Journal of Asian Earth Sciences 40 (2011) 622–635

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Journal of Asian Earth Sciences

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The effects of diagenesis on the reservoir characters in sandstones of the LateCretaceous Pab Formation, Kirthar Fold Belt, southern Pakistan

Muhammad Umar a,b, Henrik Friis c,⇑, Abdul Salam Khan d, Akhtar Muhammad Kassi a, Aimal Khan Kasi d

a Geology Department, University of Balochistan, Quetta, Pakistanb Department of Environmental Sciences, Comsats Institute of Information Technology, Abbottabad, Pakistanc Department of Earth Sciences, Aarhus University, Denmarkd Centre of Excellence in Mineralogy, University of Balochistan, Quetta, Pakistan

a r t i c l e i n f o a b s t r a c t

Article history:Received 5 November 2008Received in revised form 18 October 2010Accepted 21 October 2010

Keywords:PetrographyClay mineralsDiagenesisProvenanceIndian ShieldTurbiditesShelfal delta lobeReservoir properties

1367-9120/$ - see front matter � 2010 Elsevier Ltd. Adoi:10.1016/j.jseaes.2010.10.014

⇑ Corresponding author. Address: Department of Easity, DK-8000 Aarhus C, Denmark. Tel.: +45 8942 256

E-mail address: henrik.friis@geo.au.dk (H. Friis).

The Maastrichtian Pab Formation in the southern part of Pakistan is composed of fine- to very coarse-grained texturally mature quartz arenite and subordinate sublitharenite varieties. The sandstones haveundergone intense and complex diagenetic episodes due to burial and uplift. Diagenetic modificationswere dependent mainly on the clastic composition of sandstone, burial depth and thrust tectonics.Diagenetic events identified include compaction, precipitation of calcite, quartz, clay minerals and ironoxide/hydroxide, dissolution and alteration of unstable clastic grains as feldspar and volcaniclithicfragments as well as tectonically induced grain fracturing. The unstable clastic grains like feldsparand lithic volcanic fragments suffered considerable alteration to kaolinite and chlorite. Dissolutionand alteration of feldspar and volcanic lithic fragments and pressure solution were the main sourcesof quartz cements. Mechanical compaction and authigenic cements like calcite, quartz and iron oxide/hydroxide reduced the primary porosity, whereas dissolution of clastic grains and cements has pro-duced secondary porosity. Chlorite coatings on clastic grains have prevented quartz cementation.Coarse-grained, thick bedded packages of fluviodeltaic, shelf delta lobe and submarine channels facieshave higher average porosity than fine-grained, thin bedded and bioturbated sandstone of deepershelf and abyssal plain environments and these facies are concluded to be possible future hydrocar-bon prospects.

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1. Introduction

The Pab Formation is part of the sedimentary succession (Ta-ble 1) in the Kirthar Fold Belt in southern Pakistan. Thick marinesiliciclastic successions of the Pab Formation are well exposed inthe study area which is approximately 350 km long and 225 kmwide (Fig. 1). The formation ranges in thickness from 50 m to450 m and is thinning toward the south. The formation is mainlycomposed of sandstone interbedded with subordinate mudstoneand marl. The sandstone is commonly light brownish grey, green-ish grey, yellowish brown, medium to very coarse-grained (inplaces pebbly), moderately to well sorted, subrounded to wellrounded quartz arenite (some sublitharenites). Marl is cream whiteand very light grey, finely laminated, and thin bedded. Mudstone iscommonly brownish grey, grey and bioturbated at places. The for-mation was deposited on the tectonically controlled western mar-gin of Indian Plate. Two contrasting depositional systems were

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identified (Khan et al., 2002; Umar, 2007), as shallow marine(Northern Depositional System) and fluviodeltaic to deep marineturbidites (Southern Depositional System). Three facies associa-tions were identified as transitional from proximal to distal set-tings as shoreface, shelf delta lobe and deeper shelf faciesassociations in the Northern Depositional System. The paleocur-rent was consistently from east to west in this part. Two successivedeep marine turbidite systems, sand-rich basin floor lobes andsand-rich slope fans, were identified in the Southern DepositionalSystem (Eschard et al., 2003, 2004; Umar, 2007). The lower partof the formation is regarded as sand-rich basin floor turbidite lobesin a distal setting (north) and their corresponding bypass surface inthe proximal setting. The upper part of the formation has beeninterpreted as sand-rich slope fan turbidite systems and is madeup of stacked slope channel and associated spill-over lobe deposits.The most proximal part of this system is exposed only to east-southeast (Fig. 1), and was deposited in fluvio-deltaic environ-ments (Section 11; Fig. 1).

The Pab Formation conformably overlies and underlies theMughal Kot Formation and Khadro Formation of the Rani KotGroup respectively. The Mughal Kot Formation is composed of dark

Table 1Stratigraphic succession of the Kirthar Fold Belt, southern Pakistan.

Age Group Formation Lithology

Holocene Recent–subrecent Mixture of clay, sand and gravel

UnconformityPleistocene Dada Formation Boulders and pebble conglomerates with subordinate coarse-grained sandstonePliocene Manchhar Formation Sandstone and shale interbedded with subordinate conglomerate

UnconformityMiocene Gaj Formation Shale, sandstone with subordinate limestone and conglomerateOligocene Nari Formation Sandstone interbedded with shaleEocene Kirthar Formation Fossilifereous limestone interbedded with shale and marl

Ghazij Formation Dominantly shale with minor sandstonePaleocene Rani Kot Group Lakhra Formation Intraclastic limestone and shale

Bara Formation Sandstone and shaleKhadro Formation Sandstone, shale and marl

Maastrichtian Pab Formation Sandstone intercalated with marl and mudstoneCampanian Mughal Kot Formation Marl, arenaceous limestone, mudstone and sandstoneEarly–Late Cretaceous Parh Group Parh Limestone Biomicritc limestone

Goru Formation Micritic limestone with shale, siltstone and sandstoneSembar Formation Shale, siltstone and marl

DisconformityJurassic Ferozabad Group Anjira Formation Limestone interbedded with shale and marl

Malikhore Formation Oolitic limestone with subordinate shale and marlKharrari Formation Limestone, shale, marl and minor sandstone

Base not exposed

M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635 623

grey, greenish grey and black marine shale with minor marl andlimestone. The Khadro Formation consists of greenish grey to darkgrey marine shale and limestone.

The primary reservoir targets for hydrocarbon exploration inmost parts of Pakistan are the Sui Main Limestone (Eocene) andthe Pab Formation (Maastrichtian) e.g., in Sui, Pirkoh, Loti, Dhodak,Jandran and Savi Ragha fields (Beswetherick and Bokhari, 2000;Dolan, 1990; Kadri, 1995; Sultan and Gipson, 1995; Hedley et al.,2001; Fitzsimmons et al., 2005). The source rocks of the Pab Forma-tion are marine shales of the Sembar Formation (Early Cretaceous)and the lower part of the Mughal Kot Formation (Campanian). Boththese formations comprise of dark grey to black shale, marl andoccasional limestone interbeds of marine environments. Cap rocksare mainly shales of Rani Kot Group (Paleocene) in the study area.

