Diagenesis and reservoir quality evolution of palaeocene deep-water, marine sandstones, the Shetland-Faroes Basin, British continental shelf

ARTICLE IN PRESS

Marine and Petroleum Geology 25 (2008) 514–543 www.elsevier.com/locate/marpetgeo

Diagenesis and reservoir quality evolution of palaeocene deep-water, marine sandstones, the Shetland-Faroes Basin, British continental shelf H. Mansurbega,, S. Morada,b, A. Salemc, R. Marfild, M.A.K. El-ghalie, J.P. Nystuenf, M.A. Cajad, A. Amorosig, D. Garciah, A. La Iglesiai a Department of Earth Sciences, Uppsala University, Villava¨gen 16, SE 752 36 Uppsala, Sweden Department of Petroleum Geosciences, The Petroleum Institute, P.O. Box 2533, Abu Dhabi, United Arab Emirates c Faculty of Education at Kafr El-Sheikh, Tanta University, Kafr El-Sheikh, Egypt d Departmento Petrologia y Geoquı´mica, Facultad de Geologia, UCM, 28040 Madrid, Spain e Geology Department, Al-Fateh University, P.O. Box 13696, Libya f Department of Geosciences, University of Oslo, P.O. Box 1047 Blindern, NO-0316 Oslo, Norway g Department of Earth Sciences, University of Bologna, Via Zamboni 67, 40127 Bologna, Italy h Centre SPIN, Department GENERIC, Ecole Nationale Superieure des Mines de Saint Etienne 158, Cours Fauriel 42023, Saint-Etienne, France i Instituto de Geologı´a Econo´mica (CSIC-UCM), Facultad de Geologı´a, UCM, 28040 Madrid, Spain b

Received 12 May 2006; received in revised form 3 May 2007; accepted 9 July 2007

Abstract The Palaeocene, deep-water marine sandstones recovered from six wells in the Shetland-Faroes Basin represent lowstand, transgressive and highstand systems tract turbiditic sediments. Mineralogic, petrographic, and geochemical analyses of these siliciclastics are used to decipher and discuss the diagenetic alterations and subsequent reservoir quality evolution. The Middle-Upper Palaeocene sandstones (subarkoses to arkoses) from the Shetland-Faroes Basin, British continental shelf are submarine turbiditic deposits that are cemented predominantly by carbonates, quartz and clay minerals. Carbonate cements (intergranular and grain replacive calcite, siderite, ferroan dolomite and ankerite) are of eogenetic and mesogenetic origins. The eogenetic alterations have been mediated by marine, meteoric and mixed marine/meteoric porewaters and resulted mainly in the precipitation of calcite (d18OVPDB ¼ 10.9% and 3.8%), trace amounts of non-ferroan dolomite, siderite (d18OVPDB ¼ 14.4% to 0.6%), as well as smectite and kaolinite in the lowstand systems tract (LST) and highstand systems tract (HST) turbiditic sandstone below the sequence boundary. Minor eogenetic siderite has precipitated between expanded and kaolinitized micas, primarily biotite. The mesogenetic alterations are interpreted to have been mediated by evolved marine porewaters and resulted in the precipitation of calcite (d18OVPDB ¼ 12.9% to 7.8%) and Fe-dolomite/ankerite (d18OVPDB ¼ 12.1% to 6.3%) at temperatures of 50–140 and 60–140 1C, respectively. Quartz overgrowths and outgrowth, which post- and pre-date the mesogenetic carbonate cements is more common in the LST and TST of distal turbiditic sandstone. Discrete quartz cement, which is closely associated with illite and chlorite, is the final diagenetic phase. The clay minerals include intergranular and grain replacive eogenetic kaolinite, smectite and mesogenetic illite and chlorite. Kaolinite has been subjected to mesogenetic replacement by dickite. The K-feldspar and plagioclase grains have been albitized. Dissolution of calcite cement and of framework grain (feldspar, volcanic fragments and mud intraclasts) has resulted in a considerable enhancement of reservoir quality. r 2007 Elsevier Ltd. All rights reserved. Keywords: Diagenesis; Turbidites; Reservoir quality; Tertiary; Shetland-Faroes Basin; Sequence stratigraphy

1. Introduction Corresponding author. Present address: TOTAL, DGEP/GSR/TG/ ISS/CLAS CSTJF, Avenue Larribau, 64018 Pau Cedex, France. Tel.: +33 5 59835048; fax: +33 5 59836382. E-mail address: [email protected] (H. Mansurbeg).

0264-8172/$ - see front matter r 2007 Elsevier Ltd. All rights reserved. doi:10.1016/j.marpetgeo.2007.07.012

In the past decade, deep-water turbidite sandstones along passive continental margins have become increasingly in focus as the future hydrocarbon reservoir targets.

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The impact of diagenesis on reservoir quality of such sandstones is, thus far, relatively poorly explored in the literature. Reservoir quality, which is controlled by depositional facies and subsequent modifications by diagenetic alterations, is one of the critical aspects in understanding the basic elements of the play in the Shetland-Faroes Basin (Ebdon et al., 1995; Johnson and Fisher, 1998). The diagenetic evolution pathways of sandstones are governed by several interrelated parameters, including composition of framework grains, pore-water chemistry, tectonic setting of the basin, and burial-thermal history of the succession (Morad et al., 2000; Stonecipher, 2000). Variations in framework composition (most importantly the types and amounts of intrabasinal clasts) and porewater chemistry may be influenced by changes in the relative sea level, which occur owing to eustatic sea-level changes and/or the tectonic uplift/subsidence. Numerous successful attempts have been made to link diagenetic alteration to sequence stratigraphy of paralic and fluvial deposits (Dutton and Willis, 1998; South and Talbot, 2000; Ketzer et al., 2003; Al-Ramadan et al., 2005). Conversely, links between the mineralogical and geochemical evolution and sequence stratigraphy of deep-water marine clastic sediments are still far from fully explored in the literature, due perhaps to the uncertainty of constructing of sequence stratigraphic model for these deposits. The aim of this paper is to decipher the conditions and parameters controlling the spatial and temporal distribution of diagenetic minerals and their impact on reservoirquality evolution pathways during progressive burial of the Tertiary (Middle to Upper Paleocene) sandstones from the Shetland-Faroes Basin (SFB) within sequence stratigraphic context. Diagenetic regimes used in this paper are

515

(1) eodiagenesis (o70 1C; deptho2 km), during which pore-water chemistry is controlled by surface waters (i.e. depositional and/or meteoric waters) and (2) mesodiagenesis (470 1C; depth 42 km), which is mediated by evolved formation water and elevated temperature (Morad et al., 2000). 2. Geological setting and depositional environment The SFB is ca. 125 km wide and 600 km long and located between the West Shetland platform to the east and the Faeroe Islands to the west (Fig. 1). The basin is a major Cretaceous-Tertiary depocentre (Dore´ et al., 1997) and contains a series of NE-trending sub-basins (Fig. 2) formed by a complex tectonic history involving multiple phases of extension and volcanism (Carr and Scotchman, 2003). The area is five times that of the prospective part of the North Sea Basin, but mainly because of great water depths (ca. 2.5 km), it has been only slightly explored (Spencer and Eldholm, 1993). Fault-block rotation during the Late Jurassic–Early Cretaceous was followed by major subsidence that resulted in the accumulation of 5 km thick sediment package during the Cretaceous (Johnson and Fisher, 1998). Major continental breakup during the Late Cretaceous to Early Tertiary, accompanying the northward propagation of seafloor spreading and the opening of the North Atlantic resulted in extensive volcanic activities (White, 1989). Like for the North Sea Basin, hydrocarbon occurrences in SFB are divided into pre-, syn- and post-rift plays (Johnson and Fisher, 1998). The post-rift successions, which are the target for this study, represent deep-marine sediments that have been deposited during the Early Tertiary (middle to upper Paleocene). The main source rock in the SFB is Late

250 km

nk

tt

on

Ha

Ba

Hatton Basin

Faroe Islands

eill on yv s r W om sfe Th ran T

n ea Oc t y n ne ndar nti Co bou

Rockall Trough

B

FS

Shetland

Scotland

Rockall Bank FSB FaroeShetland Basin Water depth 0-2 km >2 km

Igneous centre South and east limits of basalt Well locations

Faero-Shetland Basin Newly designated acreage

Fig. 1. Location map of the Shetland-Faroes Basin and the wells included in this study (Modified after Brooks et al., 2002).

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516

NW Faroes UK

Schiehallion/ Foinaven play

Victory play

Clair play

SE 0

4 ?

? 6

Depth (km)

2

8

50 km

10 Lower Eocene to Recent

Jurassic

Upper Paleocene to Lower Eocene

Permian to Triassic

Paleogene basalt

Pre-Permian undifferentiated

Paleocene (study target)

Basin-floor sandstone

Upper Cretaceous Lower Cretaceous

Fig. 2. Schematic play diagram for the Shetland-Faroes Basin. The Shetland-Faroes Basin has successful plays in pre-, syn- and post-rift beds. Up-dip pinch-out of post-rift Paleocene basin-floor sandstones are the target of this study (Modified after Brooks et al., 2002). k

Pa

N

Fm

0

1000 Depth (m)

Jurassic Kimmeridge equivalent formation (Bailey et al., 1987; Scotchman et al., 1998; Carr and Scotchman, 2003) with lacustrine Middle Jurassic mudstones forming a secondary source for oil in the Foinaven sub-basin (Scotchman et al., 1998; Lamers and Carmichael 1999; Carr and Scotchman, 2003). Paleocene deposits were triggered by a major uplift of the Scottish Highlands, the Shetland Platform and presumably Greenland owing to a major underplating event (Sørensen, 2003). The burial history curve of the Tertiary deposits (Fig. 3) was dominated by rapid Early Paleocene subsidence, volcanic activities and Oligocene–Miocene compression (Carr and Scotchman, 2003). The Late Cretaceous–Early Paleocene exhumation was followed by rapid Paleocene subsidence with resultant accommodation space being filled with turbidites (Roberts et al., 1999), whereas during the Late Paleocene, the Icelandic plume resulted in basin uplift (Clift, 1999). Sequence stratigraphic studies recognized a complex basin margin, with marked basinward and landward shifts in sedimentation determining the location of sand-rich depositional systems (Mitchell et al., 1993). Ebdon et al. (1995) recognized a major sequence boundary (SB) at the base the Late Paleogene and subdivided the succession into several stratigraphic sequences, based mainly on maximum flooding surfaces (MFSs), which can be correlated with equivalent surfaces in the North Sea Basin. The Paleocene reservoirs in the SFB are comprised of submarine-fan sandstones deposited as early lowstand systems tracts (LSTs) (DTI, 2003). Factors that influenced sand distribution and trap formation in the Tertiary deposits include (Ebdon et al., 1995): (1) depositional limits of the Early Tertiary submarine fans were influenced by the underlying syn-rift, fault-block topography, which was accentuated by differential compaction of the thick, intervening Late Cretaceous mudstones, (2) fan deposition was initially aggradational (during the Paleocene), followed by periods of progradation (post Paleocene), (3) intra-Paleocene tectonic events

