A�D DEPOSITIO�AL E�VIRO�ME�TS

Facies characteristics and depositional processes

Lithofacies variations of Santonian-Campanian sequence are presented in Fig.5

and are described herein in the context of deltaic facies development.

The bottommost unit consists of sandy gravel beds that reach upto 1.5 m thick.

The clasts are pebble-gravel sized. Bulk of the clasts is cobble to boulder-sized quartzite,

granitic gneiss and older sedimentary lithoclasts (Photograph 3). Long axes of the clasts

are oriented parallel-subparallel to the bedding plane. Unsorted coarse sand and find

pebbles constitute the matrix. The beds show inverse grading. Local concentrations of

cobble-boulder clasts (Photograph 4) that form lenticular bedforms traceable upto few

tens of meters are also observed. The lower contacts of these beds are erosional and have

contact with much older Karai Formation (basinal clay deposits). These bedforms in

association with lateral extent, sedimentary structures and geomorphology indicate that

these beds might have been deposited by turbulent fluvial flows, influenced by larger

drainage area provided by sea level lowstand. In view of fault generated steep slope that

energized fluvial channels, the channels eroded granitic gneiss clasts and sedimentary

rocks alike. Presence of fresh feldspar and lithified sedimentary lithoclasts indicate

predomination of physical weathering and associated erosion. Immature nature of the

clasts suggests proximity of source, short transport of clasts and juxtaposition of source

and depocentre (Singh and Rajamani, 2001). Predominant occurrence of inverse grading

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of gravel-bounder clasts in coarse sand-fine pebble matrix suggests a cohesionless

sediment flow in which clast interactions were dominant. Undulatory upper surfaces with

large protruding clasts and their positive relief suggest prevalent plastic flow behavior

(Sohn et al. 1997). The sediment flow can be described as a cohesionless debris flow or a

modified grain flow as the sandy interstitial material must have provided supplementary

grain support forces via increasing buoyancy, viscosity and excess pore pressure. The

lateral variations of grading pattern suggest flow unsteadiness and changes in flow

rheology along the length and/or width of the flow (Kim et al. 1995). The relative

thinness and common lateral lensing of individual beds as seen in transverse sections,

suggest that the flow was of small volume and moved for short distances on the foreset

slope with an overall tongue-like geometry.

A series of repetitive thin-medium bedded sandy gravel deposits overlie the

inversely graded gravel-boulder lithoclastic beds. The facies consist of sandy gravel

deposits and contain pebble-sized clasts forming a clast-supported framework filled with

a matrix of very coarse-granule sand. These beds could be traced for many tens of meters

in the field and show pinching out nature to be overlain again by a gravel-boulder sized

lithofacies. The facies differ from that of former in terms of lateral continuity, scarcity in

larger clasts, but show repetitive occurrence of inverse grading. The contact relationships

between these facies are non-depositional and bedding surfaces are even-parallel. The

inverse grading in relatively thin and steeply inclined sandy gravel beds suggests a

modified grain flow or a cohesionless debris flow dominated by clast collisions in a

medium of viscous and buoyancy providing sandy dispersion (cf. Postma, 1986).

Next younger deposits consist of a facies type with gravel clasts, but lesser in

extent and shows massive nature and irregular gradation. It consists of beds that extend

for few tens of meters. The clasts of pebbles dominate and contain rare to insignificant

quantum of gravel clasts. The bedforms die out suddenly, grade to pebble lags and also

show fining towards downslope direction. The beds are even-parallel bedded. The clasts

are randomly oriented. The flat lying or imbricate fabric, presence of lesser quantum of

sand matrix and thin sheet-like geometry and gradual downslope transition into pebble

lags could collectively suggest tractional or near-bed movements of gravel clasts either

individually or in an assemblage, gravitationally driven on a steep slope. Although the

several grain thick, clast supported fabric in the deposits suggests grain interaction during

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deposition, lateral continuity of most gravel sheets relative to the layer thickness and the

lack of lensiodal geometry and inverse grading indicate that individual clasts were

transported in a highly dispersed state with minimal grain collision (Campbell, 1989).

