Sedimentary Facies Analysis (IAS Special Publication, No. 22)

Bude Formation subsea fans

Fig. 12. Traces in the muddy siltstone facies originally interpreted by King (1965) as traces of xiphosurid mating behaviour, but reinterpreted by Higgs (1 988) as trails left by the fins of bottom-feeding fish (Undichna). Structures have a wave length of 9 em (SS 2017 0751).

Fig. 13. Grey muddy siltstone laminations (base of section) overlain by fine-grained graded beds, then passing up into laminated beds showing 'type 3' ripple drift of Walker (1 963) (SS 201 074).

113

parallel-sided beds 5-25 cm thick. They are gen­ erally laterally continuous, but may be scoured or may themselves expand into pre-existing scours. The sole marks are predominantly linear current­ ridge moulds, flute or groove moulds. The lower part of each bed is structureless and grading is obvious only in the topmost part, where aggregates of plant material and fragments of black mudstone may be concentrated (Fig. 18). Median grain size at the base of the unit is around 2.5 ¢, decreasing to about 3.0¢ at the top. The long axes of grains (Fig. 19) in the structureless unit show slight pre­ ferred orientation in the same direction as that of current flow indicated by sole marks on the base of the bed, and, although the bed as a whole is poorly sorted, the grain-sorting at any one level is moderate to good (excluding matrix oG 0.6-0.75), although matrix content may be up to 50% (Fig. 7). The structureless silty sandstone unit is overlain by a 5-20-mm-thick unit of uniform parallel-laminated silty sandstone of slightly finer grain size and then by a similar thickness of current ripple cross-laminated silty sandstone with either single sets or ripple drift. Water-escape structures or convolution are generally not found in these beds. The cross-laminated unit is overlain by a unit of alternating light and dark parallel-laminated silts, which is, in turn overlain by structureless grey silt. Structureless sandstones are sharp-based and between 10 cm and 1 m thick (Figs 20 & 21). The most common sole marks are flute moulds and linear

114

R. V. Burne

Fig. 14. Sharp based, laminated bed showing irregular ripple-drift with stoss side erosion and convolutions due to water-escape. Note graded top of the 12-cm-thick bed (SS 1999 0493) .

Fig. 15. View to the right of Fig. 14 showing a sharp erosive top to the same bed, and an absence of convolutions (SS 1999 0493) .

current-ridge moulds. Most of the bed is ungraded and most beds are completely structureless, although some contain faint traces of lamination, often dis­ rupted into dish structures (Stauffer, 1967). They are composed of moderately well-sorted fine-grained sand (excluding matrix aG 0.85-0.9; median grain size 2-2 4 <jl; matrix content 14-25%) (Fig. 7), with grains orientated either parallel to the current direction indicated by sole marks, or bimodally about this direction (Fig. 19). The beds may have .

sharp tops or may grade up, through an abrupt and irregular contact, into either irregular lamination comparable to hummocky cross-stratification, current ripple cross-lamination (Fig. 22) or, very rarely, dune-shaped cross-sets. The beds are gen­ erally laterally continuous, but are prone to amal­ gamation. They also can be scoured or can thicken in channels. They may also be confined to channels. 'Slurried' beds (Wood & Smith, 1959) consist of sharp-based units of a poorly sorted slurry of sand,

Bude Formation subsea fans

Fig. 16. Sandstone surface bearing straight-crested ripples (SS 2014 0779).

silt and mud (excluding matrix oa 0.9; including matrix oa 2.5; median grain size 3.5 ¢, matrix content 56-70%) (Fig. 7). They are generally between 2 and 20 em thick (Figs 23 & 24). The sharp base of the bed may bear sole marks, most commonly groove or prod and bounce moulds. The lower parts of the beds show grain orientation in the same sense indicated by sole marks (Fig. 19). The beds may be ungraded or graded only in the topmost part, but they commonly show a two-part division, with a basal massive, indistinctly laminated or, rarely, cross-laminat.ed silty sandstone, and an upper, more poorly sorted 'slurried' unit often containing large contorted fragments of mudstone. In places beds have 'frozen' in the act of peeling up fragments of mudstone from the substratum (Fig. 25). Sandstone or siltstone pseudo-nodules (Fig. 18), interpreted as load balls (Kuenen, 1965), and water-escape tubes may be found in the upper part of the 'slurried' unit.

115

Fig. 17. Sandstone surface bearing cuspate-crested ripples (SS 2014 0774).

Fig. 18. Interbedded sandstone-shale facies showing graded silty sandstones (A) separated by a 'slurried' bed (B). Note load balls falling from laminated unit on top of the 'slurried' bed, and laminated top of the underlying graded silty sandstone (SS 2018 0740).

116 A

R. V. Burne

N 41 Spec. 48

B

C

Irregularly Laminated Sandstones

0

Spec. 14

Graded Silty Sandstones

N 31 5

N 30 2

Spec. 49

Spec. 2A

Structureless Sandstones

N 31 8 Spec. 2C

N 295 Spec. 2B N 3 83

N 40

Spec. 73

Spec.

6 60

1

D

E

"Slumped" Bed

"Slurried" Bed

N

I

1

N 2 87 Spec. 4

N 27 5 Spec. 30

N

N 274 Spec. 52 b

N 3 02 Spec. 31 10

15

I

20

%

Fig. 19. The orientations of long axes of elongate grains N 2

69

Spec. 58

N

=

Number of grains

Spec.

=

Specimen number (BURNE 1969b)

10

15

20

%

This unit may be overlain by a thin unit of laminated sediment consisting of either current ripple cross­ lamination or the cross-laminated deposits of small sand-volcanoes. 'Slumped' beds are 4-20-m-thick beds of disrupted sediment (Figs 26-29), similar in character to the thinner 'slurried' beds (Figs 7 & 19). The name is well established (e.g. Freshney et al., 1979), but is misleading because the beds are probably not

within representative specimens of various types of sandstone bed (data from Burne, 1969b). Measurements were made from orientated thin-sections sawn parallel to stratification. Palaeocurrent directions from sole marks or cross-laminations associated with the beds are indicated by an arrow on the rose diagram.

true slump deposits. They are of variable extent. Some appear to have a sheet-like form correlatable between sections (King, 1967), whereas others are known from only one section (Edmonds eta!. , 1979; Freshney et a!., 1979) and may pass both laterally and downward into undeformed material. The beds are composed of sediment with three separate origins; exotic muddy sandstone (excluding matrix o0 1.2; including matrix o0 2.5; median grain size

Bude Formation subsea fans

117

Fig. 20. Structureless sandstones within the interbedded sandstone- shale facies showing traces of irregular lamination, possibly due to the amalgamation of individual units. Length of extended tape is 135 em (SS 1998 061 1 ) .

3.5
Thick sandstones with few visible structures are diagnostic of the Bude Formation (Fig. 30). Some sandstones are truly structureless, but in others structure seems to have been obscured by the lack of contrast in grain-size in the beds. Irregular bedding, cross-bedding, parallel lamination, convolute lami­ nation and parting lineation can all be discerned in places. Some sandstones consist of an amalgamated series of laterally continuous beds of structureless sandstones similar to those described above. Also found in this facies are irregularly laminated sandstones, which are relatively coarse-grained (median grain size of 2-2.9
than load moulds are not common, although moulds of transverse mud ripples have been observed. They may be laterally continuous, or they may wedge out laterally. The beds are characterized by various irregular and often interfering non-parallel lami­ nations, although the internal structures are some­ times difficult to distinguish due to lack of grain size contrast. In places, uniform cross-stratification in scalloped based sets up to 20 em deep and 30 em across occur throughout the bed (Fig. 31). The beds are composed of well-sorted medium sand, and represent the coarsest sands observed in the Bude Formation. The grain fabric shows a preferred orientation at right angles to the dip of the lamination (Fig. 19). Small 1-3-m-deep scours occur near the top of some sandstones (Fig. 32). These scours are often filled by overlying beds draping into them. The fill is similar to the sediment into which the scour is cut. Deeper (> 4 m), fiat-bottomed channels also occur in this facies. The fill is confined by the channel margins and comprises only a few units of either massive sandstone or irregularly laminated sand-

118

R. V. Burne

Fig. 21. Structureless sandstones showing irregular structures possibly due to amalgamation. Length of extended tape is 70 cm (SS 2014 0775).

Fig. 22. Structureless sandstone with upper part reworked into irregular laminations (SS 199 038) .

Fig. 23. 'Slurried' bed showing three part division of graded base, centre with homogenized shale fragments, and thin laminated top with load deformation (SS 201 1 0764) .

Bude Formation subsea fans

119

Fig. 24. 'Slurried' bed with contorted fresh shale fragments in upper part of bed (SS 2014 0775 ) .

Fig. 25. Lower part o f 'slurried' bed showing incorporation o f underlying sediments into the bed as How 'froze' (SS 201 077) .

stone. The basal unit of the channel fill can contain large mudstone clasts, suggesting that the processes that deposited these units were directly connected with the erosion of the channel. In places there is evidence that the deposition of sediment was originally localized at one side of the channel, possibly as some form of point bar. Successive channel cuts can be observed within one complex.

An example of a channel complex occurs in the massive sandstone facies between King's (1967) key shales P and Q, in repeated sections between Northcott Mouth (National Grid Reference SS 202 085) and Crooklets Beach (National Grid Reference SS 202 085) (Figs 33 & 34). Mapeo & Andrews (1991) and Tanner (1992) have provided structural interpretations for these features, however, the

120

R. V. Burne

Fig. 26. 'Slumped' bed about 80 m above the Sandy Mouth Shale with sand volcano (arrowed) on surface. Hammer for scale (SS 2015 1026).

Fig. 27. Sand volcano shown arrowed in Fig. 26. Note central crater and Hank lineations (SS 2015 1026).

evidence for sedimentologically distinct channel fills and channel erosion indicate that the discordances are true channel cuts rather than being the result of early thrusting, although these bed boundaries would have provided planes for preferential movement during deformation. Facies interpretation

Black shale facies

The preservation of well-laminated mudstones of

the black shale facies indicates deposition under tranquil conditions in a basin where sediment was introduced by discrete and gentle currents, at least some of which were bottom-hugging. The rare occurrences of pelagic goniatites and bivalves can be correlated with marine bands in other Westphalian Basins and probably represent brief connections with open marine conditions during eustatic high .. stands. There is no other indication of marine conditions in the succession. The preservation of nektonic fish remains along with their faecal material in well-laminated strata lacking remains of macro··

Bude Formation subsea fans

121

Fig. 28. The Black Rock 'slumped' bed exposed north of Upton Cross. Note load balls foundering from the top of the bed and 1 . 8 m long sandstone clast (arrowed) (SS 200 053).

Fig. 29. Sandstone clast shown arrowed in Fig. 28. Note that the deformation is reminiscent of that found in load balls, and when the sandstone clast is restored to its former geometry an original lenticular cross-section is revealed, which , together with the irregular internal laminations, suggest that the feature may originally have formed as a sand volcano on the top of the 'slumped' bed (SS 2000 0529).

benthos indicates an anaerobic environment (Savrda & Bottjer, 1991), although the restricted ichnofacies dominated by Planolites, suggests dysaerobic con­ ditions (Sageman et a!. , 1991). The oxygen depletion may have been caused by the increased oxygen demand of 'surplus' organic matter (Wetzel, 1991)

brought in by the fine-grained turbidity currents in the form of finely comminuted terrestrial plant matter. Evidence for anaerobic and dysaerobic conditions is supported by the CIS ratios for the shales (Berner & Raiswell, 1983, 1984) reported by Higgs (1991). The bed-top burrows reflect

122

R. V. Burne

Fig. 3 1 . Vertical section through climbing cross-stratified sets in irregularly laminated sandstones. Arrow shows younging direction (SS 2003 0818).

Muddy siltstones facies

Fig. 30. Thick, irregularly bedded sandstones of the sandstone-dominated facies. Figure for scale, section youngs to the right (SS 281 080) .

opportunistic colonization. They do not extend above the top of the sand bed, indicating that the animals responsible died when mud-deposition resumed (Higgs , 1991) , a characteristic of turbidite successions (Seilacher, 1982) . The laminations either result from the settling of dispersed suspended material that was received in discrete pulses, possibly due to seasonal flooding of distant rivers (Burne, 1969b; Melvin, 1976; Higgs, 199 1) , or as high­ concentration sediment clouds (fine-grained tur­ bidity currents) (Burne, 1969b; Melvin, 1976; Glenn & Kelts, 1991). Melvin ( 1986) has pointed out that these laminations could be described as T3_7 and T6_7 mud turbidites using the terminology of Stow & Shanmugam (1980).

The regular lamination of the muddy siltstone facies, absence of body fossils, presence of bottom-feeding traces, and high CIS ratios indicate that this facies was probably formed below storm wave base in an oxygenated and possibly freshwater environment. Higgs ( 1991) concluded that xiphosurid tracks are known only from rocks that appear, on other evi­ dence, to have been deposited in waters no deeper than a few tens of metres. However, Goldring & Sielacher ( 1971) point out that arthropods inhabit practically all marine environments, and that deeper parts of sedimentary basins where silt and clay laminae are deposited from low-velocity currents or distal turbidity flows provide the highest fossilization potential for xiphosurid trackways. It is concluded that this facies was deposited by a series of relatively gentle but persistent bottom-hugging currents (Burne, 1969b; Goldring & Seilacher, 197 1; Melvin, 1986) , the deposits of which Melvin (1986) has compared to T0 and T023 fine-grained turbidites o:f Stow & Shanmugam (1980). Interbedded sandstone-shale facies

All the bed types encountered in the interbedded

Bude Formation subsea fans

Fig. 32. Small scour in sandstone-dominated facies (outlined) with fill of structureless sandstone beds (SS 2014 0771 ) .

sandstone- shale facies originated from short-lived bottom-hugging density currents with a higher content of suspended matter than the surrounding water, i.e. turbidity currents in the sense of Sundborg (1956, p. 60). This conclusion is supported by the high matrix content of all the beds (Fig. 7). The beds are all composed of fine-grained sediment (virtually no sediment particles larger than 1 ¢) that could be carried in turbulent suspension by turbidity currents of reasonable velocity (Lowe, 1982) and would favour the development of autosuspension (Bagnold, 1962 ) . The following detailed interpretations may be made for each of the bed types in this facies. Fine-grained graded beds are interpreted as tur-

123

Fig. 33. 'Thickening upward' succession of black shale facies, interbedded sandstone-shale facies, and sandstone­ dominated facies, the latter forming a channel fill. Structureless sandstone beds have accreted against the channel margin (accretion surfaces outlined) . Cliffs are 60 m high, section youngs to the right (SS 2014 0630) .

bidites in which the structureless graded beds result when rapid deposition of sediment outpaces the ability of a decelerating current to make stable bed­ forms (Burne, 1969b; cf. Melvin, 1986). The result is a graded unit in which grains do not have an oppor­ tunity to be preferentially orientated and are depo­ sited in a reasonably stable packing position. Laminated beds are interpreted as laminated tur­ bidites, equivalent to the Bouma (1962) Toed, Ted or Tct types. The rate of deceleration of the depositing current was low enough for the deposited sediment to construct stable bedforms beneath the waning flow (Burne, 1969b; Melvin, 1986). Graded silty sandstones are interpreted as

A

>-'

ss 20000817

�

::>;:)

:<::::

B

O:l

!::

ss 20020774 SS 201 60n2

D [:. :: : ·.] � �==j r:t//1 • � ..

� 1 m[

Sm

Structureless sandstones

Graded silty-sandstones

Cross-bedded sandstones

Grey muddy-siltstone

& laminated beds

'Slurried' beds & fine-grained graded beds

Black shales

Duplex

Fig_ 34_ Cross-sections of a channel complex in sandstone-dominated facies between key shale P and key shale Q (Tom's Cove Shale) (King, 1967). The complex consists of an earlier channel filled with beds of structureless sandstone, and a later channel filled with cross-stratified irregularly laminated sandstone. (A) Exposure in section 39 of King ( 1967), North of Earthquake Cliff. Note shale clasts at base of channel fill, apparent lateral accretion of structureless sandstones, and final cross­ bedded channel-fill. (B) The same complex 500 m south in section 36 of King ( 1 967) .

3 "'

Bude Formation subsea fans

turbidites showing the almost complete Bouma se­ quence Tabcde (Burne, 1969b; Melvin, 1986) , although the 'd' division is rarely well developed (Bouma, 1962). In the rapid deposition of the 'a' division, grains do not have an opportunity to be preferentially orientated, although they are deposited in a reason­ ably stable packing position. Deposition of the massive units of the structureless sandstones may have taken place from the bedload of a current flowing during the strongly fluctuating hydrodynamic conditions and vigorous burst-sweep cycles of large turbidity currents (Hiscott, 1994). Deposition would be either as successive sheets of sediment (under conditions of plane bed with movement) or as more irregular lenses of sediment (under conditions of antidune flow). Subsequently the escape of pore-water, trapped in the sediment during deposition, and the compaction of the unit may have destroyed this primary structure. Alterna­ tively the structureless parts of these beds may have developed if the tangential stress applied by the depositing current was high enough to maintain the dispersive stress of a bedload, the volume of which was being greatly increased by sediment falling out of suspension. This process of transport causes pro­ nounced orientation in the grain fabric and is the equivalent of the traction carpet of Dzulynski & Sanders ( 1962), but is different from the process of grain flow (Sanders, 1965 ; Stauffer, 1967, p. 502) in which the tangential stress is applied by gravity to sediment resting on a slope. Traction carpets are essentially current-impelled grain flows (Middleton & Hampton, 1976) or cohesionless debris flows (Postma, 1986). As the current waned, the tangential stress applied to the bedload would decrease to a point where the dispersive stress in the traction carpet could no longer be maintained. At this point the traction carpet would abruptly cease to move. The reworked nature of the tops of these beds results from the tail of the turbidity current con­ tinuing to shear the surface of the sediment deposited by the body of the current, and reflects the probable irregularities of that surface. , The grain size distributions of the 'slurried' beds and 'slumped' beds are very similar to that of modern mudflows (Baker, 1965) except that there are no very coarse grains. All the features of the units considered here can be explained by deposition from turbidity currents carrying a very high concentration of fine-grained sediment in suspension. Bagnold ( 1968) has demonstrated that such currents expend little energy and sediment transport might therefore

125

continue infinitely but for the fact that the gradual loss of water from the current reduces the current's volume to a point where the concentration of the suspended sediment suppresses fluid turbulence, and the current would essentially evolve into a cohesive debris flow, or mudflow, without passing through an intervening tractional stage (Postma, 1986; Postma et al. , 1988). This dense current flowing over a soft muddy bottom would tend to peel up the substrate and incorporate it into the flow in the manner described by Dzulynski & Radomski (1966) . The mud would be contorted by the flow, and would tend to disaggregate. However, it would also increase the concentration of transported solids in the flow, thereby encouraging, 'freezing' of the flow. Postma et at. (1988) concluded that a rapid transformation of flow behaviour is necessary for 'soft' intraclasts to be preserved. When the critical concentration of solids is reached the flow would stop abruptly and leave a quick bed. The coarser sediment. of the unit would not be disturbed by this and would preserve any original preferred orientation (Fig. 19). Any sediment deposited on the surface of the unit at this stage would tend to founder into it because of the lack of strength of the bed (Figs 18 & 29) . Excess water would be expelled during the compaction of the bed, forming water-escape structures, pipes of matrix-free sand, and constructing sand volcanoes on the surface of the bed (Fig. 27) . This sequence of post depositional events has been reproduced experimentally by Nichols et al. (1994). The surface of the bed may continue to be sheared by a residual current, with the result that a thin unit of current ripple cross-lamination forms on the upper surface of the bed. Sand volcanoes on the surface of 'slurried' beds may show an asymmetry in the same sense as that indicated on the sole mark of th.e bed as a result of eruption into a current (Burne, 1970) . Sandstone-dominated facies

The comparatively coarse, moderately sorted nature of the beds of the sandstone-dominated facies suggests that they have been deposited from the bedload of a transporting current. The structureless sandstones of this facies have a similar origin to those of the interbedded sandstone shale facies, but are commonly amalgamated to form thick sand­ stones. The irregularly laminated sandstones are of two types. In the majority of beds the irregular lamination cannot be related to any bedform in the lower flow regime, but strongly resemble those

126

R. V. Burne

produced by Middleton (1965, Fig. 3) in an exper­ imental study of deposition from a current forming antidunes. ln other sandstones, better organized cross-stratification can be discerned, which can be related to migrating dunes. In both these sandstone types the structures are persistent throughout the bed and there is little evidence of grading, or waning flow conditions. Burne (1969b) considered that they were deposited by single events characterized by discrete, short-lived, but relatively powerful currents of almost constant velocity.

FACIES SUCCESSION

Description

The long-term evolution of a sedimentary succession at any given place involves temporal transitions between environments, each of which is character­ ized by different depositional systems producing distinct facies. The result is a facies succession. Reading ( 1971) noted the lack of predictability of Bude Formation facies successions when compared with other cyclic Silesian successions, such as the synchronous paralic or deltaic facies succession of the Bideford Group (de Raaf et al. , 1965). The facies succession of the Bude Formation (Figs 35 & 36) shows alternation between end-member con­ ditions of tranquil mudstone sedimentation and the rapid emplacement of channelled massive sandstones with a complete gradation of intermediate environ­ ments. Burne ( 1969b) identified a composite suc­ cession in the Bude Formation consisting of a succession of black shale facies, a thickening upward succession of interbedded sandstone-shale facies, a channelled sandstone dominated facies, and then a thinning upward succession of interbedded sand­ stone-shale facies and silty mudstone facies that passes abruptly up into another black shale facies (Fig. 5). This succession was derived largely from relationships observed in the section between the key shale 0 and the top of the Tom's Cove Shale Member of King (1967) (Figs 35 & 36). However,

much of the Bude Formation does not conform to well-developed thickening-up or thinning-up suc­ cessions (Melvin, 1986). The organization of these parts of the Bude Formation is less predictable, although it is not random (Fig. 4). Interpretation

The case against a storm-influenced shelf

Higgs (1991) concluded that The Bude Formation was deposited largely, if not entirely, above storm wave base, as indicated by the fact that wave­ influenced structures occur throughout the suc­ cession, spaced no more than a few metres apart.' Melvin (1986) thought that the presence of Planolites, xiphosurid trackways and possible hummocky cross­ stratification in the succession meant that the Bude Formation had formed as a subsea fan 'within relatively shallow (shelf as opposed to abyssal) water depths' (p. 26). Higgs (1987) queried how a fan could form at such shallow depths 'where the normal configuration is a wave-cut shelf ' (p. 378). Melvin (1987, p. 382) replied that it is 'hard to imagine how 1300 m could accumulate by the passive aggradation of storm deposits on a lacustrine shelf, wherein facies changes are determined by subtle salinity variations associated with a yo-yo effect of eustatic sea level changes'. This refers to Higgs's (1986b, 1991) interpretation that each alternation of 1-10m-thick sandstone with shales of similar thickness within the Bude Formation resulted from fluctuations in depth and salinity brought about by intermittent over-topping of the sill of a silled basin. According to Higgs (1991), the sandstones were deposited from river-generated storm flows, diverted and modified by wave activity. The river-fed turbidity current and the wave-induced oscillatory current never varied to the point where the wave processes dominated distally, as would be expected in the case of a localized river underflow operating through a tempest (Brenchley, 1985). Higgs (1991) interpreted the ripple cross-lamination developed in the upper parts of these sandstone beds as 'quasi-symmetrical

Fig. 35. (Opposite.) Lateral variation in facies successions between key shale 0 and the top of key shale Q (Torn's Cove Shale). Key shales and section numbers of King ( 1967). The individual beds of the interbedded sandstone-shale facies are shown in a generalized way to demonstrate the variability of this facie. Note the tendency for thickening upward sequences to follow the black shale facies, and for the muddy siltstone facies to occur beneath the black shale facies. Note the position of the channel complex of Fig. 34 in sections 39 and 36. (Locations: Earthquake Cliff, SS 201 079; Upton Cross, SS 200 049; Lower Longbeak, SS 198 032; Widemouth Sand, SS 199 023).

Bude Formation subsea fans

127 Section 5 (Widemouth Sand)

D.. B § 0

Irregular ly laminated or structure less sandstones

Section 9 (South of Lower Long beak)

"Slurried beds" or fine grained graded beds Cross laminated beds

Graded silty sandstones

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Muddy siltstones

D

Black shales

Section 1 8 1 (Upton Cross) 1

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128

R. V. Burne

Fig. 36. A 60-m-high cliff exposure of the facies succession of section 36 (SS 201 077) as logged in Fig. 35 . Section youngs to

the right. Arrows mark the three 'thickening upward' successions from key shale 0 (left), key shale P (centre) and Tom's Cove Shale or key shale Q (right) . Highlighted in the foreground is the channel-cut shown in the upper part of Fig. 34B. This is filled with cross-stratified, irregularly laminated sandstones comparable to those shown in Fig. 31 from the laterally equivalent channel fill shown in the upper part of Fig. 34A.

ripples' intermediate between current ripples and wave ripples. However, the statistics quoted by Higgs ( 1991) in support of a wave-generated origin for these ripples are not definitive (Reineck & Wunderlich, 1968) and they could equally represent non-equilibrium bedforms developed under waning flow conditions (Baas, 1994). Unequivocal oscil­ lation ripples have not been reported from the Bude Formation (Goldring & Seilacher, 1971), neither have 'clean' sandstones characteristic of wave sorting (Fig. 7; Seilacher, 1982). The 'quasi-symmetrical ripples', described by Higgs (1983, 1984, 1987, 1991) as storm-wave influenced structures are found in beds that are quite different from the wave-generated lithotypes described by de Raaf et al. (1977), and I consider that these structures have no depth-related connotation and are more logically interpreted as the product of reworking by the tail of a turbidity current of sediment previously deposited by the

body of the same turbidity current (Middleton & Hampton, 1976). Larue & Provine (1988) illustrate similar irregular cross-laminated sets from turbidites on Barbados, which they attribute to deposition as vacillatory turbidites emplaced by multiple surging flows. Higgs ( 1991) cited the occurrence of multi­ directional tool marks and mud-draped scours as further evidence of wave action in the Bude Formation. Diametrically opposed sole marks on the same surface are a characteristic of some tem­ pestites (Seilacher, 1982), but do not occur in the Bude Formation. The multidirectional sole marks that do occur are similar to those that Middleton & Hampton ( 1976) have attributed to divergence of flow directions within a strongly lobate turbidity current head. Mud-draped scours are not definitive of wave action because Normark & Piper ( 1991) have recorded them from submarine fan-channels.

Bude Formation subsea fans

Rare occurrences of hummocky cross-stratification in the Bude Formation were cited by Higgs (1991) as evidence of storm-wave activity reworking bottom sediments. Einsele & Seilacher (1991) claim that hummocky cross stratification is not found in tur­ bidites, but Prave & Duke ( 1990) have concluded that small-scale hummocky cross-stratification does not indicate a particular flow condition or depo­ sitional environment. They interpret examples similar to those described by Higgs (1983, 1984, 1991) to have been formed by antidunes generated by standing waves at the interface between the depositing body and tail of a turbidity current and an overlying low-density layer. Similar structures in turbidites also have been interpreted as antidune structures by Walker ( 1967), Skipper (1971), Hand et al. ( 1972) and Yagishita ( 1994) . Alternatively, Larue & Provine (1988) have interpreted hummocky cross-stratification found in turbidites as being due to deposition of laminated sand on a surface modified by fluidization and current scour - a situation comparable to that inferred for some of the Bude Formation examples. Mutti (1977, fig. 16) figures thin parallel laminae in a basin-plain-facies turbidite that closely resemble the structures interpreted by Higgs ( 1983, 1984, 1991) as hummocky cross­ stratification. In the absence of associated unequi­ vocal evidence for shelf sedimentation, such as that cited by Monaco (1992), the rare hummocky cross­ stratification recorded from the Bude Formation is interpreted to represent the result of either upper flow-regime structures or traction across a quick bed during deposition from a turbidity current. Brenchley et at. (1993) have presented an in­ terpretation for the Lower Ordovician Bell Island Group of Newfoundland that is remarkably similar to that presented by Higgs (1991) for the Bude Formation, but the facies succession is quite differ­ ent. In contrast to the Bude Formation, the Bell Island Group shows more predictable facies suc­ cessions, contains no sole marks, but wave ripples, hummocky cross-stratification, herring-bone cross­ stratification, gutter casts and bioturbation are all common. The case for turbidite systems

The substantial thickness of the Bude Formation, the relative continuity of single beds, and the absence of a shallow-water benthic fauna all support a turbidite-basin interpretation rather than that of a storm-affected shelf (Einsele & Seilacher, 1991).

129

The lack of predictable vertical associations in the Bude Formation is of itself evidence against a shallow water or paralic origin for the succession, because a 'large proportion of turbidite successions cannot be assigned to any form of regular cyclicity' (Piper & Stow, 1991, p. 371). Hiscott (1981) suggested that cyclical turbidite successions, identified solely on visual criteria, could be explained by chance occur­ rences within unordered successions of turbidite beds. However, Burne ( 1969b) recognized thicken­ ing upward and thinning upward successions, which he related to subsea fan environments. Similar successions have since been described from both ancient and modern turbidite fan environments. For example, Mutti & Ricci-Lucchi ( 1972) and Ricci­ Lucchi ( 1975) described turbidite successions from the northern Appenines of Italy that contain sub­ marine fan successions or 'second-order cycles'. These either followed one another directly or were separated by 2-100 m of monotonous succession interpreted as the deposits of the open fan or basin plain. They contained thickening upward cycles that represented prograding lobes (60% of these cycles were non-channelled), and thinning upward cycles that represented a channel-fill succession (80% of these were channelled). It is concluded that the Bude Formation can also be interpreted as the deposits of a turbidite system or subsea fan (defined without reference to shape). Thus the sandstone- dominated facies may represent the deposits of an accretionary fan-valley system (O'Connell et al. , 1991). The fan-valley is on a scale much greater than that of the outcrop, but within this overall valley 'thalweg' channels occur with a scale observable in outcrop. The facies has the complex organization typical of either the fan­ channel facies (Mutti, 1977), characterized both by thick-bedded channelized units and thinner bedded units, or of the channel-mouth facies of Mutti (1977), in which beds show lensing and wedging, and massive units are overlain abruptly by dune or ripple cross­ stratification. In the silty mudstone facies the evidence of frequent though gentle turbidity currents suggests deposition either in an elevated position within a fan-valley complex or on a channel levee. This interpretation is supported by the evidence of apparent flow instability (Piper & Stow, 1991). Normark & Piper (1991) described levee deposits as containing either discontinuous sands with wavy bedding and climbing ripples or abundant thin sand­ silt laminae. Further support for this interpretation

R . V.

no

is provided by Mutti ( 1977), who found that thin, rippled sandstone units that diverged or expanded in thickness were characteristic of a channel-margin facies. There are three possible environments in which the interbedded sandstone- shale facies could form (Mutti, 1977; Normark & Piper, 1991) . In channel levee and interchannel environments, turbidite deposits will reflect the degree to which sequential channelled flows overtop the levee. They will show variability in both vertical succession and thick­ ness. Deposition on a lobe at the end of a leveed fan-valley system will show less vertical or lateral variability, but preserve evidence of continuity of aggradation. Deposits on an unchannelized open fan will tend to be composed of laterally continuous beds of rather uniform character, although reflecting currents of different sizes. The character of the interbedded sandstone- shale facies is generally not uniform and it is concluded that it represents either the deposits of levee and interchannel environments or of lobes that accumulated at the end of leveed fan-valleys. The black shale facies is composed of fine-grained rhythmic couplets typical of profundal lake deposits (Glenn & Kelts, 199 1 ) and interpreted as the deposits of open fan or basin sedimentation.

CONCLUSIONS

It is concluded that the Bude Formation was depo­ sited in a land-locked synorogenic foreland basin. Deposition occurred on the northern, inactive margin of this basin. Despite the generally unpredict­ able nature of much of the succession, periods of black shale sedimentation alternated with periods of turbidite deposition in which subsea fans were constructed. The conclusion of Melvin ( 1986) that parts of these fans did, on occasion, rise above storm wave base cannot be absolutely excluded, but it seems more likely that the environment was a pro­ fundal lake or inland-sea basin (Glenn & Kelts, 1991) in some ways comparable with the East African Lakes (Cohen, 1990; Scholz & Rosendahl, 1990; Baltzer, 1991; Scott et al., 1991) or Lake Baikal (Hutchinson et al. , 1992). Earlier alternative interpretations of the Bude Formation (Owen, 1950; Prentice, 1962; King, 1965, 1967; Edmonds et al. , 1968; Freshney et al., 1972, 1979) seem to be based, in part, on a lack of appreci­ ation of the extensive variation of structures now

Burne

known to be possible in turbidite systems (Skipper & Middleton, 1975 ; Middleton & Hampton, 1976; Postma, 1986; Postma et al., 1988; Normark & Piper, 1991; Ghibaudo, 1992) . Reading (1987 , p . 8 ) has reminded us that: 'So often a new fashion or model is developed by the subordination of alternative hypotheses or possibilities'. The model developed by Higgs (1991) for a storm-influenced shelf is based on the equi­ vocal evidence of hummocky cross-stratification and 'quasi-symmetrical' ripples (Higgs, 1983, 1984, 1987, 1988, 1991). It is confirmed neither by facies associations nor by stratification characteristics of associated sediments, and, although Melvin ( 1 987) considered that the presence of wave ripples in the Bude Formation is not in dispute, no convincing illustration of them has been published to date. The problems regarding the depositional environment of the Bude Formation all disappear when it is realized that there is no other substantial evi­ dence for shallow-water environments in the Bude Formation. The ichnofacies have no depth conno­ tation (Goldring & Seilacher, 197 1 ; Sageman et al., 1991; Goldring, 1993). Whalley & Lloyd ( 1986) note that the thickness of the Bude Formation is exceptional for the depositional environment pro­ posed by Higgs (1983, 1984), and Higgs ( 199 1 ) himself found it 'curious that such a thick 1300 m shallow water succession is totally devoid of near­ shore or emergent features'. AU the beds of the Bude Formation have features consistent with having been emplaced by turbidity currents and there is no evidence that sediment flowed into the basin by any other mechanism. What evidence is there for the form of this turbidite system? Although some have argued for a prodelta setting (Melvin, 1986) and others for a base of slope ramp (Hartley, 1991) , the deposits are more compatible with a basinal rather than a slope environment. The only constraint on the depth of this basin is the requirement that the basin should have a marginal slope sufficient to ensure either the maintenance of hyperpycnal flow or the 'ignition' (Parker, 1982) of turbidity currents (Normark & Piper, 199 1 ). Although the black shalefacies contains evidence of changes in basin salinity, marine incursion, and degree of bottom oxygenation, none of these appear to have had any direct influence on either the supply of sand to the basin or the gen­ eration of turbidity currents. The supply of sediment to the basin was probably greatest during periods of relatively low basin water level and consequently

Bude Formation subsea fans

very few powerful turbidity currents were initiated during periods of high water level. The compositions of the various turbidites indicate that they all have the same provenance, but there are marked differences in the amount of matrix in the various beds (Fig. 7). A possible explanation for these different depositional characteristics from similarly sourced turbidity currents is to view them in terms of high- and low-density flows. Low-density flows would exhibit a proximal to distal progression from full Bouma sequences Tabcde through to Tde sequences by sedimentation from fully turbulent flows. High-density flows would deposit bedload, perhaps from traction carpets, in more proximal situations, whereas distally concentrated auto­ suspensions would flow until grain concentration increased to a limiting point for turbulence and the flow would be rapidly arrested by transformation into a debris flow. It is possible that the spilling of part of the turbidity current over channel levees would suddenly dissipate energy in the channelled flow, causing bedload deposition, possibly as a result of traction carpet collapse, whereas the overbank part of the flow would carry suspended sediment only and would also dissipate energy until it increased its concentration to that of a high-density flow, eventually 'freezing' to form a 'slurried' or 'slumped' bed. These beds of contorted sediment have con­ ventionally been regarded as having originated either as slumps or debris flows (Freshney et al., 1972; Melvin, 1986; Hartley, 1991) and their interpretation as mudflows or cohesive debris flows , which are the distal deposits of high-concentration turbidity currents, is a novel concept with considerable palaeogeographical significance. Thjck and extensive 'slumped' beds interbedded with black shales may represent large muddy flows triggered by changes in water level in the basin. The thicker, more restricted 'slumped' beds probably represent the distal deposits of channelled flows that deposited massive sands in more proximal environments. These units may be compared with the chaotic silt beds recorded by Nelson et al. (1992) from the outer lobes of the Mississippi fan. The determination of whether a fan model is appropriate for the Bude Formation depends on the recognition and distinction of fan-lobe, fan-channel/ mouth-bar and interchannel deposits. Some parallels can be drawn with the fan successions described by Mutti & Ricci-Lucchi (1972) and Ricci-Lucchi (1975), though these were formed on active tectonic margins. Although the Culm Basin was synorogenic,

131

the Bude Formation was deposited on the inactive margin of the basin (Burne, 1969b; Haslam & Scrivener, 1991). The sandstones are actually richer in quartz (Burne, 1969b; Freshney et al. , 1979; Melvin, 1986) than are the sands of a modern passive­ margin fan, the Mississippi fan (Roberts & Thayer, 1985) (Fig. 7). Shanmugam & Moiola (1988) con­ cluded that interpretations derived for active-margin fans should be applied with caution to passive­ margin fans because of differences in spatial distri­ bution of turbidite facies and their associations (cf. fig. 8 of Normark & Piper, 1991). Shanmugam & Moiola ( 1988) found that characteristics of mature passive-margin taos included: low sand-to-mud ratios in their sediment supply; hyperpycnal flow is import­ ant in fan evolution; flows are uniform and of low velocity; and very large muddy flows may be triggered by sea-level changes. The result is a fan characterized by steady growth of narrow levees, progradation of channel-levee systems, and uniform aggradation of lobes - all features consistent with the facies suc­ cession of the Bude Formation.

ACKNOWLEDGEMENTS

Harold Reading introduced me to the challenges of the Bude Formation, and supervised my doctoral research. A.F. King generously made correlations of the Bude Formation available to me prior to their publication. E.C. Freshney, W. D. Gill, L.R. Moore, J.F.M. de Raaf, E.R. Oxburgh, D. Tappin, E. K. Walton, G. Whitnall, A. Wood and G. Young are thanked for discussions of various aspects of the study. Norma Burne made invaluable contributions to this research. The Chadd family of Bude provided hospitality, assistance and friendship. R.J. Korsch, J.F. Lindsay, R. Norris and T.S. Loutit commented on an early draft of this paper. R. G. Walker provided encouragement over a long period, and, together with E.C. Freshney and A.G. Plint, vigorously reviewed the manuscript. I gratefully acknowledge the support and assistance of the Australian Geo­ logical Survey in many aspects of the preparation of this manuscript. Gail Hill of the Cartographic Services Unit (AGSO) drafted the diagrams.

