FACIES-RELATED DIAGENESIS IN THE MAIN CLAYMORE …minersoc.org/pages/Archive-CM/Volume_21/21-4-479.pdf · The Claymore Sandstone Member of Turner et al. (1984) is composed of sandstone

Clay Minerals (1986) 21,479-496

F A C I E S - R E L A T E D D I A G E N E S I S IN T H E M A I N C L A Y M O R E O I L F I E L D S A N D S T O N E S

I . S. C. S P A R K AND N . H . T R E W l N

Department of Geology and Mineralogy, Marischal College, University of A berdeen, Aberdeen AB 9 1A S

(Received 8 July t 985; revised 11 November 1985)

ABSTRACT: Four major sedimentary sequences of the Triassic and Upper Jurassic of the Main Claymore Oilfield of the North Sea each contain a characteristic suite of diagenetic minerals and fabrics. (1) Triassic Skagerrak Formation fluvial sandstones contain early authigenic pore-lining smectite, together with kaolinite and chlorite which form grain replacements and pore fills. Quartz and feldspar overgrowths are minor. Ferroan dolomite forms a late diagenetic patchy poikilotopic cement. Smectite is converted to illite-smectite in a 5 m thick zone beneath the sub-Jurassic unconformity. Smectite formed early in diagenesis prior to oil migration and destroyed permeability. Thus oil is not found in these sandstones although they occur in the oilzone. (2) The Piper Formation (late Oxfordian/early Kimmeridgian) paralic deposits mainly contain authigenic, pore-lining illite-smectite, vermicular kaolinite grain replacements and pore fills. Quartz overgrowths are generally well developed. (3) The Kimmeridge Clay Formation (early Kimmeridgian/early Volgian) comprises thin marine sandstone turbidites, contained within a thick siltstone/shale sequence. In the sandstones (the 'Ten Foot Sandstone') discrete double-ended quartz crystals (1-20 #m) developed prior to quartz, K- and Na-feldspar overgrowths. Only minor kaolinite and lllite-smectite are present. Late diagenetic dolomitic occurs as a patchy poikil0topic cement and as clusters of pore-filling rhombs. (4) The Claymore Sandstone Member (early to middle Volgian) thick marine sandstone turbidites are interbedded with thin siltstones/shales. Sandstones have well-developed quartz, K and Na-feldspar overgrowths, and kaolinite and illite-smectite occur as grain replacements and rarely as pore fills. Late-diagenetic dolomite and ferroan dolomite form poikilotopic cement and clusters of pore-filling rhombs. The major factors which control diagenetic features are depositional environment and associated porewater together with original mineralogy. Burial history and textural features of the sandstones also have important influences. Reservoir quality is controlled by a complex interplay of these features.

The Main Claymore Field was discovered in 1974 and is operated by Occidental Petroleum (Caledonia) Ltd. I t is situated in Block 14/19 in the Outer Moray Fir th Basin, and is contained within a series of tilted and rotated fault blocks which trend roughly N W - S E (Fig. 1). Oil is produced from a variety of reservoir lithologies, some of which are described here to illustrate the role diagenesis has p layed in controlling reservoir quality.

The Triassic fluvial sandstones, which directly underlie the Upper Jurassic reservoir sandstones, contain no moveable oil and are generally not oil bearing, even though they occur above the o i l -water contact. The oil-bearing fine-grained late Oxford ian-middle Volgian reservoir-quality sandstones are capped and sealed by middle Volgian to early Ryazanian shales, which have a relatively high gamma- ray log response (to 180 A P I units). The s trat igraphy and facies types are briefly discussed since they have a strong influence on the subsequent diagenetic features of the varied reservoirs of this oilfield.

1986 The Mineralogical Society

480 L S. C. Spark and N. H. Trewin

S T R A T I G R A P H Y

The stratigraphic terminology used here follows that of Deegan & Scull (1977) and Turner et aL (1984). Rapid thickness and facies variation in the Piper-Claymore area will necessitate future revision of the stratigraphic nomenclature in the light of information gained from recent drilling. The scheme followed here is not necessarily that favoured by the operator.

The Main Claymore Upper Jurassic sediments belong to the Humber Group (Deegan & Scull, 1977), which here comprises the Piper Formation and Kimmeridge Clay Formation which unconformably overlie Triassic sediments of the Skagerrak and Smith Bank Formations. The Claymore Sandstone Member (Turner et al., 1984) forms the bulk of the Main Claymore reservoir.

Smith Bank Formation

The Smith Bank Formation (Deegan & Scull, 1977) comprises predominantly red silty claystones with a few thin sandstone streaks and anhydrite bands. The thickness of the formation varies considerably over the field area with 800 ft in Well 14/19-5, and 600 ft in Well 14/19-3. In Well 14/19-2 1100 ft of Trias are present, of which 700 ft could be considered a sandier facies of this formation

Skagerrak Formation

This formation (Deegan & Scull, 1977) comprises red, brown, white, green and grey conglomerates, sandstones and shales. The formation thins to the NE with 400 ft in Well 14/19-2, 300 ft in Well 14/19-3 and 100 ft in Well 14/19-4 (See Fig. 1 for well locations). The Skagerrak Formation is in part the lateral coarser equivalent of the Smith Bank Formation, and appears to be of middle-late Triassic age (Deegan & Scull, 1977).

58"30"N O0~20'W

B L O C K 14119 J WITCH GROUND

MAP LOCATION 58~ ~ ~

7000

MAIN '~ I 1"5 / CLAYMORE i ~ AREA 4 /

HALIBUT ~ MAJOR FAULTS SHELF O- WELL LOCATIONS

MAIN AREA CONTOUR ~NTERVAL 500 FT. CLAYMORE FIELD

DEPTH STRUCTURE MAP o KM. 1 ON BASE CRETACEOUS o ' MtLES ' ;

I

FIG. 1. Well locations, major faults and reservoir structure of the Main Claymore Oilfield.