These could also be called debris fall (Nemec, 1990). This facies is followed by a

bedform that consists of well-sorted, fine to coarse-grained sand and sparse pebble to

boulder size clasts whose maximum diameters outgrow the bed thickness. This bedform

reaches a maximum of 60 cm thickness and shows lateral continuity for several meters

and pinching out nature. Individual laminae of this bedform are massive and show faint

lamination. These features suggest suspension settling from turbulent flow with a very

weak tractive transport. A typical chaotic gravel-boulder deposit, at the downslope region

that spawns for many tens of meters across and few meters thick, and has scoured base,

overlies this facies. This facies contains meter scale blocks of sand deposits derived from

intraformational sediments. These beds also show syndepositional deformation structures,

namely synerasis cracks, slump folds, load casts and pillow structures (Chandrasekaran et

al. 1996). The beds have randomly oriented clasts in a mud-poorly sorted sand matrix.

This facies is interpreted as deposits of slumps and very thick debris flows that were

generated by large scale failure of the gravelly and fine-grained foresets. Poorly sorted

and muddy matrix is probably due to incorporation of mud/finer clastics from the failed

fine-grained foresets. Syndepositional deformation structures can be attributed to

differential loading by prograding foresets over the fine-grained deeper water deposits

(Hwang and Chough, 1990).

Successive younger facies consists of massive, medium to very thick-bedded,

moderately sorted fine to coarse sandstone. The beds show local development of cross

bedding and fining upward sequence. The framework grains are loose-moderately packed

and cemented by calcareous material. Grains are represented by dominant proportion of

quartz and contain sparse-significant proportion of feldspars that vary regionally.

Unsorted, articulated shell fragments, bioclasts of mollusca, brachiopoda, ostracoda,

foraminifera and peloids also constitute significant proportion of the grains regionally.

Intercalations of lithoclastic beds (predominantly intraformational lithoclasts) that show

hummocky cross bedding are also observed representing episodic storm events (cf.

George, 2000). The faunal, lithofacies and sedimentary structures indicate deposition in

regions between intertidal and storm weather wave base. This member contains abundant

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inoceramids (Radulovic and Ramamoorthy, 2000) that might have thrived in relatively

deeper water environment (Tröger, 2000). This facies forms most widespread lithology in

this basin and represents long lasted detrital influx and fairly stable environmental

conditions. This facies shows gradual transition to sandstones with significant clay

admixture and sandstone with abundant insitu serpulid colonies that were interpreted to

represent maximum flooding and slow waning of palaeosea (Chandrasekaran et al. 1996;

Ramkumar, 1997) respectively. At top, the beds of clayey sandstone and serpulid

sandstone facies show significant erosional surface associated with regression during late

Campanian.

Depositional environments and development of Gilbert type delta

From the observations presented above, a sequence of events that led to the

development of Gilbert type delta is inferred and is presented herein.

Following the deposition of Anaipadi member, there was a major regression and

reactivation of basement faults resulting in movement of the shoreline towards former

offshore regions (as could be observed elsewhere when sea level drops – Miall, 1991;

Carter et al. 1991). The period of rejuvenated significant sedimentation after this

regression was primarily under fluvial regime and the fluvial agent was turbulent enough

to transport basement rock boulders, quartzite boulders and older sedimentary rocks

(Photograph 3) in addition to unsorted coarse sand-pebble sized grains into the new

depocentre. Presence of contacts between Sadurbagam member, Archaen rocks and

Varakuppai member affirms that, on flooding, the sea covered the areas that have not

been transgressed so far. Sundaram and Rao (1986) also stated that the regions that

remained structural and topographical high since inception of this basin were brought

under marine influence as a result of fault movement during late Santonian. This faulting

movement created steep slope, energized the fluvial systems onland, by which,

development of Gilbert type delta was also initiated.