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Nature, 208, 745 -747 . H . G . (1971) Sedimentation sequences in the Upper Carboniferous of Northwest Europe. Comptes rendues 6e Congress international de Stratigraphie et Geologie Carbonifere, Sheffield, September 11 -14, 1967, Vol. 4 , pp. 140 1 - 1412. READING , H.G. (1987) Fashions and models in sedi­ mentology: a personal perspective. Sedimentology, 34, 3-9. REINECK, H .-E. & WUNDERLICH, F. ( 1968) Zur unter­ scheidung von asymmetrischen Oszillationsrippeln und Striimungsrippeln. Senckenbergina Lethaea, 49, 321 -· 354. Ricci-LUCCHJ, F. (1975) Depositional cycles in two turbidite formations of northern Appenines (Italy) . J. sediment. Petrol. , 45, 3-43 . RoBERTS , H . H . & THAYER, P . A . (1985) Petrology of Mississippi Fan depositional environments. In: Sub­ marine Fans and Related Turbidite Systems (Eds Bouma, A . H . , Normark, W.R. & Barnes, N . E . ) , pp. 331 -339 . Springer-Verlag, Berlin. SAGEMAN , B . B . , WIGNALL, P . B . & KAUFFMAN , E . G . (1991) Biofacies models for oxygen-deficient facies in epi­ continental seas: a tool for paleoenvironmental analysis . In: Cycles and Events in Stratigraphy (Eds Einsele, G . , Rieken, W. & Seilacher, A . ) , pp. 542-564. Springer­ Verlag, Berlin. SANDERS, J . E . (1965) Primary sedimentary structures formed by turbidity currents and related resedimentation mechanisms. In: Primary Sedimentary Structures and their Hydrodynamic Interpretation (Ed. Middleton, G.V.) Spec. Pub!. Soc. econ. Paleontol. Mineral. , Tulsa, 12, 192-219. SAVRDA, C.E. & BoiTJER, D . J . (1991) Oxygen-related bio­ facies in marine strata: an overview and update. In : Modern and Ancient Continental Shelf Anoxia (Eds Tyson, R.V. & Pearson, T . H . ) , Spec. Pub!. geol. Soc. London, 58, 201 -219. Geological Society of London, Bath. ScHOLZ, C.A. & RoSENDAHL, B .R . (1990) Coarse-clastic facies and stratigraphic sequence models from Lakes Malawi and Tanganyika, East Africa. In: Lacustrine Basin Exploration: Case Studies and Modern Analogs (Ed. Katz, B . J . ) , Mem. Am. Assoc. petrol. Geol. , Tulsa, 50, 151- 168. Scorr, D . L . , NG'ANG'A, P . , JoHNSON, T.C. & RosENDAHL, B . R . (1991) High-resolution acoustic character of Lake Malawi (Nyasa) , East Africa and its relationship to sedimentary processes. In: Lacustrine Facies Analysis (Eds Anad6n, P . , Cabrera, Ll. & Kelts, K.), Spec. Publ. int. Assoc. Sediment. , No. 13, 129-145. Blackwell Scientific Publications, Oxford. SEDGEWICK, A. & MURCHISON , R.I. ( 1840) On the physical structure of Devonshire, and on the subdivisions and geological relations of its older stratified deposits. Trans. geol. Soc. London, 5, 633-705. SEILACHER, A. ( 1982) Distinctive features of sandy tempestites. In: Cyclic and Event Stratification (Eds Einsele, G. & Seilacher, A . ) , pp. 333-349. Springer­ Verlag, Berlin. SHANMUGAM, G. & MoiOLA, R.J. ( 1988) Submarine fans: characteristics, models, classification , and reservoir potential. Earth Sci. Rev. , 24, 383-428. SKIPPER, K. (1971) Antidune cross-stratification in a turREADING ,

·

Bude Formation subsea fans bidite sequence, Cloridorme Formation, Gaspe, Quebec. Sedimentology, 17, 5 1 -68. SKIPPER, K. & MIDDLETON , G.V. ( 1975) The sedimentary structures and depositional mechanics of certain Ordovician turbidites, Cloridorme Formation, Gaspe Peninsula, Quebec. Can. J. Earth Sci. , 12, 1934- 1952. STAUFFER, P.H. ( 1967) Grain-flow deposits and their implications, Santa Ynez Mountains, California. J. sediment. Petrol. , 37, 487-508. STOw, D . & SHANMUGAM, G. ( 1 980) Sequence of structures in fine-grained turbidites: comparison of recent deep sea and ancient flysch sediments. Sediment. Geol. , 25, 23-42. S uNDBORG, A. ( 1956) The river Klaralven, a study of fluvial processes. Geogr. Annaler, 38, 127-316. TANNER, P.W.G. ( 1 992) The duplex model: implications from a study of flexural-slip duplexes. In: Thrust Tectonics (Ed. McClay, K . R . ) , pp. 201 - 208. Chapman & Hall, London. THOMAS, J . M . (1988) Basin history of the Culm Trough of southwest England. In: Sedimentation in a synorogenic basin complex; the Upper Carboniferous of Northwest Europe (Eds Besly, B .M. & Kelling, G . ) , pp. 24-37. Blackie, Glasgow. UNRUG, R. ( 1963) Istebna Beds - a fluxoturbidity forma­ tion in the Carpathian Flysch. Ann. Soc. geol. Pol. , 33, 49-92. UNRUG, R. (1965) Le facies de flysch a fluxoturbidites dans les Carpathes Polonaises. Carpatho-Balkan Geological Association, VII Congress, Sofia, September, Reports, Part II, pp. 301 -303. UsSHER, W.A.E. ( 1 892) The British Culm Measures. Proc.

135

Somerset archaeol. nat. Hist. Soc. , 38, 1 1 1 -2 19. R.G. ( 1963) Distinctive types of ripple-drift cross­ lamination. Sedimentology, 2, 173- 188. WALKER, R . G . ( 1 964a) Some aspects of the sedimentology of the Shale Grit and Grindslow Shales (Namurian R 1c , Derbyshire) and the Westward Ho ! and Northam Formations (Westphalian, north Devon) . PhD thesis, University of Oxford . WALKER, R.G. ( 1964b) Some aspects of the sedimentology of the Westward Ho! and Northam Formations. Proc. Ussher Soc. , 1 , 85-87. WALKER, R.G. ( 1967) Upper flow regime bed forms in tur­ bidites of the Hatch Formation, Devonian of New York State. J. sediment. Petrol. , 4, 1052- 1058. WETZEL, A. ( 1991) Stratification in black shales: depo­ sitional models and timing - an overview. In: Cycles and Events in Stratigraphy (Eds Einsele, G . , Rieken, W. & Seilacher, A . ) , pp. 508-523. Springer-Verlag, Berlin. WHALLEY, J . S . & LLOYD, G . E . ( 1986) Tectonics of the Bude Formation , north Cornwall - the recognition of northerly directed decollement. J. Geol. Soc. London, 143, 83-88. WooD, A. & SMITH , A.J. ( 1959) The sedimentation and sedimentary history of the Aberystwyth Grits (Upper Llandoverian) . Q. J. Geol. Soc. London, 114, 163- 195 . YAGISHITA, K. (1994) Antidunes and traction-carpet deposits in deep-water channel sandstones, Cretaceous, British Columbia, Canada. J. sediment. Res. Sect. A: sediment. Petrol. Process. , A64, 34-4 1 . ZIEGLER, P.A. ( 1990) Geological Atlas of Western and Central Europe. Shell Internationale Petroleum Maatschappij B . V . and the Geological Society, London. WALKER,

Spec.

Pubis

int.

Ass.

Sediment. (1995)

22,

137-156

Depositional controls on iron formation associations in Canada P HI LI P F R A L I C K* and T I M O T HY J. B A R RE T Tt *Department of Geology, Lakehead University, Thunder Bay, Ontario P7B 5El, Canada; and tMineral Deposit Research Unit, Department of Geological Sciences, University of British Columbia, Vancouver, British Columbia V6T 1Z4, Canada

ABSTRACT

Understanding Precambrian iron formation depositional processes has been hindered by the lack of precise modern analogues. However, by combining a regional basin analysis of sedimentary and volcanic rocks surrounding an iron formation with detailed examination of sedimentary structures and lithic associations within an iron formation, the depositional setting and physical processes of sedimen­ tation can be inferred. Six iron formations present in the Canadian Shield were examined using this approach. The Palaeoproterozoic Gunflint Formation consists of strand-proximal stromatolites and oolitic shoals, transitional distally to grainstones and parallel-laminated chemical muds. This succession was laid down on a wave- and tide-dominated inner shelf. In the northern Labrador Trough, the Middle Member of the Palaeoproterozoic Baby Formation represents an outer shelf to slope environment. Here massive, graded and ripple-laminated siltstone-shale couplets dominate the succession, with iron-rich chemical sediment forming fine-grained tops to some couplets; thick assemblages of parallel laminated chemical sediments also occur. In the Beardmore-Geraldton area, a submarine environment is represented by oxide facies Archaean iron formations. The chemical units accumulated where clastic mud would normally be found in coarsening and thickening upward levee and ramp assemblages. In the Terrace Bay area, Archaean abyssal plain deposits consist of black, graphitic slate with pyrite layers and zones of chert; these sediments lie on volcanics and are overlain by a submarine-ramp clastic assemblage. Archaean volcanic-associated iron formation is represented by thin interflow sediment packages in a submarine lava-plain south of Beardmore, and as a sulphide lens in a volcanic edifice south of Schreiber. In the first five examples iron formation occupies the niche usually dominated by clastic mud. The iron formation is able to form because of a reduced clastic supply reflecting some combination of the following factors: a peneplained source region; a relative rise in sea-level creating sediment storage capacity in subaerial and shallow-water regions; development of sediment bypass systems; and cessation of active volcanism near the depositional area. The facies of iron formation in a given area is dependent on factors such as water depth (both Eh and energy), degree of basin isolation from clastic sedimentation, and amount of hydrothermal input or upwelling. Proximal to submarine hydrothermal vents, iron formation accumulation can dominate without a reduction in clastic input; the mineralogy and layering are controlled directly by the temperature, Eh and drift direction of hydrothermal discharge.

INTRODUCTION

Sediments accumulating in modern environments commonly provide a basis for comparison with ancient successions, in order to reconstruct the palaeogeography and depositional processes that operated in a region. This approach is probably the most useful technique for reconstructing ancient

depositional systems, but in the Precambrian the present is sometimes not the key to the past. In the Archaean and Palaeoproterozoic, grossly different atmospheric and water chemistry, and possibly tem­ perature, led to the deposition of chemical sediments that have inadequate modern analogues . Iron

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0 137

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P. Fralick and T.J. Barrett

formations, which are chemical sedimentary rocks containing greater than 15% Fe (James, 1966), are volumetrically the most important member of this group. Modern subaqueous hot-spring activity provides clues to understanding the precipitation of Fe-rich sediments, but the mineralogy and areal extent of modern deposits is not comparable to many iron formations deposited during the Earth's early history. Theoretical chemical modelling has, to date, been unable to explain the interlayering of iron-rich sediment and chert. A broader approach is necessary when conaucting field studies of iron formation. For example, sedi­ mentary and volcanic rocks surrounding iron for­ mations provide data on depositional settings; sedimentary structures and lithic associations within iron formations give information on physical process of deposition; and regional basin analysis provides a framework within which to view iron formation genesis. Combining these techniques makes palaeo­ geographical reconstruction and depositional model­ ling feasible. Iron formations are often classified into two main groups, those of Superior type, which were deposited as thick, laterally extensive units in tectonically stable areas, with sedimentary structures and textures indicative of shallow-water environments, and those of Algoma type, which accumulated in tectonically active regions as well-layered but discontinuous units, mainly associated with volcanic rocks or deep­ water sediments (Gross, 1965, 1983). This rather arbitrary subdivision is similar to the geosynclinal theory, initially useful, but not suited to our present data base. It limits iron formation to two types, when in reality there are a multitude of types. Iron formation characteristics are controlled by depo­ sitional settings and processes that are complex multivariate systems and have little likelihood of being repeated exactly in time or space. Thus, iron formations may be similar, but will not be identical, and each is capable of providing further information on how this unusual rock type forms. Classification systems are of course useful in our nomenclature, but they must not become an end in themselves. A fuller understanding of process and depositional setting should be the primary goal of research on iron formations. This requires more detailed work on individual units. Factors involved in the deposition of iron for­ mation include the source of iron and silica, transport mechanisms, and controls on precipitation. Possible sources for the iron include weathering of iron-

bearing minerals, in either subaerial (Garrels et at., 1973; Drever, 1974) or subaqueous (van Hise & Leith, 19 1 1; Huber, 1959) environments, and hydro­ thermal ion exchange ( Goodwin, 1956; Gross, 1965; Fralick, 1987). Simonson ( 1985) has shown that lithological data from the Animikie Group and Labrador Trough do not support a cratonic source for the iron. Chemical data on trace elements, rare earths and stable isotopes do not differentiate between low-temperature weathering of mafic vol­ canic rocks and high-temperature hydrothermal alteration as possible sources for the iron. Iron released by both processes will be in solution in any ocean, and although hydrothermal input is obviously dominant in some situations, low-temperature leaching may have, at times, played an important role. In order for small quantities of Fe2+ to go into, and stay in solution, p02 must have been very low (Holland, 1973). A hydrothermal source has been proposed for the silica (Gross, 1965; Gross & Zajac, 1983; Fralick, 1987), although Cloud ( 1973) believed the silica in Palaeoproterozoic iron formations may have orig­ inated from subaerial weathering. In either case, the probable lack of silica-secreting organisms at this time would have resulted in an ocean saturated with H4Si04 (Siever, 1957; Cloud, 1973; Holland, 1973). The movement of large amounts of water into an environment with a different Eh, pH, or temperature is needed to precipitate iron and silica in the quan­ tities necessary to form an iron formation. This can be achieved through the hydrothermal venting of hot, reduced, acidic solutions (Fralick, 1987; Barrett et a!. , 1988), or upwelling of deep ocean waters with a low p02 on to a more oxygenated shelf (Cloud, 1973, 1983; Holland, 1973; Drever, 1974). However, these scenarios do not explain deep marine Archaean iron formation deposited distally to hydrothermal activity. In this situation, and with the oceanic bottom waters saturated with iron and silica, hydro­ thermal venting may produce a zone of oversatu­ ration and precipitation which extends away from the vent area for tens to possibly hundreds of kilo­ metres (Fralick, 1987). In the Palaeoprotero­ zoic, and possibly the Archaean, microorganisms influenced precipitation of iron and silica as they added 02 to the environment (Cloud, 1973), and possibly directly precipitated iron and silica as coatings and tests (LaBerge et al., 1987), although the latter process has been questioned by Oehler ( 1976). Alternatively, near-surface photo-oxidation of Fe2+ could have formed the iron-rich layers

Depositional controls on iron formation

(Cairns-Smith, 1978; Braterman & Cairns-Smith, 1987), and pressure reduction during upwelling could have caused silica to precipitate ( Holland, 1973). Studies of specific iron formation basins should attempt to integrate a number of the processes described above (cf. Morris, 1993), In this paper we provide six examples from the Canadian Shield which highlight the control that depositional setting had on iron formation character­ istics. The depositional environments described con­ stitute a succession from shallow shelf through slope, rise and abyssal plain to areas of active volcanism. It is apparent from these examples that iron formation type is, like clastic deposits, controlled by depo­ sitional setting. Of course we are not the first to emphasize this fact. Gross ( 1980), and many subsequent researchers (e.g. Ojakangas, 1983; Simonson, 1985), have stated or implied similar ideas. What we strive to emphasize here is the benefit of conducting iron formation research in a manner similar to studies of clastic depositional systems.

CASE STUDIES

Objectives

The rationale for providing this series of case studies is to emphasize the control exerted by depositional processes on iron formation attributes. The physical features that iron formations exhibit are largely the result of their depositional environment. The headings in this section are not meant to be a classification system, but are purely descriptive. We believe it is more productive to look upon each iron formation as a distinct entity reflecting the processes operative in the setting in which it formed. Shallow shelf: Gunflint Formation

Iron formation, associated with other types of chemical and clastic sediment, deposited in shallow marine settings is most common in the Palaeo­ proterozoic, although well-documented examples from the Archaean also exist. The Canadian Shield contains two extensive regions dominated by iron formation associated with shallow marine sediments: the Animikie Group and portions of the Labrador Trough. Iron formation within the Canadian portion of the Animikie Group is contained in the 2000 Ma Gunflint

139

Formation (Fig. 1). This sedimentary package was deposited on a peneplained Archaean surface forming the southwestern margin of Superior Province (the edge of the North American craton at that time). Minor mafic flows and extensive tuf­ faceous horizons within the formation indicate that the region was volcanically active at the time of basin formation and subsidence. The Gunflint For­ mation is underlain locally by a basal conglomerate; the Gunflint has a gradational upper contact with black shales and siltstones of the Rove Formation. The Gunflint Formation is divisible into two mem­ bers and averages 120 m in thickness. The lower member of the Gunflint Formation commonly has basal algal and oolitic cherts (Fig. 2). The microfossil-bearing algal structures often occur on boulders of the basal conglomerate, which pro­ vided a stable substrate. The remainder of the lower member commonly consists of grainstone beds sep­ arated by fine-grained sediment layers (Fig. 3a). Locally, this facies is termed cherty iron formation. Individual fine-grained layers are homogeneous and iron-rich, but mineralogy varies from layer to layer. Iron-oxide-chert and iron-silicate-chert mixtures are the most abundant. Sand-sized, chemical mud intraclasts forming the grainstones are commonly similar in mineralogy to interbedded fine-grained layers. Larger rip-ups of material similar to the underlying substrate are present in many beds (Fig. 3c). Internal laminations are often difficult to discern in the grainstone beds, due to the effects of dia­ genesis, but trough cross-stratification (Fig. 3d) is not uncommon and some sections exhibit hummocky cross-stratification (HCS). Small outcrop size and normal faulting in the area, make analysis of up-section trends within the Gunflint Formation difficult, and limited amounts of drill-core must be used. The section depicted in Fig. 2 exhibits an upward decrease, then increase in the thickness of grainstone beds through the lower member of the formation. Interstratified with the grainstones in some sections are units metres to tens of metres thick consisting of laminated, fine-grained chemical sediment (Fig. 3b), locally called slaty iron for­ mation, and similar in all aspects to the fine-grained material in the grainstone assemblages. The lower member is capped sporadically by breccia. In Fig. 2 this breccia is underlain by a massive layer of silt­ sized material. The upper member of the Gunflint Formation is similar in stratigraphy to the lower member. Basal algal cherts (Fig. 3e), often developed directly on

140

P. Fralick and T.J. Barrett c Ungava Bay

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top of fragments of the breccia, are succeeded by a sequence of thinning upward grainstone beds separ­ ated by chemical sediment. Regionally extensive tuffaceous shales, with accretionary lapilli present at some localities, are interbedded with the grainstone succession. A limestone unit caps the upper member and separates the Gunflint Formation from the shales of the basal Rove Formation. Lougheed ( 1983) attributed the depositional environments represented by the Animikie Group iron formations to supratidal, intertidal and subtidal zones of a broad shelf. The importance of subtidal and possibly lower intertidal sand shoals as depo­ sitional sites for the grainstone was emphasized by Simonson ( 1985). Bidirectional palaeocurrent indi­ cators at some sites indicate that tidal currents were responsible for the accumulation of some grain­ stone units (Ojakangas, 1983). Hummocky cross­ stratification present in other sequences indicates the importance of storm activity in creating and moving sand-sized intraclasts seaward (Fralick,

Fig.

I. (A) Location of the six iron formations discussed. (B) Regional geology of the area north of Lake Superior. (C) Regional geology of the northern Labrador Trough.

1988). Stromatolites and accretionary grains com-· monty appear to be limited to near-strand positions (Morey et at., 199 1). This spatial restriction is not due to browsing metazoans, as in the case of Phanerozoic stromatolites, but probably due to frequent current activity causing large areas of the offshore granular substrate to be unstable. Ojakangas ( 1983) has shown that water depth was the major control on the development of 'cherty' (grainstone dominated) versus 'slaty' (chemical mud dominated) iron formation. The grainstones were deposited at shallower depths by tide-and-wave/ storm driven currents (Fig. 4). The chemical muds accumulated during fair-weather times in shallow-· water regions, and at all times in the deeper offshore. Alternating successions of grainstone-dominated and mud-dominated iron formation packages suggest fluctuations in relative sea-level (White, 1954; Morey, 1983) and can be used to create sea-level curves. Upward-thinning and upward-thickening successions of grainstone beds also suggest, respect-

Depositional controls on iron formation

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ively, relative rise and fall of sea-level (Fig. 2). The bed thickness trends shown in Fig. 2 suggest that sea-level fluctuations controlled the cyclic nature of the lower and upper members. The positioning of the breccia and possible vadose silt unit at the lowest sea-level stand, inferred from bed thickness, suggests that these units formed through subaerial dissolution of carbonates. Increasing volcanic activity in the upper Gunflint Formation may have resulted in a

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shift from chemical sedimentation to a volcaniclastic­ dominated system. The source of the iron- and silica-rich precipitates that form the majority of the Gunflint Formation is still controversial. Simonson ( 1985) favoured a direct hydrothermal source in the offshore with coprecipi­ tation of silica and iron-rich phases in the nearshore area due to ambient high levels of silica and iron in the ocean. Dissolved hydrothermal iron and

142

P. Fralick and T.J. Barrett

Photographs of inner-shelf-associated iron formation in the Animikie Group. (a) Grainstone (intraclast)-dominated succession. The grainstone beds are arranged in lenses which may cross-cut one another. Fine-grained chemical muds (m) are interlayered with sand lenses. (b) Slaty iron formation. This chemical mud-dominated succession also contains grainstone lenses but they are less abundant and smaller than in the cherty iron formation. (c) Thick-bedded cherty iron formation with abundant rip-up clasts (r). (d) Inclined rip-up clasts in a grainstone lens lying on the backset slope of a dune. (e) Rare calcium carbonate stromatolites; most other mounds are silicified. Lens cap is 5.5 em in diameter.

Fig. 3.

Depositional controls on iron formation

143

Grainstone (reworked chemical sediment)

Fine-grained chemical sediment

Block diagram of the depositional environments in which the Gunflint chemical sediments accumulated. Deeper areas not affected by currents induced by storms and tides are dominated by laminae of silica, iron silicate, iron carbonate and iron oxide. Shallower areas affected by sporadic current activity consist of interlayered grainstone and beds of fine­ grained chemical sediment similar to that accumulating in deeper areas. Grainstone production occurred in the shallows due to erosion and abrasion of chemical muds, with the material transported offshore during times of increased current activity. Ooid shoals developed in strand-proximal zones of maximum turbulence, and stromatolites formed along the strand where the substrate had been diagenetically hardened. Karstification occurred in subaerially exposed areas.

Fig_ 4_

silica may have been carried up to shelf depths by upwelling along cratonic margins.

Outer shelf and slope: Middle Member, Baby Formation

Examples of iron formations deposited in outer shelf and slope settings are rare. Similarly, descriptions of shallow-water and deep-sea clastic units are common but outer shelf and slope secessions are not well represented. This may be due to their more restricted distribution and lower preservation potential. The Palaeoproterozoic Baby Formation of the northern Labrador Trough (Fig. 1) provides an example of iron formation deposition in an outer shelf to slope setting. The trough separates the Superior Province from the Rae Province and under­ went a major orogenic episode during closure between these two land masses. Sediments of the Baby Formation were deposited prior to the major phase of orogenic activity, probably during an earlier extensional phase (Wares & Goutier, 1990; Shulski et al., 1993). They are underlain by dolomite and overlain by tholeiitic basalt. The kilometre-thick Baby Formation is dominated by a monotonous, thin-bedded succession of fine­ grained clastic sediment. The Middle Member

departs from this trend with the upward appearance and then domination of the unit by iron formation (Fig. 5). This interval may be correlated with iron formation present in the Knob Group to the south­ west of the Baby Formation. Clastic layers within the Baby Formation are commonly 0.5- 10 mm thick and non-graded with sharp lower and upper contacts (Fig. 6a). They are composed of siltstone, silty shale and occasionally very fine-grained sandstone. Silty shale layers often contain coarser silt streaks only a few grains thick. Graded units similar in other respects to the non­ graded layers also are common. The grading may consist of either an upward grain-size decrease with no internal laminations, or alternating siltstone and silty shale laminae that thin and fine upwards. Current-ripple-laminated layers of fine sandstone are interbedded with the fine-grained clastic suc­ cession. Beds are 1-4 em thick, wavy bedded and non-graded. Current ripples in the thicker units are larger and are commonly stacked. The undulating upper surfaces of the rippled units are sharply over­ lain by silty shale, which fills troughs and covers crests. Thicker, compound fine-grained sandstone beds are also present in the succession. These are up to 20 cm thick and are composed of vertically stacked, 1-5-cm-thick, massive to parallel lami­ nated, non-graded sandstone. The layers are some-

P. Fralick and T.J. Barrett

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Fe-C, Fe-0, Fe-S

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Overburden

ss

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Fe-0, Fe-S, Si SS, Fe-C SS, Fe-C, Si SS, Fe-P

0

SS, Fe-0, Fe-S, Si

Stratigraphical section of a portion of the Middle Member, Baby Formation. All units are internally layered on the millimetre- to centimetre-scale. Chemical sediment occurs in the shale-rich tops of some clastic beds. SS, siltstone and shale; Si , chert; Fe-S, iron silicate; Fe-C, iron carbonate; Fe-0, iron oxide; Fe-P, iron sulphide.

Fig. 5.

times separated by millimetre-scale shale drapes. Rare, thicker beds in the Middle Member consist of two types. (i) Massive, metre-scale medium-grained sandstone beds with coarse sand grains scattered throughout. Reverse grading may be developed near their tops. (ii) Organized beds, with a thin, basal, current-rippled, fine-grained sandstone overlain by millimetre-thick, alternating coarse siltstone and silty shale laminae that thin and fine upwards. The upper halves of the units are composed of normally graded silty shale. Iron formation occurs interbedded with clastics and as dominantly chemical successions within the

Middle Member of the Baby Formation (Figs 5 & 6). Chemical sediment may form discrete millimetre­ scale layers either between graded silt-shale couplets, or interlaminated with the shaley top of the couplet. Iron-rich minerals may be iron silicates, oxides, carbonates or sulphides. One succession of rhythmites containing thicker (average thickness= 4 em) carbonate grainstone layers is also present. This succession is separated from a chemical sedi­ ment assemblage by a massive dolomite, 1.5 m thick. Carbonate layers in the grainstone succession contain small- to medium-scale trough cross-stratification. Chemical-dominated successions are composed of thinly to thickly laminated iron-rich, fine-grained, sediment locally interstratified with chert or jasper layers (or more rarely lenses). Carbonate, silicate, oxide and sulphide facies iron formation are all present within the Baby Formation. Interlayering of the facies is common, although all facies rarely occur in the same interval. The iron-rich layers are centimetre- to decimetre-scale in thickness, with internal millimetre- and submillimetre-scale lami­ nations caused by differences in crystal size (Fig. 6c). Microscopic examination reveals that the laminae reflect changes in the amount of clastic material and chert present. The cherts interbedded with the iron-rich units are commonly 1-3 em thick and exhibit no internal laminations except for rare submillimetre-scale iron silicate layers. The Baby Formation forms part of an assemblage that has been interpreted as a passive margin suc­ cession, thickening towards the east (Dimroth, 1981; LeGallais & Lavoie, 1982; Wardle & Bailey, 1981). The abundant fine-grained sediments, the lack of wave or tidal deposits and the rare presence of carbonate grainstones transported by bottom cur­ rents suggest that deposition took place on the outer shelf or upper slope (Fig. 7). The rhythmites were deposited from low-density turbidity currents similar to those attributed to comparable units in the off­ shore of northeastern North America (Chough & Hesse, 1980; Hesse & Chough, 1980; Stow & Shanmugam, 1980). These sediment clouds may have been raised by sporadic slumping on the slope, or as overflows from channelized bypass systems. The latter mechanism is preferred because slump scars are rare. Thin current-rippled units represent intermittent bottom currents. Rippled units of this type are rarely described from slope deposits. The thick sandstone units provide better evidence for a slope. Their massive nature combined with reverse grading indicates that downslope grainflow processes

Depositional controls on iron formation

145

Photographs of the iron formation associated with outer-shelf-slope facies in the Baby Formation. (a) Close-up of thin-bedded sediments showing lenses and continuous layers of silt (light-coloured layers), interbedded with shales. (b) Magnetite laminae assemblages (m), interlayered with lighter siltstone and shale laminae assemblages (s). (c) Laminated magnetite. Grain-size differences produce the layering. (d) Interbedded siltstone (light, s) and magnetite (dark, m); grey units (g) represent mixtures of magnetite and clastics. (e) Reflected-light photomicrograph of pyrite bands (p) interlayered with clastic material (s). Scale bar is 0.5 mm.

Fig. 6.

146

P. Fralick and T.J. Barrett Grain-flow

Channelized turbidity flows

Block diagram of the depositional environments in which the Baby Formation chemical sediments accumulated. Thinly laminated iron-rich precipitates and chert form discrete units or are interlayered with Bouma d-e turbidites in an outer shelf and slope setting. Clastic material was delivered to this environment through channel overflow, slump­ generated low-density turbidity currents, and grainflows originating at times when the shelf was dominated by quartz sand. Rare iron carbonate sands were brought into this environment when the shelf was starved of clastic material. Fig. 7.

Rainout of chemical sediment

were at least in part responsible for their emplace­ ment. The trough cross-stratified carbonate grain­ stone was obviously moved into the area by fairly strong bottom currents originating on the inner shelf, where grain production would have taken place. This places the Middle Member of the Baby For­ mation on the outer shelf to slope break, past the mud line (Stanley & Wear, 1978) but close enough to areas affected by tide-, or storm-produced currents to receive shelf sediments during unusually high­ velocity flow events. The presence of iron formation in the Middle Member requires the clastic input to be greatly diminished or chemical precipitation rates to be greatly accelerated. Not enough data are present to choose between these two alternatives. Clastic supply may be controlled by: (i) the rate at which clastic materials are supplied to the shelf; and (ii) the storage capacity of the shelf combined with the efficiency of sediment bypass systems in the outer shelf to slope area. The rate of chemical precipitation was probably controlled by upwelling rates from the deep ocean. As these variables fluctuated, the system oscillated between clastic and chemical dominance. The Baby Formation has a great variety of iron­ rich mineral phases, which probably reflects the physiographic setting on the outer shelf. There, upwelling currents would first encounter shallower waters. This may have led to large changes in Eh, pH and dissolved concentrations through time,

resulting in the varied mineralogy of the iron for­ mation. The dominance of rainout processes and the variability of sediment influx and water chemistry produced an iron formation that is both thinly laminated and laterally persistent. Submarine rise: Beardmore-Geraldton Clastic Associated Iron Formation

Archaean iron-formation-bearing successions that were probably deposited on submarine rises are present in the Rainy Lake and Spirit Lake areas (Wood, 1980), Manitou Straits district (Teal & Walker, 1977), Lake St Joseph region (Meyn & Palonen, 1980), the Abitibi greenstone belt (Hyde, 1980), and the Beardmore-Geraldton terrain (Barrett & Fralick, 1985, 1989). The latter will be used as the case example (Fig. 1B). The Beardmore-Geraldton terrane consists of three metasedimentary belts, each resting on a thick volcanic assemblage. A metavolcanic terrain lies to the north and the gneisses and granites of the Quetico Subprovince lie to the south. The northern meta­ sedimentary belt is composed of fluvial conglom­ erates and sandstones (Devaney, 1987), the central belt consists of a prograding shoreline to offshore turbidites (Devaney & Fralick, 1985), and the southern metasedimentary belt contains turbidites and iron formation (Barrett & Fralick, 1985, 1989).

Depositional controls on iron formation

The belts are kilometres wide and tens of kilometres in length, with subvertical bedding younging to the north. The assemblage represents a fore-arc basin that has been stacked tectonically (Barrett & Fralick, 1989; Devaney & Williams, 1989). Oxide facies iron formation occurs interbedded with clastics in the turbiditic portion of the assemblage (Fig. 8). The turbidites are grouped into thick-, medium­ and thin-bedded turbidite-dominated associations, and a thin-bedded iron-formation-clastic-sediment association (Barrett & Fralick, 1989). The bulk of the succession is composed of the first three associ­ ations, which forms a submarine ramp, with attri­ butes suggesting fan development at some locations. The thin bedded, iron-formation-clastic-sediment association forms packages metres to tens of metres thick, interstratified with the ramp/fan turbidites, and is itself divisible into four iron formation lithofacies associations (IFLA) (Figs 8 & 9). Their characteristics are as follows. IFLA a: domi­ nantly magnetite-rich sediment with millimetre- to centimetre-scale, graded or ungraded silt interbeds (Fig. 9a, b, d, f & g). IFLA b: centimetre-scale, graded to sharply bounded silt beds, either con-

tiguous or separated by millimetre-thick laminations of magnetite-rich sediment (Fig. 9a, b, c & e). IFLA c: sand-rich composite units up to about 1 m thick, generally consisting of thin, stacked, ungraded, laminated sand beds. These units consist of medium­ to coarse-grained sand. The composite units are separated by intervals of magnetite or magnetite and siltstone up to 15 em thick (Fig. 9a & b). IFLA d: framework-supported polymictic conglomerate beds up to a few metres thick, interbedded with sandstone and minor iron formation, or fairly thick iron for­ mation and thin-bedded sands (Barrett & Fralick, 1985, 1989). The first three lithofacies associations are commonly, although not exclusively, organized into coarsening upward successions 1-50 m thich (Fig. 8) (Fralick, 1987). Sandstone assemblages are sharply overlain by IFLA a, which is gradational upwards through IFLA b to IFLA c. This package is then overlain by ramp/fan turbidites (Barrett & Fralick, 1989). The ramp-fan system represented by the Beardmore-Geraldton turbidites provided few locations where clastic influx was low enough to allow iron formation accumulation (Fig. 10). To the

B

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IFLAa interlaminated magnetite, siltstone, and slate

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Stratigraphical sections of iron-formation-bearing units in the Beardmore-Geraldton area. (A) Outcrop section measured at the Leitch Mine (detailed location in Barrett and Fralick, 1985). (B) Outcrop section measured at Solomon's Pillars (detailed location in Fralick, 1987).

Fig. 8.

10

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Sandstone IF LAc interbedded sandstone, siltstone, slate and magnetite IFLAb interbedded siltstone, slate and magnetite

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Photographs of iron formation associated with an inner submarine ramp-fan complex in the Beardmore­ Geraldton region. (a) Magnetite layers (dark) interstratified with Bouma d-e turbidites. Clastic and chemical dominated intervals are visible (IFLA a=a; IFLA b=b; IFLA c=c). (b) Interlayered magnetite and d-e turbidites forming two thin coarsening and thickening upward successions (arrows). (c and d) Alternating magnetite-rich (dark) and clastic-rich (light) laminations (prints of thin-sections). Scale bar is 5 mm. (e) Photomicrograph showing gradation from clastic-dominated bottom to magnetite-dominated top of two thin d-e turbidites. Scale bar is 1 mm. (f) Photomicrograph of interlaminated magnetite (dark) and silt (light). Scale bar is 0.5 mm. (g) Photomicrograph of a magnetite-dominated sequence. The lighter layers are mixtures of magnetite and chert. Scale bar is 0.1 mm.

Fig. 9.

south in Quetico Subprovince, where prograding channel-fed lobes merge, iron formation is rare. In the Beardmore-Geraldton area, interchannel ramp development limits iron formation sedimentation. Chemical sediment could form only in interchannel areas where ramp progradation was limited by an insufficient sediment supply. As areas of volcanic

activity sporadically shifted in the arc to the north, sediment influx rates also shifted laterally along the multiple sediment entry points to the basin. This, and probably autocydic processes of channel switching, led to the transformation of clastic­ dominated interchannel areas into iron-oxide­ dominated environments. Either levee outbuilding,

Depositional controls on iron formation

Debris-flow

149 Fine-grained chemical sediment

Block diagram of the depositional environment in which the Beardmore-Geraldton chemical sediments accumulated. Iron oxide successions interbedded with d-e turbidites developed in clastic-sediment-starved interchannel areas not receiving ramp turbidites. These successions commonly coarsen and thicken upwards, probably due either to levee or to ramp progradation into the interchannel sites.

Fig. 10.

caused by channel re-establishment in the area, or ramp progradation, due to renewed high rates of sediment supply, overwhelmed these chemical systems, causing a gradual return to a clastic­ dominated bottom. Conglomerates interbedded with iron formation attest to an upper ramp/fan position for the chemical sediments. Conglomerates are not present in the more distal sediments to the south (Fralick et al., 1992). Abyssal plain: Terrace Bay Clastic Associated Iron Formation

Areas described as graphitic shear zones are common throughout the Canadian Shield. They are usually black, graphitic, pyritiferous slates, metres to tens of metres in thickness, which have been sheared, due to their incompetence. Depositional environments for these units vary, but the majority probably represent abyssal oceanic muds which accumulated in settings distal to active sediment sources. The area to the east of Terrace Bay (Fig. 1B) contains a number of graphitic slate assemblages. They appear to represent fault and fold repetition of the transition from sea-floor volcanics to overlying clastics (Fralick & Barrett, 1991; Eriksson etal., 1994). The Terrace Bay sediments occupy a basin bordered to the northeast and southwest by arc­ related volcanics. The sediments also overlie a volcanic pile kilometres in thickness. Thin, pyri­ tiferous black slates occur between some of the flows near the top of the pile (Fig. 12a) and a fine-grained clastic-chemical sedimentary unit, up to 40 m in thickness, directly overlies the volcanics. This unit consists of black, graphitic slate with intricately

3

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Felsic porphyry intrusive

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lnterlaminated pyritic-carbonaceous slate and pyrite

• �

Massive pyrite Chert

Representative stratigraphical section of a portion of the Kingdom iron formation in the Terrace Bay area. The 40-m-thick iron formation is underlain by altered, intermediate, pillowed volcanics and overlain by turbidites. Finely interlaminated carbonaceous slate, pyrite and chert form most of the iron formation, with occasional pyrite layers up to 20 cm in thickness (detailed location in Schnieders, 1987). Fig. 11.

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Fig. 12.

Photographs of iron formation associated with an abyssal plain assemblage in the Terrace Bay area. (a) Pyrite and chert form a zone (p) between two pillows (v). The pillow selvages have been highly silicified (si). These chemical sediments and volcanics lie directly below the iron formation (Kingdom Occurrence in Schnieders, 1987). (b) Interlayered slate (s), pyrite (p) and chert (c). Deformation has caused some disruption but the fine layering is still visible. (c) Interlayered pyrite (p) and black slate (s). (d) Turbidite succession above graphitic slate-chemical sediment zone.

laminated pyrite layers and zones of chert (Fig. 1 1; Schnieders, 1987; Barrett et a!., 1988). The pyrite occurs as massive layers (Fig. 12c) up to 25 em thick; as thin, millimetre-scale laminations in the mudrock and chert (Fig. 12b & c); as centimetre-scale spheres floating in the mudrock and chert (Fig. 12a & b); and as disseminated pyrite cubes. The cherts form zones, up to tens of centimetres thick, of either pure silica or mixtures of silica, clay and/or pyrite (Fig. 12b). The black slates are composed of quartz, sericite, carbonate, chlorite, pyrite and graphite, with scattered, floating, angular quartz and feldspar grains of possible tuffaceous origin, and rare volcanic shards. Bouma d-e turbidites are occasionally inter­ bedded with the metalliferous succession. A metasandstone-slate assemblage (kilometres in thickness) overlies the succession of metalliferous sediments. Beds are commonly graded from medium-

or coarse-grained sandstone to siltstone or slate and exhibit features typical of turbidites (Fig. 12d). The beds vary in thickness from several centimetres to several metres. Successions of turbidite beds range from a-dominated to d-e-dominated; and thinning­ fining and thickening-coarsening upward success­ ions are present, although uncommon. Deposition in the area began with subaqueous volcanism building a thick extrusive succession. Waning volcanic activity allowed sediment to accumulate between successive flows. With total cessation of volcanic activity, a blanket of fine­ grained sediment was deposited (Fig. 13). Venting hydrothermal fluids produced alteration zones in the underlying volcanics and added Si, Fe and trace amounts of other metals to the fine-grained mud raining down on the bottom. Outbuilding of a sand­ rich submarine fan-ramp complex ended deposition

Depositional controls on iron formation

15 1

Sand lobes

Block diagram of the depositional setting in which the Terrace Bay chemical sediments are interpreted to have accumulated. Intermediate to mafic volcanic sea­ floor was covered by carbonaceous muds interlayered with iron sulphides and chert. Prograding turbiditic fan/ramps built out over the chemical-clastic mixture. Fig. 13.

sediment and clay

Intermediate and mafic volcanics

of the metalliferous muds. The metal-rich deposits are interbedded with the outer-fan/ramp turbidites of this complex (Fralick & Schnieders, 1986; Schnieders, 1987). Submarine lava plain: Beardmore Volcanic Associated Iron Formation

Thick lava-plain successions are common in Superior

Province. Examples from the Abitibi Subprovince comprise major portions of cycles I and II of Dimroth et at. ( 1982), the basal komatiitic and tholeiitic portions of supergroups described by Jensen ( 1985) and, in particular, the Kinojevis Group (Jensen, 1978a, b, 198 1). Sedimentary rocks associated with the lava-plain successions are not abundant and comprise thin, interflow units, Fralick ( 1987) described a series of interflow sedimentary units

A

B

Layered siltstone and shale Massive siltstone

E2J � 1-

Layered pyrite Pyrite and layered chert Pyrrhotite and disrupted chert

f+:+l �

Felsic dike Intermediate volcanics

Fig. 14.

Fralick

\A/\ �

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Mafic intrusive Pillowed and massive flows Volcanic ash Chert lnterlaminated magnetite and chert

Stratigraphical sections representative of volcanic-associated iron formation in (A) the Schreiber area (location in 1989), and (B) the Beardmore-Geraldton area (location in Fralick, 1987).

et al.,

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present in a volcanic pile underlying a turbiditic succession in the Beardmore area (Fig. 1B). Oxide and carbonate iron formation constitute the domi­ nant lithologies present in the interflow sediment (Fig. 14B). Layers of magnetite (Fig. 15c & d) or · side1:ite (Fig. 15a & b) range from submillimetre, to centimetres in thickness, with interbedded chert (Fig. 15a, b, c & d) sometimes attaining thicknesses of metres. Laminations within beds are not common. Chlorite-rich units (with plagioclase and quartz) are often associated with the iron-rich laminae, forming bundles of alternating iron-rich and chlorite-rich layers. Clastic supply to the area was limited to volcani­ clastic ash, which forms the chloritic layers and isolated grains in the iron formation. The chemical sediment layers tend to be purer and thicker than the iron formations previously described. This is probably the result of limited clastic supply together

with proximity of the hydrothermal vent sources (Fralick, 1987; Fig. 16). The thickness of the inter­ flow sediment packages is controlled by four factors. Time duration between flows is obviously important, but the rate of hydrothermal emission, proximity to the active vents and bottom-current patterns are also major controls on unit thickness (Fralick, 1987). Submarine extrusive edifice: Morley occurrence

Submarine volcanoes produce relief of the sea­ floor. The Canadian Shield contains many such suc­ cessions, usually dominated by intermediate and felsic volcanics. Associated iron formations may be similar to those described in the section on lava plains, although they have a tendency to be inter·­ stratified with thicker successions of pyroclastic rocks due to the more explosive nature of the volcanism. Massive sulphide build-ups are also common in this

Photographs of volcanic-associated iron formation in the Beardmore area. (a) lnterlayered chert, c, and iron carbonate, a, at the Empire Mine, south of Beardmore. (b) Close-up of (a). The dark layers are siderite. (c) Interflow sediment consisting of chert, c, and magnetite, m. The thickest magnetite layer is loaded into the underlying chert. (d) Photomicrograph of interlayered magnetite-rich, m, and chert-rich, c, laminae. Scale bar is 0.5 mm.