Facies-related diagenesis, Main Claymore Field 481

Piper Formation

This formation (Deegan & Scull, 1977) in the Main Claymore Field consists of thin transgressive and thicker regressive sandstones with thin coals. Shales which mark the more extensive transgressions are commonly bioturbated. The sandstones are very coarse to granule sized, very carbonaceous and commonly contain plant rootlets. The formation varies in thickness from a maximum of 90 ft in Well 14/19-2, to 60 ft in Well 14/19-3 and 40 ft in Well 14/19-4; the variation is most likely caused by deposition on an original irregular topography, combined with differential subsidence and deposition rates.

Poor palynological evidence suggests that the Piper Formation is of late Oxfordian- early Kimmeridgian age and therefore it is likely that the Middle Jurassic is absent in the Main Claymore Field area, although reworked Callovian palynomorphs are present. If this is the case, then the basal Piper Formation unconformity is of similar age to the major late mid-Oxfordian transgressive event recognized in the Inner Moray Firth Basin by Rawson & Riley (1982). Turner et al. (1984)separated the basal part of this sequence in 14/19-4 as Middle Jurassic Pentland Formation.

Kimmeridge Clay Formation

Rhys (1974) defined the formation in the southern North Sea, and Deegan & Scull (1977) used Well 15/17-4 as a reference section in the Outer Moray Firth Basin area. The formation consists of dark grey-brown to black carbonaceous shales which commonly have a high radioactivity level, but in the reference well thin siltstone and sandstone interbeds with a much lower radioactivity level are present. The formation has a maximum thickness of 1100 ft in Well 14/19-2 but only 520 ft were recorded in Well 14/19-4 and 450 ft in Well 14/19-3; the variations may partly be due to erosion during the early Cretaceous but are mainly due to varying sedimentation rates at different structural positions on rotating fault blocks. The age ranges from the early Kimmeridgian to Volgian and possibly into the Ryazanian.

The Claymore Sandstone Member of Turner et al. (1984) is composed of sandstone units up to 140 ft thick comprising massive to thin-bedded fine grained sandstones and thin silty shales which commonly contain belemnites. The member varies considerably in thickness with 700 ft in Well 14/19-2, 300 ft in Well 14/19-3 and 270 ft inWell 14/19-4;the variations are caused by depositional thinning over fault blocks. The precise age of the Claymore Sandstone Member is uncertain but probably ranges from the late Kimmerid- gian to middle Volgian.

A lower sandstone unit within the Kimmeridge Clay Formation is informally known as the Ten Foot Sandstone (Fig. 2). This sandstone has a diagenetic history distinct from that of the Claymore Sandstone Member.

S E D I M E N T A R Y F A C I E S

Throughout the Main Claymore area seven major facies have been recognized in cores from the Triassic and Upper Jurassic. These facies have been grouped into four divisions, each of which contains a characteristic suite of diagenetic minerals and fabrics. These are: (1) Triassic fluvial sandstones of facies S (Skagerrak Formation); (2) late Oxfordian-early Kimmeridgian paralic sediments of facies PA and PB (Piper Formation); (3) early


Kimmeridgian-early Volgian marine turbidites of facies K1, K2, K3 and K4 (Kimmeridge Clay Formation); (4) early to middle Volgian marine turbidites of facies K3 (Claymore Sandstone Member). The facies subdivisions adopted in divisions (3) and (4) are similar to those of Turner et al. (1984). In the Piper Formation only two subdivisions are recognized here and are termed PA and PB to avoid confusion with the terminology of Turner et al. (1984).

Skagerrak Formation

Facies S: upward fining micaceous sandstones and conglomerates. Fades S has been studied in core from Wells 14/19-3 and 14/19-4, and occurs throughout the Main Claymore Field area. The fades comprises cycles, from 0.25 to 4 m thick, which commonly have gradational and rarely sharp erosional bases, and grade upward from mud-flake and/or mudstone-pebble conglomerate, through fine-grained, to very fine- grained micaceous sandstones. The conglomerates are poorly sorted with a fine-grained micaceous sandstone matrix. Mudstone elasts are of local origin due to reworking of fine- grained sediments. The micaceous sandstones are white, green or red, with a mean grain size of 2.9-4.1 4. Grains are angular to sub-rounded. Planar and tabular cross-bedded and plane-laminated sandstones predominate over massive and ripple-drift laminated sandstones which occur sporadically towards the top of individual cycles. Small-scale water escape and sand-intrusion structures are present in places, but no macrofossils or bioturbation were observed.

These sediments are interpreted as the distal deposits of a braided river system, on a moderate palaeoslope in a semi-arid environment. Similar sedimentary cycles from sand dominated braided rivers have been described by Elliott (1983) and Cant & Walker (1976, 1978).

Piper Formation

Facies PA: upward fining carbonaceous sandstones. In Wells 14/19-3 and 14/19-4 the facies PA forms fining-upwards cycles, 0.25-2 m thick with sharp and commonly erosional bases. The cycles grade upward from very coarse to granule-sized sandstones, into medium to fine-grained, carbonaceous sandstones with parallel and rarely ripple-drift lamination. Groups of cycles may show an overall upward-coarsening trend. Plant rootlets and vertical burrows occur, and grey-black laminated carbonaceous siltstones and mudstones form the cycle tops. In both wells, a 1 m thick coal seam occurs at the top of one of the cycles, and in Well 14/19-3 a second coal seam 25 cm thick occurs on top of a fine to medium-grained sandstone which contains abundant rootlets.

Facies PA is interpreted as representing varied coastal plain/fluvial deposits possibly of an upper delta plain environment which was subject to periodic marine invasion. Sediments from similar environments of comparable cyclicity have been discussed by Ryer (1977) and Elliott (1983). An increase in marine influence towards the top of the facies is indicated from palynological evidence, which shows that the miospore/dinocyst ratio reduces upwards.