The facies associations, bedforms and their characteristics of the Varakuppai

member suggest that the gravelly and sandy sediments near the edge of the transitional

zone between foreset and topset of the deltaic sequence might have been resedimented by

various triggering mechanisms and moved in cohesionless debris flows of various size

and volume. The cohesionless flow character might have been inherited from the

composition of the topset sediments, which were deposited in braided streams and

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shallow marine environments and are generally mud free (Hwang and Chough, 1990).

The lateral pinching out nature of gravel-boulder lithoclastic beds suggests that the flows

were generally small in volume and moved only for short distances. Therefore, sediments

of the foreset deposits are interpreted to have been recycled many times, i.e., they were

repeatedly remobilized and deposited on the foreset slope. Those deposits that were

directly derived from the topset-foreset boundary are thought to be restricted to the

uppermost part of the foresets. In contrast to the very thick, poorly sorted and chaotic

nature of the toeset deposits, the proximal part of the prodelta deposits mainly comprises

well sorted and well bedded sand deposits. There is a distinct break in grain size and

overall sedimentary features between the toeset and prodelta areas, indicative of distinct

change in depositional conditions, may be influenced by gradual smoothening of

depositional topography and energy conditions of fluvial systems.

The facies characteristics and succession as described in preceding sections,

suggest that the three lithostratigraphic members of the Sillakkudi Formation namely,

Varakuppai member, Sadurbagam member and Varanavasi member represent dominantly

fluvial, fluvio-marine and marine portions of a deltaic sequence respectively.

Depositional transition from steep fluvial channel deposit to estuarine mouth to intertidal-

fair weather wave base deposits with passage of time, associated with gradual reduction

in depositional topography and establishment of stable environmental conditions have

also been indicated from the facies succession. These three members of the Sillakkudi

Formation thus, represent deposition under fluvial and fluvio-marine environment

(Varakuppai member), marginal marine environment (Sadurbagam member) and shelf-

slope environment (Varanavasi member). As inferred through facies succession,

sedimentary structures and faunal assemblage, the entire sequence is bounded by type 1

sequence boundaries on either side. This inference, together with published data (Raju

and Misra, 1996; Raju et al. 1993), suggests that deposition of this sequence might have

been resulted during a complete third order cycle of sea level change.

SEQUE�CE DEVELOPME�T A�D ITS RELEVA�CE TO PETROLEUM

EXPLORATIO�

When a sedimentary record is analyzed for its sequence characteristics, bounding

surfaces and sub-divided into systems tracts/parasequences, based on the general

characteristics of sequences, the reservoir properties of each parasequence could be

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judged, with which applicable analogues may be traced at appropriate scale (regional,

intra and inter basinal scales). Such generalization of sedimentary record helps in

petroleum exploration (Ginsburg, 1994). The classic sequence stratigraphic studies such

as Vail et al (1977) utilized seismic data for recognition of sequences. However, the scale

of sequences recognised only through seismic survey thwart delineation of high-

resolution sequences and thus, many authors such as Grammer et al. (1996), George

(2000) and Spalletti et al. (2001) prefer outcrop-based studies. In this paper also sequence

analysis was performed based on outcrop studies instead of seismic data following the

procedures detailed in George (2000).

Sequence boundaries and parasequences

The contacts between Anaipadi member and Varakuppai member and Varanavasi

member and Kallankurichchi Formation are sequence boundaries. The three members of

the Sillakkudi Formation represent parasequences namely lowstand systems tract

(Varakuppai member), transgressive systems tract (Sadurbagam member) and highstand

systems tract (Varanavasi member). Depositional cycles of this Santonian-Campanian

sequence are thus bounded by widely traceable surfaces. The cycle boundaries are

unconformity surfaces, represented by lithostratigraphic boundaries and each depositional

cycle/systems tract fits into Walker’s (1992) definition of allostratigraphic unit. Many

authors, including Singh and Singh (1997), Sohn et al. (2001), Wignall and Newton