Fig. 15.

Depositional controls on iron formation

153

Off-axis volcanism

proximal chemical sediment

sediment and clay Fig. 16.

Block diagram representing the depositional setting of iron formation types similar to those near Schreiber and Beardmore-Geraldton. Thick successions of pyrite and chert formed in locations close to vents, probably associated with axial valleys, off-axis volcanism or rifted arcs. Iron oxides, silicates, carbonates and chert were associated with either less i ntense and/or lower temperature portions of discharging hydrothermal zones, or were precipitated distal to major hydrothermal point sources.

setting. Iron-rich massive sulphide deposits are found in clastic- and volcanic-dominated abyssal plain assemblages, but tend to be more prolific in felsic­ dominated piles. The Morley occurrence will be used as an example of this type of iron formation. It is located south of Schreiber in a volcanic pile immediately west of the area discussed for the abyssal plain association (Fig. 1B). The volcanic succession containing the iron for­ mation was deposited about 2. 7 Ga ago, and rep­ resents a large, arc-type edifice (Schnieders, 1987; Fralick & Barrett, 199 1). The Morley deposit is a lenticular chemical sedimentary unit up to 7 m thick which is underlain by intermediate flows and pyro­ clastic rocks, and overlain by thin turbidites and structurally emplaced mafic flows (Fralick et at., 1989). The lower portion of the unit (Fig. 14A, 6 -; 10 m) contains pyrite with interbanded light and dark chert: the upper half consists of bedded to laminated pyrite (Fig. 14A, 10- 1 1.5 m). Within the pyrite-rich upper part of the succession, a variety of bedding structures are developed. The pyrite layers contain delicate internal laminations of pyrite and carbonaceous chert from 0.02 to 1 mm thick. Near the tops of individual pyrite laminae the proportion of chert and disseminated clastic debris is greater. Colloform pyrite domes up to 3 em across and 2 em thick are also present. Millimetre-scale mudstone laminae thin over the small pyrite domes and thicken in flanking depressions. Discordant pyrite growth­ structures on domes, together with inclined pyrite crusts and microslumps, indicate that pyrite accumu­ lation produced an irregular microrelief on the sea-

floor (Fralick et at., 1989; Fralick & Barrett, 199 1). Clastic supply to this area was limited (Fig. 16), although some fine-grained material, of probable volcaniclastic origin, was being delivered. Each lamina in the pyrite beds probably reflects a short­ term hydrothermal injection into a stagnant bottom layer of water. Upward-thinning bundles of laminae may correspond to medium-duration hydrothermal events. The domal structures, and high carbon con­ tent of the sediment provide evidence for relatively deep-water organic mats during chemical precipi­ tation (Fralick et at., 1989).

DISCUSSION

The case studies illustrate that iron formations could form in a wide range of marine environments, given the right conditions. Shallow shelves appear to be 'all or nothing ' environments. Here sediments are distributed on the shelf commonly by non­ channelized flows. Clastics delivered to the near­ shore will be spread over large areas in the offshore, limiting iron formation development. To form iron formations, clastic delivery must be minimal, such that chemical sedimentation has the opportunity to dominate. This leads to the development of Superior Type iron formation with great lateral extent, and sedimentary structures indicative of shallow water. Limited clastic delivery can be achieved if sediment storage sites are available near the strand, thereby preserving a chemical-dominated offshore. An example of this occurs in Gunflint correlatives in

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Michigan, where clastic tidal flats with magnetite layers served as depositional sites for sands, silts and clays, leaving subtidal areas dominated by iron for­ mation and chert (LaBerge et al . , 1992). Other shelves dominated by chemical sediment display features quite different from the Gunflint succession. Carbonates and cherts of the Archaean Steep Rock Group exhibit prolific stromatolite development (Wilks & Nisbet, 1988) and lack grain­ stones. This may be due to quiet-water conditions or other, as yet unexplained, factors. Waves, storm surges and tidal currents operating on shelves, com­ bined with relative sea-level changes, variation in clastic influx , climate and mixing rates with offshore water, all govern the physical attributes of the chemical sediments deposited in this setting. Physical controls on chemical sedimentation on outer shelves and slopes are less well understood. Clastic supply is obviously a major control. Sea-level rise may be a key to limiting clastic supply as it provides more storage capacity on the shelf. It there­ fore becomes important to study the linkage in depositional response between coeval inner shelf, outer shelf and slope systems during periods of sea­ level rise. The development of sediment bypass zones can also produce interchannel areas dominated by chemical sediment. Upper slope interchannel areas are poorly described in the literature. This hinders evaluation of sediment bypass as a control on iron formation development. In submarine fan and ramp environments, devel­ opment of sediment bypass systems has been documented as a major control on formation of areas in which chemical sediment can accumulate. The development of a channel does not always prevent interchannel ramp progradation from flooding off-channel areas with clastic material (Barrett & Fralick, 1989). Changes in location of volcanic activity alter sediment supply rates along the ramp, switching depositional systems from chemical to clastic and visa versa. Bottom conditions become more unsuitable for iron formation accumu­ lation further downslope on the ramp or fan. In this area, flows become non-channelized, spreading out over the bottom. Sediment-starved areas are uncommon and likewise so is iron formation. To deposit iron formation here, major portions of the clastic supply system must be shut off. This may be accomplished through sea-level rise creating more sediment storage capacity in upslope environments or climatic change in the hinterland causing sediment transport systems to dry up.

Portions of oceanic abyssal plains removed from major clastic supply provided the most st:�ble sites for iron formation development. Here Precambrian sedimentary successions are commonly dominated by iron formation and fine-grained clastics. Factors controlling chemical sediment deposition in other areas were of little consequence on the abyssal plains. A mixture of iron formation and fine-grained clastic sediment could continue to be deposited until vol­ canism or outbuilding of a clastic pile buried this lithofacies. Depositional environments in volcanic terrains vary as much as in their clastic counterparts. Although only two examples of volcanic-associated iron formation are discussed here, they do, however, illustrate some general characteristics of chemical sediment deposited in this setting. Layering of chemical sediments in volcanic terrains reflects the dominant effect of discrete hydrothermal venting events. Layer thickness and type are inferred to be controlled by variations of temperature and com­ position of venting fluids (Fralick, 1987 ; Fralick et al., 1989). Interbedded ash layers may reflect magma recharge events with related increases in hydrothermal activity. These types of iron for­ mation should provide the best source of data on the relationship between iron formation bedding attri­ butes and hydrothermal activity. Layered poly­ metallic massive sulphide deposits are likewise an excellent source of this type of information. The above discussion highlights our relative lack of knowledge of physical controls of iron formation accumulation. Further work emphasizing the linkage between various scales of depositional process and iron formation attributes is needed. More work is also needed on chemical controls. Studies of major and trace element chemistry have been of limited value, especially where whole-rock samples rather than individual layers have been analysed. Stable isotope and REE studies have proven more interesting, although more data are needed on mono­ mineralic samples. Experimental modelling of the chemical systems responsible for Fe and Si precipi­ tation appears to be the missing link at present. In particular, we need to study the way in which progressive saturation of stable bottom-water layers by hydrothermal injections controls the nature and sequence of precipitation of iron-rich minerals. Only by integrating sedimentological, geochemical and experimental studies of iron formation can we further unravel the processes that formed these intriguing sediments.

Depositional controls on iron formation ACKNOW LEDGEMENTS

DEVANEY, J .R. & WILLIAMS,

We are particularly grateful to the Thunder Bay staff of the Ministry of Northern Development and Mines (Mines and Mineral Division) for helpful discussions and information on occurrences of iron formation in this area. Field-work in northern Quebec was assisted by Bob Wares. Useful com­ ments and suggestions on an earlier version of the manuscript were provided by Richard Hyde, Guy Plint and an anonymous reviewer. Figures were drafted by Sam Spivak and the word processing was conducted by Wendy Bourke. This research was supported by the Natural Sciences and Engineering Research Council of Canada.

REFERENCES

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GRoss, G.A. & ZAJAC , I.S. (1983) Iron formations in fold belts marginal to the Ungava Craton. In: Iron­ formation: Facts and Problems (Eds Trendell, A.F. & Morris, R.C.), pp. 253-294. Elsevier, New Y ork. HESSE, R. & CHOUGH, S.K. (1980) The northwest Atlantic Mid-Ocean Channel of the Labrador Sea: Il. D eposition of parallel laminated levee-muds from the viscous sub­ layer of low density turbidity currents. Sedimentology , 27, 697-711. HoLLAND, H.D. (1973) The oceans: a possible source of iron in iron-formations. Econ. Geol. , 68, 1169-1172. HuBER, N . K. (1959) Some aspects of the origin of the Ironwood iron formations of Michigan and Wisconsin. Econ. Geol. , 54, 82-118. HYDE, R.S. (1980) Sedimentary facies in the Archean Timiskaming Group and their tectonic implications, Abitibi Greenstone Belt, N ortheastern Ontario, Canada. Precam. Res. , 12, 161-195. JAMES, H.L. (1966) Chemistry of the iron-rich sedimentary rocks. U.S. geol. Surv. Prof. Pap . , 440 J, p. 61. JENSEN , L.S. (1 978a) Geology of Thackeray, Elliott, Tannahill and D okis Townships, District of Cochrane, Ontario. Ont. geol. Surv. Rep . , 165, 71 pp. JENSEN, L.S. (1978b) Geology of Stoughton and Marriott Townships District of Cochrane. Ont. geol. Surv. Rep . , 173, 73 pp. JENSEN , L.S. (1981) A petrogenic model for the A rchean Abitibi

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Ontario .

PhD thesis, University of Saskatchewan, Saskatoon, Saskatchewan, 520 pp. JENSEN, L.S. (1985) Stratigraphy and petrogenesis of Archean metavolcanic sequences, southwestern Abitibi subprovince, Ontario. In: Evolution of A rchean Supra­ crustal Sequences (Eds Ayres, L. D., Thurston, P.C., Card, K. D. & Weber, W.), Geol. Assoc. Can. Spec. Pap. , 28, 65-87. KRONBERG, B.l. & FRALICK, P.W. (1992) Geochemical alteration of felsic Archean rocks by Gunflint Formation­ derived fluids, Quetico-Superior region, northwest Ontario. Can. I. Earth Sci. , 29, 2610-2616. LABERGE, G.L., ROBBINS, E.l. & HAN , T.-M. (1987) A model for the biological precipitation of Precambrian iron-fonnations - a: geological evidence. In: Precambrian Iron-formations (Eds Appel, P.W . U . & LaBerge, G.L.), pp. 69-96. Theophrastus Publications, S.A., Athens. LABERGE, G.L., 0JAKANGAS , R.W. & KICHT, K. (1992) Early Proterozoic and Archean rocks of the Gogebic Range. Institute of Lake Superior Geology, 38th Annual Meeting, Part 2, Fieldtrip Guidebook. Hurley, Wisconsin. LAGALLAJS, C.J. & LAVOIE , S. (1982) Basin evolution of the Lower Proterozoic Kaniapiskau Supergroup, central Labrador miogeocline (trough), Quebec. Bull. Can. petrol. Geol. , 30, 150-166. LouGHEED, M.S. (1983) Origin of Precambrian iron formations in the Lake Superior region. Geol. Soc. Am. Bull. , 94, 325-340. MEYN, H.D. & PALONEN , P.A. (1980) Stratigraphy of an Archean submarine fan. Precam. Res. , 12, 257-285. MOREY, G.B. (1983) Animikie Basin, Lake Superior region, U.S.A. In : Iron-formation: Facts and Problems (Eds Trendell, A.F. & Morris, R.C.) pp. 13-68. Elsevier, New Y ork.

MoREY, G.B., SouTHWICK, D . L. & ScHOlT LE R , S.P. (1991) Manganiferous zones in Early Proterozoic iron­ formation in the Emily district, Cuyuna Range, east­ central Minnesota. Minn. geol. Surv. Rep. Invest. , 391, 42 pp. MoRRIS, R.C. (1993) Genetic modelling for banded iron­ formation of the Hamersley Group, Pilbara Craton, Western Australia. Precam. Res. , 60, 243-286. OEHLER, J .H. (1976) Hydrothermal crystallization of silica gel. Geol. Soc. Am. Bull. , 87, 1143-1152. 0JAKANGAS, R.W. (1983) Tidal deposits in the early Proterozic basins of the Lake Superior region. The Palus and the Pokegama Formations; evidence for subtidal­ shelf deposition of Superior-type banded iron-formation . In : Early Proterozoic Geology of the Lake Superior Region (Ed. Medaris, L. G.) , Geol. Soc. Am. Mem., Boulder, 160, 49-66. ScHNIEDERS, B. R. (1987) The geology of sulfide-facies iron formations and associated rocks in the Lower Steel River-­

Unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 196 pp. SHULSKI, T. , WARES, R.P. & SMITH, A. D. (1993) Early Proterozoic (1.88-1.87 Ga) tholeiitic magmatism in the New Quebec orogen. Can. I. Earth Sci. , 30, 1505-1520. SIEVER, R. (1957) Silica budget in the sedimentary cycle. Am. Mineral. , 42, 821 - 841. SIMONSON , B.M. (1985) Sedimentological constraints on the origins of Precambrian iron formations. Geol. Soc. Am. Bull. , 96, 244-252. STAN LEY, D.J. and WEAR, C.M. (1978) The 'mud-line': an erosion-deposition boundary on the upper continental slope. Mar. Ceo!. , 28, 19-29. STOw, D.A.V. & SHANMUGAM, G. (1980) Sequences and structures in fine-grained turbidites; comparison of recent deep-sea and ancient flysch sediments. Sediment Geol. , 25, 23-42. TEAL, P.R. & WALKER, R.G. (1977) Stratigraphy and sedimentology of the Archean Manitou Group, north­ western Ontario. Geol. Surv. Can. Pap . , 77-1A , 181-184. VAN H 1 s E , C.R. & LEITH, C.K. (1911) The geology of the Lake Superior region. U.S. geol. Surv. Monogr. , 52, 641 pp. WHITE, D.A. (1954) Stratigraphy and structure of the Mesabi range, Minnesota. Minn. geol. Surv. Bull. , 38, 92. WARDLE, R . J . & BAILEY, D . G. (1981) Early Proterozoic sequences in Labrador. In: Proterozoic Basins in Canada (Ed. Campbell, F.H.A.), Geol. Surv. Can. Pap., Ottawa, 81-10, 33 1-358. WARES, R.P. & GouTIER, J . (1990) Deformational style in the foreland of the northern New Quebec Orogen. Geosci. Can. , 17, 244-249. WILKS, M.E. & NISBET, E.G. (1988) Stratigraphy of the Steep Rock Group, northwest Ontario: a maj or Archean unconformity and Archean stromatolites. Can. J. Earth Sci. , 25, 370-391. WooD, J. (1980) Epiclastic sedimentation and stratigraphy in the N orth Spirit Lake and Rainy Lake areas: a comparison. Precam. Res. , 12, 227-255. Little Steel Lake area, Terrace Bay, Ontario.

Spec. Pubis int. Ass. Sediment. ( 1995) 22, 157-193

Facies models in volcanic terrains: time's arrow versus time's cycle G EO F F R E Y J . O RTON Department of Geology, McMaster University, Hamilton, Ontario L8S 4Ml, Canada

ABSTRACT

Vertical and lateral facies variations in volcanic terrains are abrupt owing to the sudden input of additional sediment from one or more volcanic vents. Further complications arise from the influence of volcanic activity on the subsidence history of the basin. A detailed study of one small depositional system from the Ordovician Llewelyn Volcanic Group near Tryfan Fach, North Wales is used to illustrate the problems this causes for developing facies models. The Tryfan Fach Member is made up of four facies associations characterized by their geometries, composition, and sedimentary structures. Each is assigned to a particular depositional setting: marine shelf, braided stream, floodbasin and alluvial fan. The basal portion comprises a comparatively thick , mudstone-dominated succession deposited in a quiet-water marine setting largely below fair-weather wave base. In contrast, coarse-grained, rhyolite-bearing sandstones were deposited in shallow, flood­ prone, southward flowing bedload-dominated braided streams and as unconfined sheet floods. These sandstones amalgamate to form a laterally extensive sheet c. 15m thick and at least 2 km wide, which lies with sharp contact on subjacent marine mudstones. Sandstones pass gradationally upwards into interbedded coarse and fine sandstones, laminated vitric siltstones with accretionary lapilli, and mudstones, deposited on a low-gradient coastal plain. Conglomeratic alluvial fan deposits, derived from older deposits of granite, tuffaceous sediment and rhyolite, occur along the northeast margin of the basin. Although several features can be explained by envisaging the whole succession as the product of one linked depositional system , the differences in sediment composition and palaeocurrent trends raise problems. These cast doubt on the strict application of Walther's Law to the total succession, and demand at least three genetically unrelated depositional systems. The rhyolitic braidplain to floodbasin succession is attributed to subaerial aggradation of primary and reworked pyroclastics, and had a different source from subjacent marine mudstones and the adjacent alluvial fan. The sharp basal contact of the braidplain sandstones is interpreted as an erosional unconformity, and is inferred to have resulted partly from volcano-tectonic uplift in advance of the volcanic eruption. The contemporaneous progradation of an epiclastic alluvial fan from the opposite side of the basin is related to accelerated but more differential basin subsidence during volcanism. Within the syneruptive deposits it could not be established whether facies change reflected sorting processes on the alluvial plain or progressive eruption of finer grained ash. As a consequence, vertical sequences are difficult to interpret, palaeo­ geographical maps are speculative, and the timing of relative sea-level changes could not be assessed fully.

INTRODUCTION

Vertical-sequence analysis, essentially an extension of Walther's Law, is probably the single most important advance in sedimentology in recent decades (Dott, 1983) . It provides a powerful tool with which to interpret changes in sedimentary

environments through time from two-dimensional sequences (de Raaf et al. , 1965 ; Visher, 1965). As Walther stressed, the law can be applied only to successions with gradational , non-erosive contacts between facies and/or environments (see Middleton,

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

157

158

G.J. Orton

1973). Also implicit in its application is the assump­ tion that the control on changes in grain size and/or facies largely lies within the area of deposition. External parameters (e.g. sediment supply, relative sea-level) change gradually rather than abruptly; continuity rather than discontinuity is assumed. Although modern facies models do allow for 'cata­ strophic' events (e.g. storms) , these catastrophes are assumed to recur with the same frequency in the past, present and future . And the thickness of sedimentary deposits resulting from catastrophes is normally not sufficient to alter radically the pre-existing geography and patterns of sediment dispersal. Things are not so simple in volcanic terrains. B y their very nature, explosive volcanic eruptions are catastrophic and rapidly change the supply of sedi­ ment available within a catchment area. The grain size and type (crystals, lithics, glass shards) of sedi­ ment supplied can change with equal rapidity during the course of a volcanic eruption owing to changes in the rate of magma extrusion, volatile content of magma erupted or the amount of interaction with external water in the vent (see Heiken & Wohletz, 199 1 ; White, 199 1 ) . During eruptions, patterns of sediment dispersal are not always controlled by pre­ existing topography; ash-fall deposits in particular are controlled by wind patterns and are distributed over broad areas independent of topography and , to some extent, gravity. Depositional basins around volcanoes are often infilled from multiple sources. The growth of intrabasinal volcanoes can form physiographic barriers that isolate depositional systems and limit the fetch (and hence size) of waves reaching coastlines. Relative sea-level can change quickly owing to uplift or subsidence resulting from a volcanic eruption. As a consequence, ancient volcaniclastic suc­ cessions are notoriously difficult to interpret. Two concepts jockey for recognition in every ancient succession: (i) each bed in the succession represents a unique historical event (i.e. a new eruption, a new type of eruption) with no genetic relation to the bed before; (ii) the succession represents a 'group of facies genetically related to one another and which have some environmental significance' (Collinson, 1969). The first perspective relates facies change to changing source parameters. Sedimentology essentially loses its predictive capability and ability to define environments. The second is an ahistorical perspective in which laws, processes and the pace of change remains predictable . I think these two

concepts, labelled by Gould ( 1987) as time's arrow (a linear succession of unique events) and time's cycle (recurrent patterns in a world that remains essentially unchanged), embody the major issues within analysis of every volcaniclastic succession . A small-scale but well-exposed sequence from the Ordovician Llewelyn Volcanic Group near Tryfan Fach, North Wales is used to illustrate the differences and interaction between these two concepts. At this locality, a 15-m-thick sheet of coarse-grained rhyo­ litic volcaniclastics is bounded by marine mudstones or fine-grained sandstones, and lies a few kilometres away from conglomeratic alluvial fan deposits. The main problem addressed in this paper is whether the rhyolitic sandstones are reworked pyroclastics and indicate penecontemporaneous volcanism, and whether the other stratigraphical units had the same source. Through analysis of sediment composition , it can be shown that there is no genetic relationship between vertically and laterally adjacent strati­ graphical units. Each represents a separate depo­ sitional system. The implications for facies modelling and reconstructing environments in volcanic settings are discussed.

TERMINOLOGY

Sands rich in volcanic debris are of two dominant types according to the method of clast formation . Pyroclastic material (shards, pumice, crystals) is generated by explosive volcanism , and is often contributed directly to the sedimentary record as primary pyroclastic deposits. In contrast, epiclastic volcaniclastic materials result from erosion and weathering of older volcanic rocks, including lithified tuffs, and do not usually reflect contemporaneous volcanicity. A 'grey area' within this classification concerns reworking of unconsolidated pyroclastic debris. Many authors (particularly Fisher, 196 1 ; Fisher & Schmincke, 1984) regard recycled pyro­ clastic material, redeposited by wind or water as secondary pyroclastics; others (e.g. Cas & Wright, 1987) suggest that pyroclastic particles reworked by water or other agents should be called epiclastic. In older and more deformed successions, dis­ tinction becomes even fuzzier, and one is often lucky if one can demonstrate that some particles were produced by explosive volcanism. A further complication is that many volcaniclastic deposits consist of mixtures of both syneruptive pyroclasts and epiclastic material (sensu Fisher & Schmincke ,

Facies models in volcanic terrains

1984) . The recognition of pyroclastic debris where such material is diluted by and mingled with sands of other origins is not easy. Resolving this problem , however, is far from an exercise in semantics but is crucial in determining the nature of the volcanic contribution to sedimentation. Calling all pyroclastic material that has suffered additional transport 'epiclastic' hides the fact that volcanicity may have been active at that time, and breaks the genetic lineage between deposits and their source. In this study the term pyroclastic is used where the composition, texture, variability, and geometry of the deposits all suggest that clasts were produced by a contemporaneous volcanic eruption. Pyroclastic material that was retransported by 'normal' hydro­ logical processes shortly after its initial deposition is referred to as reworked pyroclastics in order to emphasize that sedimentation was still occurring in response to volcanism.

GEOLOGICAL FRAMEWORK

During the latter part of the Ordovician, North Wales was part of an extensional or transtensional marginal basin (Campbell et al. , 1988; Kokelaar, 1988) sited on continental crust comprising accreted volcanic arcs (Thorpe , 1979). The Precambrian crust formed part of a small microcontinent, Eastern Avalonia (Soper & Hutton, 1984) , derived from continental Gondwanaland, and separated from the North American continent (Laurentia) and Baltica by the Iapetus ocean and the Tornquist's sea. Palaeo­ magnetic and faunal reconstructions record the northward movement of Avalonia throughout Early Ordovician to Early Devonian times (Cocks & Fortey, 1982; van der Voo, 1983, 1988) , with temperate southerly latitudes (c. 30-35°S) indicated for mid­ to late Ordovician (Caradoc) times (Torsvik & Trench, 1991). Alluvial fan facies suggest a humid climate (Orton, 199 1 ) . Ordovician sedimentation i n North Wales was dominated by muddy facies. Volcanism and depo­ sition of associated coarse-grained clastic material was confined largely to the Caradoc series, and confined to a NE-SW oriented rift termed the Snowdon graben. An orthogonal array of deep­ seated fractures cutting the ensialic basement (Campbell et al . , 1988; Kokelaar, 1988) controlled the position of eruptive centres within the graben. Reactivation of these and shallower fractures as faults, often in conjunction with volcanism , deter-

159

mined configuration of sub-basins, rates of fault­ block subsidence, the distribution of volcanic and sedimentary units, and depositional settings. The Snowdon graben can be divided into at least two structural basins, 10- 15 km wide, about Llanberis Pass (Fig. 1) based on the distinctive petrography of their contained sediment (Orton, 1990) and the distribution of basalt intrusions (Campbell et a!. , 1988) . The basin northeast of Llanberis Pass has been referred to as the Tryfan depocentre (Orton, 1990) . Volcanism, and the deposition of coarse clastic material, varied in time and space. It can be divided into two eruptive cycles that are reflected in two volcanic groups (Howells et at. , 199 1 ) : a lower Llewelyn Volcanic Group in northeast Snowdonia (Fig. 1) and an upper Snowdown Volcanic Group in southwest Snowdonia. Although this activity was confined to just two chronostratigraphical stages (Soudleyan and Longvillian) , about 5% of total (72 Ma) Ordovician time , its products comprise about half the total thickness of the Ordovician sequence. The average rate of sediment accumu­ lation during the Soudleyan and Longvillian approached 1 m lOOO yr-1 (Orton, 199 1 ) . The earliest volcanic activity in the Llewelyn Volcanic Group developed from at least four, partly contemporaneous , centres. Rhyolite lavas and silicic ash flow tuffs (Conwy Rhyolite, Braich tu Du Formation) , trachyandesite lavas and tuffs (Foel Fras Formation) and basaltic-andesite lavas (Foel Grach Basalts) were all erupted (Fig. 1). Extrusive rocks are associated mainly with marine mudstones and show little evidence of reworking; they were probably ponded in subsiding, fault-bounded areas of the sea-floor (Howells et al. , 199 1 ) . However, the latest activity in the Llewelyn Volcanic Group, the Capel Curig Volcanic Formation, was dominated by larger scale subaerial eruptions of silicic magma, giving rise to widespread ash flow tuff deposits. Volcaniclastic successions underlying the Capel Curig Volcanic Formation, herein referred to as the Tryfan Formation (Figs 1 & 2), reflect a compara­ tively reduced amount of volcanic activity. The Tryfan Formation can be divided into at least eight members (Fig. 2), based on differences in sedi­ mentary facies and/or environments, and the amount, petrography, chemical composition and location of any volcanicity associated with sedimen­ tation. Each member represents a distinct depo­ sitional episode with no genetic relation to the episode before. The Gwern GofTuff, and underlying fluvial- deltaic deposits were derived from a source

160

G.J. Orton

N

t

� CapeiCt.rig �"J Volcanic Formation (;:::;::::::1 Tryfan Formation

r:Mi:!;(!Jil

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Foel Fras Volcaric CC>I"\l)lex

Foel Grach Basalt Fm. � &aichTuDu � Volcanic Formation � Coowy Rhyoite Formation

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Stbvolcaric ntrusions Important faults

Fig. 1. Llewelyn Volcanic Group: distribution of formations and related subvolcanic intrusions. Box locates Fig. 2.

Intrusions in the south identified as follows: MP, Mynydd Perfedd; BC, Bwlch y Cywion; T, Talgau; CL, Carnedd Llewelyn (modified from Howells et al. , 1991). On inset map, shaded area denotes approximate position of Snowdon graben, MS refers to the Menai Straits.

to the east-southeast (Orton, 1988) . Progradation and/or transgression of coarse-grained fluvio-deltaic systems was the dominant mode of basin infill above the Gwern Gof Tuff. Each of these prograded toward the south and southeast, indicating a 'flip' in the polarity of infill of the Tryfan depocentre. The depo­ sitional package of concern here lies directly above the Gwern Gof Tuff and is referred to as the Tryfan

Fach Member (Figs 2 & 3 ) . It varies in thickness from 30 to 120 m , and has been recognized only within the Tryfan anticline (Figs 1 & 2) , over an extremely small area ( < 20 km2) . Exact information on the time-scale spanned by the Tryfan Fach Member cannot be provided. However, it was of the order of a few hundred thousand years and certainly not millions of years.

161

Facies models in volcanic terrains 68

I

I

*

fossillocalities

500

m

66

Fig. 2. Simplified geology of the Tryfan anticline (modified after British Geological Survey, 1985, unpublished 1:10000

sheets; Orton, 1990). Numbers on border refer to UK Ordinance Survey National Grid (grid square SH) and are 1 km apart. Circled letters on map refer to location of the main sedimentary sections, whereas numbers on stratigraphical log (and associated textures) denote the three basinal facies associations: 1, marine shelf mudstones; 2, Rhyolitic braided stream sandstones; 3, floodbasin sandstones .

FACIES ASSOCIATIONS

The stratigraphical record of the Tryfan Fach Member can be divided into four facies associations, each assigned to a particular depositional setting. As cleavage obscures sedimentary facies in fine­ grained lithologies (unless rhyolitic) , environmental interpretations are based largely on data collected from sedimentary rocks of medium silt grade and coarser. Marine shelf mudstones

The basal portion of the Tryfan Fach Member in the south consists of poorly exposed mudstone with sharp-based massive or massive to horizontally laminated beds (to 1 0 cm) of coarse siltstone and fine-grained sandstone. Upper surfaces of sandstone beds are either sharp or grade into overlying mud-

stone. One bed of coarse-grained lithic sandstone containing rounded rhyolite pebbles (Fig. 4, 49 m) also occurs. Rare brachiopods (Rostricellula) indicate a marine setting (Tunnicliff & Rushton, 1980). Thicker horizons (to 2 m thick) of rhyolitic fallout tuff (Fig. 3) consist of thin, 8 -1 0 cm thick, massive to graded beds of vitric siltstone. The succession generally becomes sandier upwards with bioturbated sandy siltstone giving way to inter­ stratified very fine sandstone (beds 10- 100 em) and mudstone (beds to 20 cm) . Most sandstone beds are sharp based, massive, sometimes normally graded, and rarely infill isolated-scours 30 em deep. Beds near the top of the succession , however, contain ripple cross-stratification, small-scale (to 15 em) trough cross-stratification , planar horizontal lami­ nation , and more rarely undulating lamination. Vv'here trough cross-stratification occurs, sets are often separated by discontinuous siltstone lenses.

162

G.J. Orton

Facies associations

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100

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Thinly bedded floodbasin sandstones

"

Rhyolitic braided stream sandstones

60

en u::

40

Marine shelf m udstones

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GWERN

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Fig. 3. Composite log through the Tryfan Fach Member (from the vicinity of log C, Fig. 2) showing stratigraphical relationships between facies associations. Ruled vertical bars denote position of more detailed sedimentary logs. The Gwern Gof Tuff (base of section) is about 40 m thick and only its top portion is depicted. The top of the braided stream association forms the popular rock-climbing slab at Tryfan Fach.

163

Facies models in volcanic terrains Facies interpretation 65

5cm

Small scale trough and ripple Cross stratification

? Shoreface

Wave swash on sand bar?

Low-angle planar lamination

------ -- Fairweather wave base-??

Bioturbated siltstone

55

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Alternation of rapid deposition by storms and waning alluvial floods and slow suspension sedimentation

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45

1m

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Fig. 4. Section from mid- to upper portion of the marine shelf mudstone association. Scale is in metres above Gwern Gof Tuff. From section B, east of Tryfan Fach, SH 67205995. Log is continued in Fig. 5. See Fig. 3 for stratigraphical position of log.

164

G.J. Orton

The top few metres of the association are again fine-grained (Fig. 5 ) . Mudstones are slightly tuf­ faceous and a laterally continuous comparatively pure unit of vitric siltstone (50- 100 em thick) lies directly beneath the coarse-grained rhyolitic volcaniclastic material of the braided stream associ­ ation (Figs 5 & 6) . Mudstones contain thin ( < 30 cm) tabular sandstone beds, most of which display bio­ turbation. Although the sandstone beds do not pinch and swell they could not be traced between sections only a few hundred metres apart. Indistinct parallel lamination and more rarely trough cross­ stratification were the only sedimentary structures observed. Overall, facies and fossils indicate deposition in a quiet-water muddy environment. Environmental interpretation of Caradoc fauna is facilitated by comparison within coeval strata in the nearby (< 100 km) Berwyn Hills (Pickerill & Brenchley, 1979). Here Rostricellula is associated with the Macrocoelia subcommunity of a Dinorthis com­ munity and is considered to have colonized high­ energy non-turbid, well-oxygenated environments on silty substrates under moderately high rates of sedimentation. A maximum water depth of 25 m was assigned to the Macrocoelia subcommunity based on sedimentary structures, and the presence of the fossil-boring Vermiforichnus (Pickerill , 1976) . The bottom part of the Tryfan Fach Member was dominated by deposition of mud from suspension. Coarser grained sediment was introduced period­ ically in decelerating flows. These could represent suspension clouds deposited during the waning stages of a storm (Nelson, 1982) or underflows (hyper­ pycnal effluent) from river-derived flood events. Although underflows are documented most com­ monly where river waters enter freshwater lakes, they also occur when bedload-dominated streams enter marine basins. Wright et al. (1986) confirmed that hyperpycnal underflows carrying silt occur off the mouth of the Huanghe (Yellow) River. Under­ flows carrying sand develop when floodflows from the San Lorenzo River (California) scour the river mouth and pass offshore as a plane jet (Hicks & Inman, 1987) and are suspected also to be a common phenomena on fan-deltas along fjords (Prior et al. , 1987 ) . I f storm o r river-introduced silts and sands d o not exhibit evidence of reworking by waves, it is usually inferred that beds reflect deposition below effective wave base. A comparable facies described from other Ordovician successions in North Wales was

assigned to an 'outer shelf setting' (Fritz & Howells , 199 1 ) . Care must be taken, however. As noted by Pickerill & Hurst ( 1983), if suspension clouds produced by storms or river floods contain , in addition to sand and silt-grade material, sufficient mud to provide cohesiveness, modification by later currents could be precluded even though deposition occurred above normal, even 'fair-weather' , wave base. The upper sandier part of the succession contains unequivocal evidence for deposition in shallower water. Planar lamination (Fig. 4, c. 61 m) may reflect wave swash processes. However, a similar facies , termed quasi-planar lamination, described from Lower Cretaceous lower shoreface to shelf deposits of Montana has been attributed to deposition under single-event, high-energy combined-flow conditions (Arnott, 1993 ) . Outcrop quality is not adequate to discern whether lamination remains perfectly planar or gently undulates, although the position of the lamination at the base of an upward-shoaling unit lends some credence to a combined-flow origin . Undulatory lamination, considered to form under intense oscillatory flow (Allen, 198 1 ) or combined flow (Myrow & Southard, 1991) occurs nearby at about the same stratigraphical level and provides additional evidence for wave processes. The twice observed, sandstone-filled scours are similar in size and orientation (but not abundance!) to gutter casts described by Myrow ( 1992) from shallow subtidal deposits in Newfoundland. In his model the subtidal zone , dominated by fine-grained sediment, is largely a zone of sediment bypass in which high-velocity sediment-laden flows erode shore-normal scours preserved as gutter casts. The trough cross-stratification of the sandiest deposits (Fig. 4, 63-66 m) is thought to be produced by waves and/or river mouth floods. The absence of facies such as lenticular and flaser bedding, reactivation surfaces, rhythmic bedding and mud drapes suggest that tidal processes were absent or unimportant. The small size of the cross-stratification, inter-bedding of mudstone, and absence of swaley or hummocky cross-stratification indicates low near­ shore wave power. This requires further comment. Wave power at a coastline depends on the maximum deep-water wave energy and its shallow-water fric­ tional attenuation, which is a function of the sub­ aqueous slope. Younger shorefaces of the Tryfan Formation commonly contain abundant hummocky and swaley cross-stratification, undulating lami­ nation, and large wave ripples (e.g. Orton, 1988;

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166

G.J. Orton

Fig. 6. Contact between (A) fine­

grained bioturbated sandstones of the marine shelf association (beneath hammer) and (B) trough cross-stratified coarse-grained volcaniclastic sandstones of the braided stream association (above hammer), which often contain a thin (30-50cm) bed of siliceous mudstone. General fluvial palaeocurrent is to the left and towards the reader. Hammer is 30cm long.

Fritz, 199 1 ) . As it seems unlikely that deep-water wave energy changed dramatically over a few hundred thousand years, the low wave power documented throughout the Tryfan Fach Member is related to a low subaqueous slope. At the time of emplacement of the Gwern Gof Tuf (Fig. 3) , the offshore slope must have dipped to the west­ northwest. In contrast, the rhyolitic braidplain deposits of the Tryfan Fach Member indicate a slope to the east-southeast. Sometime during the inter­ vening time period, that is when the marine shelf association was being deposited, the slope must have rotated, either tectonically or by deposition, through the horizontal. Low-gradient muddy shelves are highly efficient at attenuating wave energy (cf. Wells & Coleman, 198 1 ). On the Surinam coast­ line the total wave-energy loss through slow shoaling across the inner shelf ranges from 9 3 to 9 6% , and most waves do not reach the shoreline nor break. Rhyolitic braided-stream sandstones

A multistorey, 12- 15 m thick , laterally extensive (about 2 km) horizon of poorly sorted granule conglomerate to fine sandstone directly overlies the muddy coastal sediments (Figs 3, 5, 6 & 7). Tuffaceous siltstone rip-up clasts (to 5 em) and sub­ angular rhyolitic pebbles (to 2 cm) are common . The sandstones were deposited as broad tabular sheets 20-250 cm thick separated by thin ( < 20 cm) beds of vitric siltstone or interbedded siltstone and fine

sandstone. Fine-grained intervals are laterally continuous for distances of about 200 m, with the exception of a tuffaceous, green-coloured siltstone that extends for at least 900 m (Fig. 5). Sandstone sheets are characterized by extremely abrupt vertical changes in grain size and texture . Vertical-sided , laterally stepped erosional scours up to 30 em deep sometimes occur along their base. In thicker sheets, shallow trough cross-stratification is the dominant sedimentary structure , with rare climbing ripple cross-lamination and horizontal lamination. Palaeo­ current distribution is unimodal towards the southeast. Many thick sheets fine upwards , from conglomeratic coarse sandstone to siltstone (Fig. 8). When this occurs the set size of cross-stratification decreases upwards ( c. 20 cm to Scm) in conjunction with grain-size changes. Thinner sandstone beds display a wider range of sedimentary structures. Coarse-grained beds are commonly normally graded , fining up from conglomeratic sandstone to siltstone over short (< 50 em) vertical distances. Finer grained beds are massive, have massive bases with cross­ stratified tops, horizontal lamination, or undulating lamination. The geometry of the above sandstone packages , their internal structures, coarse texture and uni­ modal palaeocurrents indicate deposition by uni­ directional tractional currents under a broad range of flow strengths. A subaerial setting is assumed because: (i) there is no evidence (e.g. wave ripples) within fine-grained beds for marine processes ,

Facies models in volcanic terrains

167 Facies interpretation

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braided stream facies association. From Section D, east slope of Tryfan, SH 6683 5945. See Figs 3 & 5 for stratigraphical position of log. Numbers to right of log denote thickness of cross-strata. In thin beds that fine upwards the set size indicated refers only to sets at the base of the beds.

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(ii) there is no bioturbation as occurs within other facies associations, (iii) sandstones are extremely poorly sorted, (iv) tidal indicators (e.g. mudstone drapes) are absent.

Si

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The thicker sandstones are best interpreted as infill of fluvial , probably braided, channels (Miall, 1977). In relation to modern braided rivers, these sandstones resemble to some extent the South

168

G.J. Orton

Fig. 8. Channel-fill sequence from a 2-m-deep pebbly

braided stream (braided stream association). Base of sequence is just beneath the field of view. Height of cross­ stratification decreases upwards, from 20cm to 4cm. Flow of fluvial current was towards the reader.