Facies PB: very fine to medium-grained sandstones and siltstones. This facies has been recognized in a 20 ft core section from Well 14/19-3 but does not occur in Well 14/19-4; however it may be present elsewhere in the field area but is difficult to distinguish in


uncored sequences from well logs alone. Facies PB comprise units 0.25-1 m thick, with sharp bases which are rarely erosional, and consists of massive very fine to medium- grained sandstones interbedded with dark grey laminated siltstone. Carbonaceous material is concentrated in parallel laminae and rootlets are also present. Intense bioturbation occurs sporadically in both lithologies. The top of facies PB is marked by a 1 m massive fine to medium-grained sandstone bed, containing rare belemnites.

Facies PB is dominated by the deposits of an offshore shallow marine to shoreline environment which resulted from marine transgression over the coastal plain or delta of facies PA. Only rarely did emergence occur to allow plant growth.

Kimmerldge Clay Formation

Facies KI: dark grey-black finely laminated silty mudstone. Facies K1 consists of uniform dark grey-black silty mudstones which contain bclemnites, whilst body fossils of benthonic macrofauna and bioturbation structures are absent. This facies is often difficult to differentiate in core from the siltstones of facies K2; however, it is readily recognizable in well logs by having a high gamma-ray response of over 100 API units, and rarely as high as 140 API units as for example in Well 14/19-3 at the base of the Kimmeridge Clay Formation (Fig. 2). The facies commonly grades upwards into siltstones of facies K2.

Facies K2: dark grey finely laminated siltstone. The characteristic lithology of facies K2 is dark grey finely laminated siltstone, sporadically calcareous, and rarely bioturbated. Fossils include common belemnites, broken shell fragments, plant fragments, and rare ammonites, phosphatic bone fragments, fish scales, oysters and Lingula. The facies typically has a gamma-ray response of 60-100 API units, and is "best developed in Well 14/19-3 where 60 ft occur towards the base of the Kimmeridge Clay Formation (Fig. 2).

Facies K3: thickly bedded very fine to fine-grained sandstones, with minor interbedded mudstones and siltstones. Facies K3 consists of thin to very thick (up to 43 m), very fine to fine-grained, greyish brown sandstones interbedded with thinner (up to 2.1 m) mudstones, siltstones (similar to facies K 1 and K2), and siltstones which contain thin parallel laminae of very fine to fine-grained sandstone. Most beds are massive and ungraded but some are graded, and show parallel, cross and ripple-drift lamination. Many laminae are rich in carbonaceous plant fragments which also occur scattered throughout the sandstones. Load casts, flame structures, small-scale slump structures, soft-sediment faults and matrix- supported angular siltstone clasts up to 2.5 cm in diameter are locally common, whilst dish structures occur rarely within otherwise massive beds. Calcareous concretions up to 30 cm in diameter are common within the sandstones. This facies comprises the bulk of the Claymore Sandstone Member, and also forms a unit 10 ft thick (the 'Ten Foot Sandstone') within the underlying finer-grained facies of the Kimmeridge Clay Formation (Fig. 2).

Facies K4: bioturbated fine to very fine-grained sandstone. Facies K4 occurs at the top of the Claymore Sandstone Member in Wells 14/19-4 and 14/19-9 but was not recognized in core from Well 14/19-3. K4 consists of intensively bioturbated, fine to very fine-grained silty sandstones, where all primary sedimentary structures have been destroyed. Vertical, U-shaped and high-angle oblique burrows with concave-upwards backfill structures are common, with a few belemnites.

Major basin-edge faulting towards the end of the Oxfordian resulted in the formation of the rapidly subsiding Witch Ground Graben and Southern Claymore Basin which includes the Main Claymore Field area (Turner et al., 1984). The dark muds of facies K1 are


Gamma Ray Lo 9 Depth Sonic Log API UNITS (ft) mic r o l m c o ~ s per foot F l c i l s 0 100 15(:I 50

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FIG. 2. Composite log of Triassic and Upper Jurassic sections in well 14/19-3. This can be compared with the log of Well 14/19-4 illustrated by Turner et al. (1984).


interpreted as being the result of basinal hemipelagic sedimentation, whilst the siltstones of facies K2, which commonly contain broken shell and carbonaceous plant fragments, were possibly deposited by dilute turbidity flows and silt-dominated hemipelagic sedimentation close to the basin margins. The sandstones of facies K3 and K4 were introduced into the rapidly subsiding basin by a variety of gravity-flow processes. These gravity flows were probably initiated seismically as a result of movement along the basin-edge faults, as evidenced by common syn-sedimentary fracturing and dewatering structures within these sandstones. The bioturbated sandstones of K4 may represent a shallower-water facies, or merely better-oxygenated bottom conditions. The sandstones are well-sorted for gravity flow deposits, and may have been reworked from a relatively high-energy depositional environment, probably a late Jurassic shoreline or shelf system developed in the Halibut Horst (Turner et al., 1984). Facies K3 provides the reservoir quality lithologies of which the diagenetic history is described below.

Methods

Sandstones from the four major sequences described above were routinely examined using impregnated thin-sections, SEM with EDS facility, XRD and cathodoluminescence. Material was used from cores of Wells 14/19-2, 3 and 4. Mineral percentages are based on 1000 point counts per slide, and are summarized in Fig. 3. Porosity and permeability data from core analyses were utilized; porosity was not point-counted in thin-section due to the difficulty of estimating microporosity.

S kagerrak Formation fluvial sandstones (Facies S)

The sandstones are composed predominantly of quartz, and are commonly micaceous and clay-rich (Fig. 3). The counted clay mineral percentage includes a high proportion of microporosity between the clay flakes. Untwinned microcline is the most common feldspar (to 7%) with only minor albite (to 1%). Biotite is commonly altered to chlorite, and together with muscovite usually accounts for about 11% of the bulk rock. Smectite and

Stratigraphic interval

and facies

~o ~ l Sandstone

�9 ~= Calcite bE E ;~ Cemented

Sandstone m Concrotlon

Kimrneridge Clay Fm.