Nordland

Hordland

2000

Flett Lamba U. Velle

3000

150

100

50

0

Age (Ma)

Fig. 3. Burial history curve for well 204/19-1 (modified after Carr and Scotchman, 2003). The Tertiary sequences subsidence undergone during early Paleocene followed by an uplift event in the Late Paleocene and Early Eocene. It is not clear whether the uplift phase resulted in emergence deposits.

contributed to slope instability and the emplacement of large gravity slides, (4) the dominant northwesterly dip in the Late Paleocene and younger successions reflects regional thermal subsidence, which is mirrored by an oppositely dipping slope on the northwestern side of the basin, and (5) Late Tertiary structural inversion (Boldreel and Andersen, 1994; Dore´ et al., 1997), which was marked by two regionally significant compressional events (Early Oligocene and Late Miocene). The Paleocene succession reaches a maximum thickness of about 3.5 km and contains excellent quality sandstone reservoirs within the basin-floor fan, slope fan and shelf facies. In the SFB, aggradational basin-floor sandstones characterize the lower part of the Paleocene, and are likely sealed by both lowstand and highstand mudstones (DTI, 2003). The overlying Upper Paleocene to Lower Eocene

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sandstone includes a greater proportion of shelf and deltaic deposits, with a potentially high risk of imperfect seals. It has been documented that Palaeogene sedimentation in the SFB was cyclic (Mitchell et al. 1993; White and Lovell 1997). The cyclicity is related to uplift or flooding events caused by magmatic underplating and/or global sealevel changes. The stratigraphic scheme (Fig. 4A and B) shows a generalized compilation of the proposed Paleocene sequence stratigraphy of the studied part of SFB. Boundaries between major sequences or sequence sets are marked on Fig. 4B. Stratigraphic surface recognition (SB, paleosols and transgressive surfaces) is based on lithofacies analysis, changes in fining and coarsening-upward patterns, and trace fossil content. The LST was formed during a significant sea-level fall, which resulted in extensive subaerial exposure and/or widespread fluvio-deltaic deposits on the shelf (thick paleosol and/or incised valley deposits, respectively). The LST on the shelf is bounded below by a SB, which formed by subarial erosion, and

517

above by the first major marine flooding surface (FS). Transgressive systems tracts (TSTs) along the slope and basin floor are characterized by upward-deepening successions of proximal-to-distal turbidites overlain by finegrained, hemipelagic deposits, which mark the MFS. The SB can be traced along proximal parts of the basin to the shelf based on extensive paleosol development. Paleosols were identified as fine-grained, calcite-cemented mustones. High stand systems tract (HST) is overlained by a SB, and characterized by an aggradational to progradational parasequence set. It forms during the late part or a stillstand of a sea-level rise, or during the early part of a sea-level fall. The base of this systems tract is formed by the MFS over which the HST sediments prograde and aggrade. 3. Material and analytical methods One hundred and forty-five representative sandstone core samples (in terms of facies and coverage of the range of present-day burial depths) were selected from nine wells

Fig. 4. (A) Simplified sequence stratigraphic model for SFB. (B) Cross-sections and graphic logs showing the different systems tracts: LST (lowstand systems tract), TST (transgressive systems tract), HST (highstand systems tract) and sequence stratigraphic surfaces: SB (sequence boundary) and MFS (maximum flooding Surface).

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drilled in the SFB. Each sample was vacuum impregnated with blue epoxy resin prior to thin-section preparation. Modal compositions of the sandstones were obtained by counting 500 points in each thin section. Studies of occurrence crystal habits and paragenetic relationships were performed on gold-coated sample chips using a JEOL SEM instrument equipped with an energy dispersive X-ray analyser. The chemical composition of minerals was determined in 72 polished, carbon-coated thin sections using a Cameca BX50 microprobe equipped with three spectrometers and a back-scattered electron detector (BSE). Operating conditions were: 20 kV acceleration voltage, 8 nA (for carbonates and clay minerals) to 12 nA (for feldspars) measured beam current, and a 1–10 mm beam diameter (depending on the extent of homogeneous areas). Standards and count times were: wollastonite (Ca, 10 s), orthoclase (K, 5 s), albite (Na, Si, 5 and 10 s, respectively), corundum (Al, 20 s), MgO (Mg, 10 s), MnTiO3 (Mn, 10 s) and haematite (Fe, 10 s). Precision of analyses was better than 0.1 mol%. Stable isotope analysis of carbon and oxygen were performed on 28 calcite, dolomite and siderite-cemented sandstone samples; data are reported in per mill relative to Vienna Pee Dee Belemnite (V-PDB). Most of the samples contain multiple carbonate phases, which necessitated the use of sequential chemical separation (Walter, 1972). For this purpose, the bulk samples were powdered (o200 mesh) and reacted with 100% phosphoric acid at 25 1C for calcite and at 50 1C for siderite, dolomite and Fe-dolomite/ ankerite (AI-Aasm et al., 1990). The CO2 released from calcite, Fe-dolomite/ankerite and siderite was collected after 1 h, 2–3 and 6 days, respectively. Although this is not a precise method, it allows a minimum cross-contamination effect based on sequential extraction of CO2. The evolved gas for each carbonate fraction was analysed using a SIRA-12 mass spectrometer. The phosphoric acid fractionation factors used were 1.01025 for calcite (Friedman and O’Neil, 1977), 1.01060 for Fe-dolomite/ankerite, and 1.010454 for siderite (Rosenbaum and Sheppard, 1986). Precision (1s) was monitored through daily analysis of the NBS-20 calcite standard and was better than 70.05% for both d13C and d18O. Due to the presence of zonation and, in some cases, of more than one generation of the same carbonate mineral in the same sample, the isotopic data should be considered as average values. However, we have attempted to select samples containing dominant cement generation following careful CL and BSE examination. Authigenic pyrites from 2 sandstone samples were isotopically analysed. The method of the Cr reduction procedure of Canfield et al. (1986) was used to extract the pyrite-S from each of the crushed sandstone samples. The pyrite-S extracts were reacted with excess CuO at 1070 1C (Robinson and Kusakabe, 1975) and the isotope analysis of the SO, produced was determined using a VG SIRA II mass spectrometer with values reported as d34S relative to the Canyon Diablo Troilite (CDT) standard.

Fluid inclusions in authigenic carbonates were examined in 14 double-polished 100 mm thick sections using a Linkham TH600 stage calibrated for the temperature range between 100 and 400 1C. X-ray diffraction analysis of the less than 2 mm fraction in 15 of the sandstone samples were performed on air-dried samples and subsequent to ethylene glycol and heating at 550 1C. 4. Results 4.1. Composition and provenance of the sandstones The sandstones are very-fine to coarse-grained, moderately to poorly sorted, subarkoses, lithic subarkoses and sublitharenite (Table 1; Fig. 5). The quartz grains are mostly slightly undulous, monocrystalline (28–52 vol%). Polycrystalline quartz (4–12 vol%) is both of plutonic and metamorphic origins and is most abundant in the mediumto coarse-grained sandstones. Overall, K-feldspar dominates over plagioclase (1–7 vol% and o1–4 vol%, respectively). There are no pronounced variation trends in the amounts of K-feldspar and plagioclase with depth. The rock fragments are igneous (volcanic and plutonic; 1–12 vol%) and metamorphic (o1–4 vol%) in origin, with subordinate amounts of sedimentary fragments (sandstone and siltstone;o1 vol%). Mica occurs in variable amounts (1–8 vol%); overall biotite dominates over muscovite. Grains that occur in small amounts (o1 vol%) include chalcedony, chert, carbonate bioclasts, glaucony and heavy minerals (zircon, epidote, tourmaline, rutile, apatite, garnet and Fe–Ti oxides). The glaucony grains are either fresh, greenish in colour or reveal various degrees of oxidation. 4.2. Diagenetic minerals 4.2.1. Calcite cement Calcite cement attains a wide variety of occurrence habits, including concretions, scattered patches and/or extensively cemented sandstone layers, and as paleosols. Overall, calcite cement shows no systematic variation in distribution among the depositional facies, except that calcite concretions are most abundant in the shelf sandstones in the vicinity of paleosol surfaces. The concretions have spherical to oval shapes and vary in diameter from a few millimetres up to estimated 20 cm (Fig. 6). The concretions contain, in some cases, borings and burrows (commonly ophiomorphs, planolites and sophicus). The concretions are most common in poorly sorted, thinly laminated and fining upwards sandstones and in mudstones from the proximal LST turbidites. Calcite cement in the concretions is dominantly micro- to cryptocrystalline (o1–5 mm in size). Calcite crystals that coat the grains are coarser (up to 20 mm long) display a fibrous habit. Some of the concretions contain patches (ca. 2 mm across) that are occupied by fibrous calcite (ca. 200 mm long).

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Table 1 Modal and average composition of detrital and diagenetic components of representative Palaeocene sandstones Detrital and diagenetic components

Wells, depths (ft), and number of samples (n)

Monocrystalline quartz Polycrystalline quartz K-feldspar/ perthite Plagioclase Sedimentary rock fragments Igneous rock fragments Metamorphic rock fragments Biotite Muscovite Chlorite Heavy minerals Mud intraclasts Carbonate intraclasts Carbonate bioclast Opal/chalcedony Diagenetic feldspar pseudomorph Intergranular kaolinite Replacive kaolinite Intergranular illite Replacive illite Intergranular chlorite Repalcive chlorite Intergranular calcite Replacive calcite Intergranular Fedolom./ankerite Replacive Fedolom./ankerite Siderite Quartz overgrowths Discrete quartz Feldspar overgrowths Discrete feldspar Pyrite Pseudomatrix Intergranular porosity Intragranular porosity

47.2

39.3

6.3

Total

208/19-1, depth range 8036.5–8059.1, n ¼ 21

206/2-1A, depth range 11048–11103, n ¼ 17

206/1-1A, depth range 7637.5–7695.0, n ¼ 23

214/27-1, depth range 13706.4–13816.5, n ¼ 22

214/28-1, depth range 14297.0–14349.7, n ¼ 36

204/19-1, depth range 8806.2–8824, n ¼ 26

Mean S.D.