Saskatchewan River (Cant & Walker, 1978) in being dominated by trough cross-stratification. The South Saskatchewan maintains average channel depths of 3m (maximum 5 m) , in which trough cross-stratified sands accumulate by migration of sinuous crested dunes. Fining, upward sequences are generated as the channel shallows due to lateral migration. As this occurs, the trough sets are overlain and/or interbedded with stacked sets of planar cross­ stratified sand formed by cross-channel migration of shallow bar complexes. Unlike the South Saskatchewan, the Tryfan Fach Member contains no planar cross-stratification. Although puzzling this is not unique and numerous other ancient braided stream successions, such as the Jurassic Westwater Canyon Formation

(Campbell, 1976; Godin, 199 1 ) , or Pennsylvanian South Bar (Rust & Gibling, 1990a) and Boss Point (Browne & Plint, 1994) Formations are similarly dominated by trough cross-stratification. The absence of planar cross-stratification has been explained in at least two ways. Some suggest that lateral migration of deeper channels across a braided sand-flat would lead to the preferrential preservation of three-dimensional dunes at the expense of cross­ channel bars characterizing shallower channel por­ tions (Campbell, 1976). Alternatively, channel avulsion, causing rapid abandonment of the pre­ viously active tract, combined with flood-stage erosion of top-stratum would prevent development of fining upward waning channel fill sequences and result in a sheet-like sandbody dominated by trough cross-strata (Rust & Gibling, 1990a; Browne & Plint, 1994) . Several important observations suggest a differ­ ent interpretation for the trough cross-stratified 'channel' fill units of the Tryfan Fach Member. 1 There is little evidence, such as channel forms, deep scours or channel base lags, that bedforms were confined within deep channels. Instead, fluvial cycles rest on virtually planar erosion surfaces or wedge out at their base against broad, shallow channels. 2 Although sediments are poorly sorted, cosets often fine upwards with a corresponding decrease in set­ size (Figs 7 & 8) . In contrast , no clear grain size trends in the facies cycles are present within the Boss Point Formation and co-sets are inferred to represent partial-channel fills (Browne & Plint, 1994). Although trough cross-sets in the Boss Point range from 50 em to 120 em in thickness, vertical variation in set-size is not reported. 3 The set size is small enough that sandstone sheets could represent complete 'channel' fills. A broad relationship between bedform height (H) and flow depth (h) in rivers has been established (e.g. Allen, 1984, fig. 8-20). Preserved set thickness will be considerably less than the height of dunes originally present at flood stage because of relaxation during falling stage (Allen, 1 974) and truncation by sub­ sequently migrating scours. Consequently, Rust & Gibling (1990a) suggest adding 50% as a conservative estimate of the amount lost. This approach gives a flood-stage dune height of about 30 em for the largest dunes at the base of cosets and a few centimetres for bedforms at their tops. The mean relationship (H!h 0. 15) of Allen ( 1984) yields a flood-stage flow depth of about 2 m for the large basal dunes, a =

Facies models in volcanic terrains

depth comparable to the thickness of the cosets. The broad, extremely low-angle, and shallow nature of the trough cross-stratification (Fig. 8) indicates either extremely rapid flows and/or shallow flow depths (Jopling, 1965 ) . There are similarities with proglacial sandur deposits (e.g. Ruegg, 197 1 } . For all o f the above reasons, there seems iittle doubt that most of the fining upward trough cross­ stratified sandstone sheets record virtually uncon­ fined flood events. The thickness and type of sequence resulting from such flood events would depend on the depth and duration of the flood, and the amount and grain size of sediment available. Three possibilities are considered . The thickest sheets require that a range of grain sizes are available and that flow-depths change in a comparatively gradual fashion so that bedforms have time to respond to the decreasing flow strengths. These cosets represent the most prolonged (and most widespread} discharges, and may have involved some channel incision or migration. Thinner sand­ stone sheets where trough cross-stratification grades rapidly into ripple cross-stratification would develop where floods waned extremely rapidly, whereas thin incomplete (top missing) successions could result if flood waters simply ran out of sediment to deposit. Although longitudinal sections are rarely pre­ sented, some of the cross-strata (termed low-angle indistinct stratification on logs) appears similar to scour-fill bedding described from deposition of pebbly sand sheets under high-discharge , shallow flow conditions (cf. Smith , 1986, 1987, 1988) . Sand­ stone sheets with this type of cross-strata are extremely poorly sorted, rarely fine upwards, and set size is typically less than 10 em (cf. Fig. 7, top of log). These features suggest that an abundant sedi­ ment supply 'forced' rapid aggradation of low relief sinuous-crested dunes on the river bed. The thin sandstone beds without cross-stratifi­ cation contain sedimentary structures indicating high-velocity shallow-water sheetflood conditions. The massive to stratified facies, for instance, resembles hyperconcentrated flood flow deposits described from Holocene and ancient alluvial sequences of both volcanic (Smith, 1986; Maizels, 1989) and non-volcanic (Simpson & Eriksson, 1989} settings. Beds containing horizontal lamination would represent upper flow regime plane-bed con­ ditions. Following Rust & Gibling (199Gb) the beds with undulating lamination can be interpreted as vertically aggrading antidunes that may have formed in 10 -15 em of water flowing at 1 m s- 1 . These 4

169

deposits could occur on the floodplain or within channel belts at times of falling stage when water depths were locally very shallow. Facies of modern , shallow, sand-bed braided rivers differ considerably according to discharge and dis­ charge variability, depth of flood waters, and grain­ size of sediment transported . Large perennial , comparatively sluggish rivers with a finer sediment load are usually characterized by migration of large lobate linguoid bars , which deposit superimposed suites of planar-tabular cross-bedding. Examples include the Platte River of Nebraska (Crowley, 198 3) or outwash fans in Iceland (Boothroyd & Nummedal, 1978) . In contrast, with smaller, ephemeral, flood-prone streams with considerable short-term variability in discharge there is either insufficient time or insufficient depths of water for large linguoid bars to develop. This can be again illustrated by outwash fans in Iceland. Catastrophic floods caused by subglacial volcanic eruptions (jokulhlaups) have discharges many orders of magnitude larger than the normal summer discharges (Maizels, 1989) . Sandy deposits that result are structureless, horizontally laminated, or trough cross-stratified and lack stacked sets of planar cross­ stratified sand formed during normal discharges. Sinuous-crested megaripples (St) were also the most common bedform in channelized stream flood in central Australia (Williams, 197 1 ) . I n summary, I conclude that Tryfan Fach Member 'rivers' varied from a few centimetres to perhaps 2 m in depth, and were virtually unconfined. Sedi­ mentation proceeded through vertical aggradation beneath shallow ephemeral sheet-floods rather than by incision, lateral migration, and avulsion of more perennial channels. The term 'sheet braide d' (Cotter, 1978) is an apt description of the type of 'river' envisaged for the Tryfan Fach Member. This fluvial style is not common at the present day, except in arid areas where ephemeral runoff forms a network of shallow, interlacing poorly defined channels. Ancient examples that are dominated by trough cross-stratification include the Devonian Peel Sound Formation (Miall & Gibling, 1978) and perhaps distal successions of the Miocene Ellensburg Formation (Smith, 1988). Thinly bedded floodbasin sandstones

A finer grained interval of interbedded sandstones and mudstones gradationally overlies the cross­ stratified rhyolitic sandstones sheet (Fig. 9). Beds

170

G.J. Orton

Facies interpretation ��Rhizocorallium

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are sheet-like on outcrop scale with even thin ( 10 em) beds traceable for 100 m. Sandstone units vary sub­ stantially in composition and thickness. Very coarse sandstones (beds 20-50 cm thick) are either massive (Fig. lOA) or contain small-scale (lO cm) trough cross-stratification, which yields palaeocurrents toward the southeast. Coarse-grained lithologies sometimes load into underlying mudstones (Fig. lOA). Climbing ripple cross-lamination and horizon­ tal lamination are the domjnant lithofacies within the finer grained and thinner ( 1 0- 15 em) sandstone

Debris flow Reworashfalkedl Quisuspensi et-watoenr sedimentation Minorchannel fluviasl Sheetshalfllooodsw watinteor Fluvial channel Amalsheetgamatfloodsed

Fig. 9. Representative section from

floodbasin facies association. From section C, just south of wire fence on east slopes of Tryfan, SH 6696 5960. See Fig. 3 for stratigraphical position of log and Fig. 2 for location of section.

beds. The intervening siliceous siltstones display horizontal or undulating lamination and small wave ripples (Fig. lOA). Lamination in fine-grained beds often contains water-escape structures or is con­ voluted. Subhorizontal burrows in the top of the succession (Fig. lOB ) (Cruziana ichnofacies: Rhizocorallium) indicate a low-energy depositional environment, but cannot be used to infer water depths or salinity because Rhizocorallium has been described from coastal lagoon and freshwater set­ tings (Fiirsich & Mayr, 1 98 1 ) .

Facies models in volcanic terrains

171

Fig. 10. Facies o f floodbasin facies

association. Photograph is located 11m above base of Fig. 9. (A) Climbing? wave ripples, undulose and horizontal lainination within siliceous siltstones. Deposits lie about 85 em stratigraphically above position where photograph B was taken. Massive beds (M) may be bioturbated. Coarse-grained crystal-lithic sandstones at the top of the photograph represent the base of a thin (50em) debris flow deposit. Lens cap is 6 em in diameter. (B) Horizontally laminated tine-grained vitroclastic (i.e. predominantly devitritied glass shards) tine-grained sandstones and siltstones. Notice lack of grading or organization to sedimentary units. Letters denote: R, subhorizontal Rhizocorallium burrows; W, vertically-aggrading wave ripples. Coin is 3.5 em in diameter.

Facies indicate a quiet subaqueous environment. They are termed floodbasin deposits because of their fine grain-size , because there is little evidence (e.g. desiccation cracks, palaeosols) for prolonged subaerial exposure, and because they appear to have accumulated simultaneously across a basin several kilometres wide. Rare occurrences of accretionary lapilli within siltstones provide evidence for some penecontemporaneous rhyolitic volcanicity. The evenly laminated siltstones indicate episodic sedi­ mentation out of suspension. Bioturbation and wave rippling (Fig. lOB) within siltstone units indicate that they represent an amalgamation of several smaller events. These quiet-water conditions were period-

ically interrupted by input of coarser sediment. Fine-grained horizontally laminated or rippled sand­ stones, and coarse-grained massive sandstones probably represent subaqueous high concentration sheet flows (cf. Dam & Andreasan, 1990). The thin coarse-grained sandstone beds with trough cross­ stratification are not unlike beds from the previously described braided-stream association, which were attributed to waning alluvial floods and are inferred to represent a similar phenomena. Whether the floodwaters developed on an alluvial plain cr whether beds represent the subaqueous extension of the flood events is a moot point. Plausible settings for the floodbasin association

172

G.J. Orton

include: (i) an interdistributary bay, (ii) lagoons along the delta front behind barrier islands, (iii) lakes on a floodplain adjacent to an active alluvial tract and (iv) a transgressed (or flooded) low-gradient alluvial plain or braidplain. The first environmental setting is dscounted because there is no evidence (deeply incised channels, levees) for the distributary channels. If the fine-grained succession represented lagoon deposits onlapping a braidplain during a marine transgression then one would expect: (i) some evidence for landward-direct currents (e.g. from storm wash­ overs), and (ii) the lagoon deposits to be overlain by sandstones representing the protective barrier islands/ bars. Yet all palaeocurrents are seaward directed (Fig. 5) and the floodbasin association is overlain by a black laterally extensive mudstone interpreted as offshore marine (Orton, 1990) . A lacustrine origin, the third alternative, is more difficult to fully disprove. Lakes can develop on poorly drained areas of an alluvial floodplain and receive sediment from nearby river tracts. Quiet­ water suspension deposition of mud and silt occurs during normal river discharges , with sheets of coarser sediment introduced during large river floods, when crevasse splays form. The grain size of sediment within crevasse splays reflects that carried by the parental river tract, and can include coarse sands (e.g. Pollard et al. , 1982). Palaeocurrent directions from crevasse-splay deposits are usually at a high angle to the associated river deposits, and crevasse­ splay sandstones often display coarsening or thickening upward cycles according to position with respect to the active river tract (e.g. Galloway, 198 1 ) . If the floodbasin association is considered in relation to the underlying rhyolitic fluvial succession , the above lacustrine alternative seems improbable. The most important observations are: (i) the contact between the fluvial and floodbasin succession is gradational (over about 2 - 3 m), (ii) the lateral extent (perpendicular to palaeoslope) for the braided stream and floodbasin associations is the same, (iii) palaeocurrent directions from the floodbasin association are consistent with those obtained from the underlying fluvial association, and (iv) little sys­ tematic vertical change in facies or grain size could be discerned within the floodbasin sections. I therefore conclude that the floodbasin associ­ ation represents a submerged low-gradient alluvial plain. During most of the time, quiet-water sus­ pension sedimentation of silt and mud occurred. However, when flash floods and sediment gravity

flows figured prominently on a nearby alluvial plain, the alluvial input was denser than basinal waters (hyperpycnal flow) allowing flood waters to pass unchecked into the ?marine realm. Mixing between floodwaters and ambient waters appears to have been extremely limited (cf. Dam & Andreasan, 1990) and the subaqueous flood deposits are internally similar to subaerial sand and gravel sheets deposited earlier (braided stream association) on an alluvial plain. The association has many features in common with distal floodbasin/lacustrine deposits described from a rapidly subsiding Devonian strike-slip basin (Hornelen Basin, Norway) (Steel & Aasheim, 1978) . Similarities include the thickness of depositional units, occurrence of horizontally laminated, rippled and trough cross-stratified sandstones inter­ stratified with mudstones, convoluted lamination, and absence of obvious intermediate (2-20 m) or larger scale cyclicity. • Conglomeratic alluvial fan sandstones

Conglomeratic alluvial fan deposits characterize the northeast margin of the Tryfan depocentre and are well exposed between Craig yr Ysfa and Carnedd Llewellyn (Fig. 1 ) . The alluvial succession that is correlated with the Tryfan Fach Member (Fig. 1 1 ) overlies bioturbated siltstones t o very fine-grained rippled and horizontally laminated sandstones thought to reflect deposition in a shallow quiet­ water environment (Howells & Leveridge , 1980). Alluvial-fan facies rest erosively on these bioturbated siltstones, and a scour at (, :ast 50 m wide and more than 1 .5 m deep has been identified (Fig. 1 1 , 2 1 .5 m). The lowermost alluvial sediments (Fig. 1 1 , 2 1 . 5 3 6 m ) consist o f either (i) poorly sorted medium­ grained to conglomeratic sandstone with trough cross-stratification or (ii) scours infilled by normally graded , conglomeratic sandstone to fine sandstone. Palaeocurrent directions are all to the west in keeping with those obtained from younger fan successions at Craig yr Ysfa (Orton, 199 1 , fig. 4) . The trough cross-stratified sandstones suggest deposition by braided streams. Compared with the cross-stratified sandstones of the braided stream facies association these sandstones (i) have larger foresets (sets to 70 cm) ; (ii) less commonly display upwards changes in facies, grain size, or set size; and (iii) contain planar cross-stratification (Sp ). These observations imply that rivers or floodwaters were comparatively deep, that channel avulsion and/or flood-stage erosion

173

Facies models in volcanic terrains

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Distances are metres above the base of the Tryfan Formation . Note occurrence of rhyolitic ash-fall tuffs in fine-grained deposits beneath the coarse-grained alluvium .

174

G.J. Orton

of topstratum prevented the formation of fining upward cycles , and that in-channel bars (hence Sp) occasionally developed (see earlier discussions). The normally graded conglomeratic sandstones resemble hyperconcentrated flow deposits (e.g. Smith, 1986; Maizels, 1989). Deposits in the Tryfan Fach Member are either internally structureless or contain poorly developed trough cross-stratification within the finer grained upper portions. The latter relationship indicates an initial rapid fallout of suspended load and deposition by direct suspension sedimentation, followed transitionally by traction sedimentation within the more dilute low-density upper portions (see also Lowe, 1982). The inter­ calation of such debris flows within finer grained stream flow deposits indicates rapid variation in fluvial discharge conditions on the alluvial fan. The overlying alluvial fan sediments (Fig. 1 1 , 3651 m) are dominated by unstratified conglomeratic sandstone occurring in tabular beds 20 -50 cm thick. They commonly contain subvertical rhyolite pebbles (4 em size) or rip-up clasts of siliceous siltstone (30 em long). The sharp non-erosive bases, occurrence of well-defined units, lack of sorting or grading and nature of fabric are all in accord with deposition by cohesive matrix-supported debris flows. The overall succession can be interpreted in terms of increasing proximity to sediment supply. Although deposits could represent a valley-fill, with axial braided streams and laterally derived debris flows, there is no evidence for this physiography in the sedimentary record. Instead, a simpler model is favoured involving progradation of a small alluvial fan westward into the Tryfan depocentre . The main­ tenance of deep channels in a distal fan setting is not common because channel bifurcation usually results in flow expansion, progressive shallowing, and eventually unchannelized flow. It implies both an abundant supply of water and confinement of flow by channel banks. The uppermost alluvial sediments of the Tryfan Fach Member (Fig. 1 1 , 5 1 -60 m) are better sorted and finer grained. Trough cross-stratified sandstones yield palaeocurrents in a variety of directions. Thin beds of fine-grained sandstone display planar horizontal to low-angle cross-cutting lamination and contain concentrations of magnetitic heavy minerals. Together these facies suggest reworking of the fan surface , perhaps by waves, following its abandonment.

FACIES ORGANIZATION AND DISTRIBUTION

Facies relations in the Tryfan anticline

The first three facies associations are best displayed on the northwest limb of the Tryfan anticline, where four sections, spaced about 400 m apart have been measured (Figs 2, 5 & 12) . These can be correlated with a section (F) on the opposite limb of the anti­ cline, about 3 km to the southeast (after unfolding) . Along the northwest limb of the anticline , the marine shelf association thickens northeastward from 30 to 70 m. Although vitric siltstone beds are thicker and have less admixed epiclastic material in the southwest , sandstones are coarsest and most prevalent in central sections. On the southeast limb (section F) , the total succession is muddier. How­ ever, a 1 m thick, medium-coarse-grained sandstone containing moderate-scale trough cross-stratification (25-cm sets) lies 1 8 m beneath the top of the marine mudstone. Although reliable palaeocurrents could not be obtained, a fluvial origin seems unlikely given its stratigraphical position. As there is no evidence for tidal processes during deposition of the Tryfan Formation , this lends further credence to the idea that the marine shelf association represent a low­ gradient platform on which wave energy decreased shoreward. The relative thicknesses of braided stream and floodbasin components (Fig. 9) indicate that the most persistent axis of coarse fluvial deposition was along the southwest side of the braidplain. Fluvial discharge and attendant sediment/water ratios also appear to have been more variable along this side. For instance, the thickness of cross-stratified sand­ stone sheets implies that most 'channels' were quite shallow; thin debris flows and sheetfloods were an important part of aggradation on the alluvial plain. In contrast, to the northeast, successions indicate infill of several deeper (3 -4 m) 'channels' and sheet­ flood deposits are absent. Despite good exposure (> 70% outcrop) more detailed stratigraphical cor­ relation was not possible. To the southeast (section F) , the cross-stratified braidplain sandstone sheet can be correlated with an erosively based , 2.5 m thick, fining upwards suc­ cession of trough cross-stratified medium- to fine­ grained rhyolitic sandstone (Figs 12 & 13). Fine siltstone occurs as thin discontinuous lenses between individual sets, indicating numerous slack-water periods during aggradation. The cross-stratified

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G.J. Orton

176

20

Interbeddedandbiositurltsbtatoneed sandstone

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15

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Fossiliferous sandstone Mudstone

Offshore Fig. 13. Section through upper

M Si Sand sandstones overlie a 2-m-thick sharp-based suc­ cession of massive (?bioturbated) and horizontally bedded sandstones, and are overlain by interbedded fine -coarse-grained sandstones, siliceous siltstones, and interlaminated sandstones and siltstones. Overall the succession mimics the facies suc­ cessions seen on the northwest side of the anticline and interpretation of section F (Fig. 13) is made

portion of Tryfan Fach Member, southeast limb of Tryfan anticline. See Fig. 12 for location of section .

against the background of these sections . Two interpretations warrant further comment. The cross-stratified rhyolitic sandstones are similarly interpreted as representing distal fluvial currents . If the cross-sets were wave produced , their size suggests a high nearshore wave energy, yet wave power was low enough that mudstone could be deposited during interludes between input of coarser

Facies models in volcanic terrains

sediment. This does not negate the possibility, however, that some or all of the cross-stratified succession may have been deposited below sea-level. The second important difference is that the cross­ stratified sandstones lie above fine-grained sand­ stones, and rest erosively on them. On both sides of the anticline, the floodbasin association gradationally overlies the fluvial deposits, indicating a reduction in the amount of coarse sedi­ ment deposited on the braidplain. Successive sand­ stone beds, however, do not become finer grained ur-section . They merely become thinner or less common (i.e. more widely separated in vertical section). Bounding surface between the braided stream and marine shelf associations

One fundamental question concerns the position of the marine shelf association with respect to a shore­ line. At least two interpretations are possible, largely because of uncertainty concerning the amount and source of the sand along the coastline. Neither model is entirely satisfactory. The first alternative treats the sandy portion of the marine shelf association as part of a shoreface. The mudstone-dominated interval that caps the as­ sociation may represent the shoreline itself. Walker and Harms ( 1975) and Hamblin ( 1992) have de­ scribed a tideless muddy coastal deposit with some similarities. However, in the marine shelf association no indicators (e.g. change in colour or reddening of mudstones, desiccation cracks, palaeosols, beach/ foreshore deposits) for subaerial or near subaerial conditions were observed, and even the uppermost deposits appear to be fully marine. The second alternative considers the sandy succession to be a small and offshore component of an even muddier coastline. On such coastlines bed thicknesses and grain sizes first increase than decrease away from the shoreline. This pattern arises because of in situ resuspension of ambient muddy sediment with fines transported landward combined with offshore transport of coarser sediment during storms (cf. Hill & Nadeau, 1989; Myrow, 1992) . In this scenario , the uppermost mudstone-dominated interval of the marine shelf association would correspond to both muddy shoreline and nearshore subtidal deposits. Yet it contains no evidence (flaser and wavy bedding, current ripples, gutter casts) for a nearshore setting. In both alternatives the transition between the two associations appears too abrupt, with expected

177

facies successions either missing or thinner than expected. Even if the braided streams prograded into a low-wave-energy basin one would expect more widespread shoreface or delta-front deposits between the alluvial fan/braidplain and interfan/ offshore mudstones (e.g. Vos & Eriksson, 1977; Fernandez et al. , 1988; Bardaji et al. , 1990). Consequently, the abrupt juxtaposition of 'fluvial' and 'marine' along the northwest limb of the Tryfan anticline, as well as the scoured base to the more distal ?fluvial sandstones suggests a relative sea-level fall during sedimentation, with possible removal of upper shoreface/nearshore deposits by erosion. 'Uplift' appears to be greatest in the northwest, where braidplain deposits rest with sharp contact on marine mudstones. The succession to the southeast (Fig. 13) suggests more continuous aggradation, although the condensed nature and sharp base to probable shoreface sandstones indicates that relative sea-level was still falling ( cf. Plint, 1988). A similar abrupt alternation between braid­ plain and shallow water environments has been described from other strike-slip basins, including the Pennslyvanian Boss Point Formation (Browne & Plint, 1994) and English Westphalian B strata (Haszeldine & Anderton, 1980) . In the Boss Point, braidplain packages rest erosively on, and locally cut deep ( 15 m or more) scours into underlying lacustrine mudstones. A reason for the sharp contact was not given , although uplift due to episodic tectonic source area rejuvenation seems likely. Extent of the deposit

It is pertinent here to dispel the notion that the deposits described represent a small or distant component of a much larger volcaniclastic apron. In the first place, deposition of the Tryfan Fach Member occurred only to the northeast of Llanberis Pass (Fig. 1 ) . The Snowdon region was uplifted and subaerial throughout deposition of the Tryfan Formation (Orton , 1991 ) and contemporaneous marine mudstones further to the southwest belong to a separate depocentre. Although contempor­ aneous mudstone-dominated successions occur 10 km to the southeast, distal 'feather-edge' sand­ stones from the Tryfan Fach depositional system have not been recognized. In light of the rate at which the braided stream sandstones pinch out across the Tryfan anticline, correlation is not suspected. Less than 5 km to the northwest , adjacent to the Nant Peris fault (Fig. 1 ) , all Ordovician successions

178

G.J. Orton

are stratigraphically condensed (Trythall et a!. , 1987 ; Orton, 1990). The entire Tryfan Formation is about 50 m thick, compared with more than 700 m in the Tryfan anticline. It consists of mudstone or fine­ grained sandstone with no sandstones being similar in composition to the rhyolitic sandstones of the Tryfan Fach Member. These must correlate with a zone of non-deposition and/or an unconformity near Nant Peris. Ordovician successions also occur an additional 30 km to the northwest, on Anglesey (Fig. 1 , inset) . Although substantial Precambrian strike-slip move­ ment occurred along the narrow stretch of water (Menai Straits) that separates Anglesey from the Welsh mainland (Gibbons, 1983) , lithostratigra­ phical correlation of Cambrian rocks across the Menai Straits indicates little or no movement of Anglesey relative to the Welsh mainland since the Early Cambrian (Reedman et a!. , 1984). Ordovician successions on Anglesey consist of deep-marine muds and/or mass flows deposited in smaLl fault­ bounded basins (Bates, 1972). These include volcanogenic massive sulphide deposits hosted in rhyolitic volcanics and pillow basalts of probable Caradocian age (Pointon & lxer, 1980; Westhead, 1991) . Thus, although there might be some Cardocian volcanism on Anglesey, there was certainly no upstanding subaerial volcanic arc that could have sourced the Tryfan Fach Member. For the above reasons, it is concluded that the depositional system responsible for the Tryfan Fach Member affected areas within only a few kilometres of Tryfan Fach itself, and that the deposits of the Tryfan anticline represents a substantial portion of the area traversed by the system.

EV IDENCE FOR VOLCANISM

Characteristic petrographic features indicative of explosive volcanism include glass shards, pumice and scoria, bipyramidal �-quartz, and euhedral zoned feldspars. To those not familiar with volcani­ clastic deposits, it may seem an easy task to establish when volcanism occurred. However, volcanic clasts differ widely in size, shape, density and durability, with mean settling velocities varying by several orders of magnitude. During reworking, hydraulic processes can quickly remove most of the less dense or smaller ash and pumice fraction. Volcaniclastic deposits are also more susceptible to diagenesis than any other type of sand because of the chemi-

cal instability and reactiVIty of their framework grains. Glass suffers particularly, with chalcedony, opal, zeolite and clay minerals being alteration by-products (e.g. Davies et al. , 1979; Surdam &: Boles, 1979) . The end result of sorting and alteration processes in distal exposures is usually a crystal- lithic sand-­ stone, with framework grains set in a fine-grained, variably altered and/or recrystallized, non-descript matrix (e.g. Runkel, 1990; Cather & Folk , 1991) . Thus many features, including the geometry of the deposit, its sedimentary facies, the variability of its composition, shape and sorting parameters , and the extent that sediment composition differs from the 'norm' for sandstones within the basin, are needed to establish penecontemporaneous pyro·· clastic volcanism. Petrography

Alluvial fan deposits

Alluvial fan deposits contain significant mono­ crystalline quartz (24% ) , little feldspar (8%) and lithic fragments dominated by aphanitic or sparsely porphyritic quartz-bearing rhyolite (20%) , siltstone (26% ) , or granitic (3% ) lithic fragments (Table 1) . Composition indicates derivation from a mixed granite , tuffaceous sediment, and rhyolite volcanic source. Although volcanics still comprise about half of the lithic population, lithics are thought to be derived from older rhyolites of the Llewellyn Volcanic Group. Braided stream and floodbasin deposits

Sandstones contain large (to 4 mm) euhedral sani·· dine (25%) and smaller plagioclase crystals (17% ) . The percentage of monocrystalline quartz (5% ) is low compared with the 'norm' for the Tryfan Formation (Table 1) . Lithic fragments (to 1 em) are extremely angular and dominated by rhyolitic volcanics. These are mostly porphyritic two-feldspar felsic volcanics (32%) with phenocrysts of albite­ oligoclase and simply twinned alkali feldspar set in a recrystallized quartzo-feldpathic matrix that some·· times displays a trachytic texture . Quartz pheno·· crysts are rare and deeply embayed when they do occur. Other lithic clasts include aplite, ?rhyolitic tuff, or perlitically fractured glass (8% ) . Clasts are set in a quartzo-feldspathic matrix ( c. 20% of totall rock) . The texture of the matrix and lithic fragments

179

Facies models in volcanic terrains Table 1.

Composition of sandstones Tryfan Fach Member Tryfan Fach braidplain (n = 7) Mean

Monocrystalline quartz Polycrystalline quartz Chert K-feldspar Plagioclase Rhyolite Rhyolitic tuff and perlite Aplite Mafic volcanics Siltstone Granite Opaques Other heavy minerals

4.8 2.6 1.6 25.3 17.5 31.9 8.5 1.1 2.4 1.9 0.2 1.3 0.4 0.43 0. 94 0. 14

PF ratio LvL ratio QF ratio

SD

Mean

SD

2.9 3.9 1.7 15.2 9.8 14.5 8.2 1.3 3.7 2.5 0.4 2.0 1.0

23.4 10.6 11.2 5.3 2.3 10.7 Absent Absent 3.8 26.0 2.7 0.5 0.2

9.8 5.2 6.4 3.2 1.2 6.0

0. 19 0. 09 0. 1 1

9 : 43 : 47 5 : 43 : 52 5 : 89 : 8 6 : 83 : 1 1

QFR QmFLt QpLvLs LvbLvrLsp

Craig yr Ysfa alluvial fan ( n = 5)

0.33 0.51 3.3 46 : 8 : 46 24: 8 : 68 17 : 44: 39 6 : 36 : 58

2.5 12.5 3.6 0.6 0.3 0.12 0. 13 0.7

Tryfan Formation (n = 101) Mean 32.6 6. 1 7.8 14.0 6.5 7.2 1.8 1.2 1.7 14.8 0. 9 2.9 0.4 0.36 0.58 2.4

SD 15.6 5.4 7.2 1 1.4 6.5 12.3 5.8 1.5 3.9 14.8 1.6 5.8 0. 9 0.25 0.30 1.9

41 : 21 : 38 34: 21 : 45 16 : 51 : 33 4: 46 : 50

PF, plagioclase/total feldspar ratio ; LvL, proportion of volcanic lithics within lithic population; QF ratio, quartz/total feldspar ratio. Q, total quartzose grains; F, total feldspar grains; R, unstable lithic fragments; Qm, monocrystalline quartz; Lt, total lithics R + polycrystalline quartz; Qp, polycrystalline quartz; Lv, volcanic lithic fragments; Ls, sedimentary lithic fragments; Lvb, mafic volcanics; Lvr, rhyolitic volcanics; Lsp, sedimentary and plutonic lithics. =

is similar and it is often difficult to tell the two apart. Although geochemical data are not available, the crystal population (both free and within lithics) suggests derivation from a rhyodacite to rhyolite magma. The abrupt appearance of a monomict sheet consisting of feldspar crystals and angular volcanic lithics within an environment previously receiving mudstone suggests a penecontemporaneous volcanic eruption. Unequivocal evidence for explosive pyro­ clastic volcanicity is provided by the vitric siltstones containing accretionary lapilli. Although the nature of the eruption is hard to pin down, the high pro­ portion of non-vesicular lithics and crystals, low abundance of vitric-material, absence of pumice, and presence of accretionary lapilli suggest small­ scale phreatomagmatic eruption(s) (cf. Sheridan & Updike, 1975; Heiken & Wohletz, 199 1 ) . The rhyolitic deposits have a combined volume of at least 0.2 km3. As crystal content in volcanic rocks

rarely exceeds 50% (Ewart, 1979), the crystal content of sandstones must have been enriched, either by fractionation processes accompanying the eruption and deposition of the tephra and/or hydraulic sorting during fluvial reworking. The erupted volume of magma would therefore be much larger than the volume of deposits in the Tryfan Fach region, perhaps double. Although still small, the volume is still well in excess of the amount of pyroclastics (c. 0. 15 km 3) extruded in the 1980 eruption of Mount St Helens (Rowley et al. , 198 1 ) . Pre-eruptive marine mudstone association

The fine-grained sandstones intercalated with the marine mudstones are moderately sorted quartzo­ feldspathic and lithic arenites. No samples were point-counted owing to the fine grain size and dif­ ficulty of identifying components, but composition appears similar to the 'norm' for the Tryfan For-

180

G.J. Orton

mation (Table 1 ) . The anhedral nature of framework grains, the high percentages of chloritic matrix and the dominance of sedimentary lithics suggests derivation from a source of tuffaceous sedimentary rocks and mudstone. Although rhyolitic ash-fall deposits occur near the base of the succession (Fig. 3) there is no evidence elsewhere for pene­ contemporaneous volcanicity, with the exception of the vitric siltstones at the very top of the succession. Possible position of the volcanic source

In many cases volcaniclastic successions can be interpreted in terms of proximity to a central volcano (e.g. Smith, 1988) by comparison with well-studied modern examples. In the Tryfan Fach Member, the absence of primary volcanic deposits (larvas, pyro­ clastic flow deposits) , coarse breccias and debris flows is consistent with deposition in a distal setting. Similar monolithological syneruptive sands occur on the most distal ring plains that flank major central volcanoes at distances of 30- 100 km from volcanic vents. I begin by discounting these distances. Within the Llewelyn and Snowdon Volcanic Group, the absence of systematic facies variation across the entire out­ crop belt, as well as the common identification of intrabasinal vents, argues against derivation of volcaniclastic material from a line of volcanoes beyond the reach of the present outcrop. Instead the stratigraphy reflects a rather haphazard assemblage of small volcanoes, none of which had appreciable subaerial relief, erupting into and within subsiding sedimentary lowlands (Howells et al. , 1 99 1 ) . Although the Welsh basin i s typically portrayed a s a back-arc basin, this is not essential, and such low­ lying volcanic fields can also occur due to intra-arc extension (e.g. White & Robinson, 1992). In such settings, 'distal' or more accurately 'low-gradient' deposits (e.g. Tryfan Fach Member) are often locally interstratified with more 'proximal' facies (e.g. lavas, pyroclastics). In North Wales, most large Ordovician volcanic eruptions had vents within a few tens of kilometres, or less, of the sites of deposition of their primary products (excluding fallout tuffs) (Kokelaar, 1988, 1992; Howells et al. , 199 1 ) . It is not necessary to suppose that smaller eruptions had more distant volcanic sources. The absence of a correlative coarse rhyolitic unit amongst more proximal successions about 5 km away (see above 'Extent of the deposit' section) also lends support for a local source for the Tryfan Fach Member.

This remains to be located . There are several subvolcanic intrusions in the Llewelyn Volcanic Group for which comagmatic extrusive correlatives have not been recognized (Fig. 1) and which invite comparison. These lie about 5 km or more (after unfolding strata) to the west of Tryfan Fach along the northeast trending Nant Peris fault. The largest intrusion (Bwlch y Cwyion) consists of microgranite (rhyodacite/dacite) with later marginal intrusions of rhyolite (Howells et a!. , 1991) whereas smaller intrusions are rhyodacites , granodiorites and feld­ spar porphyries.

REASSEMBLING THE STRATIGRAPH ICAL P IECES

In the Tryfan Fach Member, each facies association is in some way different, in composition and/or facies, to adjoining facies associations. It will be shown that the meaning one attaches to facies changes between and within associations depends on whether the rhyodacitic sandstones are attributed to penecontemporaneous explosive volcanism. One epiclastic system: time's cycle

The simplest solution, and the one that would be adopted if the petrography of sandstones and composition of rhyolite clasts was unknown, is to integrate the four facies associations into one linked depositional system or systems tract (Fig. 14) . As noted by Hutton ( 1788) it is with 'pleasure that [we] observe order and regularity in the works of Nature, instead of being disgusted with disorder and confusion'. In this interpretation , a tributary alluvial fan at Craig Yr Ysfa would feed sediment on to an axial braidplain confined to the southwest by additional uplift along Llanberis Pass. Younger alluvial systems in the Tryfan Formation exhibit this drainage pattern (Orton, 199 1 ) . The sandier upwards marine sheU association would represent offshore to coastal plain environments reflecting deposition in front of, adjacent to , or between alluvial fan salients that were deprived of coarser sediment. Tectonic uplift along the basin margin resulted in progradation of the alluvial fan and a derivative braidplain into the marine basin. Lesser amounts of uplift within the basin resulted in a relative sea-level fall and braid­ plain sandstones rest with sharp contact on a condensed nearshore succession . Following tectonic

SEDIMENT COMPOSITION

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related. An unchanging marine transgression is assumed throughout. Place names are: TF, Tryfan Fach; CY, Craig yr Ysfa. Facies associations as in Fig.

12.

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182

G.J. Orton

uplift, a more gradual shift in the position of the aLluvial fan lobe aLlowed floodbasin waters to gradually transgress on to the fan/braidplain surface. Floodbasin siltstones therefore increase in pro­ portion upwards at the expense of coarse sandstone beds. There are many pieces that do not fit into this puzzle: (i) the composition of braided stream deposits differs from sandstones of the alluvial fan at Craig Yr Ysfa, implying derivation from a differe;1t aLluvial fan or from a more distant axial source; (ii) coarse rhyolitic sand is never present in the marine shelf association - one would expect

occasional high-velocity floods to introduce coarse material on to the shelf; (iii) the thickness and 'purity' trends in reworked ash-fall tuffs at the top of the marine shelf association , as well as grain-size trends across the braidplain (Fig. 5) suggest a source to the southwest rather than to the northeast. Together these features cast doubt on the strict application of Walther's Law to the total sequence and seems to demand three weakly or unlinked depositional systems that developed in advance of, and in conjunction with penecontemporaneous volcanism (Fig. 15).

· :-. . · · · · :·.. ·

SEDIMENT COMPOSITION

FACIES

ENVIRONMENT •

·· · �

. :;, , J:�· ·[?i;��t��)· �t:�l;-

ALLUVIAL PLAIN

BRAIDED SAND·Sf£ET

NNER SHELF

OUTER SH::LF

Fig. 15. Pyroclastic (or syneruptive) interpretation of the vertical succession. Block diagrams (also looking northeast) depict palaeogeography if Walther's Law cannot be applied to the full succession and morphological elements need to be considered separately. Two types of 'reworked pyroclastic materials' are portrayed on 'sediment composition' column in order to illustrate uncertainty about the number of volcanic events. Darker tone indicates pyroclastic materials that have been reworked more extensively (i.e. they were deposited a greater length of time after the volcanic eruption). Several (A, B and C) possible interpretations of relative sea-level change are presented for the syneruptive period. See text for discussion. See Fig. 14 for key to sediment composition.