Ten Foot Sst. K3

Piper Fm. pA, PB

Skagerrak Frn. S

Quartz

78"7-61.4 (69.3)

51.2-40.9 (45.2)

82"4-72-3 (77.8)

80"7-74.3

(78.9)

54"9- 42~

(50.2)

RANGE AND ( M E A N ) OF BULK ROCK MINERAL % P4umber of

Feldspar Mica Clays Opaques Carbonate samples analysed

26.6-13'6,12'3-0"9 12.5-3'4 3"8-0'2 2'8-0-1 28

(17.2) (5"0) (6'7) (1"2) (0'6)

10"7-5"2 5"4-0'3 7"1-0'3 1"6-0.1 46~1-3K 8

(7'9) (2"8) (2"5) (0'9) (40-7)

17'2-10"6 2"8-0"9 4"2- 1'3 I'0-0'1 3"4- 1"4 5

(14'5) (I-6) (2"9) (0"7) (2"5)

8 '3-4'2 6"1-0'2 13"8-3'8 4 '3-0.2 0"0 4 (6"7) (2'0) (10"2) (2"2)

7.4-5.3 281-1.2 3~5-145 58 -0 .5 12~-0<

(6"8) (11'3) (25'8) (2"5) (3"4) 6

FIG. 3. Composition of the Triassic and Upper Jurassic sandstones, in vol% excluding macro-pore space. Some microporosity is unavoidably included in the clay totals.


kaolinite are equally common and generally make up to 26% of the bulk rock (including microporosity). Ferroan dolomite and less commonly dolomite generally account for about 3%, but can be as high as 13% of the bulk rock. Pyrite is sporadic but can comprise 6% of the bulk rock. Tourmaline, zircon, baryte and anatase/brookite are minor.

Quartz overgrowths are generally absent, but sporadic prismatic quartz crystals up to 10 #m long (principally along the c-axis) are randomly distributed on detrital quartz grains, growth being restricted to rare pore openings between smectite clay plates which line the pores, and are similar to those described by Waugh (1970). Feldspar overgrowths are generally absent, but very rarely occur as small rhombic crystals (Fig. 4A) coating predominantly K-feldspar and very rarely Na-feldspar detrital grains. The overgrowths have similar elemental compositions to the host grains. Some detrital K- and Na-feldspar grains are variably altered to aggregates of kaolinite booklets, probably in part a result of varying degrees of grain alteration and micro-crystalline grain fracturing prior to deposition of the sandstones.

Detrital muscovite and biotite flakes are partially to wholly altered to vermicular kaolinite and chlorite flakes respectively, and aggregates of chlorite rosettes occur rarely at the margins of the chlorite flakes. Authigenic smectite clay (Fig. 4A) lines pores, normally almost completely filling the pores and bridging the pore throats. The smectite clay is fully expandable to 17 /k and the sharpness of the XRD peaks indicates a high degree of crystallinity. This feature, coupled with the delicate nature of the clay flakes (Fig. 4A), indicates an authigenic rather than a detrital origin.

A few small, tabular, well-developed crystals (up to 60 #m) of anatase or brookite have grown within the minor available pore-space between clay plates. Ferroan dolomite and rarely dolomite occur as large poikilotopic rhombic crystals randomly distributed throughout the sandstones. Baryte is present in only one sample from Well 14/19-3, as irregular, poikilotopic cemented blebs (up to 3 mm) randomly distributed throughout the sandstone. The baryte encloses, and appears to post-date, both dolomite and ferroan dolomite. Where the sandstones are brick red, hematite is probably responsible but no well-developed hematite crystals were recognized.

In the top 5 m of the Triassic below the unconformity with the Upper Jurassic deposits, a mixed-layer illite-smectite clay occurs in place of the smectite and has a similar morphology and distribution. No hematite stain occurs anywhere in this 5 m section, but pyrite occurs sporadically as well-developed, commonly twinned octahedral crystals up to 30 #m in size. Pyrite was not found in any of the red hematite-stained sandstones but only in the white and green sandstones.

The sandstones have a He porosity range of 9-22% and liquid permeabilities up to 20 md (Fig. 5). The major factor which appears to control permeability and to a lesser extent porosity is the abundance of smectite and illite-smectite clay which typically bridges the pore throats. Diagenetic clay formation and permeability destruction took place prior to oil migration so essentially no moveable oil exists in these sandstones even though they occur above the oil-water contact in Wells 14/19-1, 3 and 4. Porosity occurs mainly as microporosity between clay flakes, and to a lesser extent as open pore space free of clays towardS the pore centres. The loss of porosity caused by the formation of smectite and illite-smectite is slightly offset by the formation of a secondary microporosity created by the alteration of detrital micas and feldspar grains to kaolinite clay. The formation of later diagenetic dolomite, ferroan dolomite and minor baryte cements only locally reduces porosity and permeability.


Possible sources ofauthigenic minerals. Most detrital grains have point- or straight-edge contacts and pressure solution at these contacts is slight and therefore unlikely to be a major source of local solution for authigenic mineral formation. Cecil & Heald (1971) considered that thick clay films are effective in retarding pressure solution and secondary quartz formation. In these Triassic sandstones the smectite pore lining may be of very early formation and may possibly represent recrystallization prior to compaction of an infiltrated detrital clay. Alteration of framework grains, particularly the micas and feldspars, appears to be the major source of dements necessary for the formation of authigenic minerals. As shown by Bjorlykke (1983) early diagenesis, related to 'subsurface weathering' in the vadose zone of an arid environment, releases Na, K, Ca, Mg, AI and SiO 2 from the alteration of silicate minerals such as hornblende, augite, biotite, muscovite and feldspars, and H + ions are absorbed. The common result is the early formation of smectite, kaolinite, quartz and feldspar overgrowths, carbonate cement and iron oxides. In the Main Claymore Triassic sandstones it is clear that the alteration of feldspars, muscovite and biotite would result in the concentration of K, Na, SiO 2, A1, Ti and Fe in the porewater, with the

FIG. 4. (A) Pore-lining smectite (S) and K-feldspar overgrowths (F); Triassic Skagerrak Formation. (B) Pore-filling vermieular kaolinite (K); Piper Formation. (C) Discrete double- ended quartz crystals (Q) coating a detrital Na-feldspar grain which has an irregularly developed overgrowth (F); Ten Foot Sandstone of Kimmeridge Clay Formation. (D) Well-developed, syntaxial quartz overgrowths (Q), and K-feldspar overgrowths (F); Framework grain altered to an uncompaeted aggregate of kaolinite (K) and iUite-smectite (IS) clays; Claymore Sandstone