27.4

21.8

47.3

47.8

38.5

11.3

4.9

4.3

2.2

3.5

4.1

4.2

1.4

3.4

4.1

5.4

2.4

4.9

4.5

4.1

1.1

1.8 0.4

4.6 0.3

4.3 0.1

1.1 0.1

1.7 0.2

1.0 0.5

2.4 0.3

1.6 0.2

2.9

2.7

4.0

0.6

1.4

3.4

2.5

1.3

2.4

2.2

3.0

0.4

2.2

1.5

2.0

0.9

0.1 0.0 0.0 0.2 0.9 0.0

0.5 0.9 0.2 0.2 0.6 0.0

2.8 1.1 0.3 0.3 16.8 0.0

0.0 0.1 0.0 0.1 1.9 0.0

0.3 0.6 0.0 0.0 8.1 0.2

0.0 0.3 0.0 0.0 4.8 0.0

0.6 0.5 0.1 0.1 5.5 0.0

1.1 0.4 0.1 0.1 6.2 0.1

0.1

0.0

0.0

0.0

0.1

0.0

0.0

0.1

0.0 1.9

0.1 3.0

0.0 3.8

0.1 3.0

0.1 3.3

0.0 2.5

0.1 2.9

0.1 0.7

2.5

0.0

2.1

0.0

0.0

1.1

1.0

1.1

0.5

0.0

0.8

0.0

0.0

0.5

0.3

0.3

0.7 0.6 0.1

1.9 2.1 2.3

0.3 0.3 0.1

1.3 0.4 0.2

2.0 0.5 0.3

0.9 0.8 0.5

1.2 0.8 0.6

0.7 0.7 0.9

0.0 3.3

3.5 2.6

0.1 4.0

0.0 1.7

0.1 0.4

0.6 0.0

0.7 2.0

1.4 1.6

1.0 1.7

4.0 0.1

4.7 0.0

2.9 2.4

0.7 2.8

0.1 1.3

2.2 1.4

1.9 1.2

0.1

0.0

0.0

7.0

3.6

0.1

1.8

2.9

0.1 2.1

0.1 7.9

0.5 0.8

0.0 5.0

0.0 2.3

0.0 1.7

0.1 3.3

0.2 2.7

1.7 0.0

0.8 0.1

0.3 0.0

0.6 0.2

0.5 0.1

0.3 0.0

0.7 0.1

0.5 0.1

0.0 0.0 0.2 7.5

0.1 0.4 0.5 6.7

0.0 0.4 3.2 5.1

0.1 0.6 2.2 3.1

0.1 0.7 3.4 4.0

0.0 0.2 0.9 9.8

0.1 0.4 1.7 6.0

0.1 0.3 1.4 2.5

7.8

3.9

6.8

4.2

3.9

8.6

5.9

2.1

97.5

100.6

103.1

65.7

99.3

97.8

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The concretionary-cemented sandstones are characterized by high pre-calcite cement porosities reaching up to 45%, and up to 70% in the mudstones. Calcite cement does not reveal evidence of grain replacement in these concretions. Framework grains, particularly micas, plagioclase, and rock fragments, reveal evidence of partial dissolution and formation of intragranular pores. Sandstones hosting the concretions are either devoid of calcite cement, or cemented by variable amounts of blocky to poikilotopic calcite cement that often displays evidence of partial to pervasive dissolution. Concretions in the mudstones contain rounded sandy patches (1–3 mm across) in which the Quartz Quartzarenite 5 5 Sublitharenite Subarkose 25

25 Lithic sub-arkose

nit

Ar

re

ko

ha

se

Lit e

Lithic arkose

10 Feldspar

Feldspatic litharenite

50

10 Rock Fragment

Fig. 5. Detrital composition of the Tertiary sandstone samples plotted on McBride (1963) classification diagram. The sandstones are mainly subarkoses, lithic subarkoses and sublitharenite.

calcite cement is coarser crystalline. Some of the concretions contain thin (50 mm–1 mm), septarian fractures that are filled with blocky calcite cement and tapper towards the concretion margin. Calcite cement in the concretions engulfs, and hence post-dates, pyrite that occurs around altered volcanic rock fragments and, rarely, fibrous smectite. In interlaminated mud and sand, calcite cement occurs selectively in the sand laminae. The EMP analyses (Table 2) revealed that microcrystalline calcite in the concretions contains variable amount of Fe (0.0–7.8 mol% FeCO3), Mg (0.0–9.1 mol% MgCO3), Mn (0.3–1.7 mol% MnCO3) and Sr (0.2–0.6 mol% SrCO3). Microcrystalline and blocky concretionary calcite cement display wide variations in d13C (17.8% to +5.3%) and d18O (7.5% to 1.0%; Table 2). EMP analyses of fracture-filling calcite cement revealed a nearly pure CaCO3 composition, with trace amounts (o1 mol%) of Mg, Mn, Fe and Sr. The d13C and d18O values of one analysed fracture filling calcite are 2.3% and 7.8%, respectively (Table 2). Sandstone cemented layers (i.e. non-concretionary) by calcite vary in thickness from 30–100 cm and display homogeneous textures with weak lamination (Fig. 7). In contrast to the concretions, calcite in these sandstones occur as partial to extensive cements and has a blocky spary (30–150 mm across) and, less commonly, poikilotopic (up to 300 mm across) habits. Partial cementation by evenly scattered, blocky calcite occurs in the deeply buried sandstones (44.5 km); this calcite replaces partially, and hence post-dates, quartz overgrowths (Fig. 8A). The blocky to poikilotopic calcite cement in the sandstones, which is most common in the proximal sandstones, has replaced, partially to pervasively, the framework grains (Fig. 8A and B). Partial to complete replacement of grains has resulted in the frequent presence of oversized calcite cement patches. Calcite also fills

Fig. 6. Core photos showing the occurrence of concretionary calcite around bioturbated, siltstones. Concretions have spherical to oval shapes and vary in diameter from a few millimetres up to 20 cm.

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521

Table 2 Chemical composition from microprobe analyses (mol%) and isotopic ratios of representative carbonate cements Well, depth (ft) Carbonate type

Ca (CO3)

Mg (CO3)

Mn (CO3)

Fe (CO3)

Sr (CO3)

d18O (%PDB)

208/19-1, 8059.10 ¼

Fe-dolomite and ankerite

64.0

21.0

1.0

10.0

0.0

12.1

8.5

Fe-dolomite and ankerite

50.3

21.0

0.7

28.0

0.0

6.3

1.7

¼

Fe-dolomite/ankerite

60.9

22.2

1.0

15.9

0.0

214/27-1, 13803.5 ¼

Dolomite

55.7

44.6

0.0

0.1

0.0

Dolomite

54.9

44.3

0.0

0.2

0.0

208/19-1, 8059.10 ¼

Crypto-to microcrystalline siderite

4.8

6.5

3.7

85.0

0.0

8.6

18.6

Crypto-to microcrystalline siderite

22.1

13.3

11.7

52.2

0.7

6.6

+5.4

¼

Crypto-to microcrystalline siderite

15.6

12.1

1.5

70.8

0.0

¼

Siderite embedded in biotite

3.1

19.2

1.3

76.5

0.0

Siderite in deeply buried sandstones Siderite in deeply buried sandstones

2.3 9.8

8.6 13.3

2.3 7.2

86.6 69.7

0.0 0.0

10.2 9.8

8.2 8.2

Microcrystalline calcite in concretions

91.7

7.8

Trace

0.0

0.2

7.5

17.8

Microcrystalline calcite in concretions

81.2

9.1

1.7

7.8

0.5

1.0

+5.3

Microcrystalline calcite in concretions

98.4

0.0

1.0

Trace

0.3

Proximal turbidite Proximal turbidite Distal turbidite

Microcrystalline calcite in concretions Microcrystalline calcite in concretions Fracture filling calcite Boxwork and cone-in-cone calcite

98.9 97.2 99.0 99.6

0.0 0.0 Trace 0.0

Trace Trace Trace Trace

Trace 1.7 Trace Trace

0.6 0.2 0.2 0.0

2.3 2.2

Distal turbidite Distal turbidite Shelf sediments Paleosol

Boxwork and cone-in-cone calcite Boxwork and cone-in-cone calcite Boxwork and cone-in-cone calcite Non-concretionary blocky to poikilotopic calcite Non-concretionary blocky to poikilotopic calcite Non-concretionary blocky to poikilotopic calcite Non-concretionary blocky to poikilotopic calcite

90.4 92.6 98.0 96.8

2.7 Trace 1.0 2.1

2.7 3.4 Trace 5.5

4.0 3.4 Trace 5.0

0.2 0.1 0.2 0.5

7.5

2.1

paleosol Marine mud Marine mud Distal turbidite

96.3

Trace

1.3

2.1

0.1

15.0

+7.9

96.6

Trace

1.2

2.0

0.2

96.8

2.1

5.5

5.0

0.5

208/19-1, 8059.10 ¼ 214/27-1, 13803.5 ¼ ¼ 206/1-1A, 7595.20 ¼ ¼ ¼ 214/27-1, 13803.5 204/25a-2, 6523.4a ¼ ¼

intragranular pores in micas (Fig. 8C) and albitized feldspar (Fig. 8D and E). In granitic rock fragments, the feldspars are selectively replaced by calcite, whereas the quartz crystals are slightly replaced or unaffected. The albite overgrowths are unaffected, whereas the plagioclase cores are pervasively replaced by calcite (Fig. 8E). Calcite cement also fills primary intragranular pores in carbonate bioclasts and, in some cases, pervasively replaces the bioclast fragments (Fig. 8F). The volume of grain-replacing calcite cement is trace to 19% and of calcite filling the intergranular pores is 14–30%. BSE imaging revealed the common presence of pitted, chemically pure,

7.8 8.4

d13C (%PDB)

Sample location

Proximal turbidite Proximal turbidite Proximal turbidite Proximal turbidite Proximal turbidite Proximal turbidite Proximal turbidite Proximal turbidite Proximal turbidite Distal turbidite Distal and Proximal

Proximal turbidite Proximal turbidite

rounded calcite (probably of bioclast and echinoderm origin) is overgrown by Fe–Mn bearing euhedral calcite (Fig. 9A). In some cases, calcite that fills the central part of the intergranular pores displays brighter luminescence due to the presence of elevated Mn concentrations (Fig. 9B). The blocky to poikilotopic calcite cement has been subjected to partial to pervasive dissolution, which is evidenced by etched crystal boundaries and patchy distribution of calcite. Remnants of strongly etched calcite occur mainly within and along the grain boundaries (Fig. 10A). Calcite that has replaced or filled intragranular pores in feldspars is either preserved, or only partially

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Fig. 7. Core photos: (A) homogeneous sandstone with weak plane laminated proximal turbidites, (B) interbedded porous sandstone (proximal LST turbidites) stained with oil and siltstone that are not stained and (C) interbedded sandstone cemented by calcite (light grey) and mudstone containing lenticular siderite (distal turbidites; LST and TST).

dissolved (Fig. 10B). Secondary porosity that has resulted from the dissolution of calcite cement is fairly well connected, and hence is contributing to permeability improvement (Fig. 10C). Oil has filled some of these secondary pores. Fluid inclusions in poikilotopic calcite are relatively rare, scattered, rounded and prismatic in shape, and small in size (o3 mm). The inclusions are two-phase; measurements made on six inclusions with very small gas bubbles (i.e., high liquid/gas ratios) yielded a narrow range of homogenization temperatures (76–88 1C). The precise melting temperatures of the first and last ice crystal were not possible to obtain. The inclusions display no fluorescence under ultraviolet light. EMP analyses (Table 2) revealed that the non-concretionary, blocky to poikilotopic calcite cement is enriched to variable extents by Mg (0.0–2.1 mol% MgCO3), Fe (2.0–5.0% FeCO3), Mn (1.2–5.5 mol% MnCO3) and Sr (0.1–0.5 mol% SrCO3). The d13C values of this calcite cement range from 2.1% to +7.9%, and d18O from 7.5% to 15.0% (Table 2). Calcite cement in the paleosols occurs as boxwork and cone-in-cone habit in mottled, light grey to reddish brown mudstones in along the SB. This calcite cement contains variable amounts of Mg (0.0–2.7 mol% MgCO3), Fe (0.5–4.0 mol% FeCO3), Mn (0.3–3.4 mol% MnCO3) and Sr (0.0–0.2 mol% SrCO3). The CL imaging revealed a patchy, reddish orange luminescence colour (Fig. 9B). The d13C and d18O values of these calcite cements are 2.2% and 8.4%, respectively (Table 2).