Facies models in volcanic terrains Three unrelated depositional systems: time's arrow

Progradation of a muddy shelf and postulated volcano-tectonic uplift

An initial supply of fine-grained silts and mudstones (largely non-volcanic) was delivered to a low-wave­ energy basin, probably by small channels or sheet­ floods (cf. Wells, 1983). Sand distribution suggests that the source lay to the northwest to west (Fig. 1 5 , bottom block diagram). Reworked rhyolitic ash-fall tuffs were derived from a similar direction and may represent small eruptions precursor to the main volcanic event. The lateral variation in the amount of sand suggests that sediment input occurred from localized point sources with minimum lateral reworking at the shoreline. Widespread tectonic uplift preceding and associated with volcanicity resulted in a relative sea-level fall. Hence coarse­ grained sandstones derived from the volcanism rest with a sharp, probably erosional contact on a condensed nearshore succession . Five to 15 m of uplift would be needed to cause erosion of upper shoreface/nearshore deposits. The timing of uplift with respect to volcanism in the Tryfan Fach Member is not thought to be fortuitous and some of the uplift is related to buoyant forces associated with the intrusion and/or migration of magma. All large rhyolitic eruptions in North Wales show evidence for tens of metres of tectonic uplift in the vicinity of volcanic centres in advance of eruptions (Howells et al. , 199 1 ; Orton, 1991) and there is no reason why smaller amounts of uplift might not precede smaller rhyolitic eruptions. A volcano-tectonic control on uplift and/or development of regressive cycles has also been suspected within other ancient sedimen­ tary, particularly extensional, basins (cf. Leeder, 1983 ; Henry & Price, 1984; Watkins, 1986; Leckie & Cheel, 1989; Sloan & Williams, 199 1 ) . Our understanding of the influence o f volcanism on basin subsidence is limited by the short time span < 100 yr) over which detailed observations from modern volcanoes have been recorded and the difficulty in ancient successions to separate volcano­ tectonic from tectonic effects. Most modern data is from volcanic fields where large silicic caldera­ forming eruptions previously occurred (Newhall & Dzurin, 1988). Vertical ground movements are best recorded at calderas located by the sea, where elevation changes, forming land , are very con­ spicuous. Ground deformation in calderas occurs readily because of the existence of large shallow

183

(3 - 15 km) magma chambers, with rates of uplift ranging from 0.002 to almost 1 m per year (cf. McKee et al. , 1984; Hill et al. , 1985; Luongo & Scandone, 199 1 ) . These rates obviously are extremes or we would not notice them. The unique longer term (2000 yr) historic record from Campi Flegrei caldera (Italy) shows that episodes of unrest are often separated by decades of quiescence (Lyell, 1853, Caputo, 1979; Dvorak & Gasparini, 199 1 ) . Here, 10 centuries o f gradual subsidence ( 1 1 m) was followed by more gradual uplift beginning in the eleventh century. In 1538, a mere 0.03 km 3 of ash and scoria erupted from Monte Nuovo. This was preceded, however, by 40 years of episodic uplift (and land formation) culminating in 7 m of uplift in the two days preceding the eruption at a site (Pozzuoli) several kilometres away. As the Tryfan Fach Member was not deposited within or nearby such a caldera, it is pertinent to query whether the above data is relevant. Extrusive rocks of the lowermost Llewelyn Volcanic Group were ponded in subsiding, probably fault bound, sectors of the sea-floor (Howells et al. , 199 1 ) , the thickest accumulations (both about 1000 m) being rhyobtic lavas and ash-flow deposits of the Conwy Rhyolite and trachyandesite lavas and tuffs of the Foe! Fras Volcanic complex (Fig. 1 ) . As has been noted recently (Lipman, 1984; G.P.L. Walker, 1984) , it is a semantic issue whether the rectangular volcano­ tectonic depressions that contain these deposits should be termed caldera. The final volcanism of the Llewelyn Volcanic Group, the Capel Curig Volcanic Formation involved eruption of two rhyolitic ash­ flow tuffs (each > 20 km3) from a postulated vent in the north near Conwy, as well as smaller eruptions (c. 1 km3) from volcanic centres just north of Carnedd Llewelyn and just south of Tryfan (Fig. 1 ) . The continuity o f compositional variation within the Llewelyn Volcanic Group suggests that various intrusive and extrusive products were all originally comagmatic and related by closed-system fractional crystallization (Howells et al. , 199 1 ) . This necessi­ tates a parental magmatic system that stretched 25 km from the Llanberis Pass fault toward Conwy, and extended about 10 km in a transverse direction. This magma 'chamber' may have been an elongate mid- to subcrustal beam or dyke controlled by the dominant structural grain , a shape documented by geophysics from modern continental rifts (e.g. Cordell et al. , 1985 ) . Petrochemical and exper­ imental data (Ball & Merriman , 1989) show that rhyolites derived from this magmatic system evolved

184

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at relatively shallow crustal levels (PH,o < 500 bars) . Thus many of the conditions needed for volcano­ tectonic 'unrest', such as earlier caldera-forming eruptions, existence of large , shallow magma chambers , volcanism in an extensional graben-type structure, existed within the Snowdon graben at Tryfan Fach time. Aggradation of synerup tive pyroclastics

Sediment composition within braidplain and flood­ basin settings all indicate deposition in conjunction with a period of explosive rhyodacitic volcanicity (Fig. 15, top block diagram). The sudden increase in sediment grade and abundance resulted in a brief period of aggradation within shallow ephemeral channels. This braidplain prograded rapidly as a sheet across the pre-existing shoreline and/or exhumed marine mudflat into a low-wave-energy basin, with facies suggesting a more direct input of volcanically produced sediment loads on to the southwest portion of the braidplain. It is not certain whether sandstones represent pyroclastic material reworked from a series of small eruptions rather than a single eruption (left- and right-hand sediment composition columns of Fig. 15). Consequently, the thin siliceous mudstone beds within braidplain deposits could reflect: (i) periods of quiet-water sedimentation following lateral migration of the fluvial channels (the most 'traditional' interpretation) , or (ii) fallout of fine­ grained ash following episodes of rapid syneruptive aggradation of coarse clastic material. Fine ash may have been erupted at the same time as the underlying coarse, clastic material but initially dispersed within an eruption cloud. There are also several interpretations that can be presented for the braidplain to floodbasin transition. 1 The upwards decrease in the proportion of coarse­ grained sandstones may reflect gradual waning of the coarse sediment supply, perhaps due to erosive lowering of the volcanic edifice and/or near-vent subsidence following the volcanic eruption. In this model, the maximum basinward advance of the syneruptive braidplain is 'picked' as the top of the coarse-grained rhyolitic sandstones. Fine-grained floodbasin siltstones would largely represent matrix of the pyroclastic flows, liberated by fluvial rework­ ing, and transported to the site of deposition by low­ gradient streams. In this model, the floodbasin successions are distal and should interfinger else-

where (more proximally, or laterally) with the coarser grained braided stream deposits. If one was trying to determine relative sea-level changes, the top of the braided sandstone sheet might be picked as a marine flooding event and used to infer a rise of relative sea-level (sea-level curve B , Fig. 15). 2 A contrasting scenario results if one allows volcanism to continue and the type of volcanism to change with time. Time lines would essentially parallel bedding contacts rather than cut across them as in the first scenario . In this case, it is assumed that all vitric siltstones are primary, or only slightly reworked, pyroclastic materials. Rather than pro­ ducing coarser grained pyroclastic material, the volcano produced fine-grained ash. The occurrence of accretionary lapilli suggests that this change may have been triggered by access of water to the vent. Fluvial channels did not become more distant (i.e. distal) from the source up-section, they just became overwhelmed by aggradation of fine-grained primary and reworked fallout ash. Hence, coarse sandstone progressively formed a smaller proportion of the sediment deposited on the alluvial plain. Laminated ashfall siltstones may have been deposited in ponded depressions on the alluvial plain; a marine trans­ gression may have occurred only at the very top of the succession, after all volcanism had ceased (curve C, Fig. 15). Two observations favour the second interpret­ ation. Firstly, sandstones at the top of the floodbasin association (Fig. lOA) are as coarse as those lower down. Secondly, lamination in many of the siltstones/ fine-grained sandstone beds is often . diffuse (Fig. lOB) indicating that fine sediment from one event was still being deposited when coarsest grains from the next event arrived . Near-source ash-fall deposits commonly display such weak bedding (e.g. Branney, 1991) owing to fluctuations in the energy of the eruption, direction and strength of winds or over­ lapping of several eruptive pulses. In contrast, sedimentary laminae in lakes are commonly sharp­ based (e.g. Hamblin, 1992) because they reflect infrequent episodes (e.g. seasonal varves) of clastic input. A symbiotic alluvial fan ?

The relationship of the alluvial fan deposits to the other two genetic units is most perplexing. One might at first expect that the two stratigraphical units were not contemporaneous. However, as four

185

Facies models in volcanic terrains

coarsening upward alluvial fan sequences can be recognized in the Tryfan Formation at Craig yr Ysaf (Orton, 1988, Fig. 4) , and four braidplain -deltaic successions have been recognized in the Tryfan anti­ cline it seems reasonable to make a one-to-one correlation. None of the other alluvial fan suc­ cessions contain feldspar-phyric rhyolitic volcanics that could suggest otherwise . Rhyolitic ash-fall deposits within fine-grained 'shoreface' deposits that underlie the coarse-grained alluvial fan deposits lend further support to this correlation. The evidence suggests that the alluvial fan pro­ graded at about the same time as volcanism was active, or shortly thereafter, but from the opposite side of the basin. Tectonic changes associated with the volcanism must have produced enough relief between Craig yr Ysfa and the Tryfan depocentre to allow fan progradation. Migration of magma may have been the mechanism driving this subsidence pattern. Magma migration should have two effects: (i) it will produce relative uplift in the immediate vicinity of the volcanic vent; (ii) fault blocks that lose some of their supporting 'cushion' of magma could subside relative to surrounding fault blocks, even those more distant from the volcanic centre. In some ways this process is analogous to the movement of salt, laterally and upwards, toward the position of least overburden pressure. Magma migration is documented most commonly in extensional volcanic settings. It has been invoked in the Rio Grande rift to explain progressive increase in the gradient ( 1 m km- 1 to 7 . 1 m km- 1) of rift drainage systems from late Pliocene to Holocene times (Bachman & Mehnert, 1978) and present-day subsidence that is similar in size but opposite in sign to uplift in the Socorro area of New Mexico further along the rift axis (Reilinger & York, 1979). A more convincing illustration is provided by the 1912 eruption of Novarupta (Alaska) where a depression 600 m deep (volume 6 km3) formed at Mount Katmai 10 km away from the eruption site, implying a complex magmatic plumbing system (Hildreth, 1987). In the basaltic Bedded Pyroclastic Formation (North Wales), episodes of subsidence preceded periods of copious magmatism and rapid sediment accumulation (Kokelaar, 1992); this subsidence pattern could be explained by lateral migration of magma away from deposition sites and toward eruptive vents. Within a Variscan pull-apart basin in Spain, Besley & Collinson, ( 1991) also documented accelerated but more differential subsidence during

active volcanism and concluded that 'volcanic activity was related in a close but subtle way to the subsidence history of the basin, and thus to the sedimentology of its fill'.

THE STRA T IGRA P H ICAL RECORD OF VOLCANISM

Sedimentary sequences associated with active volcanoes are often considered to consist of two elements (G. Smith , 199 1 ; White, 199 1 ) : (i) syn­ eruptive sequences, which result from primary volcanic deposition and immediate post-eruptive reworking, (ii) inter-eruptive sequences, which record deposition without significant influence of volcanic activity where normal sediment delivery processes are dominant (Fig. 16) . The syneruptive condition is a perturbation brought about by the volcanic eruption and includes not only the time­ scale of the eruption but also the period following cessation of volcanism while hydrological and hill­ slope conditions remain disturbed. In humid cli­ mates, and without vegetation, the syneruptive condition may continue for a long time, at least one or two decades, until most loose debris from the eruption is removed completely. In the Tryfan Fach Member it includes the braided stream and flood­ basin facies associations. Syneruption sedimentation is in large part aggradational and the resultant stra­ tigraphy is an event stratigraphy with time planes essentially parallel to bedding surfaces. Around stratovolcanoes

Most conceptual models for subaerial volcaniclastic alluvium are based largely on sedimentation within volcaniclastic aprons deposited adjacent to relatively high-standing volcanic arcs (e.g. van Houten 1976; Vessell & Davies, 198 1 ; Smith, 1987, 1988). During volcanic eruptions the instantaneous production of large volumes of sediment, combined with enhanced and more variable runoff, result in sedimentation principally by aggradation of high-sediment-load flood and debris-flow processes. Inter-eruption periods are often characterized by erosion and channel incision; incision (Vessell & Davies, 1981) following syneruptive episodes was an integral aspect of the predictive facies model presented by G . A . Smith ( 1991) for continental volcaniclastic sediments. Inter-eruption streams draining the high-standing

186

G.J. Orton

volcanic highlands are commonly of gravel-bedload character. The overall geometry of the succession that results is controlled by the relative importance of syneruptive versus inter-eruptive conditions, which in turn are controlled by eruptive frequency and rates of tectonic subsidence (G . A . Smith, 199 1 ; Fig. 16a). Within a humid low-relief continental rift

The Tryfan Fach Member succession differs from the above scenario in three ways: (i) debris flow deposits are largely absent from the syneruptive sedimentary record; (although debris flows occur in a penecontemporaneous basin-margin alluvial fan, this occurrence was unrelated to high sediment loads due to volcanism); (ii) incision, if anywhere, occurred at the base of syneruption deposits; and (iii) fine­ grained muds, rather than gravels, were deposited after or between volcanic events. These differences are thought to arise largely because the Tryfan Fach sedimentary system developed within a humid­ region, actively subsiding, continental rift. Syneruptive facies and debris flows

Three reasons are given for the paucity of debris flows.

A: STRATOVOLCANOES

1 Slope was insufficient. Rhyolitic volcanic fields typically have a subdued topography. In North Wales, the stratigraphical architecture and the dis­ tribution of products from volcanic cones suggest that maximum relief was a few hundred metres (Howells et al. , 199 1 ; Orton, 199 1 ) . Walton ( 1986) also argued that the paucity of debris-flow deposits in the Tertiary rhyolitic volcaniclastic sediments of West Texas resulted from the low-gradient of source areas. At the Taupo volcanic centre , although debris flows occur in post-eruptive reworked pyroclastic deposits (R. Smith, 199 1 ) , their development depended on the moderate slopes ( 10°-20°) of the inward dipping pre-existing topography surrounding Lake Taupo. This topography is a tectonically uplifted ridge of Upper Palaeozoic- Mesozoic grey­ wackes that has nothing to do with the volcano itself (Cole , 1984). 2 Depositional site was too distal. Cross-stratified sands, not unlike those of the braided stream associ­ ation, occur within most distal volcaniclastic suc­ cessions. As volcanoes range from a few tens of metres (e.g. tuff rings) to several kilometres (e.g. some stratovolcanoes) in height, the term 'distal' has no singular meaning in terms of distance from source , with 'distal' facies developing as close as a few kilometres to more than 100 km from the volcano.

8: LOW RELIEF HUMID RIFT · ·:�_ :·:::.:s:�-:;-�-: :·: :: :};����.-: _-.=�-:: �--�_-��·:;-::-�� ��� -� .� �::.:.C.. ::.�.l.�:�;�<��-:., ;,:;.).�:� i..�:�.-:��:<; - -��- -0�_-.· .·..:; � · ·.".: ;.:� · - ·-�·:.-:·:· ::: . .·o . •:a. ,..,uo "-o.-��-o.an.: .•:. -·

f:t�: .��:�i�?:�\:�-?�_�.� �??��

,. -.

.. .

..�-

.-c

••

primary pyroclastics

SYN­ ERUPTIVE

debris flow hyperconcentrated flows braided streams

Fig. 16. The stratigraphical record of volcanicity: (A) around stratovolcanoes. Pre-eruptive and post-eruptive successions are not separated (from G.A. Smith, 1991). (B) Fluvial-deltaic successions of the Tryfan Formation, North Wales. Lower example would represent Tryfan Fach Member.

187

Facies models in volcanic terrains 3 Pyroclastic materials were reworked from a series of small eruptions rather than a single eruption. Although most small monogenetic volcanoes are short-lived , multiple soil horizons described from the 'flanks' of some volcanoes (Wells et al. , 1990) , direct observation of volcanic activity (Foshag & Gonzalez, 1956) and ancient examples (e.g. Kokelaar, 1992) indicate repeated eruptions over a long time period (decades to hundreds of thousands of years) with reuse of the same magma plumbing system. Due to the small size of individual eruptions, and time between eruptions, primary products (pyroclastic flows, debris flows) could be completely reworked into cross-stratified sandstones.

Bounding surface between syn- and post-eruptive deposits

All sedimentary basins are the result of isostatic adjustment of the lithosphere to a tectonic driving force , enhanced by subsequent sediment loading (Watts, 1989), and in this case modulated by the presence of an underpinning, buoyant magma body. The amount and rate of surface deformation is controlled ultimately by the strength of the litho­ sphere (Karner & Dewey, 1986) and gravity controls on the upwards migration and accumulation of magma (G.P.L. Walker, 1989). If the lithosphere is highly fractured, or magma is of low density and/or viscosity, deformation will be compensated locally. Unconformities can develop within any sedimen­ tary basin, their occurrence depending on rates of subsidence, rates of sediment production, and eustatic sea-level change. The only base-level scenario that can prevent syneruption deposits of the Tryfan Fach Member from being incised after aggradation is to invoke a rapid rise in relative sea­ level in conj unction with the explosive volcanism. Given the short time available, a eustatic control seems unlikely, and the sea-level rise is best related to accelerated subsidence of the underlying fault block. Although this subsidence could be purely tectonic and have no relation to the volcanism, the coincidence of the two events seems hard to explain. Several other authors have also noted that vol­ canism was associated with accelerated and/or dif­ ferential basin subsidence (Kokelaar, 1988; Besley & Collinson, 199 1 ; see above). This genetic linkage between rates of basin subsidence (e.g. relative sea­ level rise) and sediment load (e.g. an eruption) is uncommon in normal sedimentary successions. To some extent, however, particularly in strike-slip

basins, the details of the relative timing of sub­ sidence and volcanism is a 'chicken-and-egg' type of question. Did increased transtension and associated subsidence cause the volcanic eruption or did the volcanic eruption occur and allow or drive renewed subsidence? Inter-eruptive sedimentation

Inter-eruption or 'background' facies and facies geometries depend on the topography, relief, type of bedrock , climate, subsidence history, etc. of the drainage basin. Almost any facies could conceivably occur. In low-relief extensional basins fine-grained mud commonly is deposited after volcanic events on alluvial plains or within lacustrine to playa environments (e.g. Besley & Collinson, 199 1 ; Buesch, 199 1 ; Turbeville, 1991 ; White & Robinson, 1992; this study) . These environments arise because of the low topography of the basin, because syn­ eruptive deposition flattens and infills topography, because low-gradient drainage systems are easily blocked or re-routed by lava flows or debris flows, or because volcanism drives basin subsidence.

FACIES MODE L L ING : T IME ' S ARROWS V ERSUS T IME ' S CYCLE

The stratigraphical record allows us to define 'bite­ sized' pieces or facies (R. G. Walker, 1990). In facies modelling we assume 'there is a system and order in Nature' (R.G. Walker, 1992) and attempt to re­ assemble these 'bites' into a group of facies that are genetically related to each other and have some environmental or palaeogeographical significance. The two end-member conditions, time's arrow (a unilinear succession of unique events) and time's cycle (recurrent patterns in a world that remains essentially unchanged) are categories of our invention, erected here to illustrate the problems of interpreting volcaniclastic successions. Although both approaches, by themselves, are incorrect (as noted by Gould, 1987), one may become dominant. Two variables unite to determine the degree to which Walther's Law is applicable; these are the completeness of the sedimentary record (the num­ ber of time gaps), and the extent to which the sedi­ ment source changed, in terms of the amount and grain size of supply, throughout deposition (Fig. 17) . The two end-member conditions are: continu-

188

G.J. Orton

SEDIMENT COMPOSITION

STRATIGRAPHIC "BITES"

COMPLETENESS OF RECORD

WALTHER'S MARITAL STATE

t

(/) (/) w z � u None

A lot

AMOUNT SOURCE CHANGED BETWEEN SUCCESSIVE BEDS

�

PURE BLISS

DIVORCED

Fig. 17. Application of Walther's Law to syneruptive and inter-eruptive volcaniclastic successions (using ideas from Dott,

1983; Gould, 1987; Walker, 1990, 1 992). Walther's marital state is discussed further in the text.

ous sedimentation and/or an unchanging source (Walther's Law fully applicable) ; episodic sedi­ mentation and/or a constantly changing source (Walther's Law inapplicable) . As all environments, including non-volcanic ones, are to an extent influenced by episodic events (Dott, 1983 ) , the metaphor that becomes dominant depends on the time frame over which observations are made. The time's arrow metaphor embodies the entire Tryfan Fach Member. The coarse-grained cross­ stratified and thinly bedded rhyodacitic sandstones are unrelated to and lie with probable erosional unconformity on subjacent marine mudstones. The up-sequence progression of facies on this time-scale is essentially unpredictable, as pre-existing sediment dispersal systems were overwhelmed and probably

transformed by introduction of coarse-grained rhyodacitic material associated with the explosive volcanic eruption. The time's cycle metaphor is most applicable within the pre-eruptive marine mudstones and the epiclastic alluvial fan succession. Because source was comparatively constant and unchanging (or changing gradually) a facies model can be provided for the whole succession, vertical sequences can be interpreted, and palaeogeographical maps can be presented . Although unique events did occur (e.g. the ash-fall beds, flood events from the alluvial plain) , their magnitude was insufficient to radically alter pre-existing sediment dispersal systems and/or environments. Episodic hydrological and oceano­ graphic events (storms, river mouth floods) become

Facies models in volcanic terrains

'normal' over the time period (tens of thousands of years) involved and Walther's Law is fully applicable. In non-volcanic terrains , and in the epiclastic successions noted above, it is implicit (de Raaf et a/. , 1965) or explicit (Busch , 1971) that any progressive facies association with gradational facies contacts is the result of one related set of depositional conditions. However, within syneruptive reworked pyroclastic deposits, gradational grain-size and facies changes are virtually meaningless (or at least very difficult to interpret! ) in two-dimensional sequences unless the eruptive history of the volcano is known. This is rare, even in historic eruptions. Many volcanic eruptions consist of a series of closely spaced but smaller events, with the amount, grain size, density, shape and/or 'distribution of eruptive products gradually changing owing to progressive changes in the amount of magma vesiculation, the eruptive flux rate, vent parameters , and the supply rate of water. Although each bed is unique, the succession of syn­ eruptive beds may not be random but linked by gradual (?and predictable) changes in source region parameters. The problems this causes in sequence analysis are well illustrated in the Tryfan Fach Member by the braidplain to floodbasin transition, where it could not be determined whether facies change reflected sorting processes on the alluvial plain or progressive eruption of finer grained volcanic ash. Environments were difficult to reconstruct and palaeogeographical maps could not be produced. As a consequence the timing of other controls on the stratigraphy (e.g. sea-level change) could not be fully assessed . Thus although one cannot totally divorce Walther's Law from these volcaniclastic sequences it certainly must be applied with caution, with the relationship best described as an unhappy marriage (Fig. 17) .

1 89

a nearby, but unrelated, source. However, within syneruptive pyroclastics it could not be determined whether coarse sandstones were reworked from a single or several eruptions, and whether rhyolitic siltstones on the alluvial plain represent fine-grained material extracted from older pyroclastic flows during reworking or influx of ash from new volcanic eruptions. The key to unlocking the story hidden within such volcaniclastic sequences is to recognize when volcanism was active. The basic question that is addressed is whether facies changes in a vertical sequence reflect changing energy levels within the depositional setting (Walther's Law is applicable) or changing source parameters because of or during an eruption (Walther's Law does not work ) . In this regard, the study of volcaniclastic materials is still in its infancy, and is the area where much new research is needed. The above understanding is possible only if volcanism-induced eruptive and emplacement processes can be recognized and separated from 'background' conditions of sediment supply (see also McPhie & Allen, 1992). Even when the eruption is observed (e.g. Mount St Helens) , it may be dif­ ficult to distinguish between primary and reworked pyroclastic material, and to recognize individual eruptive episodes in the depositional record. These distinctions becoming increasingly difficult to make when studying the ancient record , where syneruptive sheets may be difficult to separate (G. Smith, 1991) and material from smaller eruptions becomes admixed with 'background' inter-eruption detritus (Hackett & Houghton, 1989). In these successions, it may not be possible to determine whether facies change reflects a change in type and/or energy of the environment, or a change in the type of sediment produced by the volcano, or both. We should be prepared to admit this and realize the implications for the accuracy of environmental and/or sequence interpretations .

CONCLUSIONS

Coarse-grained sandstones of the Tryfan Fach Member represent aggradation and progradation of reworked pyroclastic materials deposited in a braid­ plain setting in conjunction with penecontempor­ aneous ?phreatomagmatic rhyodacitic volcanicity. It can be shown with reasonable certainty that these deposits have no genetic relation to either muddy shelf deposits , upon which they lie with probable erosional unconformity, and/or adjacent alluvial fan deposits, which prograded at the same time from

ACKNOWLEDGEMENTS

Research formed part of a doctoral thesis ( 1990) completed under the helpful guidance of Harold Reading (Oxford University) and Roger Suthren (Oxford Brookes University) . Later visits to North Wales were financed by the Monash University Development Fund and the Australian Research Council. Over the years stimulating field discussions in North Wales with Dairmid Campbell, Ray Cas,

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Bill Fritz, Malcolm Howells, Adrian McArthur and Tony Reedman have contributed to my understand­ ing. Ideas gathered in North Wales have benefited by additional research experiences in Australia (Exhibition of 1851 Fellowship) and Japan (STA & AIST Fellowships) where similar environmental problems were encountered. Ideas presented in this paper were greatly sharpened by thorough and constructive reviews by Gary Smith, James White, and the editor, Guy Plint.

Bull. A m . Ass. petrol. Geol. , 55, 1 137- 1 154. CAMPBELL,

C .V. ( 1976) Reservoir geometry of a fluvial sheet sandstone . Bull. Am. Assoc. petrol. Geol. , 70, 1009 - 1020.

CAMPBELL,

S.D.G. , HowELLS, M.F. , SMITH , M . & REEDMAN, A.J. ( 1 988) A Caradoc failed rift within the Ordovician marginal basin of Wales. Geol. Mag. , 125,

257-266. CANT, D.J.

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pycnal plumes and plume fronts over the Huanghe

Vos, R . G .

&

ERIKSSON , K . A .

( 1 977) An embayment

model for tidal and swash deposits occurring within a fluvially dominated middle Proterozoic sequence in

&

WiSEMAN , W . J . ( 1986) Hyper­

(Yellow River) delta front . Geomarine Letters, 6, 97106.

Tectonics and Sedimentation

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

Spec. Pubis int. Ass. Sediment.

(1995) 22, 197-217

Coarse-grained lacustrine fan-delta deposits (Pororari Group) of the northwestern South Island, New Zealand: evidence for Mid-Cretaceous rifting M A L C O L M G. LAIRD Institute of Geological and Nuclear Sciences, c/o Department of Geology, University of Canterbury, Christchurch, New Zealand

A BS T R A C T

Mainly coarse-grained non-marine sediments of the Pororari Group, dominated b y breccia and conglomerate, and of mid- to late Albian age, are widespread west of the Alpine Fault in the South Island of New Zealand. They represent an unconformity-bounded sequence, resting with marked angular discordance on Early Palaeozoic sediments of the Greenland Group or non-conformably ori mid-Palaeozoic and Early Cretaceous granites, and overlain with regional discordance by coal measures of Maastrichtian age. In the southern Paparoa Range, where the Pororari Group is in excess of 2000 m thick, it can be divided into three major facies assemblages: (1) matrix- and clast-supported breccia; (2) pebbly sandstone and massive or graded sandstone; and (3) laminated dark mudstone with thin graded sandstones. Facies assemblage 1 is inferred to represent the deposits of debris flows and sheet flows on alluvial fans; facies assemblage 2, the deposits of debris flows and turbidites in a dominantly lacustrine environment; and facies assemblage 3 was deposited in a distal lacustrine environment. The three facies assemblages intertongue, and are inferred to represent elements of a fan-delta system debouching into a lake. The palaeocurrent pattern deduced from clast imbrication in breccia units, and from sedimentary structures in the lacustrine deposits, indicates a dominantly NNE-SSW direction of sediment transport. The consistent and rapid thinning of lithostratigraphical units from north to south is compatible with the palaeocurrent results. Palaeocurrent plots and clast lithologies suggests that several point sources supplied coarse sediment during· the history of deposition of the Pororari Group, each giving rise to distinct but overlapping fans. The great thickness of coarse deposits (a minimum of 1200 m for the breccia assemblage alone) and the angularity of clasts suggests that these sediments were generated by contemporaneous tectonic activity. Fluctuations in lake level also may have been controlled by active faulting at the basin margin, although climatic changes may have played a part. The fan-delta system is likely to have occupied an active subsiding half-graben. The consistently south-southwest directed sediment gravity flows suggest that the active fault was likely to have been oriented in a WNW-ESE direction, downthrown to the south, and with an implied north-northeast direction of tectonic extension. A similar extension direction for the Paparoa Range area during early. Albian times has been deduced from independent structural research on a closely associated metamorphic core complex, and is also compatible with the WNW­ E SE orientation of normal faults bounding adjacent offshore undrilled half-grabens infilled with deposits of prdbably mid-Cre��ceous age. The coarse deposits of the Pororari Group are likely to have been associated with an early period of rifting in the Albian, pre-dating by perhaps 25 Ma the separation of New Zealand from Gondwana in the late Campanian or early Maastrichtian.

INTRODUCTION

Coarse-grained non-marine sediments, o f mainly mid- to late Albian age, dominated by breccia and

conglomerate, are widespread throughout the southern and western South Island of New Zealand.

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0 197

198

M.G. Laird

Most are poorly exposed, and detailed descriptions are sparse. One exception is the Kyeburn Formation, which lies in the southeast of the South Island (Bishop & Laird, 1976; Fig. 1), and which represents talus to braided river and lacustrine deposits of mid­ Albian to at least late Cenomanian age, emplaced in a locally developed, triangular, fault-bounded depression. Another well-described exception is the Puysegur Formation of the southwestern South Island (Lindqvist, 1990; Fig. 1), which is inferred to have been deposited in terrestrial, lake margin, and lacustrine environments in a fault-controlled basin. All of these mid-Cretaceous sediments rest uncon­ formably on older rocks, and clearly post-date the Early Cretaceous Rangitata Orogeny, which marked a significant change of tectonic pattern in the New Zealand region. From the Permian to the Early Cretaceous, most rocks formed under the influence of convergent-margin tectonics and comprise incom­ plete remnants of magmatic arcs, fore-arc basins, trench-slope basins, .;:tnd accretionary complexes (Bradshaw, 1989). The later stages of this regime, in the Early Cretaceous, were accompanied by wide-

spread compressional deformation, metamorphism, uplift, erosion and calc-alkaline plutonism. Over much of New Zealand the rocks of the Rangitata Orogen are separated by a major uncon­ formity from younger, less-deformed strata. This unconformity marks the change in tectonic regime from convergent margin to extension, attributed by Bradshaw (1989) to collision with the New Zealand margin of the spreading ridge between the Phoenix and Pacific Plates. Mid Cretaceous extensional tec­ tonism is best exemplified by the post-unconformity development, in the mid Cretaceous, of half grabens throughout much of the New Zealand region, both onshore and offshore. Subsidence continued along the bounding faults in the western and southern South Island, at least in some instances, from mid­ Albian until Santonian or early Campanian times (Laird, 1993, 1994). Both the Kyeburn and the Puysegur Formations are inferred to occupy basins formed east of the Alpine Fault in mid-Albian times under the extensional tectonic regime. On the west coast of the South Island, west of the Alpine Fault, outcrops of breccia (collectively known

N

s

Kyeburn •

Fig. 1. Locality map of the South

Puysegur Point

100 km

Island of New Zealand showing ( in black) the main outcrop areas of mid-Cretaceous non-marine coarse clastic deposits. B, Beebys Conglomerate.

199

Cretaceous rift deposits of New Zealand as the Hawks Crag Breccia, forming part of the Pororari Group) are scattered but widespread over a NE-SW extent of approximately 320 km (Fig. 1). The breccia is matrix- or clast-supported with a sparse sandy matrix, and is often reddish-brown in colour. Structure is generally simple, most suc­ cessions dipping uniformly at angles up to 50°, or folded into simple, open synclines. Thickness of the Pororari Group varies considerably, but a maximum is c. 4 500m. The age of the group has been determined solely from pollen. Palynological studies (Raine, 1984) indicate that the bulk of the Pororari Group lies within the Lycopodiacides bullerensis assemblage, with the youngest rocks reaching into the lower part of the overlying Trichotomosulcites subgranulatus assemblage. Absence of angiosperms suggests that the Pororari Group lies within the New Zealand Motuan Stage (mid- to late Albian). The original upward time range of the Group in the region is unknown, being constrained only by the local occur­ rence of unconformably overlying coal measures of Maastrichtian age. However, a possible correlative of the Pororari Group, lying c. lOOkm northeast of the Paparoa Range (Beebys Conglomerate; Fig. 1) has yielded a poorly constrained microfloral age in the range late Cenomanian to early Santonian (Johnston, 1990), indicating that Pororari Group deposition may have extended locally well into the Late Cretaceous. The origin and depositional setting of the Hawks Crag Breccia has been the object of considerable speculation in the past: explanations have included glacial or glaciofluvial action (McKay, 1883), talus or alluvial fan formation (Morgan & Bartrum, 1915), or the deposits of air-lubricated landslides (Korsch & Wellman, 1988). In most occurrences, these breccia units make up the entire exposure, and the mechanism of emplacement and geometry of the bodies has been difficult to interpret because of their apparently uniform nature over large areas. How­ ever, in the Paparoa Range in the northwest of the South Island, where the most extensive areas of outcrop occur, other lithofacies are also present, and these throw light on sedimentary processes and environments of deposition of both the Hawks Crag Breccia and of the Pororari Group as a whole. This study examines the sedimentology of the Pororari Group in the southern Paparoa Range, where the breccia lithofacies is associated with fine-grained deposits (Fig. 2).

G E OL O G IC A L S E T TI N G

The Pororari Group comprises an unconformity­ bounded sequence resting on Early Palaeozoic turbidites of the Greenland Group, or on mid­ Palaeozoic and Early Cretaceous granitoids. It is overlain by the Paparoa Coal Measures of Maastrichtian age. In the southern Paparoa Range, the Pororari Group crops out in two extensive areas on the eastern and western flanks of the range, and in small exposures in Fox River to the north (Fig. 2). In the western outcrops, four lithostratigraphic units can be recognized (Laird, 1988; Figs 2 & 3): (1) a locally occurring basal unit, dominated by inter­ bedded matrix- and clast-supported breccia, massive or normally graded sandstone, and laminated mud­ stone (Watson Formation), followed gradationally by (2), thick-bedded matrix- and clast-supported breccia (Hawks Crag Breccia), followed by (3), inter­ bedded thin-bedded sandstone and matrix-supported breccia and conglomerate beds (Bovis Formation), which are in turn overlain by (4), laminated carbon­ aceous mudstone (Bullock Formation). The eastern­ outcrops lack the basal unit, and Hawks Crag Breccia rests directly on basement rocks. The landscape is extremely rugged, and covered with dense temperate rainforest (Fig. 4). This limits the accessible and usable exposure to the gorges and tributaries of the main streams, which have provided the measured stratigraphical sections (Fig. 3).

D E S C RIPT I O N OF

A N D I N T E R PR E T A TI O N

FACIES

A S S E MBLA G E S

Three facies assemblages are recognized within the southern Paparoa Range outcrops: (1) matrix- and clast-supported breccia; (2) pebbly sandstone, and massive or graded sandstone; and (3) laminated dark mudstone with thin graded sandstones. Assemblage 1: matrix- and clast-supported breccia

Facies descriptions This facies assemblage makes up the entire Hawks Crag Breccia, and also occurs within the Watson Formation, both as the basal 165m exposed in Tin­ dale Creek and as a 100-m-thick unit near the base of the formation in Pororari River (Fig. 3). Breccia units increase in frequency near the top of the Watson Formation, which passes gradationally

200

M.G. Laird

LJ Younger strata

;

N

t

Unconformity

Pororari Group Bullock Formation Bovis Formation Hawks Crag Breccia Watson Formation

Unconformity

0 Basement _.- Fault

0

2

3

4

5 km

Fig. 2. Map of part of the southern Paparoa Range showing Pororari Group outcrop and palaeocurrent plots. See Fig. 12

for key to the palaeocurrent plots.

upwards into the Hawks Crag Breccia. Maximum thickness of the facies assemblage is hard to estimate, but in the Bullock Creek area in the north, where it is thickest, it is likely to reach 1500 m.

The assemblage consists dominantly of poorly sorted matrix-supported breccia consisting of sub­ angular to subrounded clasts (often outsize) of base­ ment material in a coarse sandy matrix (Fig. 5).

Cretaceous rift deposits of New Zealand N

PAPAROA COAL MEASURES

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201

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Fig. 3. Stratigraphical columns showing relationships between the major formations and facies assemblages of the Pororari

Group.

Fig. 4. Hogback and cliffed

topography developed on Hawks Crag Breccia (Pororari Group), with steep east-facing dip slopes. The prominently bedded unit in the upper middle of the photograph is 20-30m thick. The valley of the upper Pororari River is in the foreground: crest of the Paparoa Range is in the background.

202

M.G. Laird

Fig. 5. Thick matrix-supported (upper) and clast-supported (lower) breccia beds, Fox River mouth. Beds young towards the left. Hammer for scale . Photograph courtesy of L. Homer.

Fig. 6. Succession of thin clast­ supported breccia beds, Fox River mouth. Note normal grading in the bed below the geological hammer. Some of these beds may represent the deposits of highly sediment­ charged sheet flows. Beds young upwards. Hammer for scale.

Clast-supported breccia makes up a minority of the beds (Fig. 5). Breccia beds most commonly lie in the range 1 -6 m thick, but some breccia units (probably composite) may reach 20-30 m in thickness and extend with little thickness change for at least several hundred metres (Fig. 4). Most lack internal stratifi­ cation and bedding is in many cases indistinct. Mud­ stone is absent, and thin sandstone beds up to 20 em thick occur only rarely. In rare instances clast­ supported breccia forms well-defined beds, usually

less than 1m thick but reaching a maximum of 3 m in thickness, and separated by thin (up to 10 em) beds of poorly sorted coarse to pebbly sandstone (Fig. 6). Poorly developed symmetrical grading, reverse grading, or normal grading occurs locally (Figs 6 & 7), although this is uncommon. Individual clasts reach a maximum size of 3m, and in some instances show imbrication and preferred orientation of long axes. Carbonaceous material is common, particularly at

Cretaceous rift deposits of New Zealand

203

Fig. 7. Thick, matrix-supported

breccia bed, with thin sandstone bed at base, Pororari River. Note lack of internal structure, but the presence of symmetrical grading. Beds young to left. Hammer for scale.

the interfaces between beds, and small coalified logs occur rarely. At the mouth of Fox River several thick beds of poorly sorted breccia show crude fining-upward grading into thin graded sandstone (in some instances showing synsedimentary defor­ mation structures) and then into coaly material. No root horizons have been recognized, and all plant material appears to have been transported from its source. Shallow scours or channels, up to 1m deep, occur locally. Other sedimentary structures, apart from rare synsedimentary deformation features, are notably absent. Interpretation The preponderance of poorly sorted, poorly strati­ fied matrix-supported breccia with common outsize clasts, a subordinate proportion of clast-supported breccia, the rarity of sedimentary structures within breccia beds, and the absence of cross-stratification and only minor presence of thin well-bedded sand­ stones, suggests that mass wasting events played a predominant role in the formation of assemblage 1. The major facies of matrix-supported breccia closely resembles the deposits of debris flows, which domi­ nate sedimentary processes on many modern alluvial fans, whereas the subordinate clast-supported breccia is characteristic of associated debris flow levees (cf. Pierson, 1980; Blair & McPherson, 1992). Thin, moderately well-sorted sandstone beds and thin-bedded breccia and sandstone couplets, forming a minor proportion of assemblage 1, closely resemble facies prominently represented in some modern and

Cenozoic alluvial fans (cf. Ballance, 1984; Blair, 1987), and attributed to the deposits of sheet flows emplaced during catastrophic floods. The breccia beds at Fox River mouth, which pass upwards into graded sandstone and then into coaly material, may represent the results of flash floods and the delayed settling out of fines, with transported vegetation being the last to settle. Assemblage 1 breccias are thus inferred to have been deposited on a series of alluvial fans, dominated by debris-flow depositional processes, closely similar to the humid-temperate climate, late Quaternary and modern alluvial fans that are characteristic of the mountainous regions of New Zealand (cf. Pierson, 1980, 1981; McArthur, 1987) and elsewhere (e.g. Kochel & Johnson, 1984). The abundance of debris flows, and the poor sorting and angularity of many of the clasts suggests that source slopes were steep and lay in close proximity. The common occurrence of carbonaceous material, including rare coalified logs, indicates that the source slopes were vegetated. No fine-grained subfacies potentially representa­ tive of interdistributary deposits (cf. Lewis & Ekdale, 1991) has been recognized within the assem­ blage, and fans were therefore likely to have been coalescing. A ssemblage 2: pebbly sandstone and massive or graded sandstone

Facies descriptions This facies assemblage makes up most of the Watson

204

M.G. Laird

and the Bovis Formations. It consists dominantly of matrix-supported conglomerate and pebbly sand-. stone, graded or massive very fine-grained to coarse­ grained sandstone, and alternating sandstone and mudstone successions. Thickness is variable, but an aggregate thickness of 600 m is reached in the Watson Formation exposed in Pororari River, and a maximum thickness for the Bovis Formation is approximately 400 m (in Slaty Creek}. Clasts in the conglomerate and pebbly sandstone of assemblage 2 are notably better rounded than in the breccia of assemblage 1, and all beds are matrix-supported. Conglomerate and pebbly sandstone beds, which are up to 2 m thick, lack internal sedimentary struc-

tures, and most are ungraded. Clasts commonly occupy specific horizons within each bed, but random distribution also occurs. Clasts are of basement gran­ ite or Palaeozoic sandstone, and range up to 0.5 m in size. Imbrication is relatively common (Fig. 8). Inter­ leaved very fine- to coarse-grained, moderately well­ sorted sandstones vary in thickness up to 1 m, and include both graded and ungraded beds (Fig. 9). Some have scoured bases, and show sole marks consisting mainly of grooves and less common prod marks or flutes. Primary current lineations and cur­ rent ripples are also present on bedding planes and parallel-lamination is common in the finer sandstone beds. Isolated horizons of current-rippled sandstone

Fig. 8. Imbricated pebbles of

granite and of Greenland Group sediments supported by a matrix of very coarse to gritty sandstone, Bovis Formation, Slaty Creek (at approximately 500m, Slaty Creek column, Fig. 3). Palaeocurrent direction from right to left (N-S). Beds young towards top right. Note the thin interbeds of carbonaceous mudstone. Hammer for scale.

Fig. 9. Amalgamated beds of fine

sandstone showing no internal stratification or grading, in the lower portion of the Watson Formation, Pororari River (at 120m in column, Fig. 12). Beds young towards top. Pocket knife is 9cm long.

Cretaceous rift deposits of New Zealand are common in the intervening mudstones (Fig. 10). Load casts and flame structures are common at the interface between sandstone and mudstone beds (Fig. 11), suggesting rapid deposition, and some thin sandstone layers are intersected by small low-angle synsedimentary thrust faults. The thrusting is in the same sense as the overturning of flame structures and of other palaeocurrent indicators. Widely sep­ arated burrowed horizons also occur. Some variation within the facies assemblage is evident between the Watson and Bovis Formations. The Watson Formation is characterized by inter­ leaved components of facies assemblages 1 and 2, although assemblage 2 is dominant (Fig. 12). The

Fig. 10. Mudstones of the Watson Formation containing thin massive or laminated beds of fine-grained sandstone, and horizons of isolated ripples (same locality as Fig. 9). Beds young upwards. Pocket knife is 9cm long.

Fig. 11. Load casts and flame structures in very fine-grained sandstone beds, lower portion of the Watson Formation, Pororari River (at 110m in column, Fig. 12). Palaeocurrents flowed from left to right (N-S). Beds young towards top. Pocket knife is 9cm long.