The study area (Fig. 1) is located within the Kirthar Fold Belt(part of the West Pakistan Fold Belt) (Bannert et al., 1992; Jadoon,1991) on the western suture zone of Indian and Eurasian Plates. Itprovides good outcrops of the Pab Formation which is consideredto have hydrocarbon potential and good trapping conditions. Dia-genetic studies of the Pab Formation have not earlier been carriedout in this part of Pakistan. This paper aims to provide a general ac-count of the diagenesis of sandstones of the Pab Formation andconcentrates on the following objectives:

1. the diagenetic evolution of the sandstones,2. the sequence of diagenetic processes and products with respect

to time and burial history,3. grain fracturing due to ophiolite thrusting and uplifting,4. assessing the effects of diagenesis on the composition and res-

ervoir quality of the sandstones.

2. Geological setting

The Kirthar Fold Belt has formed in response to the collision ofthe Indian continent with Eurasia (Powell, 1979; Bender and Raza,1995). Based on differences in tectonic style and stratigraphic vari-ations, the Kirthar Fold Belt may be divided into a number of smal-ler structural units (Fig. 2). Before collision of the Indian andEurasian Plates the Kirthar Fold Belt acted as passive margin untilLate Eocene time. Major continental collision was initiated in the

Early Eocene and was completed by Pliocene to Early Pleistocenetimes (Waheed and Wells, 1990). In Early–Late Cretaceous timethe western margin of the Indian Plate was separated from theMadagascar Plate, and the Indian Plate started a rapid movementtowards north (Scotese et al., 1988; Gnos et al., 1997) with anti-clockwise rotation towards northwest. During this time the Indianpassive margin was greatly affected by active normal faults andfragmentation into basins of different bathymetry.

When the Indian Plate was passing over the Reunion Hot Spotduring Late Cretaceous time, the Indian Shield area to the eastwas thermally uplifted and huge amounts of sand-rich sedimentwere supplied to the margin and deposited as the Pab Formationin a variety of tectonically controlled intra-slope basins. DuringPaleocene time a widespread transgression resulted in a reductionof the supply of coarse terrigeneous clastics to the basin, and pela-gic to hemipelagic shale and shallow marine limestone of the RaniKot Group were deposited. Emplacement of the Bela and Muslim-bagh ophiolites occurred on the western continental margin of theIndian Plate during the Paleocene (Allemann, 1979; Tapponieret al., 1981; Gnos et al., 1998). This affected the margin and a flex-ural foreland basin started developing in the west while passivemargin sedimentation continued on the eastern margin as theGhazij Shales and the shallow marine limestone of the Kirthar For-mation. The studied area subsided and deep marine clastic sedi-mentation of the Nari Formation took place in the Oligocene andpiled up additional overburden. Eventually, the collision betweenthe Indian and Eurasian plates resulted in the uplift of the Himala-yan mountain belt on the northern margin and the Sulaiman andKirthar mountain belts on the western margin. Compressionaldeformation continued until Pliocene–Pleistocene time and is re-corded in imbricated thrust sheets in the Kirthar Fold Belt (Niama-tullah et al., 1986). Then a major uplift caused exposure of theentire succession and deep erosion resulted in an unconformity be-tween the Gaj and the fluvial Manchhar Formation (Pliocene). TheManchhar Formation and the Dada Conglomerate were depositedin Pliocene–Pleistocene time in estuarine and fluvial environmentsrespectively.

The exact rate and amount of erosion and deposition above thePab Formation is difficult to estimate in outcrops because of com-plex tectonics, thrusting of ophiolites, and variations in palaeoto-

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Nari Formation (Oligocene)

Kirthar Formation/Ghazij Formation (Eocene)

Rani Kot Group (Paleocene)

Bela Ophiolites (Paleocene)

Pab/Mughal Kot Formations(Late Cretaceous)

Pahr Group (Early-Late Cretaceous)

Ferozabad Group (Late Triassic-Jurassic)

Afghan Block Lithostratigraphic units

Northern Depositional System

Southern Depositional System

Location of measured stratigraphic sections

Fig. 1. Geological map of the study area showing location of measured stratigraphic sections (modified after Bakr and Jackson (1964)).

624 M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635

pography of depositional basins with time and space. Keeping inview the above facts burial depths are estimated by the presentday overburden above the Pab Formation at various locations inthe study area based on data and maps of Hunting Survey Corpora-tion (1960) and Shah (1977). Estimated maximum burial depthsrange from 2743 m to 2916 m and 2704 m to 3252 m in the North-ern and Southern Depositional Systems respectively. Proximaldepositional settings of both systems have slightly larger burialdepths. Uplift occurred in the Paleocene–Eocene associated withophiolite emplacement. Uplift during the Miocene is marked byan angular unconformity between the Gaj and the Manchhar for-mations (Table 1). The fluvial sedimentation continued till theHolocene.

3. Methods

Twenty stratigraphic sections with continuous exposures weremeasured and sampled (Fig. 1). The description of primary andauthigenic mineralogy of the sandstones is based on the study of65 thin sections, including point counting and SEM, and X-ray dif-fraction (XRD) analyses. Polished thin sections were coated withcarbon using a Leica Emitech K950 Evaporator. Sandstone chipsfor SEI study were coated with gold. Scanning electron microscopy(SEM) was carried out using a Jeol JSM 6400, equipped with a linksystem Energy Dispersive X-ray microanalyser (EDAX). The sam-ples were examined in secondary electron (SEI) and backscatteredelectron (BSC) modes of imaging. The SEM–BSC micrographs were

Fig. 2. Map showing major tectonic features of the Kirthar Fold Belt, Pakistan, and location of the study area (modified after Bannert et al. (1992)).

M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635 625

analysed by Adobe Photoshop and stained thin sections were pointcounted to asses the sandstones porosity. XRD of sandstones andtheir clay separates was carried out using by PANanalytical X’pertPRO MPD. SEM, XRD and petrography were performed at AarhusUniversity, Denmark. The sandstone samples were disintegratedby ultrasonic treatment and tumbling. Clay separates (<2 lm) forXRD analyses were prepared by decanting and centrifuging. Afew Ml of the established >2 lm-fraction is smeared onto a glassplate and dried at room temperature. The plate-shaped clay miner-als will thus be oriented (with 0 0 1) parallel to the glass plate.These preparations are X-rayed as it is (untreated), after treatmentwith ethylene glycol vapours in a desiccator for 24 h at 60 �C, andafter heating to 500 �C for 1 h. Two-dimensional estimation ofporosity in sandstones was carried out using Adobe Photoshop ofSEM images.

4. Results

Thin section studies show that the sandstones are dominantlyfine to very coarse, moderate to well sorted, subrounded to wellrounded quartzarenites with some sublitharenites (according to

the classification of Folk, 1980). Nonundulose monocrystallinequartz is the major clastic component with minor feldspar, chertand volcanic lithic fragments. Accessory minerals in the sandstonesare glauconite, zircon, apatite and tourmaline. Opaque minerals arealso present in traces. The clastic grains are bound by cement. Car-bonate is the most common cement with subordinate amounts ofquartz overgrowths, iron oxide/hydroxide, and clay. The porosityranges from 3.5% to 15.53% with a mean value of 8.06%.