4.2.2. Siderite Crypto- to microcrystalline (o5 mm across) siderite occurs as scattered crystals and as extensive cement in concretions (a few millimetres to 5 cm across) in heavily bioturbated sandstones, which are moderately sorted, with weak horizontal lamination and deeply buried, dark grey to black mudstones (44.5 km). Concretions in the mudstones are strongly elongated parallel to the bedding plane (2–4 mm thick and 1–3 cm long), but well rounded in the sandstones. Siderite in the sandstone (trace to 5.5 vol%) occurs also in the vicinity of biotite grains that have been expanded considerably, partially dissolved and altered into poorly characterized microcrystalline, clays of probably smectite (Fig. 11A and B). Siderite occurs commonly in the palaeosols and sandstones immediately below and along the SB. Siderite is also commonly embedded in calcite concretions. EMP analyses (Table 2) revealed that siderite varies widely in composition even within the same sample, being enriched in Mg (6.5–13.3 mol%), Ca (4.8–22.1 mol%) and Mn (3.7–11.7 mol%); Sr content is usually below detection limit, but values up to 0.7 mol% have been detected. Siderite, which commonly associated with calcite cement (Fig. 11C and D), have similar ranges of chemical composition as the siderite concretions. Alternatively, these siderites have been formed by replacement of highly reactive, volcanic glass. EMP analyses (Table 2) revealed that siderite embedded in biotite is enriched in Mg (19.2 mol%) and, to a smaller extents, Ca (3.1 mol%) and Mn (1.3 mol%). The d13C values of siderite cement range

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Fig. 8. Thin section optical micrographs under X-nicols. (A) Calcite (Ca) that has partially replaced, and hence post-dates, the quartz overgrowths. (B) Blocky to poikilotopic calcite cement (arrow) has replaced, partly to pervasively, the framework grains. (C) Calcite fills intragranular pores in micas (arrow). (D) Replacement of K-feldspar by calcite. (E) Selective replacement of detrital plagioclase by calcite and not of the albite overgrowth (arrow). (F) Intragranular pores in bioclasts (arrow) filled by calcite cement.

from (18.6% to +5.4%) and the d18O values from (8.6% to 6.6%; Table 2). In the deeply buried sandstones (44.5 km), siderite occurs as small (15 mm) crystals around the framework grains and is engulfed by quartz overgrowths (Fig. 11C and D). This siderite has variable Mg (8.6–13.3 mol%), Ca (2.3–9.8 mol%), and Mn (2.3–7.2 mol%) contents; Sr content is below detection limit. The d13C values of this siderite cement are 8.2% and the d18O values range from (10.2 to 9.8%; Table 2).

4.2.3. Fe-dolomite and ankerite Fe-dolomite and ankerite occur in the sandstones both as pore-filling and as framework-grain replacive rhombs (20–80 mm in across; Fig. 12A and B). Ankerite fills mouldic pores together with albite, illite and chlorite. Such ankerite is probably originally grain replacive (Fig. 12B), and is attaining intergranular pore-fillings because the detrital grain has been dissolved and/or albitized. Evidence supporting this postulation is the common presence of secondary mouldic pores unfilled by adjacent,

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Fig. 9. (A) BSE image showing the presence of pitted bioclast that is overgrown by calcite cement. (B) CL image of poikilotopic calcite cement that fills the central part of the intergranular pores displays brighter luminescence (arrow) owing to the presence of elevated Mn concentrations.

intergranular pore-filling ferron dolomite and ankerite. Thus, replacement of framework grains by Fe-dolomite and ankerite predates grain dissolution. In some cases, Fe-dolomite and ankerite occur as large (up to 600 mm) aggregates of numerous rhombic crystals filling oversized pores. Ferron dolomite, engulfs, and hence post-dates Mg-rich calcite. Fe-dolomite and ankerite crystals are covered by, and hence post-date illite and chlorite, but are engulfed by and/or engulfing quartz overgrowths indicating a recursive origin. Low-Fe dolomite (0.0–0.2 mol%) is rare and occurs in both shallow and deeply buried rocks. Low-Fe dolomite in the shallow buried sandstones is closely associated with microcrystalline calcite in some of the concretions. In some of the sandstones, trace amounts of partially dissolved, Fe-dolomite crystals occur within intragranular pores in dissolved and albitized plagioclase, but other carbonates are absent in the intergranular pores. EMP analyses of Fe-dolomite and ankerite crystals revealed variable Fe contents (10–28 mol%). Individual crystals are either homogeneous in composition or most commonly zoned, with core containing less Fe than the rims. The Mn varies from 0.7 to 1.0 mol%, whereas Sr is not detected. The d13CVPDB values of Fe-dolomite and ankerite vary from 8.5% to 1.7% and d18OVPDB values from 12.1% to 6.3% (Table 2). 4.2.4. Clay minerals Various types of clay minerals with various textural habits occur in the sandstones. Smectite has completely replaced volcanic rock fragments, mud intraclasts, and biotite. It occurs as smooth, extensive grain coating and as meniscus crystals that may bridge pores between adjacent grains (Fig. 13A). However, the grain-coating and meniscus clays are common in the HST sandstones located immediately below the SB. Fibrous smectite has replaced angular to subangular, sand-sized grains of probable volcanic origin in some of the calcite concretions. Illitic

and chloritic clays occur as crenulated flakes that, in some cases, develop into honeycomb-like texture (Fig. 13B and C). The flakey illite crystals have hair-like filamentous terminations and are engulfed by, and thus pre-date, quartz overgrowths and discrete quartz crystals (Fig. 13D). In some cases, the illitic clays have replaced partially to pervasively biotite and fill microfractures in feldspars. Chloritic clays occur as small (0.2 mm thick and 1–5 mm across) flakes that coat the framework grains and are engulfed by quartz overgrowths. Transmission electron microscopy (TEM) indicates that most chlorite occur as Fe-rich chamosite (Fig. 14). Chlorite flakes have, together with euhedral microcrystals of authigenic K-feldspar, replaced the biotite that has been replaced by smectite and chlorite. Biotite has been expanded several times its original size and thus has often completely filled adjacent intergranular pores (Fig. 11A and B). Additionally, biotite has been squeezed due to mechanical compaction between the rigid framework grains and resulted in the development of extensive pseudomatrix. Biotite displays variable and successive degrees of alteration and dissolution, which often result in the formation of large interconnected secondary porosity (Fig. 15A and B). Kaolin occurs as booklets (5–10 mm across) and vermicular stacked pseudohexagonal crystals that have replaced the framework grains (Fig. 16A and B). Kaolin occurs as patches with well-defined, rounded to subrounded boundaries with abundant microporosity and similar sizes as the framework feldspar grains (Fig. 16C and D). Relict feldspars (primarily plagioclase) occur within the kaolin patches. Kaolin occurs also within and adjacent to severely dissolved biotite grains. The kaolin patches are commonly deformed between adjacent rigid grains such as quartz and feldspars (Fig. 16D). Based on crystal shapes, kaolin occurs as disordered kaolinite, which has, in some cases, been transformed into well-ordered kaolinite and dickite, which occur as thicker crystals (up to 6 mm) than kaolinite (o 2 mm). Kaolinite is engulfed by, and thus pre-dates, quartz overgrowths (Fig. 16A).

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sandstones that are rich in dissolved plagioclase and biotite. Kaolin is most abundant in the HST sandstones in the vicinity of and immediately below the paleosol surfaces (SB) and in proximal and distal LST turbidite facies. The XRD analyses (Fig. 18) revealed that clay minerals include mainly: (i) chlorite, which is identified based on the presence of the two typical, prominent peaks (14 and 7 A˚), and disappearance of 7 A˚ peak upon heating to 550 1C, coupled with the persistence of the 14 A˚ peak; (ii) Kaolin, which display the 7 A˚ peak that is attributed partly to the presence of kaolin and (iii) illitic clay minerals, which includes illite and mixed-layer illite/smectite and mixedlayer illite/vermiculite, which have a broad peak or numerous small peaks that extend from about 10 A˚ to about 12 A˚ in the untreated samples. Upon treatment with ethylene glycol, the illite is revealed as sharp 10 A˚ peak, while the other parts of the broad peak or the small peaks at greater values than 10 A˚ are all collapsed. Upon heating, the 10 A˚ peak is persistent or even becomes more sharpened.

Fig. 10. Thin section optical micrographs under X-nicols. (A) Remnant of strongly etched calcite within and along the grain boundaries. (B) Partial dissolution of calcite that has replaced a feldspar grain. (C) Secondary porosity, which has resulted from the dissolution of calcite cement is fairly well connected and has hence, contributed to permeability improvement.

The most dominant types of clays in the deeply buried sandstones are chlorite and illite, whereas kaolin occurs in trace amounts (Fig. 17A–C). Kaolin in the deeply buried sandstones mostly occurs as well-ordered kaolinite and dickite. Kaolin occurs in the shallow buried (o2.5 km)

4.2.5. Quartz cement Quartz cement in the sandstones occurs mainly as partial to complete syntaxial overgrowths (10–50 mm thick) around monocrystalline quartz grains. The boundary between the overgrowths and the detrital core is either poorly defined or delineated by fluid inclusions and/or thin clay coatings. Quartz overgrowths are most abundant in the deeply buried sandstones (44.5 km) where it forms up to 17%. In shallower buried sandstones (ca. 2.5 km), quartz overgrowths do not exceed 8% (Fig. 19). Quartz overgrowths are most abundant in sandstones with clean quartz grain surfaces compared with sandstones with extensive grain-coating illitic and chloritic clay. In the later sandstones, quartz overgrowths occur as isolated patches composed of small crystals (30 mm across; Fig. 20A) that are, in some cases, orientated parallel to each other (Fig. 20B). Trace amounts of quartz cement occur as discrete euhedral crystals (5–40 mm across) that fill intergranular space adjacent to quartz overgrowths (Fig. 20C), or fill secondary intragranular pores in dissolved and albitized feldspars. Quartz overgrowth is common in the distal LST sandstones as well as in the TST sandstones. Fluid inclusions in quartz overgrowths (two phase) are extremely rare, small in size (mostly o2 mm) and occur mainly along the boundary with the detrital quartz grain. Microthermometric measurements made on two inclusions in the deeply buried sandstones yielded homogenisation temperatures of 93 and 114 1C. These values were not corrected for pressure. The precise melting temperatures of the first and last ice crystal were not possible to obtain. The inclusions display no fluorescence under ultraviolet light. 4.2.6. Feldspars Diagenetic feldspars occur as albite and K-feldspar overgrowths and as numerous parallel-arranged, euhedral

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Fig. 11. Optical micrographs showing (A and B; plane polarized) Biotite grains that have been expanded considerably, forming pseudomatrix, partially dissolved and altered into microcrystalline smectitic clays. Siderite occurs in the vicinity of biotite grains (arrows). (C and D; crossed polars) Well-rounded, brown-coloured, sand-sized grains composed of cryptocrystalline siderite (s) and clay minerals. The siderite is commonly associated with poikilotopic calcite and engulfed by quartz overgrowth (arrows).