205

formation is particularly well-exposed in Pororari River and south-flowing tributaries, although its basal contact is faulted. In the adjacent Tindale Creek, the basal unit consists of clast-supported breccia consisting entirely of Palaeozoic sandstone clasts resting uncomformably on Palaeozoic strata (Fig. 3). This basal unit, and a petrographically similar 100-m-thick unit of clast-supported-breccia occurring about 300 m above the basal fault in Pororari River, are considered to represent an inter­ calation of Assemblage 1. The remainder of the formation is dominated by alternating sandstone and mudstone, including common slumped and hydroplastically deformed beds (Fig. 13). These beds

206

M.G. Laird

KEY v Sole mark

Slump u Channel

fj, 17 V

Fining-upward units Coarsening-upward units

Sequence

Palaeocurrent Plots:

• Sole mark W Parting lineation • Slumping direction [] Cross lamination --+

Imbrication direction

Lithology and Facies Assemblage

Palaeocurrents

Palaeoenvironment

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N

Alternating thin-bedded sandstone and mudstone, organized into coarsening­ upward sequences (A2)

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Fig. 12. Detailed stratigraphical column of the Watson Formation, Pororari River, showing main sedimentary characteristics, palaeocurrent plots, and inferred palaeoenvironments. Al, A2, facies assemblages 1 and 2.

Cretaceous rift deposits of New Zealand

207

Fig. 13. Slumped strata near the

base of the Watson Formation, Pororari River (at approximately 20m in column, Fig. 12). Beds young towards top of photograph. Pocket knife is 9em long.

commonly alternate with packets of thick-bedded, graded or massive sandstone. Some 460 m above the faulted base of the section, 6 m of pebbly mudstone occurs, containing indurated sandstone pebbles up to 20 em, and large rafts of thinly bedded sandstone up to 140 cm long, some showing evidence of syn­ sedimentary deformation. This unit is interpreted to be a mass flow deposit. Above this prominent slump horizon, a series of thickening and coarsening upward units occurs, varying in thickness from 15 to 110m. The base of each unit consists of laminated mudstone, passing upwards into alternating mud­ stone and thin sandstone beds, and finally up into thick, often amalgamated sandstone beds up to 1 m

Fig. 14. Thick, massive fine-grained

sandstone beds (upper one with a scoured base) in a fining upwards succession, middle portion of the Watson Formation, Pororari River (at 650m in column, Fig. 12). Note pseudonodule horizon below the lower sandstone bed. Beds young towards top of photograph). Pocket knife is 9em long.

thick (Fig. 14). Small scours or channels up to 0. 5 m deep, infilled with fine-grained sandstone, in some instances slumped, occur at rare intervals in the succession (Fig. 15). A change in lithofacies occurs at about 800 m above the base of the Pororari River section. At this horizon, 5 m of cross-bedded sandstone occurs, the cross-bedded sets averaging 10-20 em in thickness. Some sets have suffered hydroplastic deformation. At about the same level, thin beds of matrix- or clast-supported graded breccia, the clasts consisting of granite fragments, start to appear, commonly with erosive bases (Fig. 16). Units of alternating sandstone and mudstone, although still common,

208

M.G. Laird

Fig. 15. Channel cut in mudstone and infilled with slumped sandstone, Watson Formation, Pororari River (at 440m in column, Fig. 12). Beds young towards top of photograph. Hammer for scale.

Fig. 16. Thin graded breccia beds in the upper part of the Watson Formation (at approximately 820m in column, Fig. 12). Note slumped laminated mudstones between the breccia beds. Beds young to the left of the photograph.

become less frequent upwards. The top 10-20 m of the Watson Formation is dominated by matrix­ supported conglomerate, carbonaceous sandstone, and thin coaly layers in a series of fining upward units 1. 5-3 m thick. Each unit comprises a basal conglomerate consisting of moderately well-rounded pebbles of basement rocks, predominantly granite, up to 15 em in diameter, commonly resting on a scoured surface, passing upwards into carbonaceous parallel- or cross-bedded very coarse sandstone, and then into thin beds of carbonaceous siltstone or coal seams up to 20 em thick (Fig. 17). The fining upwards units pass with rapid gradation upwards into thick

massive matrix-supported breccia beds compnsmg the basal beds of the Hawks Crag Breccia. In the Bovis Formation, assemblage 2 forms part of one systematically fining upwards succession, with only minor intercalations of other facies associations compared with the Watson Formation. There is an upward gradation from matrix-supported breccia of assemblage 1 into metre-bedded pebbly coarse sand­ stone of assemblage 2, containing scattered sub­ angular and subrounded clasts of basement rocks up to 1m (Figs 8 & 9), with minor interbeds of decimet:re to metre-bedded medium- to coarse-grained massive sandstone. The proportion of conglomerate to sand-

Cretaceous rift deposits of New Zealand

209

Poorly-stratified matrix-supported conglomerate, scoured at base Cross-bedded carbonaceous sandstone

Poorly-stratified matrix-supported conglomerate, scoured at base

Carbonaceous siltstone(lenses)

Poorly-stratified matrix-supported conglomerate, scoured at base, and fining upwards

Carbonaceous sandstone Coal Carbonaceous sandstone

Poorly-stratified matrix-supported conglomerate, scoured at base

Fig. 17. Sketch of conglomeratic

fining upwards units at the top of the Watson Formation, Pororari River (940m in column, Fig. 12).

stone beds decreases upwards, and in its upper portions the succession consists almost entirely of massive or normally graded sandstone beds separ­ ated by thin layers of carbonaceous mudstone (Fig. 18). Some of the massive sandstone beds contain pebble trains or scattered clasts, the clast size decreasing upwards in the succession. The top few tens of metres of the Bovis Formation consists of centimetre to decimetre graded very fine to fine­ grained sandstone alternating with mudstone. Initially thin mudstone interbeds between sand­ stones thicken upwards until near the top of the unit, the predominant lithofacies is alternating mud­ stone and fine-grained graded or massive sandstone. Flame-structures are commonly developed at the base of sandstone beds, and show a uniform direction of overturning. Sole structures, mainly grooves, but a few flutes, also occur. Burrows occur locally.

Poorly-stratified sandstone

Interpretation The high proportion of mudstone throughout assem­ blage 2 in the Watson Formation and in the upper portions of the Bovis Formation, preponderance of sediment gravity flow deposits and scarcity of traction current deposits, and the occurrence of rare bur­ rowed horizons suggests sedimentation in a standing water body. The highly carbonaceous nature of the deposits and the absence of shelly fossils suggests that this is likely to have been a lake. The very poor sorting of the matrix-supported breccia beds, the presence of megaclasts in some beds, and absence of sedimentary structures suggests deposition of these beds by subaqueous debris flows. Common loads and flame structures at the interface between sand­ stone and mudstone beds suggest rapid deposition of the coarse sediments, and the increasingly common

210

M.G. Laird

Fig. 18. Thin, massive to poorly

graded fine- to medium-grained sandstone beds with thin partings of carbonaceous mudstone, near the top of the Bovis Formation, Slaty Creek. Beds young towards top of photograph. Hammer for scale.

presence of normally graded bedded sandstones upwards in the Bovis Formation suggests deposition by turbidity currents. The common occurrence of slump deposits in the Watson Formation suggests the presence of a significant slope throughout much of its depositional history. The thin fining upwards successions near the top of the formation in Watson Creek are inferred to have resulted from a high-.energy braided stream system developed at the toe of a Hawks Crag alluvial fan, which prograded over a vegetated coastal flat between the advancing alluvial fan and the standing water body. Miall's (1978) 'Scott-type' braided stream model, considered typical of proximal braided stream deposits on alluvial fans, is prob­ ably a close analogue. The Scott River was also represented as typically containing minor interbeds of cross-bedded sand and small fining upwards cycles. A similar environment has been inferred by Hamblin (1992) for parts of the Lower Carboniferous Horton Group of Nova Scotia.

Slaty Creek. Near the base, assemblage 3 contains mudstone interbedded with decimetre-bedded fine­ to very fine-grained, parallel laminated sandstone. Sandstone beds are commonly normally graded, and have sharp bases and tops. This facies passes rapidly up into finely laminated, dark, carbonaceous mudstone (Fig. 19). Sole marks, consisting mainly of grooves, but also with rare prod and flute casts, occur on the base of some sandstone beds. Isolated vertical burrows occur at some horizons. Rare slumped units up to 1. 5 m thick also are present near the base of the assemblage. In Slaty Creek, a horizon of coal 0. 5 m thick, bounded above and below by dark mudstone with prominent vertical burrows, occurs a few metres above the gradational contact between assemblages 2 and 3. The coal could not be traced more than 5 m laterally and is almost certainly lenticular. No root horizons are present, and the coal is inferred to represent allochthonous material. Interpretation

Assemblage 3: laminated mudstone

Facies descriptions This facies assemblage, which makes up the bulk of the Bullock Formation, consists dominantly of dark brown to grey, highly carbonaceous laminated mud­ stone. It passes upwards with rapid gradation from alternating thin sandstones and mudstones of assem­ blage 2, attaining a thickness in excess of 500 m in

The ubiquitous presence of highly carbonaceous mudstone, absence of traction current deposits, and absence of marine shelly fossils but presence of burrowed horizons, suggests deposition in a standing body of water, inferred to be a lake. The rare occurrence of sediment gravity flow deposits and slumps suggests that deposition occurred distal to the sediment distributary system, and probably well out on the floor of the lake basin.

Cretaceous rift deposits of New Zealand

211

Fig. 19. Laminated carbonaceous

mudstone, Bullock Formation, Slaty Creek. Beds young towards top of photograph. Hammer for scale.

D EPO SIT IO N A L

SYS T E M

O F THE PO R O R A R I G R O U P

The three facies assemblages intertongue, and are obviously closely related portions of a common depositional system. The Watson, Hawks Crag Breccia, and Bovis Formations of the western Paparoa Range, and their associated facies assem­ blages also show rapid systematic thinning from north to south (Fig. 3). This, together with the nature of the interrelationships between the facies assemblages supports the interpretation of the Pororari succession as a lacustrine fan-delta system. The alluvial fan deposits of assemblage 1 are inferred to have fed either directly , or via a narrow braid­ plain, into a lake, where deposition continued on a fan consisting of assemblage 2 proximal deposits of the Watson and Bovis Formations. Lacustrine, mainly quiet-water conditions, are represented by the assemblage 3 mudstones of the Bullock Forma­ tion. A simplified reconstruction showing the inferred relationship of assemblages to each other, and the evolution of the fan delta system, is given in Fig. 20. The most informative profile is that given by the Pororari River and its tributaries (Fig. 12). Here there is clear evidence of a fluctuating lake shoreline, with irregular progradation of the subaerial portion of an alluvial fan into the lake, and subsequent periods of transgression. The presence of 165m of basal breccia in Tindale Creek suggests that depo­ sition started with an alluvial fan. Subsequent

inundation by a relative rise in lake level and depo­ sition of proximal lacustrine sediments is inferred to account for association 2, with a brief re-establish­ ment of an alluvial fan represented by association 1 deposits from 32 5 to 425 m in the Pororari River section (Fig. 12). The lacustrine deposits of the Watson Formation were increasingly encroached on and finally overwhelmed by the alluvial fans making up the Hawks Crag Breccia, which represented the major alluvial fan development of the Pororari Group. The end of Hawks Crag Breccia times shows increasing incursion and swamping of the fan by first proximal and then distal lacustrine associations in all areas of outcrop. The thickening and coarsening upward successions in the middle part of the Watson Formation may have been autocyclic, i.e. represented periodic diversion of the main feeder channels of a lacustrine delta and the build-up of a new pro-delta lobe, with accommodation provided by continued subsidence of the basin. However, the repetitive nature of the successions, and the evidence for major changes in base level outlined above, suggest that a more likely cause was fluctuations in lake level. The mudstone at the base of each cycle is inferred to represent a flooding event caused by a rise in lake level, the subsequent development of the thickening upward succession caused by lobes of sediment prograding into the lake during relative high-stand. The fluctu­ ations in lake level could have been brought about by a variety of mechanisms, including tectonic activity on a fault or faults controlling the depo-

212

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making up the Pororari Group. a-a' shows the approximate location of the Pororari River measured section (Fig. 3).

sitional basin, or by climatic cycles. Cyclic suc­ cessions present in the dominantly lacustrine early Mesozoic Newark Supergroup of eastern North America have been attributed to climate-controlled high-frequency fluctuations in the depth of rift lakes (e.g. Olsen, 1986, 1991).

P A L A E O C U R R E N T PAT T E R N

Palaeocurrents (summarized in Fig. 2 ) were deter­ mined from a wide variety of sedimentary features, including grooves, prods, flutes, primary current lineation, channel margins, and cross-lamination, and palaeoslopes from slump fold axes in the lacus­ trine facies associations 2 and 3. Palaeocurrent indicators are relatively common in the Watson Formation, and in the Bovis and Bullock Formations of the western Paparoa Range. Although no suitable bedding surfaces were exposed for measurement of palaeocurrent directions in the Bovis Formation of the eastern Paparoa Range, numerous overturned flame structures, and local occurrence of imbrication in pebbly beds indicate a consistent general direction

of transport. Only a small number of directional features were recorded in the Bullock Formation in the eastern Paparoa Range. The palaeocurrent and palaeoslope trend in the Watson Formation is generally from north to south. However, in the thick, well-exposed succession in the Pororari River and its tributaries, plotting palaeocurrents in segments according to facies groups clearly shows that significant variations occur within the succession (Fig. 12). The lower, poorly exposed 300 m, consisting of alternating sandstones and mudstones of association 2, has palaeocurrents directed towards the south or south-southwest. By contrast, the similar facies higher in the column between 430 and 800 m in the section has palaeocur­ rents directed towards the south-southeast. The sole imbrication measurement recorded in the inter­ vening 100-m unit of assemblage 1 breccia (between 330 and 430 m, section Fig. 12), although statistically insignificant, is also directed to the south-southeast, suggesting a source in common with the overlying sediments, but differing from the source providing the earlier sediments. The 100-m thick breccia unit is thus inferred to represent an abrupt incursion of

213

Cretaceous rift deposits of New Zealand an alluvial fan into the lake. The monolithological nature of clasts, which consist of indurated basement sandstone, and their high degree of angularity, suggests a very proximal source. The same fai'!_. prob­ ably continued to feed clastic detritus in 'a 0uth­ southeasterly direction into the lake up until:a�:'reast the 800 m mark (Fig. 12), an inference supportetl by the fact that the clasts in the 6-m bed of slumped pebbly mudstone 20m above the breccia (at 460 m, Fig. 12) also consist largely of indurated basement sandstone. The upper part of the Watson Formation, between 800 and 950 m in Fig. 12, although dominated by assemblage 2 facies, includes thin beds of breccia and conglomerate containing granitic clasts, indi­ cating a different provenance. This mixture of sources is reflected by the wide variation in current directions recorded, ranging through southeast to southwest. Southwest and west-southwest directions are indicated by clast imbrication on two surfaces in the overlying Hawks Crag Breccia (Fig. 12), and a south-southwest direction of transport is recorded overall for the Hawks Crag Breccia in the study area (Fig. 2). Palaeocurrent data were not available from the Watson Formation in the adjacent Tindale Creek, the only other area of major outcrop, thus these changes of palaeocurrent trend could not be confirmed as a regional phenomenon. However, the inference is that at least three distinct sources and fans influenced sedimentation: the earliest from the north or north-northeast; the next from a Palaeozoic area to the north-northwest, and the upper part of the Watson Formation is increasingly influenced by debris derived from a granitic source to the north­ east. Derivation of indurated sandstone clasts in the second fan from the north-northwest is consistent with the presence of Early Palaeozoic turbidites of the Greenland Group cropping out to the west of the Pororari Group. However, basement rocks cur­ rently lying to the northeast of the Hawks Crag breccia and extending for many kilometres consist entirely of foliated gneiss of the Charleston Meta­ morphic Group (Nathan, 1978; Laird, 1988), a lith­ ology almost absent from the Hawks Crag Breccia, which is dominated by granitic clasts. This problem of derivation is discussed below under Tectonic Implications. Clast imbrication is the only usable palaeocurrent indicator in the alluvial fan assemblage represented by the Hawks Crag Breccia, which is dominated by granitic clasts. Suitable surfaces for recording imbri-

s

cation were relatively uncommon however, and only a few measurements were obtained. Although there is some variation, a dominant NNE-SSW, or NE-SW trend is apparent. In the Bovis Formation, palaeocurrents are widely variable (Fig. 2), but still show a generally southerly trend. The information from sole marks and other features is strongly reinforced by the direction of overturning of flame structures, which occur at the base of many sandstone beds. Sole marks from the Bullock Formation show a closely constrained and consistent trend from north-northeast to south­ southwest (Fig. 2). The consistent thinning of the Watson Formation, Hawks Crag Breccia, and Bovis Formations from north to south in the western Paparoa Range also supports the inference of a transport direction in a generally N-S direction.

T E C T O NIC

1M PLIC A T I O N S

I t has been widely recognized that fan-deltas are preferentially located in areas of active tectonism, and are commonly controlled by an active fault or faults (e.g. Ethridge & Wescott, 1984). The great thickness of the clastic wedge (a minimum of 1200 m for the Hawks Crag Breccia alone), and the relative angularity of clasts, suggests that they were gener­ ated by an active fault system in close proximity. The evidence suggests that the fan-delta system occu­ pied a half-graben, downthrown to the south (Fig. 20). The facies relationships, with thick coarse sub­ aerial deposits concentrated in the north, and rapid thinning and fining to the south, i.e. downcurrent, also closely fits the model proposed by Surlyk (1978) for fan-delta sedimentation in a half-graben. Move­ ment on the inferred bounding fault to the north continued during much of the deposition of the Pororari Group, and was responsible not only for generating the coarse facies of the Hawks Crag Breccia and the associated lacustrine sediment grav­ ity flows, but also for the necessary subsidence to provide accommodation for the great thickness of the deposits. Contemporaneous fault movement may also have been a factor in causing the fluctuations in lake level noted earlier, although climate also may have played a role. Judging from the consistently s-ssw directed sediment gravity flows, both sub­ aerial and subaqueous, the active fault scarp was likely to have been oriented in a westerly or west-

214

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northwest direction with an implied N-NNE direc­ tion of tectonic extension (Fig. 21). As noted previously, the similarity of Lithofacies, stratigraphical position, and palaeocurrent pattern of the Hawks Crag Breccia and, in particular, the overlying lake beds in the now widely separated outcrops on the eastern and western flanks of the Paparoa Range suggests that they were formed in the same depositional basin. However, the thick­ nesses of the Hawks Crag Breccia and the overlying

Fig. 21. Map showing generalized

palaeocurrent trends and inferred alignment of the mid-Cretaceous fault bounding the half-graben containing the Pororari Group.

lacustrine strata in the outcrop area on the eastern flank of the Paparoa Range, which lies south­ southeast of that on the western flank, i.e. downcur­ rent, do not conform with the southward thinning pattern of units clearly shown by the western out­ crops. The thickness of the Hawks Crag Breccia of the eastern outcrop is similar to the maximum thick­ ness in the north of the western unit, but the over­ lying strata of association 2 are more than twice the thickness of the western maximum. This is hard to

Cretaceous rift deposits of New Zealand explain by invoking a change in sedimentation pat­ tern, as the palaeocurrent pattern is essentially indis­ tinguishable between the two outcrop areas. A more likely explanation is that transcurrent movement along a north-northeast trending fault (Hawera Fault?) subsequent to deposition has offset the eastern part of the half graben to the south by at least 10-15 km. It was noted previously that the clasts making up the Hawks Crag Breccia consist almost entirely of granite , inferred to have been derived from a proxi­ mal source to the northeast. The basement rocks to the northeast and north, however, consist for many kilometres of foliated gneiss of the Charleston Meta­ morphic Group, a lithology that occurs only rarely in the Hawks Crag Breccia. An added apparent complexity is the evidence provided by Tulloch & Palmer (1990) that granite samples typical of the bulk of the Hawks Crag Breccia at the Fox River mouth and from several large exposures in Bullock Creek are similar to the Buckland Pluton, the closest outcrop of which lies 20 km to the northeast. The clasts are distinct from the Meybille Granite , which crops out to the immediate south of Fox River mouth. The inference drawn by Tulloch & Palmer (1990) is that, at the time of their deposition, either the Fox River and Bullock Creek deposits were situated 20-2 5 km to the northeast, or the eroded upper parts of the Buckland Pluton had a consider­ ably greater horizontal extent. Tulloch & Kimbrough (1989) inferred that the Greenland and Pororari Groups, together with the granite from which the latter had b'een derived, were part of an upper plate that detached from a metamorphic core complex (Charleston Metamorphic Group) which comprised a lower plate. On this hypothesis, dislocation of the Hawks Crag Breccia from its original source is possible. A possible analogue of the southern Paparoa Range half-graben has been recognized in seismic profiles offshore from Greymouth, 30 km to the southwest (Bishop, 1992; Fig. 21). This half-graben is also aligned WNW-ESE, and is inferred to be infilled largely with Pororari Group sediments. It is clearly bounded to the north by a southward-dipping normal fault, and to the south by a pinchout. It is divided by major north-northeast striking transfer faults into several segments. The orientation of the half-graben and bounding normal fault clearly sup­ ports the direction of tectonic extension suggested for the southern Paparoa Range half-graben. The inferred E-W to SE-NW oriented half-graben

215

infilled by the Beebys Conglomerate (Johnston , 1990; Fig. 1) is also likely to have been formed during the same period of north-northeast extension. These data fit well with the north-northeast exten­ sion direction for the Paparoa Range area during late Aptian to early Albian times deduced from structural research on the basement noted above. Tulloch & Kimbrough (1989) provided evidence that stretching lineations associated with low-angle normal detachment faults separating the upper from the lower plate defined a regionally consistent north­ northeast extensional trend, which coincides with that inferred from the half-graben. It has been suggested (Laird, 1993, 1994) that half-graben formation and synchronous infilling were associated with a New-Zealand-wide early period of extension beginning in the mid-Albian, pre-dating by approximately 2 5 Ma the separation of New Zealand from Gondwana and sea-floor spreading in the Tasman Sea.

CONCLUSIONS 1 The Pororari Group of the southern Paparoa Range represents a fault-controlled lacustrine fan­ delta complex. The Hawks Crag Breccia and equiv­ alent matrix- and clast-supported breccias represent deposits of humid-temperate climate alluvial fans, while intertonguing and overlying finer grained sedi­ ments were deposited predominantly in a lacustrine environment. 2 Mass-flow processes dominated both alluvial fan and lacustrine sedimentation, suggesting that tec­ tonic activity continued throughout most of the period of deposition of the Pororari Group, dying away only towards the end. The abundance of debris­ flow deposits in the alluvial fan association, and the poor-sorting and angularity of many of the clasts suggests that source slopes were steep and lay in close proximity. The ubiquitous occurrence of debris-flow deposits, turbidites and slumps in the proximal lacustrine succession indicates the presence of a significant slope on the subaqueous fan, which may account for the limited presence of deposits transitional between the alluvial and sublacustrine fans. 3 Fluctuations in lake level are indicated by episodic transgressions and regressions of the lake with res­ pect to the adjacent alluvial-fan system, and by the presence of repetitive thickening and coarsening upwards successions in the proximal lacustrine

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deposits of the Watson Formation. These fluctuations may have been caused by active faulting at the basin margin, by climatic changes, or by a combination of both. 4 The fan-delta system is inferred to have occupied an actively subsiding half-graben. The consistently south-southwest directed sediment gravity flows and rapid southward thinning of the sedimentary wedge suggests that the active basin margin fault was likely to have been oriented in a WNW-ESE direction. The implied NNE-SSW direction of tectonic exten­ sion compares well with a similar extension direction determined for this area from structural research on basement rocks, and with that of a probable mid­ Cretaceous half-graben recognized on offshore seis­ mic profiles. The orientation of the half-grabens is essentially parallel to the future spreading axis associ­ ated with the opening of the adjacent Tasman Sea. This opening occurred at about 80 Ma, suggesting that initial rifting and half-graben formation occurred 20-25 Ma before sea-floor spreading began.

ACKNOWLEDGEMENTS

I am grateful to D . Lewis, J . Lindqvist and A. Tulloch for stimulating discussions, both in the field and in the office, which helped to clarify my ideas on processes of emplacement and geological setting of the Pororari Group. Thorough reviews by T. Astin, S. Flint, D. Lewis, J. Lindqvist and G. Plint greatly improved the manuscript and are appreciated.

REFERENCES

P.F. (1984) Sheet-flow-dominated gravel fans of the non-marine middle Cenozoic Simmler Formation, central California. Sediment. Geol. , 38, 337-359. BISHOP, D.J. (1992) Extensional tectonism and magmatism during the middle Cretaceous to Paleocene, North Westland, New Zealand. N.Z. J. Geo/. Geophys. 35, 81-91. BISHOP, D.G. & LAIRD, M.G. (1976) Stratigraphy and depositional environment of the Kyeburn Formation (Cretaceous), a wedge of coarse terrestrial sediments in Central Otago. J. R. Soc. N.Z. , 6, 55-7 1. BLAIR, T.C. (1987) Sedimentary processes, vertical stratifi­ cation sequences, and geomorphology of the Roaring River alluvial fan, Rocky Mountain National Park, Colorado. J. sediment. Petrol. 57, 1 - 18. BLAIR, T.C. & McPHERSON, J.G. (1992) The Trollheim alluvial fan and facies model revisited. Geol. Soc. Am. Bull. , 104, 762-769. BRADSHAW, J.D. (1989) Cretaceous geotectonic patterns in the New Zealand region. Tectonics, 8, 803-820.

BALLANCE,

F.G. & WESCOTT, .W.A. (1984) Tectonic setting, recognition and hydrocarbon reservoir potential of fan­ delta deposits. In: Sedimentology of Gravels and Con·· glomerates (Eds Koster, E.H. & Steel, R. J . ), Mem. Can. Soc. petrol. Geol., Calgary, 10, 217-235. HAMBLIN, A.P. (1992) Half-graben lacustrine sedimentary rocks of the Lower Carboniferous Strathlorne Formation, Horton Group, Cape Breton Island, Nova Scotia, Canada. Sedimentology, 39, 263-284. JOHNSTON , M.R. (1990) Geology of the St Arnaud District, Southeast Nelson (Sheet N29). New Zealand Geological Survey Bulletin 99. New Zealand Geological Survey, Lower Hutt, New Zealand, 119 pp. KocHEL, R.C. & JOHNSON, R.A. (1984) Geomorphology and sedimentology of humid-temperate alluvial fans, central Virginia. In: Sedimentology of Gravels and Con­ glomerates (Eds Koster, E . H. & Steel, R.J.). Mem. Can. Soc. petrol. Geol., Calgary, 10, 109-122. KoRSCH , R.J. & WELLMAN, H.W. (1988) The geological! evolution of New Zealand and the New Zealand region. In: The Ocean Basins and Margins, Vol. 7B (Eds Nairn, A.E.M., Stehli, F.G. & Uyeda, S. ), pp. 411-482. Plenum, New York. LAIRD, M.G. (1988) Sheet S37 Punakaiki. Geological map of New Zealand 1 : 63 360. New Zealand Geologi­ cal Survey, Department of Scientific and Industrial Research, Wellington. LAIRD, M.G. (1993 ) Cretaceous continental rifts. New Zealand Region. In: Sedimentary Basins of the World. South Pacific Sedimentary Basins (Ed. Ballance, P.F. ), pp. 37-49. Elsevier, Amsterdam. LAIRD, M.G. (1994) Geological aspects of the opening of the Tasman Sea. In: The Evolution of the Tasman Sea Basin (Eds van der Lingen, G.J., Swanson, K. & Muir, R.J.), pp. 1-17. Balkema, Rotterdam. LEWIS, D.W. & EKDALE, A.A. (1991 ) Lithofacies relation­ ships in a late Quaternary gravel and loess fan delta complex, New Zealand. Palaeogeogr. , Palaeoclimatol. , Palaeoeco/. , 81, 229-251. LIN DQVIST, J.K. (1990) Puysegur Group: a mid Cretaceous lacustrine fan-delta complex, Balleny Basin, southwest Fiordland. Geol. Soc. N. Z. Misc. Pub/. , SOA, 83. McARTHUR, J.L. (1987) The characteristics classification, and origin of Late Pleistocene fan deposits in the Cass Basin, Canterbury, New Zealand. Sedimentology, 34, 459-471. McKAY, A. (1883) On the geology of the Reefton District, lnangahua County. Geol. Survey Reports during 1882 , Vol. 15, pp. 142-144. Wellington. MLALL, A.D. (1978) Lithofacies types and vertical profile models in braided river deposits: a summary. In: Fluvial Sedimentology (Ed. Miall, A.D.), Mem. Can. Soc. petrol. Geol., Calgary, 5, 597-604. MORGAN , P.G. & BARTRUM, J.A. (1915) The geological and mineral resources of Buller-Mokihinui Subdivision, Westport Division. Geological Survey Bulletin No. 1 7 (new series) . New Zealand Department of Mines, Geo­ logical Survey Branch, Wellington. NATHAN , S. (1978) Sheet S31 & Part S32 Buller- Lyell. Geological Map of New Zealand 1 : 63 360. New Zealand Geological Survey, Department of Scientific and Indus­ trial Research, Wellington. OLSEN, P.E. (1986) A 40-million-year lake record of early Mesozoic orbital climatic forcing. Science, 234, 842-848. ETHRIDGE,

Cretaceous rift deposits of New Zealand P.E. ( 1991) Tectonic, climatic, and biotic modu­ lation of lacustrine ecosystems - examples from Newark Supergroup of eastern North America. In: Lacustrine Basin Exploration. Case Studies and Modern Analogs (Ed. Katz, B.J.) , Mem. Am. Assoc. petrol. Geol. , Tulsa, 50, 209-224. PIERSO N , T.C. ( 1 980) Erosion and deposition by debris flows at Mt Thomas, north Canterbury, New Zealand. Earth Surf Process. , 5, 227-247. PIERSON, T.C. ( 1981) Debris flows. An important process in high country gully erosion. 1. Tussock Grasslands Mountain Lands Inst. , Rev. , 39, 3- 14. RAINE, J.l. ( 1984) Outline of a palynological zonation of Cretaceous to Paleogene terrestrial sediments in West Coast region, South Island, New Zealand. New Zealand OLSEN,

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Geological Survey Report 109 Department of Scientific and Industrial Research, New Zealand. SuRLYK, F. ( 1978) Submarine fan sedimentation along fault scarps on tilted faultblocks (Jurassic-Cretaceous bound­ ary, East Greenland). Gr¢n. geol. Unders. Bull. , 128, 108 pp. TULLOCH , A.J. & KIMBROUGH, D.L. ( 1989) The Paparoa Metamorphic Core Complex, New Zealand: Cretaceous extension associated with fragmentation of the Pacific margin of Gondwana. Tectonics, 8, 1217- 1234. TuLLOCH, A.J. & PALMER, K. (1990) Tectonic implications of granite cobbles from the mid-Cretaceous Pororari Group, southwest Nelson, New Zealand. N.Z. 1. Geol. Geophys. , 33, 205 -217.

Spec. Pubis int. Ass. Sediment. (1995)

22, 219-236

Sedimentation and tectonics of a synrift succession: Upper Jurassic alluvial fans and palaeokarst at the late Cimmerian unconformity, western Cameros Basin, northern Spain N IG E L H . P L A T T Geco-Prakla Schlumberger, Schlumberger House, Buckingham Gate, Gatwick Airport, West Sussex RH6 ONZ, UK

ABSTRACT

The Upper Jurassic (late Kimmeridgian to Berriasian ?) Senora de Brezales Formation of the western Cameros Basin, northern Spain, comprises a laterally variable succession of continental conglomerates, sandstones and pedogenetic carbonates deposited in wadi-type channels, alluvial fans, sandflats and palaeosols in a semi-arid environment. This succession rests on a complex unconformity surface that developed in response to relative falls in sea-level and strong Late Jurassic extensional faulting. Footwall uplift resulted in truncation of underlying strata at the crestal axes of fault blocks. The clastic rocks of the Senora de Brezales Formation represent the erosion products of the Jurassic anq older strata. Lateral facies variations reflect the changing lithology beneath the unconformity surface, whereas strong lateral thickness variations record fault control on sedimentation. In areas of limited erosion, karst surfaces developed on subaerially exposed lower Kimmeridgian limestones. Where these were remo ved, erosion of upper Oxfordian marginal marine sandstones led to the deposition of red continental sandstones. In other areas, erosion led to the incision of channels into a pediment of Middle Jurassic carbonates. The channels were filled with conglomerates largely deri ved from erosion of the Jurassic marine limestones. The location of channels was strongly influenced by NE-SW faults. Laminar and nodular calcretes formed directly on the unconformity surface, in interchannel areas, and within the o verlying clastic succession. The Senora de Brezales Formation is an outcrop analogue for Upper Jurassic clastic successions present in the subsurface of basins to the west of Britain on the northern Biscay margin . Field studies highlight the complex lateral facies variations within this synrift succession and underline the control of seismic-scale and subseismic-scale faults on erosion and sedimentation patterns at the late Cimmerian unconformity.

INTRODUCTION AND GEOLOGICAL SETTING

Many of the Mesozoic basins bordering the North Atlantic display complex unconformities and thick continental successions recording major rifting events during the Late Jurassic to Early Cretaceous (Tankard & Balkwill, 1989; Hiscott et a!., 1990). Improved understanding of the stratigraphy and tectonic evolution of rifts bordering the Bay of Biscay has highlighted the similarities between the offshore and onshore basins of Spain (Garcfa-Mondejar

1985; Platt, 1989a, 1990; Platt & Pujalte, 1994), Ireland (Petrie et a!., 1989; Shannon, 1991) and southern England (Kamerling, 1979; Chadwick, 1985; Evans, 1990; Ruffell & Coward, 1992). Upper Jurassic clastic rocks are reported from the offshore basins to the southwest of the UK and Ireland on the northern Biscay margin (Millson, 1987; Evans, 1990; Shannon, 1991; Moore, 1992), but there has been limited sedimentological study of

et a!.,

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

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these subsurface successions to date. In northern Spain, on the southern Biscay margin, Upper Jurassic-Lower Cretaceous continental clastic de­ posits are well-exposed, permitting detailed sedi­ mentological study and mapping of lateral facies and thickness variations within the Upper Jurassic synrift succession (Platt, 1986; Clemente & Perez-Arlucea, 1993; Gomez Fernandez & Melendez, 1994). This paper presents the results of a sedimen­ tological study of the Upper Jurassic Senora de Brezales Formation of the western Cameros Basin (Fig. 1), outlining the tectonic controls on alluvial fan deposition and palaeokarst development at the late Cimmerian unconformity.

the modern Florida Everglades or to the marshlands of southern Iraq (Platt, 1989b, c; Platt & Wright, 1991, 1992). The Rupelo Formation is overlain by oncoidal limestones, sandstones and mudstones of the ?Valanginian-Barremian Hortigtiela For·· mation (Fig. 2), and a thick ?Barremian-Aptian series of alluvial clastic rocks comprising the sandy-· muddy Piedrahita de Mufi6 Formation and the conglomeratic-sandy Salas Group. This succession is capped by alluvial sandstones and conglomerates of the upper Albian to lower Cenomanian Utrillas Formation.

LITHOFACIES STRATIGRAPHY

The stratigraphy of the western Cameros Basin, described by Platt (1986, 1989a, b) and Platt & Pujalte (1994), is outlined in Fig. 2. A thick suc­ cession of Early-Middle Jurassic marine carbonates is truncated by a complex series of unconformities (Mensink & Schudack, 1982; Schudack, 1987; Platt et al., 1991). In the south, Bathonian­ Callovian limestones are discordantly overlain by (?Callovian-) Oxfordian sandstones and cal­ carenites of the San Leonardo Formation (Wilde, 1988) and Oxfordian to basal Kimmeridgian car­ bonates of the Talveila Formation (Dfaz et al., 1983; Platt, 1986; Thalmann, 1989). Sedimentary cycles within this complex and heterogeneous succession of marginal marine carbonates and littoral clastics (Platt, 1986) may be compared with those present in the Corallian of southern Britain (Talbot, 1973; Sun, 1989). The middle Oxfordian to Kimmeridgian shallow­ to marginal-marine rocks are absent in many areas, reflecting erosion prior to deposition of the Senora de Brezales Formation, which is a laterally variable succession of continental deposits present through­ out the western part of the basin (Platt, 1989a). The Senora de Brezales Formation comprises poly­ genetic conglomerates, red sandstones and calcareous palaeosols, and is probably of Kimmeridgian­ Berriasian age (Platt, 1986; Wright et al., 1988). The Senora de Brezales Formation reaches a maximum thickness of 75 m to the north of Espej6n (Figs 1 & 3), and is overlain by lacustrine-palustrine lime­ stones of the Rupelo Formation (Berriasian). The Rupelo Formation is interpreted as a succession of freshwater carbonates deposited in a low-gradient, low-energy system of lakes and swamps similar to

The Senora de Brezales Formation shows strong lateral variability (see discussion below), but in all cases comprises three main facies: sandstones, con­ glomerates and carbonate rocks. Figure 3 presents a representative logged sedimentological section from the Senora de Brezales Formation. Red sandstones

Red sandstones reach a maximum of 40 m in thick­ ness and are present: 1 forming the lower part of the Senora de Brezales Formation to the southwest of the San Leonardo Fault (Fig. 1); 2 above basal Senora de Brezales Formation lime­ stone conglomerates in the northeast of the study area. The sandstones are typically dark brick-red in colour and are of medium grain size. They have subrounded grains, are generally quartzose, and contain scattered quartz pebbles 1-4 mm in dia­ meter. The sandstones are generally structureless (Fig. 4A), although trough cross-bedding, thin pebbly lags and plane lamination are locally discern­ ible (Fig. 4B). White-red mottling and/or 1-2 cm diameter burrows are developed at a few localities. Vertical cylindrical structures 5-10 em in diameter are also present locally, as are 2-cm-diameter tubu­ lar structures parallel to bedding. In the west, at Hortezuelos and Mamolar (Fig. 1), the sandstones are cut by vertical and horizontal cracks. Brecciated, nodular carbonate horizons 0.5-1 m thick are also present. These are commonly mottled and display floating quartz grains, angular, branching spar-filled cracks and grey carbonate stringers 1 mm in width. These carbonate facies are described in detail below.

221

Alluvial fans, palaeokarst and tectonics

A

1. (A) Palinspastic reconstruction of the North A tlantic and Bay of Biscay region for Late Jurassic times showing the location of major Mesozoic basins. The Cameros Basin is in central Iberia, to the south of the Biscay rift. (B) Map of the western Cameros Basin showing major structural elements and localities mentioned in the tex t: C, Castrovido; H, Hortigiiela/Valparaiso; Ho, Hortezuelos; HR, Huerta del Rey; J, Jaramillo Quemado; LG , La Gallega; M, Mamolar; Mo, Moncalvillo; ML, Mambrillas de Lara; PB, Pinilla de los Barruecos; PM, Pinilla de los Moros; Q, Quintanilla de las Vinas; R, Rupelo; SI, Salas de los Infantes; SL, San Leonardo de Yagiie; T, Talveila; QCF, Quintanilla­ Castrovido Fault; JCF, Jaramillo­ Covarrubias Fault; SLF, San Leonardo Fault. (C) Inset showing area covered by map B . Fig.

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thrust

transfer fault

Interpretation

The homogeneous character of the sandstones is consistent with strong bioturbation. Primary sedi­ mentary structures are preserved in only a few

places, but the rare presence of lags and mesoscale cross-bedding suggests that the sandstones were waterlain. Mud is absent; flows were prol::ably short­ lived with no deposition from suspension. Thin sec­ tion observation shows that the red colour is the

W CAMEROS

BASIN- STRATIGRAPHY

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Albian

Aptian

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Sandstone

Organic-rich limestone

Conglomerate

2. Upper Jurassic to Lower Cretaceous stratigraphy of the western Cameros Basin. Fig.

Limestone

223

Alluvial fans, palaeokarst and tectonics

z 0

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limestone and marl

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3. Logged sedimentological section from the Senora de Brezales Formation, section near chapel of Nuestra Senora de Brezales, 2 km north of Espej6n.

Quartz conglomerate Charophyte-gastropod

0

Massive red sandstone

j���� Marine limestone

Fig.