4.1. Clastic mineralogy

The clastic mineralogy was determined by point counting (Ta-ble 2). The average clastic composition of the sandstones is95.58% quartz, 0.88% feldspar and 3.53% total lithic fragments (in-cludes chert, volcanic and sedimentary fragments). Sandstones ofthe Northern Depositional System contain on average 97.57%quartz, 1.01% feldspar and 1.4% lithic fragments (no volcanics)and are classified as quartz arenite. Sandstones of the SouthernDepositional System contain on average 93.09% quartz, 0.68% feld-spar and 6.21% lithic fragments (mostly volcanics) and are classi-fied as quartz arenite and sublithic arenite. The amount of lithic

Table 2Petrographic composition of the Pab Formation. Clastic grains: Qm = monocrystalline quartz; Qp = polycrystalline quartz; Kf = alkalifeldspar; Pl = plagioclase; Lv = volcanic lithicfragments; Ls = sedimentary lithic fragments; Ch = chert; Hm = clastic heavy minerals; Cements/replasive: Cal = calcite; Ioc = iron oxide/hydroxide; Qc = quartz cement.

Sample No. Qm Qp Kf Pl Lv Ls Ch Calc Ioc Clay Qc Hm

1-19 74.2 4.4 0.2 0 0 0 0.6 0 3.0 4.8 12.4 0.42-1 77.8 3.4 2.4 0 0 0 2.0 0 1.4 2.4 10.2 0.42-2 76.6 2.0 0.8 0 0 0 3.2 0 1.2 3.4 12.4 0.42-5 80.2 3.2 0.4 0 0 0 1.4 0 1.0 3.4 10.0 0.42-6 80.2 0.2 0.6 0 0 0 4.2 2.0 Trace 3.0 9.8 02-8 76.6 2.2 2.4 0 0 0 2.2 0 1.2 3.0 11.6 0.82-10 81.2 2.0 0.8 0 0 0 0.8 0 1.0 2.8 9.4 2.03-3 70.0 0.2 0.2 0.6 0 0 0 0 18.4 6.0 3.2 1.43-4 89.4 2.0 1.0 0 0 0 0.8 0 Trace 2.6 4.2 03-6 61.8 0.2 0.4 0 0 8.6 0 0 20.6 2.6 5.2 0.64-1 74.4 0.6 1.4 0 0 0 0 0 18.8 1.2 2.6 1.05-1 66.8 0 0.4 0 0 0 0 30.2 Trace 1.4 1.2 Trace5-3 77.2 0 1.4 0 0 0 0.4 16.4 Trace 3.8 0.8 Trace5-5 60.8 0 2.8 0.2 0 0 0.8 31.4 Trace 1.8 1.8 0.45-6 87.0 0.4 0.4 0 0 0 0.4 2.0 Trace 1.6 7.6 0.66-2 60.0 0 0.6 0 0 5.2 0.2 30.8 0 1.4 1.8 06-4 76.6 0.6 0.2 0 0 0 0.2 18.4 0 2.0 2.0 06-5 67.4 2.4 0.6 0 0 0 0.4 24.8 0 2.6 1.6 0.26-6 87.2 0.2 0.2 0 0 0 0 7.0 0.2 1.2 4.0 Trace6-8 84.8 0 1.2 0 0 0 0 5.2 Trace 0.8 8.0 07-1 65.2 0 0.6 0 0 0 0 30.8 Trace 1.0 2.4 Trace7-4 82.2 0 0.6 0 0 0 0 15.8 0 0.2 1.2 Trace7-7 76.2 0 0.8 0 0 0 0 16.6 2.2 2.2 2.0 07-9 71.8 0 0.2 0.2 0 0.2 0.2 24.2 0.2 1.4 1.4 0.27-10 70.6 2.2 1.2 0 0 0 0.6 20.4 0.6 1.6 2.6 0.27-11 80.2 0 0.8 0 0 0.2 0 16.6 Trace 0.8 1.0 0.47-14 72.2 2.6 0.6 0 0 0.4 0 18.6 0 2.4 3.2 08-3 74.4 0.4 0.4 0 0 0 0.4 16.8 Trace 4.4 2.8 0.48-4 72.0 1.4 0.2 0 0 0 0 21.8 Trace 2.2 2.4 Trace8-6 69.0 1.2 0.2 0 0 0.6 25.2 0.4 1.6 1.8 Trace8-7 76.6 0.6 1.2 0 0 0 1.4 16.2 1.2 1.0 1.2 0.68-9 79.6 1.0 0.4 0 0 0 0.6 15.0 0.8 0.8 1.8 Trace8-11 63.8 0.4 0.6 0 0 0 1.2 21.4 9.8 1.2 1.6 Trace9-3 69.6 3.6 0.6 0 0 0 0.4 20.2 0 3.8 1.2 0.69-4 58.8 5.2 0.4 0 0 0 0.8 27.4 0 5.6 1.4 0.49-6 59.4 3.4 0.2 0 0 0 0.6 29.4 0 5.2 1.2 0.610-3 56.2 20.2 0.8 0 0 0 1.2 9.82 3.4 2.4 4.6 1.411-1 71.2 2.2 0.2 0 0 0 0.6 2.0 20.0 2.2 1.4 0.211-2 42.2 3.4 0.8 0 0 0 0.4 23.6 25.2 1.4 2.6 0.411-3 75.2 1.2 0.6 0 0 0 0 7.6 11.2 1.6 1.8 0.811-5 66.2 0.8 0.6 0 17.6 0 0 10.6 1.2 1.2 1.8 Trace13-2 61.6 1.4 0.2 0 10.6 0 0 20.2 0 1.4 4.6 Trace13-3 86.0 0.2 0.2 0 7.2 0 0.2 3.4 0 0.2 2.6 Trace14-7 72.2 1.4 0.4 0 7.2 0 0 15.2 1.2 1.2 1.2 Trace14-9 64.8 2.6 0.2 0 13.4 0 0.8 14.4 0.8 1.4 1.6 Trace14-10 67 0.4 0.6 0 7.6 0 0.2 19.6 1.4 2.2 1.0 Trace15-4 84.2 3.2 0.4 0 0 0 0 7.2 1.4 1.2 2.4 Trace15-5 63.6 0.2 0.6 0.6 27.6 0 0.2 3.8 1.4 0.4 1.6 Trace15-6 80.2 0.2 0.8 0.2 9 0 0 1.4 0.8 0.2 7.2 Trace16-2 84.2 1.8 0.4 0 0 0.2 0 0 4.8 6.0 2.2 0.416-5 63.2 2.6 0.6 0 0 0.4 0 28.2 0.4 0.8 3.6 0.216-7 88.2 0.2 0.4 0 1.6 0 0 2.4 0.6 2.6 4.0 Trace16-8 61.2 0.2 1.2 0.2 10.4 0 0 20.6 0.6 2.4 3.2 Trace16-11 52.8 0.4 1.2 0.4 11.2 0 0 30.0 0.2 0.6 3.2 Trace17-1 65.2 2.0 0.4 0 0 0 0.2 24.4 1.8 1.6 4.0 0.417-3 84.2 1.2 0.2 0 0 0 0.8 4.6 0 2.4 5.6 1.017-4 75.8 1.2 0.4 0 0.6 0 0 16.6 0 1.2 4.2 Trace17-5 83.0 0.4 0.4 0.6 3.6 0 0 2.6 0.4 2.2 6.0 0.818-1 76.2 2.6 0.2 0 1.6 0 0.2 13.2 0 1.6 3.4 1.018-3 75.0 0.4 1.2 0 3.6 0 0.4 16.6 0 0.8 2.0 Trace18-6 83.0 4.8 0.4 0 0 0 1.2 0 1.2 6.6 1.8 1.019-1 76.8 4.0 0.2 0 0 0 0 13.2 0 1.6 3.4 0.819-2 82.8 4.6 0.2 0 0 0.8 0 0 0 6 4.2 1.420-1 68.0 5.2 0.2 0 0 0 0 20.4 0.4 1.2 4.2 0.420-2 75.2 5.6 0.4 0 0 0 0 10.2 0 3.4 4.8 0.4

626 M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635

fragments (volcanic rock fragments) increases in the upper part ofSouthern Depositional System.