Fig. 12. Rhombs of Fe-dolomite and ankerite occur as pore-filling (A) and framework-grain replacive (B).

albite crystals within partly to extensively dissolved plagioclase (Fig. 20D), which resemble albitized feldspar described by Morad (1990). The albitized grains are commonly coated with chlorite that also partially fills intragranular pores and engulfs, and hence post-dates, the albite crystals (Fig. 20E). Albite overgrowths display twinning and develop on partly to completely albitized grains. K-feldspar overgrowths occur as rare, small discrete (5–10 mm across) or coalesced adularia-like crystals (Fig. 20F) around detrital microcline grains that are coated with illitic or chloritic clay. The overgrowths are engulfed by, and hence predate, quartz overgrowths that have been developed on adjacent quartz grains. K-feldspar occurs

also as discrete etched microcrystals within dissolved and chloritized biotite grains. EMP analyses (Table 3) revealed that diagenetic albite and K-feldspar are nearly pure NaAlSi3O8 and KAlSi3O8 end-member composition (99 mol%; Ab and Or, respectively). Detrital K-feldspar contains variable amounts of albite solid solutions (Ab 2–14 mol%) and detrital plagioclase contains variable amounts of Ca (CaO 0.2–3.7 wt%; An 1–18 mol%). 4.2.7. Other diagenetic minerals Minor diagenetic minerals include pyrite and anatase. Pyrite is most abundant (up to 10 vol%) in the calcite

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Fig. 13. (A) Smectite completely replacing a framework grain (volcanic rock fragment, biotite or mud intraclast). (B) Grain-coating, illitic and chloritic clays occur as crenulated plates that develop into honeycomb-like texture. (C) Chlorite and/or mixed-layer chlorite-smectite with honeycomb texture (D) Grain coating illitic clay that is engulfed by quartz overgrowths and by discrete authigenic quartz crystals.

coarse-crystalline, intergranular cement, as scattered framboids, and by pervasive replacement of biogenic calcite. Anatase occurs as small (10–30 mm across) euhedral crystals that have replaced detrital Fe–Ti oxide grains and titanite. 4.3. Porosity

Fig. 14. Chemical composition of chlorite based on TEM analysis showing that chlorite is mainly chamosite.

concretions that occur in mudstones. The sulphur isotope analysis of this pyrite indicate a considerable enrichment in 32 S (d34S ¼ 12%). Pyrite occurs in the sandstones as local

Porosity of the sandstones is of primary and secondary origins. However, the distinction between intergranular porosity of primary and secondary origins is problematic. Sandstones buried at depths o2.5 km are characterized by elevated intergranular (up to 27%), and moldic porosity (up to 18%). Porosity decreases sharply with depth, and thus the deep reservoir (44 km) are characterized by low (o7%) inter- and intragranular porosity (Fig. 21A and B). A plot of the total volume of intergranular volume (IGV) versus volume of cement % (Fig. 22) reveals that compaction has been far more important in porosity destruction than cementation. The plot does not include sandstones that are cemented by concretionary carbonates, which have lost their intergranular porosity due to pervasive, pre-compactional cementation. Secondary intragranular porosity has been developed by partial to pervasive dissolution of detrital biotite, plagioclase, volcanic rock fragments and mud intraclasts. Moldic pores have commonly been formed from the complete dissolution of very-coarse (up to 1.5 mm) to very-fine sand grains, presumably plagioclase and biotite grains. Dissolved biotite grains have been expanded into adjacent open pore-forming pseudomatrix (Fig. 11A and B).

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Fig. 15. Plane polarized optical micrographs (A and B) showing variable and successive degrees of alteration and dissolution of detrital biotite and the formation of large secondary porosity.

Fig. 16. (A) SEM micrograph showing dissolved and partially dickitized kaolinite that is engulfed by quartz overgrowths. (B) SEM micrograph showing a close up view of blocky dickite that has replaced kaolinite. (C) Plane-polarized optical micrograph showing dissolved biotite grain and in the vicinity of which kaolin has precipitated. (D) Plane-polarized optical micrograph showing an aggregate of kaolin booklets that have been deformed into pseudomatrix-like clay.

The coarse plagioclase grains are more pervasively dissolved than the small plagioclase grains, being frequently associated by the formation of kaolinite (Fig. 23A and B). Moldic pores that have been formed by the dissolution of plagioclase are outlined by chloritic or illitic clays (Fig. 23C). Dissolution of plagioclase grains is also commonly associated with variable degrees of albitization (Fig. 23D). Overall, the detrital K-feldspar grains are fresh or only slightly etched. The dissolution of smectitic clays that have replaced the volcanogenic rock fragments and mud intraclasts is fairly extensive. The secondary intergranular and moldic pores are, thus, contributing to a substantial enhancement of reservoir quality of the sandstones. The intergranular pores are overall fairly well interconnected except in sandstones containing abundant

biotite grains that have expanded and blocked adjacent pore throats and resulted in poorly connected macropores, and thus low-permeability reservoirs. The dissolved biotite and clay minerals are extremely rich in microporosity. There is no single diagenetic cement that is solely or mainly controlling the pattern of porosity evolution of the sandstones (Fig. 24A–C). Instead, it appears that the main types of cements (carbonates, clay minerals and quartz) as well as compaction have collectively controlled the reservoir quality evolution of the sandstones. 5. Discussion The diagenetic evolution and impact on the spatial and temporal distribution of reservoir quality of the Tertiary sandstones were accomplished during eo- and

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mesodiagenesis. Although it is difficult to obtain precise timing and duration of the diagenetic processes, an overall paragenetic sequence for diagenetic alterations was constructed based on petrographic textural relationships, isotopic composition of the diagenetic minerals, fluid inclusion and burial history curve (Fig. 25).

6000

Depth (ft)

8000

10000

5.1. Eogenetic alterations 12000

14000 r = -0.005 16000 0

4

12

16

Chloritic clay (%) 6000

Depth (ft)

8000

10000

12000

14000 r = -0.6 16000 0

2

4

6

Illitic clay (%)

6000

8000

Depth (ft)

529

10000

12000

14000

r = -0.7

16000 0

2

4

6

Kaolinite (%)

Fig. 17. Plots showing variations in the amounts of chloritic clays (A), illite (B) and kaolin (C) with depth. Kaolinite is mainly abundant in shallow sandstones. Chlorite is not correlated with present depth. Illite is overall increasing with increasing present depth.

The eogenetic alterations of Palaeocene turbidites have been controlled by several parameters such as changes in the relative sea level, pore-water chemistry, depositional facies, and detrital composition. The eogenetic alterations have exerted a critical control on the pattern of mesogenetic evolution and related reservoir quality modifications. The eogenetic alterations resulted in the cementation by carbonate, dissolution and kaolinitization of detrital feldspars and micas, and formation of smectite. 5.1.1. Carbonates High pre-carbonate cement porosities encountered (up to 45% in the sandstones and 70% in the mudstones) implies a pre-compactional, eogenetic precipitation of concretionary, microcrystalline calcite cements. The concentration of the eogenetic, concretionary cements around borings and burrows, points out the role of local carbonate alkalinity increase due to the decomposition of organic matter. Furthermore, bacterial sulphate reduction in the marine porewaters resulted in the precipitation of eogenetic, framboidal and coarse-crystalline pyrite (Bottrell et al., 2000). The close association of pyrite with eogenetic carbonate cements suggests that the alkalinity increase due to oxidation of organic matter is related to bacterial sulphate reduction coupled with the reduction of iron oxides/oxyhydroxide (Curtis et al., 1986). The fairly strong enrichment of this pyrite in 32S indicates that bacterial sulphate reduction occurred under semi-closed conditions with respect to the overlying seawater (Raiswell, 1982). The d18O values (7.5% to 1.0%) of the eogenetic concretionary calcite suggest temperatures of (15–50 1C; Fig. 26A), if precipitation is assumed to have occurred from marine porewaters with unmodified oxygen isotopic composition (1.2% relative to SMOW; Shackleton and Kennett, 1975). The elevated Mg (up to 9 mol%) and Sr concentrations (up to 5000 ppm) in the microcrystalline calcite support precipitation from marine porewaters (Morad, 1998). However, the elevated Fe content (8 mol%) is fairly unusual at these temperatures but, does indicate that calcite precipitation occurred under reducing conditions. The reduction of Fe3+ occurs in the sub-oxic, bacterial sulphate reduction, and methanogenesis zones below the sea floor (Froelich et al., 1979; Berner, 1981). The d13C values of the eogenetic concretionary calcite (17.8% to +5.3%) indicate derivation of carbon from the bacterial sulphate reduction and from microbial methanogenesis, respectively (Irwin et al., 1977).

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K > I > Chl

I-S >KI > Chl

EG OA H

K+Chl

EG OA H

Q I-S

Chl

K

Q

I

A 0

B 5

10

15 2 Theta

20

25

0

5

EG OA H

I-S >KI > Chl Q

10

15 2 Theta

20

EG OA H

K+Chl

K>Chl>I

25

I-S I

Q K

Chl

I Chl

C 0

D 5

10

15 2 Theta

20

25

0

5

10

15 2 Theta

20

25

Fig. 18. Clay fraction XRD patterns for a representative Tertiary sandstones. The runs include: (i) air-dried (ii) ethylene glycol treatment and (iii) heating to 550 1C. The X-ray diffraction patterns in A, B, C and D show chlorite, illite, kaolinite and mixed-layer illitie–smectite, respectively.

6000

Depth (ft)

8000

10000

12000

14000

r = 0.3 16000 0

4

8 12 Quartz overgrowth (%)

16

20

Fig. 19. Plot showing variation in the amounts of quartz overgrowth with depth. Quartz overgrowth in many samples do not display any significant correlation with present depth.