Member, Kimmeridge Clay Formation.


subsequent authigenic formation of smectite, iron oxides, anatase or brookite, and quartz and feldspar overgrowths. Dolomite most likely precipitated from the porewaters, which in arid envrionments are frequently enriched with Mg and carbonate, due in part to dissolution of ferromagnesian minerals such as augite and hornblende (Bjorlykke, 1983). The later ferroan dolomite was probably deposited from upward-migrating solutions rich in Mg, and carbonate derived from the solution of underlying Permian dolomites. The present shallow burial and preservation of smectite shows that the ferroan dolomite of the Skagerrak Formation is not the result of illitization reactions within the Triassic rocks. The Ba for baryte could similarly have been derived from the underlying Permian.

The later diagenesis involving illite-smectite was influenced by the incoming of marine porewaters derived from the overlying unconformity from the late Oxfordian onwards. A reduction zone was established and the earlier-formed iron oxide was reduced to pyrite, and smectite was converted to illite-smectite. This diagenetic phase is clearly linked with that of the Piper Formation (described below) where subsurface reducing conditions and abundance of sulphate resulted in pyrite formation. The maturation of organic material (mainly plant debris) in the Piper Formation during later burial also probably supplied organic acids to the porewater which invaded the adjacent Triassic sandstones.

The sequence of authigenic mineral formation is difficult to reconstruct because of the rarity of some of the minerals and because the various authigenic minerals are seldom observed in contact. However, since all the authigenic minerals, with the exception of kaolinite and chlorite, enclose the smectite, the proposed phases of authigenic mineral formation are as follows: (1) kaolinite, chlorite, and smectite; (2) quartz and feldspar overgrowths, anatase or brookite; (3) dolomite followed by ferroan dolomite and baryte; (4) pyrite and illite-smectite developed beneath the sub-Jurassic unconformity. The sequence of authigenic mineral formation (Fig. 6), is very similar to that in other red-bed Triassic sandstones, such as those described in greater detail by Burley (1984).

Piper Formation paralic deposits (Facies PA and PB)

The sandstones of facies PA and PB are mineralogically similar with predominant quartz (to 80 vol%). Untwinned microline is the commonest feldspar (to 8%) with rare albite (to 0.1%). Micas, mainly biotite and lesser muscovite, form ~2% of the bulk rock. Clay minerals, mainly kaolinite with minor illite-smectite, form ~10%, and include a relatively small proportion of the micro-porosity of the bulk rock. Pyrite is very common in many samples (to 4%), locally as cement.

The sandstones of both facies have undergone a similar diagenetic history, except that syntaxial quartz overgrowths of facies PB generally tend to be more fully developed than those of facies PA. Detrital K- and Na-feldspar grains are usually partially to wholly altered to aggregates of kaolinite booklets, but more commonly to vermicular kaolinite growths (up to 700 pm long) which are commonly pore-filling (Fig. 4B). Illite-smectite sporadically occurs as a thin rim-coating on detrital framework grains, but is minor in comparison to kaolinite which makes up the bulk of the clay mineral assemblage of the rock (Fig. 3). Muscovite and biotite are very rarely altered to vermicular kaolinite, the crushed and bent mica flakes being more prone to alteration than the physically undeformed flakes. Feldspar overgrowths were not recognized in any of the samples studied. Syntaxial quartz overgrowths are common, poorly to well developed, and appear to have been partially restricted in some cases by the clay minerals which are partially or

Facies-related diagenesis, Main Claymore Field o

3 0 %

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20 %'..~�9149 �9149 " �9 "�9149 �9 ; �9

15

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PERMEABILITY md.

FIG. 5. Relationships between porosity and permeability in Well 14/19-3. Other wells generally show lower permeability. (A) Triassic sandstones facies S. The points circled are those from the top 5 m of the Triassic below the sub-Jurassic unconformity where smectite has been altered to illite-smectite. (B) Piper Formation facies PA and PB. (C) Claymore Sandstone Member (dots) and the Ten Foot Sandstone (circles) to show higher preserved porosity in the Ten Foot Sandstone. The three other data points with high porosity are from plugs at the base of the

Claymore Sandstone Member adjacent to shaly rocks.

wholly engulfed by quartz overgrowths. Pyrite forms aggregates of well-developed, pore-filling pyritohedra, which locally develop into a slightly replacive patchy cement. Carbonate cement was not recognized in any of the samples.

Porosity ranges from 12 to 25% and permeability from 0.2 to 800 md in the sandstones from both facies (Fig. 5). The two major factors which appear to control the porosity and permeability of the sandstones are the formation of kaolinite and, to a lesser extent, pyrite. The higher values of porosity and permeability are from sandstones having ~4% kaolinite and ~0.2% pyrite, and in which the kaolinite occurs as feldspar grain replacements and only rarely as pore fills. In contrast, the lower recorded values of porosity and permeability


are from sandstones having ~ 14% kaolinite and ~4% pyrite. In these examples the kaolinite also occurs as feldspar grain replacements and also commonly as pore fills. In the clay-rich sandstones the porosity is mainly micro-porosity between individual clay platelets, booklets and vermicules. The effect of quartz overgrowth formation on the reduction of porosity and permeability is very minor in comparison to that of kaolinite and pyrite formation.

Possible sources of authigenic minerals. These sandstones have a high organic content (up to 2% of the bulk rock) of carbonaceous plant fragments. Marine-brackish porewaters probably resulted in the early diagenetic formation of pyrite by the reaction between iron-bearing minerals and H2S produced from sulphate in the seawater by sulphate- reducing bacteria (Berner, 1970). According to Bjorlykke (1983) kaolinization of feldspar is favoured in acid porewaters, here produced by humic acids and CO 2 from the degradation of organic plant material.