The most abundant clastic grains are quartz. Both monocrystal-line (42.2–89.4%) and polycrystalline (0–20.2%) quartz types are

present. The monocrystalline quartz is clean and shows domi-nantly nonundulose extinction, although in thin sections undulose(more than 50) extinction is also noted. The undulose quartz grainsdo not show common orientation, thus indicating that strain was

30 um30 um

100 um100 um

40 um40 µm

200 um200 µm

A

DC

B

Qtz

Qtz

QtzKfs

Qtz

Qtz

Qtz

Cal Cal

Kfs

Fig. 3. Clastic feldspar alteration. (A) Physically fractured potassium feldspar (Kfs) grain. Qtz is clastic quartz. SEM–BSC micrograph. (B) Kaolinitization of feldspar grain.Arrows pointing to kaolinite booklets. SEM–BSC micrograph. (C) Calcite replacement of rounded clastic feldspar grain. Arrows point to remnant parts of the clastic feldspar.Cal is calcite cement, Qtz is clastc quartz. SEM–BSC micrograph. (D) Partly dissolved clastic alkali feldspar (Kfs). Calcite cement (Cal) does not fill the secondary porosity alongcleavage planes (arrows). SEM–BSC micrograph.

10 um10 um 100 um100 um

BA

K

ChDol

Cal

Fig. 4. Clastic volcanic rock-fragment alteration . (A) Replacement by Fe-rich chlorite (Cl). SEM–BSC micrograph. (B) replacement by kaolinite (K) and marginally withsyntaxial dolomite (Dol) on calcite cement (Cal). SEM–BSC micrograph.

M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635 627

inherited from the source area. Polycrystalline quartz types withtwo, and with more than two crystals per grain are both seen inthe thin sections. But polycrystalline quartz with more than twocrystals is more frequent. The subgrain size is variable, even withina single composite grain of polycrystalline quartz. Most subgrainsare fine to very fine sand size.

Feldspar is present as a minor component, maximum up to 3.0%(Table 2). The observed varieties of feldspar include orthoclase,microcline and plagioclase. SEM and thin section studies show thatfeldspar may be highly fractured during compaction (Fig. 3A). Mostfeldspar grains show intense diagenetic alteration and replacement(Fig. 3B–D). The major part of the feldspar grains were altered tokaolinite (Fig. 3B) or chlorite. Calcite has also replaced feldspar par-tially or completely (Fig. 3C). In some cases the feldspar grainswere partially coated by clay rims. These grains have no calcite ce-

ment in their internal porosity, which may represent dissolutionalong cleavage planes (Fig. 3D).

Volcanic rock fragments are common only in upper part of theSouthern Depositional System. They constitute up to 27.6% in theuppermost part of the Pab Formation. They have undergone in-tense episodes of diagenetic alteration and dissolution. Volcanicfragments were altered to chlorite (Fig. 4A) or kaolinite (Fig. 4B).Some of them contain albite crystals. They probably represent dia-genetic albitization of the original Ca-rich plagioclase of basalticlithic fragments.

4.2. Compaction

The sandstones of the Pab Formation were subjected to intensemechanical and chemical compaction during its progressive burial

mm1 mm1

BA

0.5 mm

Fig. 5. A) Photomicrograph of compacted fabric of sandstone with little calcite cement. Arrows point to tight contacts between clastic quartz indicating some pressuresolution of quartz. The sandstone is strongly cemented by quartz (12%). Porosity of sample is 8%. Crossed polars. (B) Calcite cemented sandstone with limited compactionalloss of intergranular volume. Light grey is calcite cement (25%); dark grey is clastic quartz grains; black is porosity. SEM–BSC micrograph.

628 M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635

as evidenced from much loss in primary porosity of the sandstones.The compaction effect in the sandstones is also evidenced bystraight, concavo-convex and sutured contacts of neighbouringclastic grains (Fig. 5A). During compaction clastic grains were slid-ing past each other and packed into a tighter configuration. Thegrains penetrated into one another with increased force of over-burden and chemical compaction. Primary porosity in well sortedsands is in the order of 45% (Atkins and McBride, 1992). Initiallyall sandstones were subjected to some mechanical compaction tillcalcite cementation occurred. Massive calcite cementation ceasedthe effect of mechanical compaction (Fig. 5B). But the mechanicalcompaction continued in sandstones with little or rare calcite asindicated by their long and sutured grain contacts (Figs. 3A and5A).

4.3. Authigenic components

A variety of authigenic minerals were observed in thin sectionand SEM microscopy including quartz, feldspar, calcite, dolomite,

Table 3Relationship of clay mineral assemblage with estimated burial depths, porosity, clay mine

SN BD (m) n% K C I I–S

1-19 2893 7.42 p p2-2 2916 8.70 p3-5 2891 p p5-2 2792 7.84 p5-5 2745 10.61 p6-2 2812 8.04 p p6-5 2747 8.59 p7-1 2825 7.63 p8-2 2892 9.96 p p8-8 2810 9.25 p9-2 2743 2.77 p10-1 2824 10.11 p10-3 2770 p p11-3 3252 10.30 p p p11-5 3185 10.30 p p p13-2 2705 p13-3 2754 9.47 p14-7 2704 6.98 p p p15-4 2764 p p p p16-4 2720 p p16-5 2765 6.74 p p16-7 2809 p17-4 2724 p p17-5 2791 p p18-1 2777 8.83 p p p p19-2 2760 3.29 p p p20-1 2780 6.65 p

SN = Sample number; BD = estimated burial depth; n% = mean porosity in percentage;minerals present and IGV = Intergranular volume.

kaolinite, chlorite, illite/mixed layer illite–smectite, iron hydroxide,anatase, hematite and pyrite. Calcite, clay minerals, quartz and ironhydroxide are the main cement types identified in the sandstoneswith only minor dolomite in few samples.

The sandstones are mainly composed of monocrystallinequartz, which may carry quartz overgrowths (Fig. 7A). Quartz is acommon cement in the sandstones and it constitutes up to 12.4%of the whole rock volume as determined by point counting (Ta-ble 2). The amount of quartz overgrowth may be under-estimatedbecause of the small size of some overgrowths and because ofunrecognizable boundaries between some overgrowths and theirdetrital core. In porous sandstones quartz overgrowths are gener-ally euhedral, whereas they have compromise borders in stronglycemented sands (Fig. 5A). Quartz overgrowths are commonthroughout the formation, laterally and vertically. Large euhedralquartz overgrowths were formed in clean sands, preferentially insandstones of shoreface depositional setting (Umar, 2007), wherethe amount of calcite cement is low. The detrital cores may beclearly outlined by dust rims (Fig. 7A). The presence of authigenic

rals and depositional settings.