The close association of siderite with mud intraclast and micas that have been expanded considerably owing to dissolution and alteration into microcrystalline smectite and kaolinite suggest that these mud intraclast and micas acted as initial source of needed irons for siderite formation. Altered mica might also have provided suitable micro-geochemical conditions, such as elevated pH needed for the formation of siderite (Boles and Johnson, 1984). Eogenetic siderite has precipitated at near-surface conditions in association with pedogenic calcite concretions along the SB in the shelf areas. Although, the obtained two values of d18OVPDB and 13 d CVPDB of siderite cement preclude a precise elucidation of its origin, some clues to their precipitation conditions can be achieved. Using the d18OVPDB values of siderite (8.6% to 6.6%) in the paleosols, the fractionation equation of Carothers et al. (1988), and assuming nearsurface temperature 20–30 1C, precipitation would have occurred from pore water with d18OSMOW ranging between 11% and 6.5%, which are equivalent to mixed marine meteoric and dominantly meteoric pore waters (Fig. 26B). Meteoric water in the North Sea region during

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Fig. 20. (A) SEM images of quartz cement showing isolated quartz overgrowth crystals. The inhibition of more extensive quartz overgrowths is attributed to the presence of chlorite coatings around the detrital quartz core. (B) Partly coalesced quartz overgrowth crystals oriented parallel to each other. (C) Discrete euhedral crystals of quartz cement that fill intergranular space adjacent to quartz overgrowths. (D) Numerous parallel-arranged euhedral albite crystals within partly dissolved and albitized plagioclase. (E) Albitized grain, which is coated with chlorite that partially fills intragranular pores and engulfs, hence post-dates the albite crystals. (F) Authigenic K-feldspar crystals that occur as overgrowths on detrital K-feldspar coat illitic and chloritic clays.

the Tertiary was inferred to be about 12% (Fallick et al., 1985). The high carbon isotopic composition of siderite (+5.4%) suggests precipitation under condition of microbial methanogenesis (Irwin et al., 1977; De Souza et al., 1995). The formation of siderite in the Paleocene sandstones was favoured by low concentration of SO2 4 in the meteoric influenced HST and proximal LST sandstones (Postma, 1982). Calcite with cone-in-cone and boxwork habits in the pedogenically sandstones along and immediately below the SB, has relatively elevate d13C value (2.2%) indicating derivation of dissolved carbon from C4 plants (cf. Cerling, 1984) and/or mixed marine-meteoric pore waters. Using

the d18O value of this calcite (8.4%), the assumed precipitation temperature of 20 1C and the fractionation equation of Friedman and O’Neil (1977), the oxygen isotope value of the pore water from which calcite has precipitated was 7% relative to SMOW, which corroborates with a mixed marine and meteoric water origin. 5.1.2. Silicates Among the earliest eogenetic processes is the replacement of highly reactive, volcanic grains and mud intraclasts into smectite, and formation of grain coatings and pore bridges smectite as well as dissolution of silicate grains such as feldspars, micas, rock fragments, and mud intraclasts and formation of kaolinite. However, the formation of

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Table 3 Oxygen microprobe data and mineral formulas of detrital and diagenetic feldspars Well

Depth (ft)

Spot

Wt%

Mol%

SiO2

Al2O3

CaO

Na2O

K2O

Ab

Or

An

214/28-1

8361.7

Diagenetic Detrital Detrital Detrital Detrital Diagenetic Detrital Detrital Diagenetic Detrital Detrital Detrital Detrital Detrital Detrital Detrital

63.6 67.1 69.0 64.9 66.0 68.7 66.9 65.5 69.3 67.8 65.2 66.8 67.9 63.3 63.9 64.0

19.1 20.2 19.4 21.2 20.1 18.9 19.4 21.4 19.2 19.6 21.1 19.2 19.3 17.7 17.9 18.0

0.01 1.0 0.2 2.6 1.3 0.1 0.6 2.3 0 0.4 2.3 0.7 0.5 0 0 0

12.0 11.1 11.6 10.0 11.1 11.6 11.3 10.5 11.7 11.4 10.2 11.3 11.4 0.3 0.3 0.7

0.03 0.04 0.03 0.2 0.1 0.03 0.02 0.1 0.01 0.07 0.1 0.07 0.07 15.9 16.2 15.8

99.77 95.04 98.9 86.4 93.57 99.4 97.08 88.51 99.9 97.72 88.39 96.31 97.03 2.79 2.71 5.95

0.18 0.23 0.15 1.03 0.55 0.18 0.09 0.66 0.07 0.41 0.53 0.42 0.39 97.21 97.29 94.05

0.06 4.73 0.94 12.56 5.88 0.42 2.83 10.84 0 1.88 11.08 0 2.57 0 0 0

214/28-1

8362.9

Diagenetic Detrital Detrital Detrital Diagenetic Detrital Detrital Detrital

68.4 66.2 68.4 64.3 68.4 65.4 63.7 63.2

19.5 20.5 19.0 18.8 19.8 20.7 18.4 22.5

0.2 1.3 0.1 0 0.1 2.2 0 3.7

11.7 10.4 11.5 1.1 11.4 10.4 1.0 9.7

0 0.1 0.05 15.83 0.04 0.1 15.6 0.02

99.28 93.04 99.08 9.29 99.4 89.38 9.0 82.49

0 0.32 0.28 90.71 0.25 0.36 91.0 0.12

0.72 6.64 0.64 0 0.35 10.26 0 17.39

214/29-1-1Z

4901.5

Detrital Detrital Detrital Detrital Detrital

64.1 64.1 67.5 66.5 65.1

22.4 22.3 20.2 21.1 21.3

3.7 3.6 0.8 1.7 2.3

9.6 9.9 11.4 10.7 10.3

0.2 0.1 0.2 0.1 0.1

81.32 82.59 95.76 91.96 88.59

1.34 0.65 0.82 0.4 0.56

17.34 16.77 3.42 7.63 10.84

214/29-1-1Z

2564.53

Detrital Detrital Diagenetic Detrital

64.4 64.9 68.4 64.9

18.9 18.8 19.2 18.2

0 0 0 0

1.7 0.5 11.6 0.2

14.9 15.9 0.02 16.8

14.23 4.52 99.88 1.84

85.77 95.48 0.12 98.16

0 0 0 0

204/25a-2

6523.4a

Detrital Detrital Detrital Detrital Detrital Detrital Detrital Diagenetic Diagenetic Detrital Detrital

64.4 64.7 63.9 67.7 62.3 64.5 69.2 68.0 67.7 64.5 64.8

18.8 18.4 18.9 20.0 23.3 18.1 19.5 19.5 19.6 65.8 18.6

0 0.01 0 0.6 5 0 0.3 0.2 0.1 0.3 0.02

0.5 0.4 0.3 11.6 9.3 0.3 10.9 11.6 11.2 0.9 1.1

16.4 16.6 16.9 0.1 0.1 17.0 0.1 0.01 0.02 16.0 15.5

3.89 3.38 3.05 97.07 76.94 2.22 98.32 99.21 99.29 7.09 8.92

96.11 96.6 96.95 0.43 0.52 97.78 0.25 0.05 0.11 91.74 90.99

0 0.02 0 2.5 22.54 0 1.43 0.74 0.6 1.17 0.09

kaolinite is usually attributed to the incursion of meteoric waters in continental and paralic siliciclastic sequences. Such meteoric-water incursion and kaolinite formation in deep-water, marine turbidites are suggested to be unlikely to occur (Longstaffe, 1984; Bjørlykke and Aagaard, 1992; Wilkinson et al., 2004a, b). Thus, the origin and chemical composition of the waters that are responsible for the formation of kaolinite in such deposits remains difficult to constraint (Morad et al., 2000). Although, several workers suggested that diagenesis in turbiditic sandstones was most

likely influenced by descending meteoric water circulation (Ma´tya´s and Matter, 1997; Carvalho et al., 1995), the mechanism of meteoric-water flux into deep-water, turbiditic sandstones is enigmatic but could have occurred as a consequence of (i) major fall in the relative sea level during formation of LST (Meisler et al., 1984; Wilson et al., 1999), (ii) hyperpycnal flow (Plink-Bjo¨rklund and Steel, 2004) and (iii) telodiagenesis during basin uplift (Morad et al. 2000). The influx of meteoric water and formation of eogenetic kaolinite in the deep-water, turbiditic sandstones probably

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Cement (%) 0

10

20

30

40

40

0 10

30

14000

0

10

50 sit y

(% )

20

%

ul

ar

po

ro

20

75

an

12000

30

gr

10000

25

In te r

Intergranular volume (%)

Depth (ft)

8000

10

90

-0.4 16000 4

12

16

Intragranular porosity (%)

0

25

50

75

100

Original porosity destroyed by cementation (%)

Fig. 22. Plot of intergranular volume (IGV) versus volume of cement (Houseknecht, 1988, modified by Ehrenberg, 1989) of the Tertiary sandstones showing that the intergranular porosity has been reduced mainly due to compaction rather than to cementation. Carbonate concretions are not included.

6000

8000

Depth (ft)

100

0

0

Original porosity destroyed by mechanical compaction and intergranular pressure solution (%)

6000

10000

12000

14000 -0.5 16000 0

4

12

16

Intergranular porosity (%) Fig. 21. (A)–(B) Cross plots showing variations in the amounts of intergranular and intragranular porosity with depth. Porosity is lower in deeply buried sandstones.

occurred as a consequence of major fall in relative sea level, which occurs during deposition of the LSTs. Major fall in relative sea level exposes the shelf and the shoreline rapidly migrates basinward, leading to the enlargement of the meteoric recharge areas (Morad et al., 2000; Ketzer et al., 2003; Hayes and Boles, 1992; Carvalho et al., 1995). Hyperpycnal flow, which occurs when river effluent transfers into a sediment gravity flow and enters the sea as dense underflow is responsible of transporting big masses of meteoric water into deep-sea sediemts. The likelihood of hyperpycnal flow increases, when rivers and distributary

channels reach the shelf edge, and their sediments are delivered directly onto a slope. In hyperpycnal flows, fresh river water is captured as the ambient fluid in the sediment gravity flow (Plink-Bjo¨rklund and Steel, 2004). However, the impact of hyperpycnal flow on kaolinitie formation is fraud with uncertainties owing to the episodic nature of this flow, which precludes continous flushing of the sand deposits. The formation of kaolinite during telodiagenesis is unlikely because the burial history curve of the SFB does not show that the studied rocks have been emerged during uplift phase (Carr and Scotchman, 2003). However, Smallwood and Gill (2002) have provided evidence that some parts of the SFB might have been exposed to the surface during uplift. Furthermore, there are two lines of petrographic evidence, which suggest that kaolinite is of eogenetic rather than of telogenetic origin, include (i) considerable expansion texture of kaolinitized micas to fill adjacent pores and vermicular habit, which typically imply formation in poorly compacted sandstones prior to significant burial (Ketzer et al., 2003; Wilkinson et al., 2004a, b) and (ii) the engulfment of kaolinite by quartz overgrowths and the transformation of kaolinite into dickite (Fig. 16A and B). 5.2. Mesogenetic alterations The mesogenetic alterations, which include the processes that have occurred at burial depths greater than ca. 2 km, were controlled by increasing temperature (470 1C) and

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H. Mansurbeg et al. / Marine and Petroleum Geology 25 (2008) 514–543

Fig. 23. Optical micrographs (A and B; plane polarized) showing variable degrees of dissolution of plagioclase grains and the formation of secondary porosity. Plagioclase grains are frequently associated wit the formation of kaolinite. (C) Optical micrograph (plane polarized) showing completely dissolved plagioclase, which resulted in mouldic pore that is outlined by chloritic or illitic clay minerals. (D) BSE showing dissolution of plagioclase, which is associated with variable degrees of albitization.