(2001) and Hiscott (2001) found that lithostratigraphic boundaries often tend to be

sequence boundaries. The lower sequence boundary is separated by an unconformity

surface created over shelf sands, followed by initiation of fluvial system advancement

and deposition of fluvial sediments over former offshore regions and thus, it is interpreted

to be Type 1 sequence boundary (cf. Sarg, 1988; 2001). In other words, as sea level might

have crossed shelf break [as inferred from change of depositional environment from shelf

sands (Anaipadi member) to fluvial channel and estuarine mouth deposits (Varakuppai

member)], it is inferred as sequence boundary type 1. The cycle for such a type 1

sequence is defined as starting at a relative lowstand with sea level located below the

shelf edge; the accompanying lowstand systems tract (LST) includes incised fluvial

valleys, shelf-margin deltas, sediment gravity flow deposits, submarine fan, slope and

lowstand shoreline sediments (Anderson et al. 1996; Carter et al. 1991). In the

sedimentary record under study, fluvial and estuarine mouth sediments of Varakuppai

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member represent it. This member was deposited consequent upon base level fall causing

fluvio deltaic channels to incise into delta plain and emergent shelf, thus increasing the

length of the alluvial feeder systems. The amount of incision was controlled by the

magnitude and rate of base level fall, the difference in gradient between the fluvio-deltaic

and shelf profiles and proximal to distal variations in stream power and sediment flux.

Enhanced channel incision occurred in proximal locations where the alluvial gradient was

less than the shelf gradient. Subsequently, during sea level rise, the shoreline crossed the

shelf break and migrated landwards across the shelf, accompanied by deposition of

shoreline-shelf facies of the transgressive systems tract [(TST) - Sadurbagam member, a

pebbly sandstone unit deposited along shoreline between supratidal-subtidal regions of

the advancing sea]. Finally, at the termination of sea level rise, the shoreline-shelf facies

of the highstand systems tract [(HST) - Varanavasi member that contains medium-coarse

shelf sandstones] prograded seawards across the maximum flooding surface (MFS).

These lowstand, transgressive and highstand systems tracts reach up to maximum

thickness of 45 m, 80 m and 270 m respectively. As could be commonly observed

elsewhere (Walker, 1990), sedimentary record of HST is thicker than TST (Pekar et al.

2000; Spalletti et al. 2001) in this basin too. Grammer et al. (1996) stated that cycle

thickness could be an indicator of amplitude of sea level change, based on which, the

marine flooding during highstand systems tract has been interpreted as long lasted steady

increase of relative sea level in this basin. Alternatively, gradual subsidence and

concurrent sea level rise could also be inferred. As the thickness of parasequences shows

increase from bottom to top, gradual increase in accommodation space created by these

processes is inferred.

Influence of sea level changes and tectonics on sequence development

The sea level curve of the Cauvery basin constructed independently by faunal data

(Raju and Ravindran, 1990; Raju et al. 1993), lithofacies and geochemical data

(Ramkumar et al. 2002; Stüben and Ramkumar, 2003) shows the presence of global sea

level peak at 85.5 MY (±1; Early to Late Santonian) within a third order sea level cycle

during Santonian-Campanian. The depositional history of the study area as enumerated in

previous sections also corroborates with the views of these authors. While the duration of

this 3rd order cycle (Santonian-late Campanian -- ~13 my.) and independent inference of

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sea level variations through faunal and depositional environments indicate deposition

under the influence of sea level fluctuations [which in turn were influenced by eustatic

control (Carter et al. 1991) as is also verified by the presence of global sea level peak in

these strata (Raju et al. 1993; Hart et al. 2000)], record of faulting at the dawn of

Santonian raises an iota of doubt over initiation of this having been influenced by

faulting. Cloeting (1988) and Miall (1991) also raised doubts regarding 3rd order cycles

such that they might be the consequence of tectonic forces. However, present

observations in terms of field survey coupled with data from earlier studies (e.g.