IGV Depositional settings

20.2 Shoreface facies association17.0 Shoreface facies association28.4 Shelfal delta lobe facies association24.0 Shelfal delta lobe facies association35.0 Shelfal delta lobe facies association34.0 Shelfal delta lobe facies association29.0 Shelfal delta lobe facies association34.2 Deeper shelf facies association25.2 Shelfal delta lobe facies association19.0 Shelfal delta lobe facies association25.2 Deeper shelf facies association20.2 Deeper shelf facies association22.2 Deeper shelf facies association22.2 Fluvio-deltaic facies association14.8 Fluvio-deltaic facies association26.2 Submarine slope channels-fan lobe facies association6.2 Submarine slope channels-fan lobe facies association24.2 Submarine slope channels-fan lobe facies association12.2 Submarine slope channels-fan lobe facies association29.9 Submarine fan lobe facies association33.0 Submarine lobe facies association9.6 Submarine slope channels-fan lobe facies association22.0 Submarine slope channels-fan lobe facies association11.2 Submarine slope channels-fan lobe facies association15.2 Submarine slope channels-fan lobe facies association10.2 Submarine basin floor and slope fan lobes and channels, turbidites.26.2 Basin floor fan lobe facies association

K = kaolinite; C = Chlorite; I = Illite and I–S = Illite–Smectite mixed layer, p = clay

0.5 mm

A

40 um40 um200 um200 um

DC

CementCement

CementCementCementCement

GrainGrain

GrainGrain

GrainGrain

GrainGrain

GrainGrain

GrainGrain

GrainGrain

GrainGrain

GrainGrain

GrainGrainB

Fig. 6. Microfractures in clastic grains. (A) Calcite in fractured clastic quartz grain (arrows). Photomicrograph, crossed polars. (B) Sketch of (A) alignment of fractures. Arrowsindicated inferred direction of maximum stress. (C) Strongly compacted and fractured sandstone with calcite cement in remnant porosity and inside fractures (arrows). Lightgrey is calcite cement; dark grey is clastic quartz. SEM–BSC micrograph. (D) as (C) with calcite cement enclosing booklets of kaolinite (arrows). SEM–BSC micrograph.

20 um20 um10 um10 um

A B

Fig. 7. Quartz cement. (A) Euhedral quartz overgrowths (arrows) embedded by calcite cement. Photomicrograph, crossed polars. (B) Euhedral quartz overgrowth, growthpartly being retarded by grain coating clays (arrows). SEM-SE micrograph.

M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635 629

clays, such as kaolinite, chlorite, and mixed layer illite–smectitemodified the overgrowth habit of quartz. Quartz commonly nucle-ates on clean, clay free parts of the clastic grain surface and thengrows outward and laterally to form overgrowth. Thus, irregularquartz cementation (Fig. 7B) has resulted where the presence ofclay obstructed the complete quartz overgrowth. Early calcite ce-ment reduced the primary porosity and prevented further quartzcementation as indicated by low quartz cementation in somesamples.

Albite is found in the interior of volcanic fragments as plagio-clase laths as determined by SEM–BSC and EDAX. Albite is alsodocumented by its 3.18 Å reflection in samples with high amountsof volcanic lithic fragments (Fig. 4) and is thought to be formed byalbitization of Ca-rich plagioclase within volcanic fragments. Nosuch albite was observed in samples without volcanic fragments.

Calcite is the most dominant cement in the sandstone and itranges from traces to 31.4% (Table 2). It is mainly poikilitic andmay replace clastic grains (Fig. 3C). Calcite is detected by the pres-ence of a strong 3.03 Å reflection on X-ray diffractograms in allsamples (except sandstones of shoreface facies), covering alldepths. Calcite occurs as intergranular cement (Fig. 5B) and ascomplete or partial replacement of clastic components. Calcite re-placed feldspar and quartz grains partly or completely, mostly attheir margins. Locally these grains were severely attacked and re-placed by calcite (Fig. 3C), and replacement even penetrated intothe cores of grains. In many sandstone samples authigenic calcitehas filled most secondary porosity after dissolution of feldspar.

Minor dolomite rims (Fig. 4B) were formed around and insidealtered volcanic lithic fragments. The dolomite crystals have arhombic shape. XRD shows a 2.9 Å peak confirming the presence

d-spacing (Å)40.0 20.0 6.0 0.40.8

4

3

2

500

CPS

1

B

A

B

B

B

C

C

C

C

A

A

A

Chl

orite C

hlor

ite

Chl

orite

Chl

orite

Qua

rtz

Qua

rtzQ

uartz

Kao

linite

Anat

ase

Kao

linite

Kao

linite

Kao

linite

Anat

ase

Cal

cite

Goe

thite

Chl

orite

I/S 0

01

I/S 0

02

I/S 0

03

I/S 0

05

Dol

omite

IIllit

e/sm

ectit

e

Illite

/sm

ectit

e

Illite

Fig. 8. X-ray diffractograms of < 2 lm fraction, air dried (A), ethylene glycollated (B) and heated to 500 �C (C). 1 is dominated by kaolinite with minor amounts of quartz,anatase, calcite and dolomite (Section 11, sample 11-5); 2 is dominated by chlorite with a minor amount of quartz (Section 13, sample 13-2); 3 is interstratified illite/smectitewith minor amounts of quartz and goethite (Section 10, sample 10-3); 4 is irregular interlayered illite/smectite, illite and minor kaolinite, with quartz and anatase. showingpeak positions of feldspar, mixed illite–smectite (Section 20, sample 20-2).

630 M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635

of dolomite. A small amount of siderite has been identified in fewsamples by its 2.8 Å reflection on X-ray diffractograms.

Authigenic clay minerals are kaolinite, chlorite, mixed layer il-lite–smectite and minor illite. Their occurrence is shown in Table 3.Diagenetic clay minerals are found in a wide variety of morpholo-gies in SEM. The presence of such clay minerals is confirmed bySEM, BSC, EDAX and XRD. Authigenic kaolinite is present in moststudied sandstone samples. It occurs as booklets and vermicularaggregates of stacked platelets as indicated by SEM (Figs. 3B, 4Band 6D). Kaolinite is the most frequent clay mineral, which is iden-tified by its 7.17 Å, 3.58 Å peaks on XRD (Fig. 8-1 and 4). The peak

10 um10 um

A

Fig. 9. A) Grain coating diagenetic mixed layer illite/smectite. SEM-SE micrograph. (B) Bmicrograph.

position was not influenced by treatment with ethylene glycol;they were reduced in size after heating to 500 �C but did not com-pletely disappear, thus confirming the presence of kaolinite. Kao-linite has formed along cleavage planes in clastic feldspar(Fig. 3B). It also formed by full dissolution of feldspar and occupiesoversized, irregular or elongated pores which may still containskeletal remnants of partially dissolved feldspar. Both K-feldsparand plagioclase are observed to be altered to kaolinite. Delicateeuhedral booklets, vermicular texture, high intercrystalline micro-porosity within patches of pore filling kaolinite indicate an in situdiagenetic origin of the kaolinite (Hurst and Nadeau, 1995). Kaolin-

6 um6 um

B

I

I

I K

K

KK

rushy and hairy illite (I) growing on partly modified kaolinite booklets (K). SEM-SE

50 um50 um3 um3 um

8 um 3 um3 um

A

DC

B

Fig. 10. A) SEM–BSC micrograph showing well developed anatase crystals (arrows). (B) Goethite with rosette structure. SEM-SE micrograph. (C) Randomly oriented goethite.(D) Hematite clusters. SEM-SE micrograph.