mediated by evolved formation waters (Morad et al., 2000) and by the types and distribution patterns of eogenetic alterations. Evolution of formation waters, which is usually resulting in increase in the amounts of total dissolved solids, occurs owing to sediment–pore-water interaction in the Paleocene deposits as well as in adjacent lithologies and probably owing to the flux of exotic formation waters (Wycherley et al., 2003). 5.2.1. Silicates Paleocene turbiditic sandstones display a variety of mesogenetic alterations including albitization of K-feldspar and plagioclase, conversion of kaolinite into dickite, chloritization of smectite as well as illitization of kaolinite and of grain-coating infiltrated clays and quartz cementation. The precise timing and mechanism of extensive plagioclase dissolution and albitization are uncertain. Plagioclase albitization occurs at shallower burial depths and temperatures than albitization of K-feldspars (Morad et al., 1990), which explains the co-existence of fresh Kfeldspar and dissolved/albitized plagioclase. The minimum temperature of plagioclase albitization is still poorly delineated in the literature. Based on a study of Tertiary to Jurassic reservoir sandstones from the North Sea, offshore Norway it has been suggested by Saigal et al. (1988) that feldspar albitization is commenced at about 60 1C (2 km depth). In the Snorre oilfield, Triassic sandstone (present-day burial depths 2.5–3.0 km;

temperature 75–100 1C) reveal evidence of pervasive plagioclase albitization, whereas the detrital K-feldspar is relatively fresh (Morad et al., 1990). Partial to pervasive dissolution of smectitic clays, mud intraclasts and formation of abundant secondary porosity require an interaction of the sandstones with acidic fluids (Milliken, 1989; Wilkinson et al., 2001). These fluids could have been of meteoric origin or are probably carboxylic and carbonic in composition, derived from the burial diagenesis of organic matter in the interbedded mudstones and in the source rocks prior to oil generation (Surdam et al., 1984; Surdam and Crossey, 1985). Extensive dissolution of these alumino-silicates requires a mechanism able to complex and transport the immobile aluminium (Wilkinson et al., 2004a, b). Surdam et al. (1984) argued, based on experimental work at 100 1C, that the presence of difunctional carboxylic acid anions at concentrations similar to those observed in formation waters can increase the solubility of aluminium from alumino-silicate dissolution by more than three orders of magnitude compared with inorganic solubility of gibbsite and even more with respect to kaolinite. This increase in solubility of Al is attributed to the Al-chelating characteristic of these organic acids (Wilkinson and Haszeldine, 1996). However, estimates of organic acid requirements for Al-chelating reactions are fraught with uncertainty about the structure of the complexes at reservoir conditions (Lundegard and Land, 1986) and about the very low concentration of

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535

2km 70

20 0

Eodiagenesis

Mesodiagenesis

Quartz over growth (%)

16 Kaolinite

Dickite

Smectite

12

Siderite

Siderite Pyrite

8 Non-ferron dolomite Concretionary calcite

4

Ferron dolomite and ankerite Pore-filling and grain replacive calcite

Mechanical compaction

0 0

4

12

16

Intergranular porosity (%)

Grain dissolution

Calcite cement dissolution Albite Ferron dolomite / ankerite

16 0

Chlorite Illite

Carbonate cement (%)

12

Pressure dissolution Quartz overgrowth k-feldspar

8

Discrete quartz cement

4

0 0

10

20

30

40

Intergranular porosity (%)

Fig. 25. A simplified paragenetic sequence of the diagenetic processes in the Tertiary sandstones. Eogenetic processes include: formation of kaolinite, smectite, eogenetic siderite, pyrite, non-feron dolomite and concretionary calcite. Mesogenetic alterations include: pore-filling and grain replacive calcite, dissolution of calcite cement, albitization of feldspar, ferron dolomite/ankerite, mesogenetic siderite and formation of illitic and chloritic clays.

Intergranular clay minerals (%)

10

8

6

4

2

0 0

4

12

16

Intergranular porosity (%)

Fig. 24. Cross plots of porosity versus amounts of quartz cements, carbonate and clay minerals showing that no single diagenetic cement solely controlling the pattern of porosity evolution of the Paleocene sandstones.

dissolved Al (o1.0 mg/l; Morton et al., 1981; Kharaka et al., 1986) in formation waters. The common association of pervasively dissolved and albitized plagioclase with fresh K-feldspar grains is in line with the observations of albitized feldspars described by Morad et al. (1990). The dissolution and albitization of K-feldspar occurs at greater burial depths and temperatures (ca. 100–130 1C) when the sandstones enter the diagenetic zone of extensive kaolinite illitization, which imposes a master control on K+ budget in the diagenetic system. The studied sandstones reveal evidence of only incipient illitization of kaolinite. Instead, kaolinite is partially to pervasively transformed into the dickite, which has a better-ordered crystal structure, and hence is stable at elevated temperatures towards illitization than kaolinite (Ehrenberg et al., 1993; Morad et al., 1994; McAulay et al., 1994; Worden and Morad, 2003; Fialips et al., 2003). The transformation of kaolinite into dickite occurs when the formation waters are characterized by low aHþ /aKþ ratios (Morad et al., 1995), which can be established owing to the

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150

100

-8

calcite concretions

siderite

-13

-7

-11

-5

-9

-12 Temperature °C

-4

50

100

0

50

0 -10

0

10

-10

0

10

150

150 poikilotopic Calcite

Fe-dolomite /ankerite

-14 -16

-11 -9

-10

-7

Temperature °C

-5 100

-6

100

50

50 -10

0 δ18OSMOW‰

10

-10

0

10

δ18OSMOW‰

Fig. 26. Curves of oxygen isotope fractionation between calcite, siderite and dolomite/ankerite and water as a function of temperature. The shaded fields illustrate the possible ranges of precipitation temperatures if waters involved were of marine origin (d18OSMOW ¼ 1.2%; Shackleton and Kennett, 1975) and moderately evolved brine (d18OSMOW ¼ +2.0%).

incursion of acidic fluids, the origin of which is, so far, poorly constrained. Using the average EMP analyses, the albitization reaction of plagioclase can be envisaged as follows: Na0:76 Ca0:26 All1:23 Si2:77 O8 þ H2 O þ 0:52Hþ ¼ plagioclase 0:76NaAlSi3 O8 þ 0:24Al2 Si2 O5 ðOHÞ4 þ 0:26Caþ 2 albite kaolin This reaction reveals that: (i) albitization reaction is not equi-volume, but instead involves a volume reduction by 26% that accounts for the presence of abundant intragranular porosity in albitized plagioclase grains, and (ii) sodium ions for the formation of albite can be derived internally, and hence no external source of sodium-rich fluids is required. Chloritic and illitic clay coatings have been formed through successive transformation of smooth, grain-coating smectitic clays as evidenced by gradual crenulations

and detachment of smectite. A portion of the chlorite is authigenic in origin, and has been precipitated around the grains and in the intragranular pores. Ions needed for the formation of authigenic chlorite have presumably been derived from the dissolution of Fe–Mg silicates, such as biotite and volcanic rock fragments. Grain-coating chlorite and illite have prohibited the precipitation of more extensive quartz overgrowths (e.g., Thomson, 1959; Dixon et al., 1989; Ehrenberg, 1993). Limited fluid inclusion data obtained for Tertiary sandstones, which indicate precipitation temperatures at the range (93–114 1C). This range of temperature is in agreement with the fluid inclusion data of Parnell et al. (1998) for quartz overgrowths and the data farther north in the North Sea (e.g., Walderhaug, 1994). The deeply buried (44.5 km) Tertiary sandstones are rich in quartz overgrowths and poor in clay grain coatings. The grains in these sandstones are instead partially coated by microcrystalline siderite that is engulfed by the quartz overgrowths. Chemical compaction in these sandstones is

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manifested by intergranular pressure dissolution and no stylolitic surfaces have been detected. 5.2.2. Carbonates Total IGV values in sandstones (less than 30%) cemented by blocky and poikilotopic calcite indicates precipitation after considerable compaction and loss of intergranular porosity. The d18O values of this calcite (7.5% to 15%) indicate precipitation at temperatures of 50–100 1C (Fig. 26) if the unmodified marine porewaters with d18OSMOW of 1.2% are assumed and by using the fractionation equation of Friedman and O’Neil (1977). This range of calculated temperatures does not fit well with the calculated temperature from the fluid inclusions (i.e., 125 1C) of these calcite cements. More realistic temperatures are obtained (70–140 1C; Fig. 26C) assuming that calcite precipitation occurred from evolved marine porewaters with a d18OSMOW of +2% (e.g., Lundegard and Land, 1986). The wide range of d13C values of calcite cement (10.5% to +7.9%) suggests variable sources and/or processes for the derivation of dissolved carbon. The uppermost values suggest derivation of carbon from microbial methanogenesis of organic matter, whereas the lowermost values could be attributed to a variety of sources and/or processes, such as the derivation of carbon from thermal maturation of organic matter (Irwin et al., 1977; Surdam et al., 1984). The latter postulation is supported by the calculated precipitation temperatures (70–140 1C; Fig. 26C) using the d18O values of the calcite cement (7.5% to 15%). The weak positive correlation between d18O and d13C values 10

δ13CV-PDB‰

5

0

-5

-10

r=6 -15 -20

-16

-12

-8

-4

0

δ18OV-PDB‰

Fig. 27. A plot showing correlation between d18OVPDB versus d13CVPDB of calcite, siderite and Fe-dolomite/ankerite. The weak trend of positive correlation, indicating that there was a small increase in the input of light carbon (12C) with the increase in temperature.