Marine Jurassic

product of iron oxide grain coatings, typical of con­ tinental 'red bed' deposits. This interpretation is consistent with the sedimentological evidence of subaerial exposure. This includes cracks, interpreted as sheet cracks developed during desiccation, mottles and tubular structures of varying morphology and size, interpreted as 1-2-cm-diameter burrows and 2-10-cm-scale roots. The brecciated, nodular and mottled fabrics of the carbonates are typical of pedo­ genetic horizons, favouring an interpretation as cal­ cretes. The presence of floating quartz grains indicates the in situ growth of replacive calcite. The red sandstones are thought to represent ephemeral stream deposits. The climate was prob­ ably semi-arid, with sporadic rainfall permitting only short-lived stream flow; in the intervening periods the sandstones were homogenized by burrowing and root action. The occasional rainfall permitted the growth of hardy scrub vegetation and periodically allowed the establishment of high or perched water tables, favouring the development of palaeosol pro­ files and pedogenetic carbonates (see also below). Desiccation and strongly oxidizing conditions were responsible for the red colour of the sediments and

prevented the preservation of organic material so that the roots are preserved only as moulds. Conglomerates

Conglomerates are common in the Senora de Brezales Formation. These are poorly sorted, with subangular to subrounded clasts, mostly of carbonate and generally varying in size from 2 to lOcm. Locally, the clasts reach 30-100 em, as at Jaramillo Quemado (Fig. 4C). The matrix consists of medium­ grained red sandstone with 1-2-cm rounded quartz pebbles, which are more abundant in the southwest. Conglomerates commonly occur in channels. The channels are 10-50 m across, have erosive bases and are 2-3 m deep (Fig. 5). Coarse conglomerates at the channel bases pass up into red, trough cross­ bedded sandstones. Many channels are cut directly into marine Jurassic carbonates, as in the northeast of the study area. At Castrovido (Fig. 1), channelized conglomerates rest on lower Callovian limestones, passing up into interbedded conglomerates and red sandstones. Channelized conglomerates also rest unconformably on lower Callovian marine

224

N.H. Platt

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4. Lithofacies and sedimentary structures within the Senora de Brezales Formation. (A) sandstones- massive. Huerta del Rey. Person for scale. (B) sandstones- cross-bedded. Huerta del Rey. Hammer for scale. (C) conglomerates- very coarse conglomerates containing locally derived clasts of lower Callovian limestone. Jaramillo Quemado. Lens cap for scale. (D) conglomerates- stacked, fining up sheet-flood conglomerates. Arrows mark bases of sheet-flood units. La Gallega (Navas road) . Hammer for scale. (E) carbonates- cross-cutting laminar calcrete present within conglomerate, Hortezuelos. Lens cap for scale. (F) carbonates- stacked, horizontally bedded laminar calcretes in sandstone matrix, Mamolar. Lens cap for scale. (G) karst surface- karstic cavity developed in Lower Kimmeridgian limestones, filled by limestone breccia and red internal sediment, Talveila. Hammer for scale. (H) karst surface- red sandstone, marked by arrow, piped down into lower Callovian limestone surface, Pinilla de los Moros. Hammer for scale. Fig.

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5. Channels of polymict conglomerate cut into red sandstones and marls, Espej6n. The arrows mark the bases of three distinct channels. These channels are not fault-bounded. Bedding is vertical, younging away from camera: The photograph was taken standing on Jurassic marine limestones and looking southwards. The prominent ridge running left to right across the photograph above the Senora de Brezales Formation channels is composed of non-marine limestones of the o verlying Rupelo Formation (R). The village and cornfields beyond lie on Valanginian-Cenomanian clastic rocks (KL), whereas the mountain on the horizon consists of Upper Cretaceous marine limestones (KU). Fig.

Alluvial fans, palaeokarst and tectonics

carbonates at Quintanilla de las Vinas and Hortigiiela in the northwest, an area where red sandstones are absent. In the southwest, conglomerates occur inter­ bedded with sandstones or with palaeosols some metres above the base of the Senora de Brezales Formation. Many channels incised into the basal unconformity are bounded on one or both margins by minor NE­ SW trending faults. Lateral variations in conglomer­ ate thickness also commonly occur across these faults. In contrast, channelized conglomerates pre­ sent at higher levels within the Senora de Brezales Formation are not fault-bounded. At Espej6n {Fig. 1), a number of individual channels are spectacularly exposed. The channels are asymmetric, with strongly erosive bases. The coarsest conglomerates fill the base of the channel, passing upwards and towards the channel margin into red, trough cross-bedded sandstones. Sheet-like conglomerates are also common; these show planar erosive bases and typically occur in 2-5-m beds (Fig. 4D). Several pulses of conglomer­ ate occur at Senora de Brezales; these are traceable laterally for 20-50 m. At Hortezuelos (Fig. 1), conglomerates near the base of the succession are monomict, dominated by angular 2-4-cm clasts of grey micrite. Elsewhere, the conglomerates are polymict ('polygenetique' of the French usage; Salomon, 1982). Clast lithologies include: 1 fossiliferous yellow-orange packstones containing 0.5-mm quartz grains and brachiopod fragmentsthese are derived from the marine lower Callovian; 2 pink to grey micrites, commonly partially recrystallized- these may represent subaerially weathered clasts of marine carbonate from the Talveila Formation, Middle Jurassic, or Liassic; 3 pink oolitic grainstones- these are especially common in the southwest of the study area, and were probably derived from Middle Jurassic shallow water carbonates; 4 rarer, small ( < 1 em) clasts of red sandstonethese may be derived from red sandstones in the Senora de Brezales Formation, or from Triassic red­ beds. The conglomerates show few sedimentary struc­ tures, and are mostly clast-supported. Matrix­ supported conglomerates occur at a few localities, as at Senora de Brezales and La Gallega (Navas road), but even in these cases there is little mud and the matrix is comprised of sandstone.

227

Interpretation

The channels provide evidence of powerful, strongly erosive flows which were probably intermittent events associated with sporadic heavy rainstorms. The wadi-type channels were incised into the uncon­ formity surface on uplifted blocks, and locally, as at Espej6n, into finer grained deposits of the Senora de Brezales Formation. Away from the fault scarps, flows spread out into sheet-floods, depositing tabular units of conglomerate and sandstone. Carbonate rocks

Pedogenetic carbonate horizons are interbedded with the red sandstone facies of the Senora de Brezales Formation (see above). Carbonate horizons also rest directly upon marine Jurassic limestones in interchannel areas. Laminated limestones are present in the north of the study area, lying either directly on the uncon­ formity surface (as at Castrovido and Mambrillas de Lara; Fig. 1), interbedded with or cross-cutting the sandstones and conglomerates (as at Hortezuelos and Mamolar; Fig. 4E & F) or developed as exten­ sive sheets resting on conglomerate bodies (as at Quintanilla de las Vinas). The laminae are irregular and undulating, consist­ ing of 1-mm alternations of light and darker grey (rarely black) micrite. Many of the laminated hor­ izons display a fabric of horizontal 1-mm pale beige, concentrically laminated carbonate tubules 0.5 mm in diameter filled with sparry calcite. At Senora de Brezales and at Espej6n (Fig. 1), a succession of marly, brecciated and heavily mottled carbonate horizons is interbedded with sheet-flood and channelized conglomerates. These sediments display a complex reticulate fabric produced by horizontal 2-20-mm layers of laminar calcrete and vertical tubular structures 1-10 em in diameter and up to 1 m in length. Some of the vertical structures are composed of homogeneous micrite, but others show an internal fabric similar to that of the laminar horizons. Between these structures, the mottled marly, or sandy, matrix commonly contains abun­ dant fine carbonate strings and spar-filled angular, branching cracks. At Senora de Brezales, stacked laminar horizons form mound-like structures 5 m across with a relief of 0.5 m (fig. 9a in Wright et a/., 1988).

228

N.H. Platt

Interpretation

The carbonate horizons within the Senora de Brezales Formation are interpreted as calcareous palaeosols. The laminated limestones are similar to laminar calcretes described from modern exposure surfaces in Florida by Multer & Hoffmeister (1968). Although similar in some respects to laminar stro­ matolites, an absence of diagnostic algal structures, such as calcified filaments (Read, 1976), and their association with pedogenetic features, suggests inter­ pretation as biogenic laminar calcretes or 'rhizolites' (Wright et a!., 1988). The fine spar-filled tubules are interpreted as root structures (rhizoliths). The beige micrite coating may either be a replacement of the root tissue itself (internal septae commonly form the characteristic 'alveolar texture' of Esteban, 1974), or may be made up of layers of concretionary car­ bonate precipitated around the root. A more detailed description of the laminated horizons is to be found in Wright et a!. (1988). The thickness and abundance of laminar calcrete horizons towards the top of the formation point to long periods of root colonization, both on the surface and beneath it. These times were punctuated by occasional input of fine-grained distal alluvial clastic material, producing an alternation of laminar cal­ cretes with mottled sandstone or marl. Desiccation and the action of larger roots were probably respon­ sible for most of the observed brecciation of the carbonate horizons. Brecciated, mottled carbonate horizons are also common within the basal part of the overlying Rupelo Formation, and in places there may be an upward transition between the Rupelo and Senora de Brezales formations. Although Platt ( 1986) con­ sidered these to be two essentially separate units, they may be partial lateral equivalents.

GEOMETRY OF LATE JURASSIC UNCONFORMITIES

The study of Upper Jurassic facies distributions is complicated by the presence of at least two import­ ant unconformities, one beneath the San Leonardo Formation (probably late Callovian to middle Oxfordian in age) and a later one beneath the Senora de Brezales Formation (probably intra­ Kimmeridgian). Further discontinuities may exist within the San Leonardo Formation. The superpos­ ition of these unconformities causes problems in

unravelling the depositional history of the study area. Effects of local faulting were certainly signifi­ cant in the localized preservation of strata beneath the unconformities. The Senora de Brezales Formation rests on the most important intra-Mesozoic unconformity in the study area. This unconformity records a period of uplift, faulting and erosion that began some time after the early Kimmeridgian. It marks a major change in basin configuration, and a change from relatively uniform, gentle subsidence and generall.y shallow marine conditions from the Rhaetian to early Kimmeridgian, to the more rapid, strongly localized subsidence and continental sedimentation that characterized the latest Jurassic and Early Cretaceous. The combined effects of tectonic activity and ero­ sion during this period resulted in the formation of a faulted surface with significant relief. Thus the Senora de Brezales Formation lies unconformably on a variety of different strata, from Bathonian-­ Callovian shallow-marine limestones, to upper Oxfordian to lower Kimmeridgian marginal marine sandstones and limestones of the San Leonardo and Talveila Formations. Where the Senora de Brezales Formation rests on Kimmeridgian and Callovian limestones, it reslts variably upon karst surfaces on palaeohighs and on an erosive channelled base in the lows. Where sand­ stones of the San Leonardo Formation are truncated by the unconformity, the basal contact is planar without the development of palaeokarst or significant relief (Platt, 1986). Karst surface

The Senora de Brezales Formation lies on karstified lower Kimmeridgian limestones at Talveila (Mauthe, 1975; Fig. 40), Pinilla de los Barruecos and Horte­ zuelos (Fig. 1). At all these localities, red internal sediment fills cavities and joints within the underlying carbonate rocks. At Hortezuelos, the karstified sur­ face is cut by a small fault with a throw of 1 m. Here, deeper karstification and/or partial erosion of the fault scarp led to the local deposition of a 1-2-m lens of coarse chaotic conglomerates. At Castrovido in the north of the study area (Fig. 1), a karst surface is developed on marine Callovian limestones in interchannel areas. The marine Jurassic carbonate surface is locally dis­ coloured and brecciated and laminar calcretes rest on the unconformity. A similar situation occurs

Alluvial fans, palaeokarst and tectonics

at Mambrillas de Lara (Fig. 1), where the lower Callovian limestone surface is draped by laminar calcretes penetrating downwards for 30 em along joints and cracks (see Wright et al., 1988). At Pinilla de los Moros (Fig. 4H), a karstified surface of lower Callovian limestone is penetrated by vertical cracks 1 em across filled with red sandstone. At these localities, subaerial exposure and erosion prior to deposition of the Senora de Brezales For­ mation resulted in karstification of underlying lime­ stones. Red internal sediment was piped into joints and karstic cavities; laminar calcretes formed at the surface and along joints and cracks beneath. Planar contact

To the north of Hortezuelos, and at Mamolar and La Gallega (Fig. 1), red sandstones of the Senora de Brezales Formation rest on yellow sandstones of the San Leonardo Formation and paleokarst is not devel­ oped. A S-cm-thick, dark red, ferruginous-cemented sandstone layer is locally present at the contact. The poorly defined contact between the red sand­ stones and the underlying yellow sandstones suggests that the red sandstones may have been at least partially derived by reworking of the San Leonardo Formation. Channels

Channels filled with polymict conglomerates occur at the base of the Senora de Brezales Formation in the north of the study area, where the San Leonardo and Talveila Formations are absent. The channels record powerful stream erosion of the limestones beneath. The siting of channels on the downthrown sides of NW-SE trending faults (Fig. 6) suggests that they formed in structural lows produced by the Late Jurassic faulting.

TECTONIC CONTROLS ON FACIES DEVELOPMENT

The style of sedimentation within the Senora de Brezales Formation strongly reflected the pattern of fault blocks and the lithologies exposed. Where the sandstones of the San Leonardo Formation were eroded, planation of the contact ensued and rework­ ing led to the deposition of red sandstones. Karst surfaces developed on fault-bounded structural highs, where Middle Jurassic and lower Kimmerid-

229

gian marine carbonates were exposed. Erosion of these highs acted to equalize the pediment relief. A schematic NW-SE stratigraphical section for the Senora de Brezales Formation in the north of the study area (Fig. 6) demonstrates the tectonic control on facies that is evident after restoration to the base of the overlying Rupelo Formation. The presence at Jaramillo Quemado of coarse basal con­ glomerates with clasts of lower Callovian limestone up to 1 m in size (Fig. 4C) probably records erosion of a faulted high to the west of the Jaramillo­ Covarrubias Fault, a major NE-SW transfer struc­ ture (Platt, 1990). At this locality there may be a gradation to karst conglomerates, where brecciation was followed by only limited transport, so that the locally derived clasts were deposited more-or-less in situ, infilling an irregular karstic relief. Red sand­ stones are restricted to the area to the east of the fault, whereas immediately to the west, between Valparaiso and Mambrillas de Lara, karst surfaces are developed on the marine Jurassic, and the Senora de Brezales Formation is thin or absent. Moving further to the west, conglomerates are present firstly on isolated fault blocks and eventually in fault­ controlled channels. Many of the conglomerate-filled channels are bounded on one or both sides by NE-SW faults (Fig. 6); good examples occur at Castrovido, Quintanilla de las Vinas, and between San Leonardo de Yagiie and Hontoria del Pinar (Fig. 1). These minor faults are parallel to the Jaramillo-Covar­ rubias Fault. At La Gallega, a NE-SW fault with a throw of 25 m on the western side of the Navas road bounds a depression filled with Senora de Brezales Formation conglomerates and sandstones. Many more minor faults occur, typically at a 100-m spacing. These faults have throws of only 2-10 m, but still controlled lateral variations in the subcrop to the unconformity and exerted strong local influence on facies development and thicknesses within the Senora de Brezales Formation. Changes in facies and thickness also occur across NW-SE faults. Although the exposure is not con­ tinuous in the dip direction and the basin structure only allows comparison between outcrops on parallel anticlines several kilometres apart, general contrasts are evident. Thick successions with red sandstones up to 40 m thick occur near the southwestern basin margin, whereas thinner successions are present to the north. In the north of the study area, the Senora de Brezales Formation is thicker to the south of the Quintanilla-Castrovido Fault than immediately to

230

N.H. Platt

NW

SE

m ottled I brecciated lim estone

l

MARINE JURASSIC PEDIMENT JCF Quintanilla de las Vinas

KEY

Mambrillas de Lara

?Berriasian

Valparaiso Hortiguela

Jaramillo Quemado

Rupelo Formation (Las Vinas Member)

Pinilla de los Moros

Castrovido

non-marine limestone & marl

J.+.�+----············-·······-·-············-··············

...............................................................................................................

red sandstone

?Kimmeridgian - Berriasian

Senora de Brezales Formation

..................................................•....•............•....•••.................................•

Lower Callovian

top of Marine Jurassic

conglomerate

laminar calcrete

ee;;r;<;i;l·

.................•....................................

marine limestone

Fig. 6. Cartoon strike section between Quintanilla de las Vinas and Castrovido, illustrating tectonic control on lateral changes in facies development within the Senora de Brezales Formation. Section restored to the base of the Rupelo Formation to eliminate the effects of post-depositional faulting of Lower Cretaceous and Tertiary age (see Platt, 1990). Horizontal distance between Quintanilla de las Vinas and Castrovido is 30 km; maximum vertical thickness of the Senora de Brezales Formation is 50 m.

the north. Red sandstones are well-developed to the east of the Jaramillo-Covarrubias Fault between Jaramillo Quemado and Castrovido but are very thin or absent immediately to the northwest of the fault. These lateral variations suggest that both NW-SE and NE-SW fault sets were probably active during deposition of the Senora de Brezales Formation, defining a complex mosaic of fault blocks with differing subsidence histories and different detailed stratigraphies.

EROSION, CLAST DERIVATION AND DEPOSITIONAL HISTORY

The composition of the conglomerates within the Senora de Brezales Formation strongly reflects the subcrop to the unconformity (Fig. 7). The irregular

clast shapes and abundance of coarse carbonate clasts suggests limited clast transport; clast derivation was chiefly from erosion of the marine Jurassic limestones of underlying or adjacent fault blocks. The vertical succession may therefore record the progressive erosional history ('unroofing') of the source areas. 1 Locally, as at Hortezuelos, a basal conglomerate is present, containing clasts only of grey micritic limestone. These clasts appear to have been derived from laminar calcretes deposited in interchannel areas and from Talveila Formation lagoonal carbonates. 2 Red sandstones above consist mainly of quartz, probably including a component reworked from the sandstones of the San Leonardo Formation. 3 Polymict conglomerates above these contain abundant reddened oolitic limestone clasts, indi-

231

Alluvial fans, palaeokarst and tectonics

CLAST DERIVATION

DEPOSITIONAL HISTORY Valanginian Cenomanian

4

quartz sandstones and conglomeratesprogressive unroofing and erosion of

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···············

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Formation

.....................

Senora de Brezales Formation

Talveila Formation:

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-

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San Leonardo Formation:

.....'!��-��!��!-�-��!�-�-�-���-�!��!:�---Middle Jurassic: shallow-marine limestones

limited clastic supply

--------------------- --·-············-·----------------·

Lower Jurassic: marine limestones and marls

polymict conglomerates reworking of marine Jurassic, especially Middle Jurassic red sandstones - reworking of San Leonardo Fm

monomict basal conglomerate ��:-y-���!�-�-?!.!.����i!� .F.�.: .................

--------------------------------

Triassic: continental, mostly sandstones

.

7. U pper Jurassic-lower cretaceous clast derivation . Clastic rocks- vertical succession illustrating likely derivation of clasts from the successive 'unroofing' and erosion of underlying stratigraphical units. See text for explanation. The schematic section shown is a composite based on outcrops on the northeastern flank of the Hortezuelos anticline.

Fig.

eating erosion and reworking of Bajocian-Callovian shallow marine carbonates. Their deposition may therefore date the 'unroofing' of the underlying Middle Jurassic. Similar polymict conglomerates containing abun­ dant coarse limestone clasts occur in a number of areas, in each case recording local clast derivation from the marine Jurassic. Examples occur in the north of the study area, as described above (Fig. 6). The San Leonardo Formation is absent in the north, but the local presence of red sandstones, as at Castrovido, may record transport of eroded material elsewhere from the San Leonardo Formation or from basement lithologies. In the southwest of the study area, the presence of quartz pebbles may record erosion of quartzitic material from the San Leonardo Formation or from basement rocks on the Duero block to the south. However, widespread erosion of Triassic sandstones and Palaeozoic quartzites probably did not com­ mence until much later (Platt, 1989a). The low clastic content in the freshwater carbonates of the overlying Rupelo Formation suggests that this phase of peneplanation was complete by the Berriasian, although the catchment area geology was still appar­ ently dominated by carbonate lithologies. The first wide-spread erosion of the Triassic and Palaeozoic is probably dated by the appearance of sandstones within the Barremian-Aptian Hortigiiela and Pied­ rahita de Mufi6 Formations (Figs 2 & 7). These rocks do not contain Jurassic limestone clasts-

this may reflect lengthier transport in more perma­ nent river systems developed under more humid conditions.

ENVIRONMENTAL SYNTHESIS

Figure 8 presents a schematic depositional model for the Senora de Brezales Formation, showing the main environments represented and the control of the unconformity subcrop and of tectonics on facies development. Sedimentation took place under a semi-arid climate. Small-scale wadi-type channels were incised directly into the underlying marine limestone pediment, but much of the sandstone and conglomerate was deposited from laterally extensive sheet-floods spreading out from these channels. Pedogenetic carbonates developed on the pediment in interchannel areas, and also formed as surficial layers and nodular brecciated horizons on or within the exposed alluvial clastic deposits. Bioturbation and root penetration acted to destroy primary sedimentary structures, particularly within the sandstones. The association of channels and sheet-floods is consistent with deposition in small-scale alluvial fans. Basin-bounding NW-SE faults and NE-SW trans­ fer faults both influenced sedimentation, which acted to infill fault-bounded depressions; clastic supply came dominantly from the erosion of neigbouring uplifted blocks.

N.H. Platt

232

pedogenesis - laminated crusts

& cross-graben wrench or "transfer" faults

laminated crusts on interchannel areas of pediment

\

paleosols on alluvial substrates

small fans sourced at fault intersections

channels cut into pediment located along transfer faults

basin-bounding extensional normal faults

_.U I

2

Oxfordian arenites

erosion of Mid. Jurassic sequence - polymict conglomerates localised erosion of Lr Kimmeridgian limestones - monomict conglomerates

8. Schematic depositional model for the Senora de Brezales Formation, illustrating sedimentary environments represented and the influence of tectonics and the underlying geology on facies development and clast derivation.

Fig.

REGIONAL PERSPECTIVE

Upper Jurassic clastic rocks

The Kimmeridgian-Tithonian saw a major re­ gression in Iberia and across much of northwest Europe. In the Lusitanian Basin of Portugal (Wison et al., 1989), tectonic movements led to an acceler­ ation of subsidence; the lacustrine-lagoonal car­ bonates of the Oxfordian Caba'
Brezales Formation occur at the base of the laterally equivalent Aguilar Formation (see Platt & Pujalte; 1994). Here, too, the subcrop to the Late Jurassic unconformity was dominated by marine Jurassic limestones developed in facies similar to those of the Cameros. Faulting and continental erosion then permitted the deposition of locally derived, tec­ tonically controlled conglomerates of similar composition. Upper Jurassic continental clastic sediments are also reported from several basins to the west of Britain on the northern Biscay margin, where they form potential hydrocarbon reservoir targets. Exam­ ples occur in the Celtic Sea Basins (Millson, 1987; Petrie et at., 1989; Shannon, 1991) and the Brittany Basin of the Southwestern Approaches (Evans, 1990). in these areas, a major period of Late Jurassic extension led to fragmentation of the basins into fault blocks, which were tilted and eroded. These Late Jurassic fault block structures form a major exploration target in the Porcupine Basin of offshore

233

Alluvial fans, palaeokarst and tectonics

Ireland (MacDonald et al., 1987). In the Porcupine, the Kimmeridgian saw a change in tectonic regime, as widespread broad subsidence gave way to fault­ controlled subsidence and sedimentation. This evol­ ution is parallel to that recorded by deposition of the Senora de Brezales Formation in the western Cameros Basin. Moore (1992) recognized an Upper Jurassic synrift to post-rift transition succession up to 500 m thick within faulted sub-basins of the Porcupine Basin. These rocks are characterized on seismic data as mounded, fan-style deposits resting on a faulted unconformity, passing laterally and upwards into parallel, onlapping reflectors of a ?marginal marine succession. Moore suggests that these fans may have significant reservoir potential, particularly in the sand-prone proximal-mid- and mid-fan regions. Proximal areas of the fan contain poor reservoir­ quality unsorted conglomerates (Shell well 34/19), whereas the distal part is expected to contain fine siltstone and shale. The clastic fans were bounded by faults on one side, passing basinwards into finer grained distal fan deposits and onlapped by Lower Cretaceous shales. Although this succession was deposited in a marginal marine setting, the facies geometries developed are essentially identical to those inferred for the alluvial clastics of the Senora de Brezales Formation (Fig. 9) and the overlying low-gradient lacustrine deposits of the Rupelo Formation. In the North Celtic Sea (Millson, 1987), inter­ bedded marginal marine carbonates and littoral and continental clastics of the mid- to upper Oxfordian

Clastic fans

Timoleague Formation show similarities with the heterogeneous marginal marine succession devel­ oped in the San Leonardo and Talveila Formations. Like the western Cameros, the North Celtic Sea saw a change to continental conditions during the Kimmeridgian. Here, the coarse- to fine-grained, locally conglomeratic sandstones of the Clonakilty Formation may correspond with the clastics of the Senora de Brezales Formation in the Cameros. This interpretation is further supported by Millson's (1987) descripton of thin, concretionary, silty micritic carbonate horizons, possibly representing calcretes, within the Clonakilty Formation. Late Jurassic unconformities

Many authors have recognized a Late Jurassic to Early Cretaceous unconformity in northwest Europe (see e.g. Ziegler, 1990). This 'late Cimmerian' unconformity, which is variably dated as earliest Cretaceous or Volgian-Tithonian, is known from the North Sea (Rawson & Riley, 1982), Porcupine Basin (Croker & Shannon, 1987; Tate, 1993), North Celtic Sea (Millson, 1987; Shannon, 1991; McCann & Shannon, 1993), and Cantabria (Platt & Pujalte, 1994). Accurate dating of this surface is commonly hindered by problems of stratigraphical resolution in marginal/restricted Upper Jurassic facies; for exam­ ple, in the study area, the basal Senora de Brezales Formation unconformity can be dated only as post­ lower Kimmeridgian and pre- lower Berriasian. Superposition of several unconformities means that the time gap represented by the unconformity is

Onlapping infill

Clastic fans

Basal unconformity

9. Sketch showing inferred stratigraphical relationships in the Upper Jurassic Senora de B rezales Formation of the western Cameros Basin. Figure is modified after a cartoon by Moore ( 1992) depicting seismic relationships in the Upper Jurassic synrift to post-rift transition of sub-basins within the Porcupine Basin (Moore, 1992). The two successions share a common tectonic evolution and are thought to show e ssentially identical geometry. Fig.

N.H. Platt

234

of variable extent in different areas. In basinal areas of the North Sea and the Wessex Basin, the 'base Cretaceous unconformity' is represented by a con­ densed section of Volgian-Ryazanian (Tithonian­ Berriasian) age (Rawson & Riley, 1982; Ruffell, 1992). However, on structural highs and in areas adjacent to the Biscay rift, Lower Cretaceous rocks may rest directly upon Liassic or Triassic rocks, as occurs in western parts of the Celtic Sea basins (Millson, 1987), the Western Approaches (Evans 1990) and northern Cantabria (Garcia-Mondejar et a!., 1985; Hines, 1988). Study of the Cameros Basin indicates that the complex association of superimposed unconformities in the areas around Biscay is the product of vari­ ations in relative sea-level (notably falls in the early Oxfordian and ?mid-Kimmeridgian and rises in the late Oxfordian and ?late Kimmeridgian to Berriasian), coupled with widespread Late Jurassic extensional block faulting. The effect of these, reflected in the basal contact, deposition and deri­ vation of the Senora de Brezales Formation, was to produce truncation during times of lower base level, particularly on the footwall crests of extensional fault blocks, and complex sedimentary onlaps during infilling of local accommodation space created by faulting or rises in relative sea-level.

were derived from localized reworking of lower Kimmeridgian limestones, red sandstones from erosion of upper Oxfordian sandstones, and polymict conglomerates were derived from erosion of the entire Jurassic succession, in particular the Middle Jurassic. The Senora de Brezales Formation offers a useful outcrop analogue for Upper Jurassic clastic suc­ cessions in basins bordering the northern Biscay margin to the west of Britain. A depositional model developed for the Senora de Brezales Formation (Fig. 9) serves to illustrate the importance of seismic­ scale and subseismic-scale fault movements in con­ trolling sedimentation and facies distributions and highlights the strong lateral variability so character­ istic of synrift deposition.

ACKNOWLEDGEMENTS

My sincere thanks go to Dr Harold Reading for support and encouragement. I am grateful to the students from Oxford and Bern Universities who helped with detailed mapping and logging. Helpful reviews and comments were provided by Guy Plint, Alastair Ruffell, Bruce Sellwood and Pat Shannon. My special thanks go to Adrian Forrer for discussion and good humour in the field.

CONCLUSIONS

The Senora de Brezales Formation is a laterally variable Upper Jurassic succession of continental red-beds deposited under a semi-arid climate. Con­ glomerates and sandstones were deposited in wadi­ type channels and by sheet-floods in small-scale alluvial fans. Bioturbation, desiccation and root action led to homogenization of the waterlain sand­ stones during protracted subaerial exposure. Pedo­ genetic carbonates represent palaeosols and laminar calcretes formed within and at the base of the clastic succession. The Senora de Brezales Formation rests on the late Cimmerian unconformity, which truncates a thick succession of marine Jurassic carbonates. Lateral variations in thickness in the Senora de Brezales Formation reflect Late Jurassic extensional block faulting; many channels are fault-bounded and the coarsest conglomerates occur adjacent to major faults. The deposition of the Senora de Brezales Formation recorded erosion of the under­ lying marine Jurassic; monomict conglomerates

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Spec. Pubis int. Ass. Sediment. (1995) 22, 237 - 2 64

The use of geochemical data in determining the provenance and tectonic setting of ancient sedimentary successions: the Kalvag Melange, western Norwegian Caledonides RODMAR RAVNAS

and

HAR ALD FURNES

Geological Institute, Univers ity of Bergen, Allegt. 41, N-5007 Bergen, Norway

ABSTRACT

The inferred Silurian Kalvag Melange is a thick (> 2500 m ) bedded olistostromal melange, which formed due to gravitational resedimentation driven by tectonic activity along the steep submarine slope of a basin margin and was deposited in the lower slope or base-of-slope environment. The melange consists of olistoliths, more than 2.8 km in length and up to 0.5 km in thickness, of shallow-marine deposits, deep-marine turbidites, bedded cherts, pebbly mudstones, calcareous rocks, lava-flow basalts and andesites, rhyolites and rhyolitic ignimbrites embedded in a sheared olistostromal 'groundmass'. The latter comprises chiefly debris-flow deposits. Interfingering channel-fill conglomerates and gully-fill turbidites with erosional contacts to the olistostromal 'groundmass' are present. The channel-fill conglomerates are dominated by debris of lithic-volcanic arenites and volcanic rocks. Olistoliths of volcanics and the volcanic debris in the channel-fill conglomerates were derived from alkaline, calc­ alkaline , !AT, MORll and boninitic sources. Geochemical and isotope data from the volcanic rock olistoliths reflect an evolution from magmatism associated with initial rifting to incipient back-arc spreading, and finally to volcanic arc magmatism. Geochemical analyses of conglomerate debris and sedimentary-rock olistoliths suggest derivation from an evolved, mature island arc, and an ophiolitic source. The volcanic arc is thought to have formed by rifting of accreted terranes along an active continental margin, and the MORJJ-type and boninitic debris most likely represent derivation from an older, accreted ophiolitic terrane. The sedimentological and geochemical data jointly suggest that the deposition of the Kalvag Melange took place in a back-arc basin , formed by progressive cannibalization of the volcanic arc apron . Based on the present knowledge of the development of marginal basins within the outboard terranes of the Caledonian orogen, the basin-evolution history recorded by the Kalvag Melange can be correlated with the history of other late Ordovician to early Silurian marginal basins.

INTRODUCTION

The primary control that tectonics exerts on the production, d ispersal, accumulation and lithification of sediments has long been recognized (Pettijohn et al., 1987). Provenance regions, and especially those forming parts of the overriding plate at destruc­ tive plate margins, can u ndergo significant and rapid uplift. Much information is lost due to subsequent erosion, deformation and/or tectonic displacement, and a complete in s itu record of the source rocks is rarely preserved. The sedimentary record of adjacent sedimentary basins can provide some constraints toward redressing this preservational bias (e.g.

Dorsey, 1988; Smith et al., 1988; Haughton & Halliday, 1991; Marsaglia et al., 1992). Intercalated volcaniclastic rocks and clasts in coarse conglom­ erates can be characterized and sometimes dated, providing supplementary geochemical and isotopic information about the hypothetical provenance area and the original plate tectonic setting of the depo­ sitional basin. The mineralogical and geochemical composition of terrigenous sediments and sedimentary rocks depend chiefly upon provenance, relief and climate of the source region(s), weathering, transport

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

237

238

R. Ravnas and H. Fumes

processes, basin geometry and morphology, depo­ sitional e nvironment and diagenesis (e. g. Bhatia, 1983; Basu, 1985; Helmold, 1985; McBride, 1985; Mutti, 1985; Normark, 1985; Ricci Lucchi, 1985; Dorsey, 1988; Johnsson & Stallard, 1989; McLennan et a/., 1990). Plate-tectonic setting has been advo­ cated as the primary control on sediment com­ position (Blatt et a/., 1980; Pettijohn et a/., 1987) and the link between plate-tectonic setting and modal framework of modern sediments (mainly sands and sandstones) appears empirically robust (e. g. see Crook, 1974; Schwab, 1975; Dickinson & Suczek, 1979; Ingersoll & Suczek, 1979; Dickinson & Valloni, 1980; Valloni & Mezzardi, 1984; Dickinson, 1985; Valloni, 1985). The nature of the primary framework grains can, however, be modified significantly, obliterated or destroyed during post­ depositional diagenesis, burial, deformation and/or metamorphism. The geochemical composition, element ratios and isotope composition of sediments seem less depe ndant upon these post-depositional processes (Hiscott, 1984; Robertson & Henderson, 1984; Argast & Donnelly, 1987; Larue & Sampayo, 1990), and as such have proved powerful and more reliable in provenance studies and discrimination of plate-tectonic setting of ancient sedimentary successions (Bhatia & Taylor, 1981; Bhatia, 1983; Bhatia & Crook, 1986; Roser & Korsch, 1986; McLe nnan et a/., 1990). The present paper discusses the provenance and tectonic setting of an olistostromal melange, the inferred Silurian Kalvag Melange of the western Norwegian Caledonides (Fig. 1), on the basis of data obtained from the melange's ' groundmass', olisto­ liths (of volcanic and sedimentary rocks) and cobbles and boulders of interbedded polymictic conglom­ erates. Secondly, an attempt is made to further detail the reconstruction and probable evolution of the melange's provenance region and depositional setting by integrating the present knowledge of the island arc/ophiolite terranes and marginal basins of the Scandinavian Caledonides.

GEOLOGICAL SETTING

The area between Solund and Bremanger (Fig. 1) forms part of the westernmost Norwegian Caledo­ nides. Three tectono-stratigraphical units are recognized: the lower, middle and upper tectonic units (Brekke & Solberg, 1987; Andersen et a/., 1990). The lower tectonic unit comprises Pre-

cambrian migmatitic orthogneisses (Jostedalen Complex), paragneisses and associated augen­ gneisses, amphibolites, eclogites and meta­ anorthosites (Fjordane Complex), and low- to medium-grade metamorphic sedimentary, volcanic and plutonic rocks (Askvoll Group) of the Western Gneiss Region. A quartz diorite from the Askvoll Group has yielded a middle Proterozoic magmati.c U-Pb age of 1640. 5 ± 2.3Ma (Skaret a/., 1994). The middle tectonic unit is separated from the lower by the extensional Kvamshesten Detachment Zone. It consists of the Dalsfjorden Suite, composed of various types of syenitic to charnockitic ortho­ gneisses, granites and gabbros (Kolderup, 1921), and the unconformably overlying pre-Silurian and Silurian metasediments of the H�yvik and Herland Groups, respectively. The H�yvik Group experienced polyphase deformation and associated metamorphism prior to deposition of the u nconform­ ably overlying Herland Group (Brekke & Solberg, 1987; Andersen et a/., 1990). The upper tectonic unit comprises the Solund­ Stavfjord Ophiolite Complex and its cover of meta­ sediments and metavolcanics (Fumes eta/., 1990). It is separated from the middle tectonic unit by a composite terrane, the Sunnfjord Melange, inter­ preted to have developed during obduction of the ophiolite complex upon an e arly to middle Silurian continental margin represented by the Herland Group (Berg, 1988; Andersen et a/., 1990; Alsaker & Fumes, 1994). The Solund-Stavfjord Ophiolite Complex has yielded a U - Pb age of 443 ± 3 Ma (Du nning & Pedersen, 1988). Single zircon U-Pb datings from its cover sediments have yielded both Archaean, Proterozoic and early Ordovician ages (Pedersen & Dunning, 1993). The Kalvag Melange is separated from the gneisses, metasediments and metamorphic plutonic and volcanic rocks of the lower, middle and upper tectonic units to the north by a major thrust zone (Hartz et al., 1994). The metasediments on Bremanger have bee n correlated with the H�yvik Group (Fumes et a/., 1990; Hartz et a/., 1994). The metamorphic igneous rocks include a number of small serpentinite bodies, amphibolites, green­ stones, gabbros and schists (Bryhni et a/., 1981; Hartz et a/., 1994), which may possibly correlate with the Sunnfjord Melange and/or the Solund·­ Stavfjord Ophiolite Complex and its cover. More­ over, Hartz et a/. (1994) recognized major E-W striking, southward dipping extensional shear zones along the northern and central parts of

TECTONOSTRATIGRAPHY Devonian

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Detachments between Lower. Middle and Upper Plates

Fig. 1. Simplified geological map and tectono-stratigraphy of the So lund- Bremanger area, compiled from Fumes eta/. ( 1990), Hartz eta/. (1994) and Osmundsen & Andersen ( 1994) , and simplified geological map of Frs;lya, modified from Bryhni & Lyse ( 1985).

t;5

\0

240

R. Ravnas and H. Fumes

Bremangerlandet (Fig. 1) that separate major plates, analoguous to the tectono-stratigraphy of the Sunnfjord area (Brekke & Solberg, 1987; Andersen et al. , 1990; Andersen & Jamtveit, 1990; Osmundsen & Andersen, 1994). The Kalvag Melange has been interpreted in terms of a tectonized olistostrome formed in an island arc or accretionary prism setting (Bryhni & Lyse, 1985; Fumes et al., 1990), and was tentatively included in the upper tectonic unit by Fumes et al. (1990). The melange succession is intruded by two plutons of gabbronoritic/dioritic and granodioritic composition (Fig. 1). The older gabbronorite/diorite has yielded a Sm-Nd age of 380 ± 26 Ma (Fumes et al., 1989), thus providing a minimum age for the melange formation. The uppermost stratigraphical level of the region is represented by early to middle Devonian clastic sediments, deposited in late to post-Caledonian collapse basins (Hossack, 1984; Norton, 1986; Seranne & Seguret, 1987; Andersen & Jamtveit, 1990). In the Bremanger area these sediments rest unconformably upon the pre-Devonian rocks as well as on the granodioritic pluton.

THE KALVAG MELANGE

The Kalvag Melange forms a more than 16-km-wide rock unit on the island s of Fqllya and Bremanger (Fig. 1). On Fr¢ya, the Kalvag Melange makes up a minimum 2500-m-thick succession consisting of clasts and large detached blocks (olistoliths) of native and exotic rocks (sensu Raymond, 1984) d ispersed in a pelitic 'groundmass' , and interbedded conglom­ erates and sandstones (Figs 1 & 2) (Ravnas, 1991). Graptolites and tabulate corals of probable Ordovician-Silurian age were reported by Reush (1903) and Kolderup (1928), but have not been confirmed by later studies. The melange succession shows no pervasive shear fracturing, but local zones of pervasively sheared matrix occur. Strong tec­ tonic deformation, including folding, faulting and boudinage, is limited to discrete, isolated zones (fault zones and shear fracture zones). Primary structures within olistoliths and the fragmental fabric of the melange 'groundmass' are generally well preserved within the contact aureoles of the two plutons, allowing distinction between different types of olisto­ liths and recognition of their primary depositional environments. The melange as a whole can be considered in

terms of its two principal components: the bedded olistostromal 'groundmass', and the olistoliths that are scattered in this groundmass. The olistoliths, which may exceed 2. 8 km in length and up to 0. 5 km in thickness, comprise shallow-marine (offshore to shoreface) sandstones and mudstones, deep-marine sandy turbidites, bedded cherts, pebbly mudstones, calcareous rocks, subaerial and subaqueous lava·· flow basalts and andesites with interstitial calcareous rocks and intercalated sediments, rhyolite and rhyolitic ignimbrite (Fig. 2A-E). The olistoliths are interpreted to have been emplaced as large slides and slumps. The melange 'groundmass' comprises mainly matrix-rich sedimentary breccias and gravelly mudstones (Fig. 2F & G) that are inferred to represent chiefly cohesive debris-flow deposits. The interbedded conglomerates and sandstones (Fig. 2H), interpreted to have been deposited from gravelly non-cohesive debris-flows and high-density turbidity currents and sandy high- to low-density turbidity currents respectively, commonly have sharp, erosive basal contacts, locally displaying large loadcasts. They form broad, lenticular bodies that locally interfinger laterally with the olistostromal facies of the melange's 'groundmass', and are inter­ preted to represent submarine slope channel- and gully-fill deposits. Black shales and subordinate fine·· grained turbidites, the latter probably representing overbank deposits of the submarine channels, form minor constituents of the melange succession. Locally the 'groundmass' is intruded by fine- to medium-grained sandstone dykes and sills. The Kalvag Melange is interpreted to have been formed due to gravitational resedimentation driven by tectonic activity along the steep submarine slope of a basin margin, and deposited in the lower slope to base-of-slope environment (Ravnas, 1991). Several points of evidence suggest that the entire melange succession chiefly is of sedimentary origin. On the island of Fr¢ya the melange d isplays a pervasive, although subtle, bedding defined by textural and compositional variations. The general stratigraphical way-up direction is recognizable, and there seem to be some large-scale thickening/ thinning upwards trends. These trends are correla­ tive throughout the island area, marked predomin­ antly by the stratigraphical thickness distribution of the m ass-flow facies (Ravnas, 1991). The slide and slump units (olistoliths) stand out as isolated, out­ sized peaks on this background, and are correlative (Fig. 3) only when associated with the maxima of the mass-flow facies thickening. The olistoliths that are

Geochemical data for determining provenance

encased in the bedded 'olistostromal' matrix show primary depositional contacts, and are in some cases draped by black shales. In some places, gradations down the inferred palaeoslope from sedimentary­ rock olistoliths of slide/slump origin into debris-flow deposits can be seen. The matrix of the melange is thought to have been derived from the fine-grained interlayers of the shallow marine sediments in addition to slope mud (see below). Partly similar clast lithologies are present in the interfingering channel-fill conglomerates and the olistostromal 'groundmass' or debris-flow deposits. Locally the overbank and channel-fill deposits show evidence of slumping prior to deposition from the subsequent flow, and the olistostromal 'groundmass' deposits locally are found as olistoliths and conglomerate debris. The spatial distribution of facies associations and olistoliths (Fig. 3) suggests that the melange developed during a series of successive resedimen­ tation events (Ravnas, 1991). Basal and internal penecontemporaneous, 'soft­ type' deformation (folding, faulting and boundinage) of sedimentary-rock olistoliths is attributed to their emplacement by sliding and slumping while still in a semi-consolidated state. The volcanic-rock olistoliths show little internal deformation, apparently due to their inherently rigid character. Sandstone dykes show cross-cutting relationships with the sheared olistostromal 'groundmass' and locally are found to intrude the deformed basal part of sedimentary-rock olistoliths. High-pressure-low-temperature miner­ alogy has not been observed, neither in olistoliths nor in the sheared matrix. The melange has been divided broadly into a lower part comprising mainly sedimentary slides, slumps and cohesive debris-flow deposits, and an upper part comprising mainly cohesive debris-flow deposits and subordinate slumps, fine-grained turbi­ dites and resedimented coarse conglomerates. The lower part contains few or no volcanic-rock olisto­ liths, whereas the upper part shows an upward increase in the content of volcanic-rock olistoliths.