100 um100 um400 um400 um

A B

Fig. 11. Late stage iron-hydroxide impregnation along fracture zones (arrows) in calcite cemented sandstone. SEM–BSC micrograph. (B) Early pyrite (arrows) in intrafossilvoids. SEM–BSC micrograph.

M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635 631

ite dominates in sandstones from the Northern Depositional Sys-tem and may be the only diagenetic clay mineral in these sand-stones (Fig. 1; Table 2).

Chlorite occurs as grain-coating rims, rosette aggregates and insecondary pores left by the dissolution of unstable grains like vol-canic rock fragments and feldspar (Fig. 4A). It also partly replacesauthigenic kaolinite. Quartz grains with complete chlorite coatingsshow poor quartz overgrowths. Chlorite may be totally engulfedwithin later quartz overgrowth. Presence of chlorite is confirmedby its 14 Å, 7 Å, 4.7 Å, 3.5 Å and 2.83 Å reflections (Fig. 8-2). The7 Å and 3.5 Å are strong peaks with the 3.5 Å peak being the stron-gest. Chlorite peaks were not affected by treating with ethyleneglycol but heating 500 �C caused slight reduction of the 7 Å and3.5 Å peaks. In the heated sample, there is a distinct separationfrom the remnant kaolinite peaks, which have slightly higher Å-values than those of chlorite. Chlorite is strongly dominating in

samples from the upper part of the Southern Depositional System,and its occurrence is mainly coinciding with the occurrence of clas-tic volcanic rock fragments.

Mixed layer illite-smectite clays showing honeycomb structure(Fig. 9A) are also observed in few samples in the deeper shelf faciesassociations. In some samples, the mixed layer clay is a well or-dered 1–1 illite–smectite (Fig. 8-3). The basal reflections of theair dried sample at 25.5 Å, 12.5 Å, 4.97 Å and 3.16 Å expand on eth-ylene glycol treatment to 27.5 Å, 13.1 Å, 9.1 Å5.3 Å 3.3 Å. On heat-ing, it collapses fully to a 10 Å mineral. Other samples exhibit lessordering and may contain both illite and poorly ordered illite-smectite (Fig. 8-4). Illite is represented by three different texturetypes: hair like radially disposed crystals, fibrous and membra-neous (Fig. 9B), which occur as pore filling and pore lining cementas described by Cocker et al. (2003). Pore-filling illite occurs amongclastic grains and is partly included in quartz overgrowths.

632 M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635

Tabular crystals of anatase (Fig. 10A) are observed in most sand-stones. Anatase occurs in trace amount and is found in the porespace between clastic grains. Anatase is probably also recordedin many of the clay fraction separates by its reflection at 2.51 Å,although it cannot be verified by other peaks (Fig. 8-1 and 4).The amount of anatase in the clay fraction depends on the crystalsize.

The iron oxide and iron hydroxide cements vary from trace to25.2% (Table 2), although most sandstones contain only smallamounts. It occurs as interstitial pore filling and in some cases itmay completely fill most pores and strongly reduce porosity. Goe-thite is a common iron hydroxide mineral in the clay fraction of thesandstones as determined by its 4.18 Å reflection, which disap-peared on heating to 500 �C (Fig. 8-3). It may form rosettes(Fig. 10B) and is locally concentrated in clusters or distributed ran-domly (Fig. 10C). Minor hematite occurs in small clusters (Fig. 10D)and in the intergranular space among clastic grains. Hematite isidentified by its 2.7 Å reflection, which does not react on glycoltreatment or heating to 500 �C. These cements are quite rich inthe proximal depositional setting of the Southern Depositional Sys-tem from fluvio-deltaic environments. Some of the iron oxide andhydroxide was formed in oxidizing environmental conditions asearly diagenetic constituents. During Eocene–Pleistocene time,the area was strongly affected by uplift, folding and faulting. Suchuplift caused migration of dissolved iron resulting in precipitationof late stage iron oxide/hydroxide cements (Fig. 11A).

Pyrite is less commonly present in the sandstones as framboidaland cubic crystals (Fig. 11B). Scattered distribution of pyrite cubesin calcite cement within a shell indicates the formation of pyriteahead of precipitation of calcite cement.

4.4. Grain fracturing

In sandstones located close to the Bela Ophiolite thrust (Fig. 1),a large proportion of clastic grains are strongly fractured (Fig. 6A).This also applies to strongly cemented samples with shallowestburial depths. The major fractures were parallel to stress axis(Fig. 6B) with some diagonal, irregular microfractures which re-sulted from brittle behaviour of quartz. The stress direction is dif-ferent from stratigraphically up–down. Sandstones located awayfrom the ophiolite thrust are unfractured, even those with rare/lit-tle calcite and high remnant porosity, and buried at great depths.

5. Discussion

Sandstone reservoir quality is largely determined by diageneticprocesses that either reduce or enhance porosity. In the Pab Forma-tion, the most important diagenetic processes are compaction,cementation by quartz, calcite, chlorite, kaolinite and iron hydrox-ides. Compaction, both mechanical and chemical, has a large influ-ence on the intergranular volume. The intergranular volume ispartly correlated to the amount of diagenetic calcite, indicatingthat porosity reduction continued in sandstones with little or nocalcite cement, whereas massive cementation by calcite preventedlater compaction and preserved a larger intergranular volume.Mechanical compaction and authigenetic cements reduced theporosity and permeability, whereas dissolution of unstable clasticgrains and soluble cements increased porosity in sandstone. Thepresent porosity of sandstone varies from 2.7% to 15.5%, with anaverage of 8.0%. The primary porosity of the sandstones is reduceddue to intense mechanical compaction and due to filling of earlyauthigenic cements such as calcite, clay minerals, quartz and ironoxides. Sultan and Gipson (1995) who studied the Pab Formationin the Sulaiman Fold Belt (NNE of the study area), obtained slightlyhigher estimates of porosity (0–16%, mean approximately 9%) than

recorded in this study. They also indicated the same causes ofporosity reduction in sandstones in their study. They pointed outthat primary porosity was originally much higher but has beengreatly reduced by successive phases of cementation and mechan-ical compaction during diagenesis, and that porosity reductioncontinued due to mechanical compaction in sandstones with littleor rare calcite.

Long and sutured contacts between neighbouring clastic grainsof the sandstones resulted from intense mechanical and chemicalcompaction and is more intense in sandstones with rare or poorcalcite cement (Fig. 3A). The successive burial was responsiblefor reduction in porosity due to compaction. In some samples themesogenetic calcite cement reduced or prevented further compac-tion and the resulting close packing of clastic grains. On the otherhand the calcite cement reduced the porosity in sandstones appre-ciably (Fig. 5B).

Kaolinitization and chloritization also played an important rolein either reduction or preservation of porosity. Large amounts ofauthigenic kaolinite and chlorite have reduced the permeabilitybut preserve porosity as intercrystalline porosity (Figs. 3B, 4A, Band 9B). Grain-coating rims of chlorite (Fig. 7B) in some cases re-tarded or prevented in the formation of quartz overgrowths andthus preserved porosity.