537

(r ¼ +0.6; Fig. 27) precludes systematic increase in the input of dissolved carbon from thermal maturation of organic matter during progressive burial depth and increase in temperature. The Fe-dolomite and ankerite, which occur mainly in deeply buried sandstones, have been formed through direct precipitation from porewaters and the replacement of eogenetic, ferroan and non-ferroan dolomite and calcite. Using the d18OVPDB values of Fe-dolomite and ankerite (range between 12.1% and 6.3%), the fractionation equation of Fisher and Land (1986) and considering their dominantly deep, mesogenetic origin, precipitation from an evolved brine (d18OSMOW ¼ +1%), suggest precipitation temperatures of 73–128 1C (Fig. 26D). Somewhat higher temperatures (80–140 1C) are obtained if the formation waters were isotopically more evolved to have a d18OSMOW of +2% (Fig. 26). Lower temperatures (57–104 1C) are obtained if precipitation is assumed to have occurred from unmodified marine pore waters (d18OSMOW ¼ 1.2%). The d13CVPDB values (8.5% to 1.7%) indicate a low to moderate input of 12C from the oxidation of organic matter. 5.3. Reservoir quality evolution Reservoir quality of the Palaeocene sandstones has been subjected to overall successive deterioration with increase in burial depth, although there is a wide variation in porosity within each depth interval (Fig. 21). Deterioration in reservoir quality has been mainly controlled by mechanical compaction and, to a smaller extent, by cementation. Although, the presence of grain-coating chlorite has inhibited more pervasive quartz cementation, there is no anomalous porosity preservation in the deeply buried sandstones (44 km). Reservoir quality of the sandstones has been enhanced to various extents owing to secondary porosity development as a result of dissolution of framework silicate and calcite cements. Dissolution of detrital grains and cements during eodiagenesis is presumably related to flushing by meteoric waters (Morad et al., 2000; Ketzer et al., 2002), whereas during mesodiagenesis is poorly constrained, but possibly related to organic acids that were capable of complexing and transporting the immobile Al from dissolving silicates (Surdam et al., 1984; Wilkinson and Haszeldine, 1996). Dissolution of detrital grains, which include volcanic rock fragments, mud intraclasts and feldspars, and thus porosity enhancement, is more common in the LST and HST sandstones especially, below the SB owing to the more effective circulation of the meteoric waters into these sandstones compared with TST sandstones. The dissolution of these detrital grains has resulted in oversized intragranular and moldic pores that display variable degrees of connectivity. The dissolution of calcite cement has also resulted in considerable portion of open, well-connected secondary intergranular pores. The preservation of secondary moldic

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porosity was enhanced by the presence of small amounts of evenly distributed patches of quartz and carbonate cements. However inter- and intragranular moldic porosity decrease systematically with increasing burial depth (Fig. 28). Despite the formation of secondary porosity by grain and calcite cement dissolution, diagenesis has resulted in overall considerable destruction of reservoir quality. The impact of cementation by quartz overgrowths on reservoir quality reduction is greater in distal LST and TST sandstones owing to the more intense compaction and cementation than in the proximal LST and HST sandstones. The influence of compaction on porosity reduction was presumably more important in the distal LST and TST

B

C

plagioclase

K-feldspar

D

E

F

Mesodiagenesis

70˚ C

Eodiagenesis

A

sandstones owing to their finer grain size and relatively more abundant mica contents than in the proximal LST and HST sandstones. Abundant mica would induce greater enhancement to pressure dissolution of quartz grains (i.e., chemical compaction; Thomson, 1959, Gjeldsvik and Bjørkum, 1984). Additionally, the silica released has been frequently re-precipitated as syntaxial quartz overgrowths in the vicinity of the sites of intergranular pressure dissolution, causing further deterioration to porosity of the sandstones (Tada and Siever, 1989). Quartz overgrowths are most abundant cement in the deeply buried sandstones (44.5 km) and have precipitated over a range of burial depths and temperatures (93–114 1C). In the shallower buried sandstones

quartz

kaolinitized mica dickite

siderite

mud intraclast

mica albitized feldspar

microcrystalline calcite

VRF altered to chlorite VRF altered to smectite

illitized mica

qua rtz cement pressure dissolution

volcanic rock fragment (VRF) porosity

bioclast

dolomite cement infiltrated clay

illitized infiltrated clay

Fig. 28. Schematic diagram showing the evolution pathways and distribution patterns of diagenetic alteration for Tertiary sandstones within a sequence stratigraphic sequence. Meteoric water influence in the LST shelf and proximal sandstones is evidenced by extensive dissolution of detrital silicate grains and replacement by kaolinite.

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(o2500 m), the grain-coating illite and chlorite has strongly inhibited the precipitation of quartz overgrowths and possibly deterioration of reservoir quality. The lack of well developed and extensive clay–mineral coatings in the deeply buried sandstones renders it difficult to evaluate their role on inhibition of quartz cementation in these reservoirs and a subsequent impact on reservoir quality evolution. However, other workers indicate that chlorite coating is controlling the reservoir quality of Paleocene turbiditic sandstones (Sullivan et al., 1999). Other diagenetic alterations that have had little impact on reservoir quality of the sandstones include the formation of small amounts of predominantly, grain replacing, Fe-dolomite, ankerite, siderite and pyrite. 5.4. Summary model for the diagenetic and reservoir-quality evolution pathways of the turbiditic sandstones The spatial and temporal distribution of diagenetic alterations and related reservoir evolution pathways of the turbiditic sandstones are controlled by detrital composition, depositional facies, changes in the pore-water chemistry owing to changes in the relative sea level and maximum burial depth reached by the sandstones. HST and proximal LST sandstones have been subjected to greater extent of dissolution and kaolinitization of feldspar, mica and mud intraclasts than the finer-grained, distal LST sandstones. Kaolinitized mica has been expanded to completely fill adjacent pore space, and hence induced deterioration to sandstone permeability. In addition to the role of proximity to the basin margin, kaolinitization was also controlled by the amounts of mica, mud intraclasts and the chemically most unstable feldspars, such as Ca-rich plagioclase and structurally disordered K-feldspars. Variations in detrital composition and/or in the eogenetic alterations of the sandstones within HST, proximal and distal LST resulted in variations in the reservoirquality evolutions, and hence in considerable reservoir heterogeneity (Fig. 28A–F). The types and extent of eogenetic alterations exerted a profound impact on the mesogenetic and thus related reservoir-quality evolution of the sandstones. The formation of illite occurred primarily by illitization of eogenetic grain-coating, infiltrated clays and mica in which its distribution was facies controlled, being most extensive in the pedogenically influenced HST than the finer-grained proximal and distal LST sandstones. The formation of illite in the sandstones suggest minimum burial temperatures of about 90–100 1C (Ehrenberg and Nadeau, 1989; Giles et al., 1992). Considering the burial depth of 3 km of the Tertiary sandstones and assuming a geothermal gradient of 30 1C/km and a surface temperature of 20 1C, would suggest that the minimum temperature reached by the sandstones is about 110 1C. This temperature could account for the formation of illite and extensive quartz overgrowths (Lander and Walderhaug, 1999). Quartz cement occurs as overgrowths and less commonly,

539

as pore filling cement. Quartz cement is most abundant in the deeply buried sediments in the distal LST turbiditic sandstones. The silica needed for quartz cementation was probably derived from the intergranular, mica- and clayinduced pressure dissolution of detrital quartz and from stylolitization (Houseknecht, 1988). Quartz overgrowth is commonly associated with poikilotopic, mesogenetic calcite cement in contrast to the pore-filling quartz cement, which is not associated with poikilotopic calcite cement. Dissolved potassium needed for the formation of illite was presumably provided by concomitant albitization of K-feldspar. Illite increases systematically in abundance with increasing burial depth of the sandstones (Fig. 17B), which is partly due to replacement of eogenetic kaolinite and to the pervasive replacement of unstable framework silicates and clay pseudomatrix (Worden and Morad, 2003). Dickite, which was formed by the replacement of eogenetic kaolinite has not been illitized. The lack of illite in some of the studied sandstones is not because of low temperatures reached but is rather linked to the lack of internal source of dissolved potassium, which was in turn, caused by pervasive to complete dissolution and kaolinitization of detrital mica and K-feldspars. This postulation is obviously evidence for the importance of diffusion rather than potassium supply by large-scale advection (Morad et al., 2000). Another reason for the formation of small amounts of illite in many of the sandstones is the earlier, wide-spread conversion of kaolinite into dickite. Dickite is more resistant towards illitization than kaolinite (McAulay et al., 1994; Morad et al., 1994, 2000) due to smaller degree of crystal structural disorder. Other important diagenetic alteration is the formation of smectitic clays. Smectite formed during eodiagenesis as a result of dissolution of volcanic rock fragment, micas and infiltrated clay coatings. (Wilson and Stanton, 1994). The rapid alteration of volcanic rock fragments and micas associated with significant expansion of these detrital components cause significant reduction of permeabilities. Fe–Mg smectitic volcanic rock fragments, infiltrated clay and micas (mainly biotite) facilitated formation of chloritic clays during mesodiagenesis (Ehrenberg, 1993). Chloritic clays are most abundant in the proximal HST facies, which is attributed to the presence of the suitable precursors such infiltrated smectitc clays. Another important diagenetic alteration that has profound impact on reservoir heterogeneity and quality is cementation by calcite. Calcite in Tertiary sandstones occurs as concretions within the proximal HST. Such a distribution pattern of calcite cement within HST facies is attributed to the increased amounts of carbonate fragments (South and Talbot, 2000; Ketzer et al., 2002, 2003) and/or to the long residence time of the sediments at shallow depths below the sea floor (Taylor et al., 1995). The microcrystalline calcite cement is presumably precipitated during eodiagenesis, and continued as mosaic and poikilotopic crystals during mesodiagenesis. In many studied sandstones, mesogenetic calcite along with other silicates

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(feldspar, mica and volcanic rock fragments) are subjected to extensive dissolution and considerably enhanced the porosity and permeability of the sandstones. In addition to the formation of kaolinite, infiltrated clays, smectite, illite, chlorite, quartz and calcite other diagenetic alterations that contributed to quality and heterogeneity of reservoirs include mainly mechanical and chemical compaction. Mechanical compaction includes grain rearrangement and pseudo-plastic deformation of mica and mud intraclasts. Chemical compaction occurred in sandstones rich in mica and grain-coating illite, which are known to enhance the pressure dissolution of quartz grains (e.g. Heald, 1955; Oelkers et al., 1992). Diagenetic alterations that have had little impact on reservoir quality of the sandstones include the formation of small amounts of predominantly, grain replacing, Fe-dolomite, ankerite, siderite and pyrite. 6. Conclusions The petrological and geochemical study of the Paleocene lowstand, transgressive and highstand systems tract turbiditic sandstones, Shetland-Faroes Basin have revealed that the reservoir quality of the sandstones is controlled by depositional facies, detrital composition, influx of the meteoric waters and burial depth reached by the sandstones. Diagenetic and related reservoir-quality evolution pathways have been accomplished during eo- and mesodiagenesis. Pervasive dissolution and kaolinitization of feldspars, mica and mud intraclasts in the turbiditic sandstones are attributed to flux of meteoric waters, which is attributed to major fall in the relative sea level below the shelf edge. There is no single diagenetic process that is solely or mainly controlling the pattern of porosity evolution of the sandstones. Instead, it appears that the main types of cements (carbonates, clay minerals and quartz) as well as compaction have collectively controlled the reservoir quality evolution of the sandstones. Grain dissolution and kaolinitization of the silicate grains was more extensive in the shelf and proximal than distal sandstones owing to their proximity to basin margin and hydraulic head. Mechanical and chemical compaction was far more important than the cementation in porosity destruction. Chemical compaction was most extensive in the distal turbidites owing to the presence of mica and illitic clays. Illitization of kaolinite and smectite occurred concomitantly with albitization of plagioclase and K-feldspars, which acted as internal source of potassium. Considerable reservoir heterogeneity in turbiditic sandstones is attributed to various facies-controlled, diagenetic evolution pathways in the sandstones. Thus, the porosity enhancement is greatly related to the original framework composition of the sandstones. The dissolution of these grains has resulted in oversized intragranular pores that display variable degrees of connectivity. The dissolution of calcite cement has resulted in considerable open,

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