Silica released during feldspar alteration probably formed quartz overgrowths in these sandstones. Another possible source of quartz for overgrowth formation may be from the mixing of fluvial surface waters, containing an average 13 ppm dissolved SiO2, and marine porewater with an average 6 ppm dissolved SIO2, resulting in the precipitation of quartz as outlined by Blatt (1979). The alternation of marine and freshwater deposition gives the potential for frequent mixing of porewaters of variable chemistry early in diagenesis. Pressure solution at grain contacts is insignificant and therefore unlikely to be an important local source of SiO z for overgrowths. Some of the illite-smectite clay may have originated from infiltration of detrital clay during or just after deposition of the sandstones. Diagenetic illite-smectite formed by direct precipitation from the porewaters, where the source of K + may have been from the kaolinization of feldspar, and/or from the partial dissolution of feldspars in the acidic porewaters. Feldspar dissolution would also mobilize AI and Si for clay mineral formation (Bjcrlykke, 1983). The proposed sequence of authigenic mineral formation (Fig. 6) is as follows: (1) ?illite-smectite (by infiltration), pyrite; (2) kaolinite, illite-smectite (by kaolinization or dissolution of feldspars); (3) quartz overgrowths.

Kimmeridge Clay Formation; Ten Foot Sandstone turbidites (Facies K3)

Samples from the Ten Foot Sandstone turbidite unit (facies K3) enclosed in siltstones (facies K1 and K2) of the Kimmeridge Clay Formation (Fig. 2) are composed predominantly of quartz (to 82%). Feldspar occurs as untwinned microcline (to 16%) with only minor albite (to 1%). Micas are rare, with muscovite and biotite accounting for only about 3% of the bulk rock. Generally clay minerals are very rare, illite-smectite being the commonest (to 4%) mainly as squashed detrital clay pellets and rarely as a thin pore lining, whilst kaolinite is very rare as feldspar grain replacements. Pyrite, dolomite, zircon and tourmaline are minor.

Aggregates of tiny (1-10 /zm) double-ended quartz crystals (Fig. 4C) coat all the framework grains and are partly engulfed by later quartz and feldspar overgrowths. Syntaxial quartz overgrowths are common and usually poorly to moderately well developed their growth having been inhibited by the double-ended quartz crystals. This has resulted in many randomly distributed, but similarly orientated, prismatic quartz crystal overgrowths on individual detrital quartz grains. Rarely, detrital quartz grains have very well developed euhedral quartz overgrowths which, in thin-section, show two generations of growth. These grains have been reworked from a previously quartz-cemented sandstone.


Major Sedimentary

Sequence

Claymore Sandstone Member

turbidites K3

Auth igen ic M inera l

Quartz over(~lrowt hs feldsi~overgrowths

kaolinite illite ,;smectite

pyrite Fe / dolomite

calcite

Early Oiagenetic Sequence

Late

m

m

Kimmeridge Clay Formation Ten Foot

Sandstone turbidites

K3

Piper Formation PA, PB

quartz overgrowths quart z crystals

teldsp overgrowths kaolinite

i l l i t (P.smecti te pyrite

dolomite

q~rtz overgrowths kaolinite illite ~smectite

pyrite - ?

m

m

m a n

i

Skagerrak

Formation

S

quartz overgrowths feldsp overgrowths ~kaolinite

smectite

m /

m

I I I i m - X -

? '~ '1 i m - X -

FIG. 6. Sequences of authigenic mineral formation in the Triassic and Upper Jurassic sandstones of the Main Claymore Oilfield. Thick bars indicate the most important diagenetic phases connected with porosity and permeability preservation and destruction. Diagenetic sequences in the Jurassic probably developed at about the same time and pre-date oil migration. Diagenetic phases in the Triassic sandstone were formed prior to deposition of the Piper Formation with the exception of the phases indicated by asterisks which formed beneath the sub-Jurassic

unconformity. It is also possible that the ferroan dolomite and baryte are of later formation.

Abundant moderately to well-developed K-feldspar and rarely Na-feldspar overgrowths occur, with host grain and overgrowth having similar elemental compositions. Again, the formation of the overgrowth appears to have been inhibited by authigenic double-ended quartz crystals which coat the detrital feldspar grains. (Fig. 4C). Feldspar overgrowths are partially enclosed by quartz overgrowths, which continued to develop after the cessation of feldspar overgrowth development. Dolomite occurs as a minor patchy poikilotopic, slightly replacive cement, and as slightly replacive pore-filling clusters of small (to 100 gm), well-developed rhombic crystals. Both forms of dolomite enclose the tiny double-ended quartz crystals and partially replace the quartz and feldspar overgrowths and host grains. Minor framboidal pyrite (to 60/Jm diameter) occurs in pore space and is enclosed by quartz overgrowths. Micas (muscovite, and lesser biotite) are unaltered and normally have large flat contacts with framework grains. Intense dissolution of the quartz and feldspar framework grains occurred against mica with almost complete removal of some grains. This intense solution of framework grains also occurs where they are in contact with carbonate shell fragments, and carbonaceous plant fragments.

Porosity ranges from 25-30% and permeabilities from 350-1400 md in Well 14/19-3 (Fig. 5). The major textural features influencing porosity and permeability of these sandstones are the abundant discrete double-ended quartz crystals, overgrowth development, grain size, and sorting. The sandstones are moderately to well sorted and grain size


varies from a mean ~ value of 2.71 to 3.01; the finer grained, poorer sorted sandstones have lower porosities and permeabilities. The rare, widely spaced distribution of illite- smectite and dolomite cement results in only local reduction of the porosity and permeability of the sandstones. The early deposition of the small quartz crystals on grain surfaces was an important factor in porosity preservation. These crystals provided cement when primary porosity was high, but only occupied a small proportion of the porosity. With further burial the crystals inhibited the deposition of larger pore-tilling quartz and feldspar overgrowths.