An appreciable amount of dissolution of unstable grains tookplace during diagenesis of the sandstones. Most commonly dis-solved clastic grains are feldspar and volcanic fragments. Dissolu-tion can occur at shallow depths by meteoric groundwater(Bjørlykke, 1984; Mathisen, 1984) or at greater depths by fluidsproduced during diagenesis of organic matter or clay minerals (Sei-bert et al., 1984; Surdam et al., 1984). The dissolution of unstableclastic grains began in early stages of diagenesis which producedan empty pore volume for secondary cementation; some poreswere protected by clay coats and remained open. Another episodeof dissolution occurred in late stages and dissolved authigenic min-erals, mainly calcite, thus producing secondary porosity.

The Pab Formation is mainly sourced from the Indian Shield(Umar, 2007). When the Indian Plate was passing over the ReunionHot Spot during Late Cretaceous time, the Indian Shield started tosupply volcanic rock fragments from the Deccan Volcanism. Thiscaused a change in sandstone composition. The Northern Deposi-tional System does not contain volcanic rock fragments, whereasthe Southern Depositional System contains an upwards increasingamount of volcanic rock fragments (Table 2). Since the volcanicrock fragments are very labile components this shift in provenancehas a significant influence on diagenesis.

5.1. Paragenetic sequence

The inferred paragenetic sequence of the Pab Formation isshown in Table 4. The sandstones have undergone intense andcomplex episodes of diagenesis, including eogenesis, mesogenesisand telogenesis due to influence of depositional environment, deepburial and uplift. The paragenetic sequence is inferred with respectto time by SEM, XRD and thin section studies. The major diageneticevents include early mechanical compaction, dissolution of unsta-ble clastic grains like feldspar and volcanic fragments, kaolinitiza-tion, chloritization, quartz, carbonate and iron oxide/hydroxidecementation, chemical compaction, illitization, pyrite, hematiteand anatase formation.

The diagenesis of the Pab Formation sandstones was initiatedby early mechanical compaction. Compaction is documented by atight grain supported fabric of the sandstones. The compactioncontinued in all sandstones till the precipitation of massive calcitein some sandstones ceased further compaction and continued upto late stages in sandstones with little/rare calcite cement.Mechanical compaction continued from early to late diagenetic

Table 4Paragenetic sequence of sandstones of Pab Formation.

Relative TimingEvent Early Late

Mechanical Compaction Dissolution of feldspar and volcanic fragmentsKaoliniteChloriteChemical compactionQuartz cementIlliteCalciteDolomiteIron OxideGrain fracturingCalciteDissolution

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M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635 633

stages particularly in sandstones from shoreface facies deposi-tional environment (proximal settings) in the northern part ofthe study area. Mechanical compaction can be observed by physi-cal breakdown of feldspar grains (Fig. 3A). The dissolution andalteration of unstable grains such as feldspar and volcanic frag-ments to kaolinite (Figs. 3B and 4B) was the second important dia-genetic event in terms of relative timings. Partial to completekaolinitization of feldspar took place. The kaolinite is not replacingany authigenic minerals and is absent in secondary microfracturesin clastic grains. It is in situ at the former position of feldspar grainsin most cases, but exceptions do occur where partially dissolvedfeldspar grains have not been replaced by kaolinite. Chlorite wasformed in the next stage of diagenesis. Chlorite is mainly foundin samples with a large content of volcanic fragments (Fig. 4A),which indicate that it is an alteration product of such fragments.It formed earlier than quartz cement because it obstructed thequartz overgrowth.

Quartz cement had started to precipitate in open pores beforethe massive calcite cementation. This is demonstrated by the pres-ence of well developed euhedral quartz overgrowths, which areembedded in calcite (Fig. 7A). Calcite cement also corroded the ear-lier formed quartz overgrowth in such samples. The massive calcitecementation stopped further quartz cementation and mechanicalcompaction. But quartz overgrowth is very well developed in sam-ples with little or rare calcite cement (especially in Sections 1 and2) and shows its continuation to later stages. The proportion ofquartz overgrowth is moderate, generally only a few percent (Ta-ble 2) in the available pore space (primary porosity plus secondaryporosity provided by the dissolution or kaolinitization and chlori-tization of feldspar and volcanic fragments), and its growths ispartly influenced by the presence of earlier formed kaolinite andchlorite (Fig. 7B). Because chemical compaction was active at thattime, pressure solution may have provided the silica for quartzcementation. The presence and growth of illite on kaolinite indi-cates its later origin than kaolinite. During diagenesis sedimentswere subjected to different conditions which might activatesources of silica for quartz cementation, such as dissolution of feld-spar (Hawkins, 1978), pressure solution (Bjørlykke et al., 1986;Houseknecht, 1988; Dutton and Diggs, 1990; Bjørlykke and Ege-berg, 1993; Dutton, 1993; Walderhaug, 1994), replacement ofquartz and feldspar by calcite (Burley and Kantorowicz, 1986)and transformation of clay (Hower et al., 1976; Boles and Franks,1979). For every mole of K-feldspar altered to kaolinite, two molesof silica are released and made available for cement (Siever, 1957).The extensive dissolution of feldspar and volcanic lithic fragments,kaolinitization, chloritization and chemical compaction were thesuggested sources of the silica for quartz cementation.

Calcite is the most abundant cement. The calcite cementationtook place in two stages, i.e., prior to and after grain fracturing.The first stage of calcite cementation started slightly later thanquartz cementation as indicated by euhedral quartz overgrowthswhich are embedded and partially replaced by calcite (Fig. 7A).In samples with pervasive calcite cement further developmentof quartz overgrowth stopped, whereas quartz cementation stillcontinued in samples with little or rare calcite. Massive calcitecementation also ceased the compaction as demonstrated bythe large remnant intergranular volume (Fig. 5B). In contrast,samples with little or rare calcite cement are much strongercompacted (Fig. 3A). On the other hand, calcite cement also fillsporosity and microfractures formed by late stage ophiolitethrusting (Fig. 6C and D) and this indicates a second, late stagecalcite cementation.

During the alteration and dissolution of volcanic fragments, Mgions were liberated and reacted to precipitate dolomite over-growth on calcite cement into the dissolved grain void (Fig. 4B).

Chlorite precipitated through the alteration (Anjos et al., 2003)and dissolution of volcanic fragments (Klass et al., 1981) and alter-ation of kaolinite (Burton et al., 1987). Volcanic fragments mayhave contributed Mg, Fe and Si ions to the precipitation of chlorite.

Illite may be formed diagenetically by a number of processes.Among these the important ones are replacement of feldspar andvolcanic fragments particularly along thin cleavage planes and ill-itization of kaolinite. These mentioned two causes were responsi-ble in the formation of illite.

The titanium oxide content of pure quartz arenite is small. It issuggested that the titanium ions needed for the formation of theauthigenic titanium oxides are mainly derived from in situ alter-ation of clastic Fe–Ti oxide grains (Weibel and Friis, 2007). Biotitewas also a possible source of Ti ions (Mader, 1980, 1981; Turner,1980), for the formation of anatase in sandstones.