Sundaram and Rao, 1986; Bandel, 2000) clearly indicate that a sequence of events such

as regression at the end of Coniacian→hiatal period→establishment of fluvial systems

and initiation of fluvial deposition→faulting→energizing of fluvial systems and

increased detrital influx→slow increase of sea level →stable environmental conditions

upto the end of deposition of Sillakkudi Formation. This observation also indicates that

after initial tectonic movement, tectonic forces were either inactive or less active and thus

has not imparted any imprints over sedimentation. Hence it could be inferred that, the 3rd

order cycle has been a consequence of eustatic sea level cycle with coeval faulting at

early stages of the sea level cycle. This inference is in accordance with the views of Sarg

(1988) who stated that the influence of sea level fluctuations over marine sedimentation

systems overwhelm the influences of other variables such as tectonics, sediment influx

and climate. Spalletti et al. (2001) and Hiscott (2001) opined that it is a common

occurrence to have relative sea level changes are also contributed by local tectonics.

Implications of this study on hydrocarbon exploration

The deposition of Gilbert type delta during Santonian-Campanian, as enumerated

above, exhibits three levels of lateral variability in terms of reservoir quality properties

(porosity, grain size, packing and porefills and lateral and vertical extent etc). Three

systems tracts define these three levels. Among which, the lowstand systems tract shows

high lateral variability of reservoir quality while the highstand systems tract shows more

or less monotonous facies characteristics (George, 2000; Wignall and Newton, 2001) and

is also widespread in areal extent. The rapid transgression that took place during lowstand

systems tract, developed myriad varieties of facies successions, owing to which, when

reservoirs are located in this systems tract, high lateral variability of reservoirs could be

13

expected (George, 2000). Similarly, identification of lowstand systems tract signifies

possibility of lenses of porous and permeable down-slope gravity deposits and to

increased porosity and permeability produced beneath the subaerial unconformity on the

shelf or platform top (Ginsburg, 1994). Recognition of key stratal surfaces in coeval

strata of this basin thus would help predict stratigraphic positions and also aid in reservoir

characterization (George, 2000). Recognition of type 1 sequence boundary is important as

generation of this type of sequence boundary triggers substantial loss of sediments from

exposed terrains. Another process that occurs during formation of type 1 sequence

boundary is regional migration of fresh water lens in a seaward direction and consequent

upon this, precipitation of abundant meteoric cement will occur deeper in the phreatic

zone. Thus, primary reservoir properties of downslope regions might have been altered

owing to type 1 sequence boundary formation (Sarg, 1988).

Although presence of hydrocarbons in the Cauvery basin has been indicated

earlier, exploration activities have shown only a small sized reservoirs confined between

stratigraphic traps (Raju and Misra 1996). The exploration activities in this basin have

also indicated that, while significant source rocks were identified in Albian, significant

reservoir rocks are available in Campanian and Eocene (Govindan et al. 2000). Govindan

et al. (2000) have stated that siliciclastic deposits associated with regressive cycles form

bulk of reservoir facies of this basin. These authors have also suggested that generation of

hydrocarbon in this basin commenced during Santonian that reached its peak during early

Eocene. Integrating this information with the generalized reservoir characteristics based

on sequence analysis indicates that, the Varanavasi member (HST), deposited during

middle-upper Campanian could serve as potential reservoir for hydrocarbon. Govindan et

al. (2000) also arrive at the same conclusion and suggested to target this member for any

future intense exploration activities. This member could be recognised in addition to the

facies characteristics presented in this paper, also through its faunal content (as described

in terms of foraminiferal and ammonite zones) from bore well (onland and offshore)

samples too.

CO�CLUSIO�S

a. The depositional sequence represents more or less complete stratigraphic record of

hitherto less known Santonian-Campanian ages of the Cauvery basin, and the record

14

is represented by development of Gilbert-type delta under the influence of third order

sea level cycle.

b. The initial faulting during the beginning of Santonian had provided adequate slope to

establish/energize fluvial systems that brought large quantum of detrital sediments.