Possible sources of authigenic minerals. SEM observation of leaf fragments from the surrounding very thick siltstones and shales which sandwich the Ten Foot Sandstone (Fig. 2) shows that, within the pore space inside the leaves, minute (up to 10 am) authigenic double-ended quartz crystals occur. The only possible source of silica for the development of these crystals within the pore space of the leaf fragments is from the siltstones and shales which enclose them. Fuchtbauer (1983) showed that quartz cementation within individual sandstone beds was greatest near the contacts with shale beds and suggested that silica, derived from pressure solution of quartz grains within the shales, had migrated by diffusion from the shales into the sandstones with the subsequent precipitation of quartz as overgrowths.

Since the size and form of the discrete quartz crystals both in the leaf fragments and sandstones are similar, the enclosing shales are considered to be the source of silica for the discrete quartz crystals in the adjacent sandstones. The fact that discrete crystals rather than syntaxial quartz overgrowths developed, suggests that the porewater within the sandstones was supersaturated with respect to silica. Since feldspars and micas are essentially stable in marine porewaters (Bjorlykke, 1983), alteration of the framework grains is not an important source of elements for authigenic mineral formation in the Ten Foot Sandstone. Within the sandstones, pressure solution of the framework grains during burial, with subsequent reprecipitation as quartz and feldspar overgrowths, seems the most likely source for these authigenic phases.

Since the development of the quartz and feldspar overgrowths post-dates the formation of the tiny discrete quartz crystals, and assuming that the sources of these authigenic minerals are as suggested above, pressure solution within the siltstones and shales may have occurred at shallower depths of burial than within the sandstones. Belemnites and carbonate shell fragments in the siltstones and shales have undergone intense pressure solution during burial which, coupled with Mg from the porewater within the sandstones, may possibly be the source of the authigenic dolomite cements. Usdowsky (1968) showed that Mg 2+ cannot enter the carbonate mineral lattice at low temperatures in marine porewater, a feature which supports the late diagenetic origin of the dolomite.

The sequence of diagenetic mineral formation is interpreted as: (1) pyrite, minor kaolinite and illite-smectite; (2) discrete double-ended quartz crystals; (3) feldspar and quartz overgrowths; (4) dolomite. The minor dolomite cement and pore-tilling aggregates of small dolomite rhombs post-date the formation of all the other authigenic minerals; they were not observed in contact and are included together in the authigenic mineral sequence which is summarized in Fig. 6.

Claymore Sandstone Member turbidites (Facies K3)

The sandstones of the Claymore Sandstone Member (Fig. 2) are predominantly


composed of quartz, are feldspathic and rarely arkosic (Fig. 3). The feldspar is mainly untwinned microdine with minor (to 2.5%) albite. Muscovite and biotite, roughly in equal proportions, form ~5% of the detritai fraction. Kaolinite and illite-smectite, roughly equally account for 7% of the bulk rock. Dolomite, ferroan dolomite, pyrite, carbonaceous material, tourmaline and zircon are minor. The mineralogy of the calcite-cemented concretions is similar except that calcite comprises up to 40% of the bulk rock.

Authigenic clay minerals are rare in these sandstones with only sporadic aggregates of pore-filling kaolinite booklets, and a sparse rim-cement ofillite-smectite. Most pores remain free of clay minerals and the sandstones have a clean texture. The bulk of the clay in these sandstones occurs as compacted and distorted detrital clay pellets, and as altered feldspar grains consisting of uncompacted aggregates of kaolinite and illite-smectite (Fig. 4D). Rarely, detrital feldspar grains have undergone severe dissolution and occur as skeletal framework grains. Detrital quartz grains generally have very well developed euhedral, syntaxial quartz overgrowths (Fig. 4D), but in the few cases where illite-smectite rim cement and pore-filling kaolinite booklets are in contact with the quartz grains, only poorly developed overgrowths occur which engulf the clay minerals.

Both detrital K-feldspar and much rarer Na-feldspar grains have well-developed overgrowths (Fig. 4D), with the host grain and its overgrowth having similar compositions. The feldspar overgrowths are commonly partially enclosed by the quartz overgrowths, as in the Ten Foot Sandstone. The feldspar overgrowths generally have a slightly different extinction angle from the host grain, and twinning does not extend into the overgrowth.

Detrital muscovite and biotite flakes are normally unaltered and uncrushed and nearly always form straight contacts with other framework grains. Intense solution transfer has occurred at these contacts, and in some cases nearly the whole framework grain has dissolved. Similar intense solution of the quartz and feldspar framework grains has occurred at the contacts with shell fragments and carbonaceous plant fragments. Pyrite forms rare framboids, engulfed by quartz overgrowths, and also highly replacive patches up to 1 mm in diameter. Minor dolomite and ferroan dolomite (<3%) occur as patchy replacive poikilotopic cement and as clusters of pore-filling rhombic crystals (to 100 gm).

Calcitic concretions, up to 30 cm in diameter, occur within the sandstones, in which calcite forms a highly replacive poikilotopic cement. Cathodoluminescence studies of these calcite-cemented sandstones shows clearly that the feldspar overgrowths (which are non-luminescent), and their host detrital feldspar grains (with bright blue luminescence), are both highly replaced by the calcite cement (with bright yellow luminescence). Both the dolomite and ferroan dolomite cements, (with red luminescence), are completely enclosed by the calcite cement. Quartz overgrowths and host grains are both partially replaced by the calcite cement. The above evidence clearly shows that the calcite concretions are of late diagenetic formation. For further interpretation of the carbonate sequence, isotopic data are required but the small proportions of dolomite would prove difficult to separate.