Pyrite probably was an early diagenetic product and probablyformed during the sulphate reduction phase. Hematite and pyritehave no significant influence on reservoir characters of sandstoneof the Pab Formation.

The iron oxide and hydroxide were precipitated in the laststages of diagenesis. Sources for iron oxide (hematite)/hydroxide(goethite) cements might be the liberation of iron by alterationof Ti–Fe oxides and volcanic fragments. Telogenetic iron hydroxidepartially replaced nearly all older cements in sandstones, so is laterstage cement (Fig. 11A).

The late diagenetic events such as clastic grain fracturing, disso-lution of grains and calcite cement and formation of iron oxide ce-ment were caused by uplift and thrusting of Bela Ophiolites. Thegrain fracturing (Fig. 6A–C) was not a result of burial compactionas indicated by least grain fractures in sandstones with great burialdepths 2916 and 3252 m, at greatest distance from the ophiolitethrust (Sections 1–3 and 11; Fig. 1). The obduction of the BelaOphiolites was responsible for various thrust faults present in thestudy area. The Bela Ophiolites were thrusted over the PaleoceneRani Kot Group and the Maastrichtian Pab Formation in the wes-tern most (distal) part of the study area and buried all the youngerrock units under the thrust. As the distance from the ophiolites in-creases, the intensity of grain fracturing reduces, even in sand-stones with high remnant porosity and buried at great depths.This indicates the tectonic origin of grain fracturing in thesandstones.

5.2. Potential reservoirs

Reservoir characters of sandstone depend on a variety of factorssuch as texture, composition, diagenesis, facies, thickness of beds,their lateral extent, vertical connectivity and bed geometry. Thestudied sandstones are medium to coarse-grained moderately to

634 M. Umar et al. / Journal of Asian Earth Sciences 40 (2011) 622–635

well sorted, with rounded to well rounded clastic grains. Feldsparand lithic fragments were the factors which retain and enhancedporosity. Lithic fragments are only present in the Southern Deposi-tional System (Table 2). These fragments are partly transformed toillite-Smectite and chlorite, which may form grain-coating rimswhich may retard quartz cementation (Figs. 7B and 9A). Thus, itmay be expected that sandstones from upper part of the SouthernDepositional System have preserved more porosity.The sand-rich,amalgamated, and laterally extensive packages of the successionwith good vertical connectivity are thought to have good potentialfor hydrocarbon occurrences. Reservoir characterization is largelyconcerned with the regional distribution of attributes, such as totalthickness and percentage of sandstones together with the internalarchitecture, heterogeneity and geometries of major sand bodiesand their constituent facies. Coarse-grained, thick sandstonesshowing massive, trough cross bedded, hummocky cross stratified,submarine channel fills and slope fan lobe facies show higher val-ues of porosity. The porosity range of shelf delta lobe, submarinechannel fill, deeper shelf and shore facies is 7.8–10.6%, 6.98–9.4%,2.7–10.1% and 7.4–8.7% respectively (Table 3). Fluvio-deltaic sand-stone show the highest average value as 10.3% (Table 3). Averageporosity of fine-grained, thin bedded, bioturbated facies of abyssalplain and deeper shelf environments are 3.2% and 2.7% respectively(Table 3). In the studied sections, the highest cumulative sandthicknesses are encountered in the submarine slope channels-fanlobe facies associations (Sections 13–17 and 18; Fig. 1) and in theshelfal delta lobe facies association (Sections 5, 6 and 8) of thesouthern and northern depositional systems respectively. Thehighest net to gross values occur in the thinner Sections 1 and 2shoreface sequence (94% sandstone), followed by Sections 5, 6and 8 shelfal delta lobe facies (87%, 89% and 89% respectively)and Sections 13–18 submarine slope channels-fan lobe succession(88.06%). The lowest sand percentages (46% and 35%) are observedin the deep shelf and abyssal plain fan lobe facies associations seenin Sections 7 and 20. However, the nature and geometry of thesandstones vary greatly among these facies associations. In theshoreface facies in the Sections 1 and 2, the thickest sand bodiesare cross bedded, highly lenticular and mutually erosive units, orstrongly bioturbated, semi tabular packets. In the mid-shelf andsubmarine slope channels-fan lobes sequences of Sections 5, 6, 8,13–17 and 18, the major sand bodies are stacked with maximumlateral dimensions (scores to hundreds of M) probably perpendic-ular to palaeoflows. These are likely to offer good vertical and lat-eral connectivity. However, pervasive bioturbation, especially inSections 4 and 10 has obscured the original boundaries of someof these compound bodies (and also has introduced significant tex-tural heterogeneity). In the Southern Depositional System, Sections13–18 the dominant sand bodies are deep marine channels andslope fan lobes, traceable laterally for several hundreds of M. How-ever, good vertical connectivity can be anticipated only in thestrongly amalgamated sandstone-rich packages, especially inthickening upward packages in the upper part of the sections. Interms of reservoir properties, thick sand bodies, composed of mas-sive and hummocky cross stratified sandstone, found in the mid-shelf sequences of the Northern Depositional System appear toprovide the most favourable external geometries, internal architec-ture and connectivity. Lateral and vertical intercalation of thesebodies with more fine-grained shelf sediments also offers addi-tional potential for fluid trapping and sealing, augmented in thebasal parts of these sequences by local channeling into the fine-grained substrate (Khan et al., 2002).

It is concluded that coarse-grained, well sorted, amalgamatedand thick packages of sandstones are more porous and have a goodlateral and vertical connectivity. Specifically, the packages of fluvi-o-deltaic, submarine channel and shelf delta lobe facies associa-tions are thought to contain potential reservoirs, whereas the

mud-marl dominated and bioturbated sandstone facies have poorreservoir characteristics.

6. Summary

Diagenetic signatures observed in the sandstones of the PabFormation include compaction, cementation, grain fracturing anddissolution. Sandstones composition, burial depth and upliftingwere the factors which influenced the diagenetic modifications.Major authigenic cements are calcite, quartz, iron oxide and clayminerals. The paragenetic sequence is identified with relative dia-genetic timings. The feldspar and volcanic fragments were severelyaffected during diagenesis as shown by their intense dissolutionand alteration to clay minerals. Early–Late dissolution of clasticand authigenic minerals/cements has created secondary porosityin sandstones. Compaction and calcite cementation were the maincauses of deterioration of porosity in sandstones. The Bela Opiolitethrusting was responsible for grain fracturing.

Cementation is most pronounced in proximal facies. Coarse-grained, well sorted, amalgamated and thick packages of sand-stones are more porous and have a good lateral and vertical con-nectivity. The deep water facies, mainly turbidite channels sandsand shelfal delta lobe sands have retained a high porosity andare expected to contain the best reservoir units within the PabFormation.

Acknowledgements

We acknowledge the partial financial assistance of the PakistanScience Foundation for this research. The principal authoracknowledges University of Balochistan for granting funds for a 6months split PhD. We appreciate the cooperation of Professor Ras-mussen, Anne Thoisen, Charlotte Rasmussen, Laboratory staff fortheir assistance in XRD analyses in Aarhus University, Denmark.The cooperation of Laboratory staff of Keele University, UK, andGeoscience Laboratory, Pakistan are appreciated for making goodthin sections and polished thin sections. An anonymous journal re-viewer is acknowledged for his very constructive and helpfulcomments.

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