The initial transgression, followed by steady increase and subsequent fall of sea level,

that resulted in gradual reduction in depositional slope and deposition of gravel and

boulder deposits, followed by pebbly-gravely deposits to be overlain by coarse

sandstones. All these represent a complete deltaic sequence developed under a third

order sea level cycle of glacio-eustatic origin.

c. Bounding of type 1 sequence boundaries on either side of the Santonian-Campanian

stratigraphic record and recognition of parasequences that have typical characteristics

of such depositional cycles during this study, sheds light on probable reservoir

qualities of the study area. It would aid in exploration activities and field

development, when potential reservoirs are identified in this part of the Cauvery

basin.

ACK�OWLEDGEME�TS

Comments and modifications suggested by the reviewers have immensely helped the authors for lucid presentation of this paper. Prof.Dr.Jutta Winsemann, University of Hannover, Germany, is thanked for reading this manuscript and making suggestions that helped improve the style and content of this paper. Shri.T.Sreekumar, Department of Geology, National College, Tiruchirapalli, India, is thanked for his assistance during field survey. MR acknowledges the financial assistance by Alexander von Humboldt Foundation, Germany. Council of Scientific and Industrial Research, India extended financial support during fieldwork component of this study.

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19

Pla

te I

20

Explanations for photographs of Plate I

1. Field photograph showing insitu thalassinoid burrow. The burrows are found to

occur at upper portion of the Varakuppai member indicative of shallow marine

conditions of deposition. They also signify gradual transition of fluvial channel

deposits into shallow marine conditions under the influence of increasing sea

level that had submerged river mouths. It is also to be noted that except these

burrows, no other fossils are found in this member.

2. Massive featureless sandstones of Varanavasi member. Note feeble fining upward

sequences that are interspersed with pebble layers.

3. Lithoclastic boulder drawn from erosion of older sedimentary rocks found in the

Varakuppai member. Consolidated nature of these boulders indicate significant

time gap between cementation of the original rocks and erosion of them.

4. Field photograph showing upward coarsening gravel-boulder beds of the

Varakuppai member. In the foreground, much older Karai Formation clay

deposits could be noticed over which the boulder beds rest with a pronounced

erosional angular unconformity. Top of the photograph shows planar view of the

gravel-boulder bed.

21

TABLE A�D FIGURE CAPTIO�S

TABLE

Table �o. 1 Cretaceous-Tertiary lithostratigraphy of the Cauvery basin (after Ramkumar et al. 2002). Santionan-Campanian part of this comprehensive stratigraphic table forms the focus of this paper.

FIGURES

Figure �o.1 Sedimentary basins of India and location of the Cauvery basin (after Chandra, 1991). Present study is confined to the onland exposures found to occur in and around Tiruchirapalli district, located ~300 km south of Chennai.

Figure �o.2 Spatio-temporal variations of lithostratigraphic units along the traverses drawn in figure No.3. In the study area, oldest rocks are located at western margin of the basin that are aligned in left side of the figure No.3. In the figure No.2, oldest rocks are at bottom of the diagram. The Arabic numbers agains each box of the legend refer to the indexes of different members as given in figure No.3. Present study is concerned with 16 (Varakuppai member), 17 (Sadurbagam member) and 18 (Varanavasi member). Note the pinching out nature of Varakuppai member within short distance compared to all other stratigraphic units of the Cauvery basin and its contact with much older strata. The roman numbers at bottom of each lithological column refer to concerned traverselines drawn in figure No.3 in such a way that, the western part of the traverse is located at bottom of the figure No.2. Scale bar indicates the vertical exaggeration. Horizontal direction is not to scale.

Figure �o.3 Geology of the study area. Figure �o.4 Tectonic map of the Cauvery basin (after Prabhar and Zutchi, 1993). Figure �o.5 Stratigraphic succession, parasequences and relative sealevel of the

Santonian-Campanian strata of the Cauvery basin. Note that although the succession was deposited within a complete third order sealevel cycle, within the cycle, there are few perturbations which may be attributed to local causes such as tectonics, sediment inflow, subsidence, sealevel oscillation, etc. Within Varanavasi member (Highstand systems tract), there were periodic high-energy conditions that produced storm beds with higher grain size.