Core analysis of the sandstones, excluding the concretions, gives a porosity range of 13-25% and a permeability range of 3-1000 md, with most data clustered in the range 17-25% porosity and 20-800 md permeability (Fig. 5). The lower values of porosity are recorded from sandstones containing abundant carbonaceous plant fragments. These rocks have suffered a high degree of chemical compaction due to solution at grain contacts. The local production of carboxylic acid during the maturation of the organic material present may be responsible for the observed grain solution. Since a strong framework was not present, solution by organic acids resulted in local porosity reduction, rather than


recognizable secondary porosity (Surdam et al., 1984). Sandstones having recorded porosities and permeabilities higher than 20% and 100 md, respectively, generally contain no carbonaceous material and were less affected by organic acids of local origin. The main factors which affect porosity and permeability of these marine sandstones are the degree of quartz and feldspar overgrowth development and the grain size and sorting of the sediments. The mean grain size varies from 2.4 to 3.7 4, and sorting from moderately well to well sorted. The higher values of porosity and permeability are recorded from the well sorted, coarser grained sandstones. The clay minerals, dolomite and ferroan dolomite cements, are rare and of widely spaced distribution, with the result that porosity and permeability of the sandstones is only locally reduced. In the basal metre of the member, higher porosities and permeabilities (Fig. 5) indicate a localized diagenetic history similar to that of the Ten Foot Sandstone.

Possible sources of authigenic minerals. The main sources of the clay minerals are detrital clay pellets, infiltrated detrital illite-smectite clay and alteration, particularly of detrital Na-feldspar. However, most detrital feldspar grains are unaltered, suggesting that altered grains were more strongly weathered and micro-fractured prior to deposition of the sediments. Solution at grain contacts, particularly with detrital muscovites, biotite, carbonaceous plant fragments and carbonate shell fragments, appears to be the most likely source for authigenic quartz and feldspar overgrowths. Compaction fluids derived from intervening thin siltstone/shale beds may also be a source of silica for quartz overgrowths.

Pyrite formed during early diagenesis, similarly to that described for the paralic deposits of facies P. The source of the minor late diagenetic dolomite is probably similar to that described for the Ten Foot Sandstone turbidite unit. The ferroan dolomite is typical of late carbonate cements (Franks & Forester, 1984). If the Fe and Mg originated from illitization of smectite as suggested by Boles & Franks (1979), solutions must have been derived some distance from more deeply buried strata since smectite occurs in the underlying Trias. The Fe may be derived from within the formation porewater, and incorporated in the dolomite due to higher temperature of formation during burial. Usdowsky (1968) showed that Fe cannot enter the carbonate lattice at low temperatures in marine porewater, but only at elevated temperatures and pressures. Any Fe remaining after early formation of pyrite could be utilized in the formation of ferroan dolomite later during burial diagenesis. Since the proportion of ferroan dolomite is so small (<3%), the required amount of Fe could have a variety of origins, including minor alteration of iron-bearing detrital minerals such as biotite. This is supported because ferroan dolomite is late diagenetic, replaces quartz and feldspar overgrowths, and commonly encloses the early diagenetic clay minerals.

The calcite cement of the concretions appears to be the last authigenic mineral to form, since it encloses dolomite. Calcite formed rather than dolomite or ferroan dolomite possibly because the porewaters within the sandstones were sufficiently depleted in Mg and Fe by the earlier formation of the two latter minerals thus allowing only calcite to precipitate. Isotopic data might help interpret the origin of this late stage concretion cement. The proposed sequence of authigenic mineral formation in the Claymore Sandstone Member is" (1) pyrite; (2) kaolinite and illite-smectite; (3) feldspar and quartz overgrowths; (4) dolomite and ferroan dolomite; (5) calcite (Fig. 6).

C O N C L U S I O N S

Several factors caused considerable variation in the diagenetic history of the facies


described, with subsequent strong influence on the reservoir characteristics of the sandstones as reflected in their porosity and permeability. In both the Triassic fluvial (facies S) and Jurassic paralic deposits (facies P), the environment of deposition strongly controlled the early diagenetic history. Alteration of framework grains was due to circulation of acid-neutral porewaters in the Triassic (facies S) sandstones, and mixing of acidic terrestrial porewaters with alkaline marine porewaters in the Jurassic paralic sandstones. The early formation of clay minerals, particularly smectite and illite-smectite in the Triassic fluvial deposits, and kaolinite in the Jurassic paralic deposits, appears to have inhibited pressure solution of the framework grains later during burial diagenesis. In the Triassic sandstones smectite effectively destroyed permeability, but the kaolinite in the Jurassic sandstones was less damaging to permeability.

In the two examples of marine turbidite deposits (facies K3), alkaline porewaters of marine origin prevented the alteration of framework grains during early diagenesis since micas and feldspars are essentially stable under alkaline conditions (Velde, 1983; Bj~rlykke, 1983). Later, during burial diagenesis, dissolution at grain contacts with subsequent reprecipitation controlled the formation of authigenic minerals.

The pH of the original porewaters, itself a factor of climate, original depositional environment and rock composition, strongly controlled early diagenesis. Dissolution of the framework grains was influenced to a large extent by the presence or absence of earlier formed diagenetic clay minerals and detrital grains such as muscovite, carbonaceous plant fragments and carbonate shell fragments. Local variations in cementation can be linked with organic content and are probably due to solution effects due to carboxylic acid production during organic maturation. Local variation in quartz cementation was related to dewatering of adjacent shales, and produced an early generation of small, double-ended quartz crystals prior to quartz overgrowth formation. The original detrital mineralogy, organic content and depositional environment strongly influenced the later burial diagenesis of the sandstones. Within each facies described the coarser, better sorted, sandstones have the best reservoir characteristics, except where late diagenetic carbonate cementation has taken place.

ACKNOWLEDGMENTS

Occidental Petroleum (Caledonia) Limited and its partners in the Claymore Oilfield kindly allowed access to the core material and data used in this study and permitted publication of this paper. The Production Geology Group of Occidental in Aberdeen have provided much useful discussion on the topic. The technical facilities of the Department of Geology and Mineralogy, University of Aberdeen are gratefully acknowledged. This work was completed during the tenure by I.S.C.S. of a Natural Environment Research Council studentship. N.H.T. thanks Occidental for generous grants to enable purchase of the SEM and ancillary equipment used in this study.

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FACIES-RELATED DIAGENESIS IN THE MAIN CLAYMORE …minersoc.org/pages/Archive-CM/Volume_21/21-4-479.pdf · The Claymore Sandstone Member of Turner et al. (1984) is composed of sandstone