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).
2
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
10
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
11
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
12
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.