GEOCHEMISTRY

Chemical analyses of samples from olistoliths, conglomerate debris and the olistostromal 'ground-· mass' were carried out by X-ray fluorescence. The glass bead technique of Norrish & Hutton (1969) was used for major element oxides, and pressed powder pellets for the trace elements, using inter-

241

national basalt standards and the recommended values of Govindaraju (1984) for calibration. Possible post-depositional alteration of basaits, vol­ caniclastic materials and sediments have been dis­ cussed in a number of studies. The elements Ti, P, V, Zr, Y, Nb and Cr, with the addition of Si, Fe, Mg, La, Th, Ce and Co in sedimentary rocks, are considered relatively stable during low-grade meta­ morphism (e.g. Cann, 1970; Coish, 1977; Shervais, 1982; Bhatia, 1983; Staudigel & Hart, 1983; Hiscott, 1984; Robertson & Henderson, 1984; Bhatia & Crook, 1986; Cas & Wright, 1988; Larue & Sampayo, 1990) and are used systematically through­ out this study. Ca, Na and K in sedimentary rocks show increased mobilization with increasing meta­ morphic grade (Larue & Sampayo, 1990) and thus are less useful as indicators of primary composition. The melange succession, forming mainly a biotite­ quartz hornfels (Bryhni & Lyse, 1985), shows prograde contact metamorphism towards the two plutons, especially the gabbronoritic/dioritic pluton, which has a border zone up to 150 m thick of migma­ tized and foliated host rock (Fumes et a!., 1989). To minimize the effect of remobilization of elements due to the increase in metamorphic grade, most samples were collected well away from the two intrusive bodies. To avoid possible systematic com­ positional variations of sediments related to vari­ ations in grain size (Roser & Korsch, 1986; Argast & Donnelly, 1987), chiefly medium- to fine-grained arenites (greywackes) were selected for analyses. Conglomerate debris of channel-fill conglomerates

Clasts of well-rounded igneous and well-rounded and angular sedimentary rocks are present, rep­ resenting gabbro, diorite, metabasalts, trachyande­ site/dacite, granite, rhyodacite/rhyolite, bedded chert, quartzite, subarkose, lithic arenite, lithic­ volcanic arenite and wacke, black shale and calcareous rocks. The chemical analyses of rep­ resentative samples are presented in Table 1. The metabasalts can be divided into three different types based on their geochemical characteristics (Fig. 4, Table 2). These are MORB (mid-ocean ridge basalts), tAT (island arc tholeiites) and rocks approaching boninitic composition. All types show negative Nb anomalies and exhibit either a flat pattern or show slight or pronounced depletion of P, Zr, Ti andY in the MORB normalized, multi-element variation diagrams (Fig. 4A), typical of back­ arc basin basalts, tAT and boninites respectively

242

R. Ravnas and H. Fumes OLISTOLITH COMPOSITION Alkaline potassic meta basalts Subalkaline MORB-Iike metabasalts Calcareous rocks Bedded chert Transitional type metabasalts Large olistoliths of native and exotic blocks

Ignimbrite and IAT metabasalts Deep-marine sandy turbidites

E

8 "'

'Background' sedimentary mass of bedded olistostrome (gravitational melange); ---contacts sedimentary, non-ophiolitic matrix and rock clasts/blocks

N

Shallow marine siliciclastic sediments

Alkaline potassic metabasalts, Ignimbrite and bedded chert

Shallow marine siliciclastic sediments

Fig. 2. (Above. ) Composite profile through the Kalvag Melange, compiled from measured sections along the eastern and

western coast of Frll\ya. (Right. ) Locations of field photographs are indicated by letters. (A) Flow-folded, welded ignimbrite. Lens cap is 6cm in diameter. (B) Massive andesitic and rubbly aa-type basaltic flows. Length of hammer is 40cm. (C) Shallow-marine (shoreface) sediments; interbedded inferred wave-formed ripple cross-laminated and parallel stratified sandstones. Scale bar is 5 em. (D) Normally graded, sandy Bouma-type Tabc and Tab turbidites. Lens cap is 6em in diameter. (E) Evenly bedded, turbiditic ribbon chert. Length of hammer is 40 cm. (F) Disintegration of olistolith of shallow-marine deposits into (G) clast-rich, debris-flow deposits of the melange 'groundmass'. Scale bars of (F) and (G) are 20cm and Scm respectively. (H) Clast-supported conglomerate representing inertial, non-cohesive, high-density gravelly debris-flow. Note polymictic clast population. Length of hammer is 40cm.

Geochemical data for determining provenance

243

t

B A

c ----

-

----

2000

? --

--

. ..:.; ��-�:-<�:

_.·����:�Jg�6�c��IC RO��§ts !D

---<< . ...:-:-:�.::.:�:�

· : -. ·�:-: !J . ·. : .- :·- : : : ·(::-: :. ·� .. . ·. .· . ·. · · : ?. · . · ·

1500

. ·

. ·

-

+---- 2.0km ___

2.5km

•

?

.

1500:

.

.

Stratigraphic height

.

.. .

.

uncertain

�)1 --��-�-� /

CHANNEL-/GULLEY FILL DEPOSITS ..,

1500

... ... ... . . . . . . .

.--:--=--.� �.-.-,.

------.

___

:;_,

_.,..-,""'":"-:- .-.-:

.

E

: :- ;,·,

_.;./ · Stratigraphic\..·...:....:...:_>­ height

o o"';--

2.8km -------!�

'/ /

<:... . .

oo

"�-·����OLI TH OF DEEP-MARINE SANDY TURBIDITES

� �

--<;_.:.-·� .

D

-o--

:

!:;·

·

"' ;:s "'-

500

uncertain

·

·� : �: :;i�::.:::.�::::-:-:::.-:-::.:-6

("

.� : 17 & & .::

Fig. 3. Spatial distribution of olistoliths within the Kalvag Melange. Possible correlations are marked by dashed lines . Sections A and B are generalized logs

interpreted from the map, whereas sections C, D and E are simplified versions of measured sections described by Ravnas (1991).

:;_,

"' ""' ;:s

� � � �

Table 1. Representative geochemical analyses of conglomerate debris Sample no. R8-157 R8-163 R8-164 R8-166 R9-17 R9-21

Location

Debris type

Si02

MORB

49,60

MORB

51,37

Mulen

MORB

50,52

Mulen

MORB

Mulen Mulen

Mulen Mulen

Ti02

Al203 Fe203" MnO

1,65

14,43

1,18

16,58

1,47

13,65 15,13

.50,23

1,55

MORB

49,00

1,91

MORB

49,82

MORB

14,34

1,63

14,19

MgO 6,92

CaO 9,01

12,60

0,29

9,34

0,23

5,79

10,62

12,98

0,24

6,59

9,23

11,62

0,26

6,53

9,80

13,05

0,34

6,77

9,93

12,51

0,30

7,21

9,07

Na20

1<20

P205

LOI

Total

v

�

�

�

�

�

�

&

y

li

0,79

99,29

232

203

54

45

20

41

157

35

103

175

308

39

0,60

98,8�

237

257

56

62

3

109

78

6

72

20

1�

38

0,34

99,65

221

346

50

101

3

101

7

88

0,88

99,05

269

314

54

51

23

117

195

39

114

288

114

167

53

11

295

170

43

35

36

149

48

18

22

1,00

0,13

2,84

0,32

0,10

2,76

0,66

0,12

3,59

0,44

0,17

2,48

0,25

0,10 0,21

1,66

98,66

2,86

2,65

1,12

0,13

0,30

0,70

98,66

99,34

234

48,47

2,30

14,35

13,87

0,28

6,63

7,65

2,13

R9-16

Mulen

MORB

53,74

1,40

14,88

10,22

0,23

5,09

8,51

2,73

1,12

0,32

0,13

0,95

98,19

195

R8-155

Mulen

IAT

53,22

0,79

22,86

4,83

0,08

2,32

7,13

4,62

1,07

0,11

1,87

98,89

99

3,81

1,41

1,20

99.40

6,91

4,87

3,71

2,70

'o.o5

1,49

98,47 100,36

R9-15

R8-156 R8-162

Mulen

Mulen Mulen

51,87

0,55

IAT

49,71

0,58

IAT

0,26

3,91

4,80

0,03

17,06

14,48

18,05

10,23 10,10

0,24

5,88

10,10

3,21

0,25

0,13

0,55

0,23

9,04

3,74

1,62

1,34

0,03

3,96

98,42

0,19

6,23

4,13

1,57

0,03

1,91

98,57

6,30

3,98

3,89

1,48

0,07

2,56

7,93

0,03

R8-165

Mulen

IAT

53,64,

0,98

15,28

R9-18

Mulen

boninitic

51,33

0,33

13,58

13,23

52,58

0,28

17,71

10,51

0,17

11,32

3,55

12,56

3,59

76,13

0,29

0,16

12,00

2,21

75,03

0,22

11,36

3,22

71,06

0,27

12.55

5,52

ign./rhyol.

76,09

0,22

11,37

dac./tr.and.

69,24

0,29

14,37

67,05

0,43

16,52

3,77

100,60

0,02

0,35

98,51

0,_03

98,00

0,05

1,39

R8-161 R9-30 R8-167

Mulen Mulen Mulen Mulen Mulen

R8-168

Mulen

R8-169

Mulen

granite (co) granite (co) granite (m) granite (m) 1gn./rhyol.

R9-38

Mulen

dac.ltr.and.

R8-170

Mulen

chert

R8-171

Mulen

chert chert

0,33

0,04 0,06

2,89

26

47

4

374

11

28

27

13

328

165

73

163

10

24

29

14

M8

313

10

24

0,04

0,55

99,86

40

0,07

1,16

99,65

38

5,08

0,80

0,08

0,59

99,82

50

3,98

0,40

0,03

1,44

99,50

33

0,14

0,75

5,20 5,51

18

2,37

0,10

1,44

2,34

2,66

0,04

0,39

2,88

2,56

0,06

0,94

4,55

5,27

0,40

47

1,74

4,50

4,18

0,42

0,23

1,02

98,56

0,07

99,93

69

45

0,27

0,01

0,04

0,08

0,22

0,01

O,Q2

0,23

101,83

0,53

0,02

0,02

0,30

102,05

M H

� U

0,13

101,44 99,76

�

�

0,90

101,53

30

16

0,51 0,63

0,02

0,04

0,02

0,12

1,04

0,08

0,04

0,01

0,26

0,96

1,24

0,07

0,02

0,16

1,52

2,03

0,18

R9-28

Muten

99,65

0,03

0,34

0,01

R9-36

Mulen

chert ang.

92,54

0,07

3,67

0,61

R8-160

Mulen

subarkose?

92,06

0,07

4,26

0,32

0,53

0,50 0,51

0,04

0,01 0,02

0,53

0,30

101,05

29 25

53 23

51

71

48

251

57

22

14

201

27

14

145

32

14 42

12

87

10

39

220

7

18

21

95

24

42

506

13

8

20

306

24

38

161

9

10

216

15

3

47

213

19

29

18

35

174

13

20

35

71

24

43

310

10

202

18

79

19

35

105

9

25

96

11

91

4

23

31

67

17

26

24

231

5

63

18

5

29

26

8

15

17

13

5

12 13

§­

38

186

5

[

10

141

20

38

62

13

231

12

150

11

78

47

82

33

70

16

255

18

218

14

144

69

118

56

10

3

28

11

3

143

55

5

47

9

4

75

48

6

84

6

47

59

23

54

57

17

16

14

�

922

215

12

6

66

13

"' 0 (") � "'

4

1192

23

4

C)

16

12

14

13

8

15

12

15

5

21

21

23

7

25

40

46

Th

35

8

67

10

17

82

18

20

10

12

2,30

1,17

4,32

43 41

11

27

0,58

Mulen

0,59

2,99

9

154

0,04

0,03

R9-27

chert

0,97 1,09

66

78

0,69

0,38

0,22

72,48

R9-24

0,13

66

158

0,44

4,57

75,36

R9-23

77,45

349

7

0,64

granite (co)

232

15

4,65

1,51

Mulen

17

51

1,85

11,24

R8-159

24

42

0,39

2,76

43

100,02

0,01

12,41

25

100,12

0,03

0,18

312

0,70

0,32

0,25

72.69

2

1,15

5,81

54,89

granite (co)

73

0,08

2,88

diorite

Mulen

75

277

0,04

0,61

Mulen

R8-158

40

197

0,63

0,03

R8-154

651

33

0,75

81

480

98,80

0,03

836

1

0,90

1,13

28 17

0,02

3,12

18

36

38

1,98

6,77

25

32

32

98,30

3,04

'5,36

39

10

34

0,58

7,40

0,17

150

550

-47

32

170

0,16.

9,37

566

21

144

9,18

16

15

163

13

116

17,04

345

53

51

161

96,29

0,30

4

12

28

99

110

185

98,59

50,23

260

57

1,30

gabbro

19

34

�

233

128

22

Mulen

147

95

52

R8-153

260

17

130

54

0,20

21

23

61

173

9,95

6

18

24

27

99,23

16,93

43

37

24

63

27

0,26

61

160

71

83

49

120

53,61

45

705

68

52

boninitic

42

171

�

19

178

62

Mulen

2

-�

=n

�

31

175

57

R8-172

17

8

�

250

199 1050

boninitic

14,45

66

48

177

3,94

Mulen

R9-22

60

88

•

5

27

44

39 41

11

50

14 15

11

R9-29

Mulen

subarkose?

91,18

0,11

4,59

1.24

0,30

1,04

1,84

0,21

0,02

0,47

101,02

M

�

17

14

10

39

111

10

R8-151

Mulen

p. mudst.

55,95

0,77

17,42

6,98

0,09

3,08

4,40

4,54

1,54

0,15

3,54

98,47

234

196

31

89

39

130

72

261

24

156

11

36

30

517

13

17

73,74

0,47

10,46

3,83

241

15

89

4

19

26

432

17

10

2,38

99,24

37

0,06

2,49

69

1,51

0,11

62

2,73

1,00

53

2,78

2,69

14

2,05

2,90

151

0,04

1,53

171

5,52

0,04

98,03

231

149

21

69

33

95

65

133

22

103

5

28

32

363

22

9

0,07

1.62

1,94

2,52

1.22

0,09

0,75

100,18

89

54

16

23

84

51

93

24

1529

43

52

47

265

22

26

140

180

38

45

29

105

69

262

34

255

14

90

40

727

29

129

187

33

36

26

97

55

220

29

242

11

72

45

627

30

Mulen

bl. mudsh.

R9-33

Mulen

bl. mudsh.

68,44

0,50

12,01

R8-175

Mulen

x-lam arenite

76,70

1,62

8.81

4,84

R9-31

Mulen

lit.-volc. aren.

R8-152

R9-32

Mulen

tit.-volc. aren.

58,92

0,97

16,04

8,17

0,08

3,38

2,76

4,45

1,63

0,10

1,75

98,24

63,31

0,94

13,96

7,86

0,10

3,12

3.02

3,31

1,34

0,09

1.18

98,22

146

65

f} 1i)

�· s �'1::3

\3

12

39

!:)

'Cj' ....

"" "' ;:s !:> ;:s

�

10

LOI, loss on ignition; Fe.o; , total iron as Fe.03 + 1 . 1 x FeO.

�

A

8 COMENOITE

CD

100

PANTELLEAITE

\ PHONOLITE '

10

0,1

10

- - -� --- � I BASALT / . ., . / --------

0,01

.... .. . . -

;;_�// -��2�-----

' '

............

0

Zr/Ti02

®

RHYOLITE

0

0,1

100

I I

ANDESITE

····-......

ALKALI­ BASALT

SUB-ALKALINE BASALT

�

0,001 /

-l-------+--+---1 ®

1000

10

0,1

0,01

Nb/Y

0,1

,,,,

I I I I

0,01

®

100

ppm Cr

I I I I I I

100

\

\,.rJl

10 ID a: 0 :IE

>< 0 0 a:

D

[]

10

0,1

0

100

IAT

I

\

WPB \

I I I\ I I \ I \ \ I \ :

I\

+------+-'�-���--_L-�

100

10

®

10

0,01

\ \ \

t, t: II

ppmY

'

Trachyandesilic/dacilic debfis

10

'

\

WITHIN PLATE Ti02(%)

I

I \

\

'

1

' I I

0,1

+---..C.:..---1----_____,

0,1

®

100

Granitic and rhyolite debris 10

10

100 ppmZr •

a •

0,1

o

Sr

K

Rb

Ba

Nb

Ce

p

Zr

Ti

y

Cr

•

MORB melabasatts Boninilic melabasalts and gabtxoic rocks tAT metabasatts granitic and rhyolitic debris trachyandesitesl dacites

1000

247

Geochemical data for determining provenance

Table 2. Classification of metabasaltic and gabbroic conglomerate debris according to the discriminant plots of Fig. 4B

Discriminant plot Nb/Y-Zr/Ti02 Y-Cr Zr-Ti02 (%) Total

A

B

c

D

Metabasaltic debris

Metabasaltic debris

Metabasaltic debris

Gabbro/diorite debris

Subalkaline andesite/ basalt MORB/WPB MORB (LATfWPB)

Subalkaline andesite/ basalt LAT (MORB) LAT

Subalkaline andesite/ basalt LAT Boninite

Subalkaline andesite/ basalt !AT Boninite

MORB

LAT

Bonini tic

Boninitic

MoRB, mid ocean ridge basalts; WPB, within plate basalts; LAT, island arc tholeiites.

(Saunders & Tarney, 1979; Basaltic Volcanism Study Project, 1981; Hickey & Frey, 1982; Cameron et al., 1983; Thompson et al., 1984; Crawford, 1989; Wilson, 1989). This tripartite subdivision is also suggested by the discriminant plots of Fig. 4B. The debris of phaneritic, mafic volcanic rocks (gabbro and diorite) show a similar chemical composition to the boninitic samples, as suggested by their nearly identical MORB normalized, multi-element variation diagram patterns (Fig. 4A) and the discriminant plot of Fig. 4B. The gabbro and diorite are interpreted accordingly to represent the plutonic equivalents of the boninitic metavolcanics (see further discussion below). Although the value of geochemical analyses of acidic volcanic rocks has been questioned (e.g. see discussion in Cas & Wright, 1988, and references therein), the chemical composition of the trachy­ andesite/dacite cobbles testifies to their alkaline affinity (Fig. 4) and demonstrates their incompati­ bility with the granitic and rhyolitic rock debris. The rhyolitic and fine- to coarse-grained granitic clasts display a similar mineralogy (particularly diagnostic is the presence of blue quartz in both types) and nearly identical chemical compositions, and are interpreted to represent extrusives and their intrusive equivalents. Their geochemical signature (e.g. pro­ nounced negative Nb, P and Ti anomalies, see Fig. 4A) indicates derivation from a depleted tAT parent (Pearce et al., 1984; Holm, 1985), and suggests erosion of low-K granitic rocks, typical of island arcs (Ewart, 1982). Moreover, the rhyolitic debris exhibits a mineralogical and chemical composition

close to that of the rhyolite and rhyolitic ignimbrite olistoliths, thereby suggesting a common protolith (see interpretation below). Bedded chert occurs as large angular blocks up to 1. 7m across. Angular cobbles of 'quartzite' are, based on mineralog.ical and geochemical criteria, interpreted to represent disrupted chert beds derived from the ribbon chert. The subarkosic debris are geochemically similar to chert, but differ with respect to their coarser grain size and higher proportion of plagioclase. However, detrital sodic plagioclase locally is present in some chert beds of the larger bedded chert olistoliths, suggesting that a common parent cannot be ruled out. Well-rounded quartzite clasts are recrystallized completely, and their protolith (whether representing vein quartz, quartzitic sandstone or chert, see below) cannot be determined. Well-rounded cobbles of lithic arenites differ mineralogically and geochemically from the angular debris of lithic-volcanic arenites. The lithic arenites consist, in decreasing order, of quartz, lithic sedi­ mentary grains (quartzite/quartz schist), sodic plagioclase, alkali feldspars, ilmenite, zircon and minor apatite and opaques. Modal ap.alyses indicate either a continental craton or recycled orogen/ collision orogen as provenance (Dickinson, 1985; Valloni, 1985), as is also suggested by their chemical composition. Lithic-volcanic arenites, on the other hand, are texturally, mineralogically and geochemi­ cally similar to the olistoliths of shallow-marine sandstones and deep-marine turbidites (see description and interpretation below), and are

Fig. 4. (Opposite.) (A) MORB normalized, multi-element variation diagrams of conglomerate debris of volcanic rocks.

(1) MORB metabasalts, (2) LAT metabasalts, (3) boninitic metabasalts and debris of gabbroic rocks, (4) trachyandesites/ dacites, and (5) granites and rhyolites. The order of the elements and the values of the normalizing constants are from Pearce (1983), except for Cr, which is from Pearce (1980). (B) Discriminant plots for conglomerate debris of ( 1) volcanic rocks, and (2) and (3) mafic volcanic rocks. Discriminant plot (B . 1 ) is from Winchester & Floyd (1977) , whereas discriminant plots (B.2) and (B.3) are from Pearce ( 1980).

248

R. Ravnas and H. Fumes

accordingly interpreted to have been derived from such deposits. Angular clasts of pebbly mudstone and black shales are geochemically similar to the olistoliths of pebbly mudstone and interbedded black shales of the melange succession, respectively. Furthermore, black shales are present only in conglomerate deposits resting with erosive basal contacts on black shale, thus suggesting the intrabasinal provenance of this clast type. Well-rounded pebbles of calcareous rocks have not been analysed geochemically. Although an extrabasinal origin of this clast type cannot be excluded, they may have been produced by erosion of lava-flow basalts with calcite-filled interstices now present as large olistoliths in the melange. Olistoliths of volcanic rocks

Geochemical analyses of samples from the different olistoliths of lava-flow basalts and andesites (Table 3) show considerable variations, allowing subdivision into alkaline, mildly alkaline to subalkaline MORB­ like and island arc basalts (Fig. 5 & Table 4). In addition there are some transitional types. Isotope analyses were performed on samples from olistoliths of ignimbrite and alkaline and subalkaline MORB-like basalts/andesites, and the results are presented in Table 5. The Es, and ENct values are calculated for several ages (470, 450 and 430Ma) based on the interpretation that the Kalvag Melange forms part of the Caledonian island arc/marginal basin terrains (see discussion below). The high E5, values of the lava-flow basalts/andesites indicate contamination, possibly from (a) syn- or post-lithification hydro­ thermal solution(s) from which the calcitic infills of vesicles and interstices precipitated (Ravnas, 1991). As the volcanic-rock olistoliths occur scattered in the olistostromal groundmass and do not form parts of a continuous succession, the geochemistry of the different volcanic-rock olistolith types are described separately, followed by a brief discussion of their possible primary tectonic setting. Alkaline basalts comprise subaerial compound flows of pillowed and scoriaceous, porphyritic and vesicular, greenschist facies metabasalts. Phenocrysts are augite/diopside, saussuritisized plagioclase, primary sanidine and microcline pseudomorphs, probably after leucite. Calcite-filled amygdules are common. The chemical composition, characterized by high content of alkali/alkaline earth metals, Ti02, Zr and Nb, denotes a strong alkaline affinity of these

metabasalts. The high K content, together with the inferred primary texture and mineralogy jointly suggest that these rocks belong to the potassic ig­ neous rock suite (Foley et at., 1987; Wilson, 1989), probably representing shoshonites transitional into leucite-tephrites (Bates & Jackson, 1987). The original tectonic setting of these metabasalts is difficult to assess. The high contents of Ti02 and Nb, and to a lesser extent Zr, typical of ocean island basalts (O!B), argue against a subduction-related setting (e.g. Perfit et al., 1980; Morris & Hart, 1983). However, high Ti02 and Nb values are reported from alkaline basalts of back-arc, intraplate settings undergoing extension (e.g. Thorpe et a/., 1984; Briggs et a/., 1990) and subduction-related ultrapotassic rocks (Thompson, 1977; Edgar, 1980; Thompson eta/., 1984). The MORB normalized, multi­ element variation diagrams resemble closely the pattern shown by continental flood-basalts (CFB) (e.g. Thompson et a/., 1984; Wilson, 1989), and differ from those of OtB by their lack of pronounced positive Nb anomalies. Provided these metabasalts represent potassic rocks, the chemical composition indicates a transitional character between lamproites and potassic/ultrapotassic rocks of active orogenic zones (i. e. group I and III potassic rocks of Foley et at. (1987)), whereas the content of Nb relative to Zr strongly suggests a subduction-related origin (see discussion in Thompson & Fowler, 1986; and their Fig. 9). The calculated ENct value (Table 5) coincides with values reported from both OtB, potassic rocks, island arc and active continental margin volcanics., as well as some rift-related continental flood-basalts (e.g. Faure, 1986; Wilson, 1989; and references therein). Accordingly, rift-related, continental within-plate or back-arc, or alternatively subduction·· related potassic volcanics provide the best analogues for these alkaline metabasalts. Subalkaline MORB-like basalts are present as com·· pound accumulations of greenschist facies aphyric to slightly phyric, vesicular, subaerial aa and pahoehoe, and subaquatic, pillow lava-flows. The flow-units range in composition from basaltic to andesitic, with the pahoehoe- and pillowed-type flows representing the former. Saussuritisized plagioclase and abundant calcite-filled amygdules are present in all flow types. Despite some variations in geochemical signatures between different olistoliths (see Fig. 5 and dis·· cussion below), all metabasalts classify as sub·· alkaline MORB-like (Fig. SB, Table 3). The MORB normalized, multi-element variation diagram pat­ terns (Fig. SA), however, indicate a slight alkaline

Table 3. Representative geochemical analyses of volcanic rock olistoliths Si02_

Ti02

Al203 Fe203"

Alk.basalt

47,25

2,88

21,65

9,31

0,18

3,88

7,91

Alk.basalt

51,83

2,47

19,68

5,16

0,18

2,35

13.75

Sleneset

Alk.basalt

52,36

2,49

19,67

5,49

0,23

2,19

13,69

0,96

R9-13

Sleneset

Alk.basalt

52,60

2,59

19,97

4,96

0,20

2,42

9,92

R9-14

Sleneset

Alk.basalt

53,62

2,33

18,05

5,23

0,18

2,52

10,33

R9-52

Kalvegj.n.

MORB-Iike

53,74

2,30

23,19

6,27

0,06

3,63

4,49

2,94

R9-53

Kalvegj.n.

MORB-Iike

54,56

2,16

22,84

6,27

0,06

3,42

4,40

3,49

R9-54

Kalvegj.n.

MORB-Iike

58,44

1,96

19,64

8,74

0,13

3,99

0,00

0,88

R9-55

Kalvegj.n.

MORB-Iike

53,44

2,17

22,65

7,71

0,12

3,69

2,74

3,19

3,86

R9-56

Kalvegj.n.

MORB-Iike

57,57

2,00

20,74

7,00

0,21

3,47

2,33

2,97

3,23

RB-48

Botnan.

MORB-Iike

50,13

1,60

18,15

12,54

0,16

7,02

6,17

1,84

RB-50

Botnan.

MORB-Iike

52,64

1,58

19,60

10,12

0,08

5,90

3,93

4,99

RB-51

Botnan.

MORB-Iike

51,15

1,59

18,55

12,04

0,11

8,40

3,31

4,04

0,64

RB-54

Botnan.

MORB-Iike

52,62

1,64

20,32

10,58

0,20

4,30

3,04

3,51

3,50

R9-58

Botnan.

MORB-Iike

50,33

1,72

19,39

11,05

0,10

8,12

4,06

3,77

R9-59

Botnan.

MORB-Iike

51,90

1,53

18,49

9,50

0,10

6,51

6,29

4,08

R9-60

Botnan.

MORB-Iike

54,05

1,63

19,80

8,94

0,06

5,69

4.97

3,56

Sample no.

Location

Rock type

RB-8

Sleneset

RB-9

Sleneset

R9-12

MnO

MgO

CaO

Na20

v

�

�

N

�

239

�

�

&

v

�

m

45

70

35

21

100

42

87

32

125

57

106

26

89

13

101

25

72

12

93

158

101

209

47

376

49

119

41

371

37

104

83

146

127

265

36

113

17

121

98

591

38

129

18

120

76

514

42

105

�

�

154

143

680

56

567

30

175

37

34

519

29

174

257

47 35

74

456

31

202

37

101

409

34

180

36

�

�

70

34

841

40

54

637

20

34

36

550

33

56

41

870

36

38

41

905

14

17

26

34

17

29

41

167

14

17

23

29

546

10

20

26

45

510

17

17

46

307

26

13

309

32

24

178

K20

P205

LOt

Total

1,80

4,61

0,54

1,00

100,01

272

0,95

3,36

0,49

1,13

100,24

193

188

2,53

0,51

1,93

100,11

199

178

1,71

5,02

0,53

1,94

99,92

230

198

0,89

6,41

0,63

1,52

100,18

199

169

3,08

0,30

7,31

100,01

2,46

0,31

6,69

99,98

286

439

30

128

5,86

0,38

9,40

100,01

255

455

41

160

0,45

7,80

100,02

268

409

34

178

0,50

12,54

100,03

238

430

28

224

2,27

0,14

3,26

100,03

283

28

52

21

20

123

40

316

36

95

0,96

0,21

3,70

100,00

210

647

52

277

84

107

20

407

29

87

0,20

4,70

100,03

229

633

60

266

74

114

12

338

26

90

29

12

113

0,31

7.87

100,01

249

437

49

239

86

122

78

404

38

102

11

34

38

431

1,24

0,21

3,42

99,99

243

742

59

306

52

124

25

380

29

98

10

36

20

220

9

1,39

0,21

1,33

100,01

198

560

47

212

56

99

29

388

26

87

35

20

254

10

0,85

0,41

6,50

99,96

216

574

41

247

67

108

16

446

30

84

9,

10

19

27

160

247

292

440

34

38

35

111

241

15

0

!}

Mulen

Trans. bas.

51,24

1,12

18,87

7,55

0,35

3,78

15,61

0,75

0,12

0,28

3,84

99,66

156

27

41

38

88

3

454

26

48

4

20

28

61

20

Minnet

Trans. bas.

50,60

1,91

18,59

12,07

0,24

9,17

2,78

3,22

1,11

0,31

2,33

100,00

229

302

59

93

28

104

40

446

29

157

10

70

11

787

11

R9-89

Mulen

Trans. bas.

55,44

0,99

15,78

10,71

0,29

5,10

6,53

4,60

0,25

0,27

1,99

99,94

189

257

47

85

48

110

5

26

81

6

34

27

264

12

R9-90

Mulen

Trans. bas.

53.15

1,21

18,05

10,09

0,26

5,64

6,06

4,29

0,87

0,29

1,07

99,91

203

290

43

56

88

89

24

221· 361

27

103

35

27

1251

0,53

0,25

0,72

100,01

168

235

47

94

16

104

15

367

25

64

30

15

117

17

384

51,00

1,03

tAT

49,09

1,09

tAT

50,45

1,01

Sjeneset

tAT

50,63

1,14

Sjeneset

tAT

49,63

1,04

tAT

RS-19

Sj0neset

RB-20

Sjeneset

RB-21 R9-1

!.:)

� ..., � �

� ;:;: �-

·

5,45

9,69

2,95

0,24

6,35

4,27

1,89

4,00

0,19

6,71

100,03

262

77

53

32

63

141

111

447

20

55

24

0,26

7,13

5,82

1,76

2,78

0,28

4,53

100,06

242

54

58

26

39

131

82

329

19

44

36

21

264

8

11.42

0,22

6,56

7,22

3,13

1,26

0,27

5,37

100,09

239

50

49

25

50

130

32

475

22

51

23

16

195

13

11,39

0,22

6,74

7,05

2,49

1,84

0,21

4,19

100,02

234

64

50

31

44

140

78

248

17

48

27

30

448

1

1,22

2,15

0,01

2,70

98,62

28

2

18

201

63

140

94

400

59

47

1742

22

11

682

10,67

0,30

21,03

11.90

17,66

12,93

18,25 19,41

18,14

� s. �

[

RS-95

Sjeneset

CJ

::!

RB-223

RB-18

Th

3,21

0,28

0

�

;j

Mulevik

lgn.matrix

68,52

0,24

15,37

4,89

0,03

RS-103

Mulevik

lgn.matrix

70,92

0.23

14,61

3,45

0.03

1.95

0,65

4,70

1,18

0,02

1,55

99,30

27

4

13

155

45

174

4

50

49

24

18

RS-104

Mulevik

lgn.matrix

70,87

0,23

14.53

3,52

0,03

2.00

0,64

3,96

1,18

0,02

1,49

98,47

28

1

10

5

101

43

146

62

383

5

42

46

719

17

11

RB-105

Mulevik

lgn.matrix

70,05

0,22

14,14

4,91

0,02

3,16

0,28

2,76

1,24

0,02

2,08

98,87

30

2

18

4

161

35

215

115

390

10

70

58

841

22

R9-84

Mulevik

lgn.matrix

71.70

0,20

12,73

4,84

0,02

3,24

0,97

4,81

0,67

0,02

1,83

101,40

29

8

18

6

179

16

378

81

348

5

25

12

61

37

48

217

RB-102

R9-82

Mulevik

Rhy. core

RB-40

Mulevik

Rhyolite

RB-98

Mulevik

Ash-layer

R9-85

Mulevik

Ash-layer

83,54

0,13

7,98

2,56

0,02

0,41

0,14

4,20

0,02

0,01

0,56

99,57

73,18

0,23

14.18

2,02

0,02

0,31

0,46

7,59

0,04

0,02

0,65

�8.70

66,58

0,22

15,07

6,14

0,05

3,63

0,35

1,40

2,87

0,02

2,56 - 98,88

65,47

0,24

16,36

6,04

0,05

4.07

0,29

1,12

3,74

0,00

2,75

100,20

4

47

353

49

48

395

16

19

40

20

11 27

65

73

352

68

66

16

28·-

24

155

80

101

57

392

54

34

3319

10

28

23

159

106

87

72

414

46

42

4180

4

7

30

86

LOI, loss on ignition; F�o; , total iron as Fe203 + 1 . 1 x FeO; Listed values of major oxides of mafic volcanic rocks are recalculated values to 100% at LOI

=

"" � ;::s !:) ;::s

�

11

0.

�

250

R. Ravnas and H. Fumes

A

100

10

Q)/>;�

/'/-

. .

.

•

�� c.< '< c�

COMENOITE

B

Olislolilh Slenesel- Alkaline

PANTELLERITE

I i

\PHONOLITE

\

'

' '............

0.1

- ... ... .. ZrfTI02

0.1

100

®

0,01

Olistolith Kalvegjerdsneset

•

A NDESITE

Subalkaline MORB-Iike (KO)

--�

I BASALT
-- -,t- --

� ., ,� �

ALKALI· BASALT

SUB-ALKALINE BASALT 10

0.001

+------+---1---i

O.Ql

10

0.1 Nb/Y

®

1000 0.1

100

®

,/'"'

I

Olislolilh B otnanesesl- Subalkallne MORB-Iike (BO)

, , , ,

10 ppmCr

I I I

\I

100

� •

.t.6

1

' '

m a: 0 :::E

100

I ' I I I I

'

0.1

-" ()

0 a:

I

G)

10

IAT

f, 1: h

WPB

I\

I\ • I \

\ I I

I

I I ', I \ :I

+-------��---'�-�--�---i '

100

10 10

ppmY

®

10

----- ......

100

'

\ I

I

0.1

® Olistolilhs of IAT metabasahs

TI02(%)

\ \ \ \ I

1

'

0.1 100

10

ppmZr

•

0

01. alkaline metabasalts 01. subalkaline/MORB-Iike metabasalts (KO) 01. subalkaline/MORB-Iike metabasalts (80)

0.1

O.Ql +--+---+--+--1----'< Ti y Cr p Zr Rb Be Nb Ce Sr K

o

01. transitional metabasalts

•

01. IAT metabasalts

1000

2S1

Geochemical data for determining provenance

Table 4. Classification of metabasalts of mafic volcanic rock olistoliths according to the discriminant plots of Fig. 58

Discriminant plot

B

Alkaline

Subalkaline basalt

MORB/WPB WPB (MORB)

MORBIWPB MORB/WPB

WPB

MORB?

Nb/Y- Zr/Ti02 Y - Cr Zr-Ti02 (%)

c

A

Total

D

E

Subalkaline andesite/ Subalkaline andesite/ Subalkaline andesite/ basalt basalt basalt !AT MORB/WPB MORB/WPB !AT (MORB) All MORB (!AT) MORB?

!AT

Transitional

A, Large olistolith of alkaline metabasalts exposed at Sleneset; B, large olistolith of subalkaline MORB-Iike metabasalts exposed at Kalvegjerdsneset (KO); C, large olistolith of subalkaline MORB-like metabasalts exposed at Botnaneset (BO); D , smaller olistoliths of transitional metabasalts; E , smaller olistoliths of !AT metabasalts. Table 5. Isotope data from volcanic-rock olistoliths

470 Ma Sample AI kaline basalt MORB-like (80) basalt MORB-like (BO) basalt andesite Ignimbrite matrix

450Ma

430 Ma

ENd

Esr

EN d

Esr

ENd

Esr

7.08 8.13 4.14 7.08

71.26 1 87 . 8 6 69.43 -50.81

6.82 7.99 4.06 6.9 6

72.38 1 87 . 87 70. 65 -45.55

6.5 6 7.86 3.98 6.74

73.49 1 87 . 89 7 1 . 88 -40.3

BO, large olistolith of subalkaline MORB-like metabasalts exposed at Botnaneset.

character, more akin to om or CFB. The lack of positive Nb anomalies and in som e cases slight relative depletion ofNb (Fig. SA), and the relatively low values of Ti02, P205, Zr and Y jointly favour a CFB-type rather than an OIB (Thompson et al., 1984; Wilson, 1989 ) . Low values of Ti, Nb, Zr and Y are generally attributed to crustal contamination or reflect subduction-related magma genesis (Perfit et al., 1980; Saunders et al., 1980; Cox & Hawkes­ warth, 1984; Thompson etal. , 1984). Moreover, the subalkaline MORB-like metabasalts also display geo­ chemical characteristics similar to subduction-related potassic and high-K calc-alkaline basalts (Foley et al. , 1987; Wheller et al. , 1987; Stolz et al. , 1990). The isotope data from a basaltic pahoehoe (ENct = 8) and an andesitic aa (ENd = 4) flow-unit are indicative of MORB/back-arc basin basalts and oceanic island arc/some continental rift-related (northern Basin

and Range) basalts respectively (Faure, 1986; Wilson, 1989) . The lower ENd value of the latter may, how­ ever, indicate that the more felsic volcanics have experienced higher degrees of crustal contamination compared with the basaltic types (Faure, 1986 ) . Transitional basalts chiefly consist of scoriaceous and subordinate massive, aphyric or porphyritic aa metabasalts and basaltic meta-andesites. Rare phenocrysts are saussuritisized plagioclase and FeTi­ oxides. Calcite- and minor quartz-filled amygdules are common, and locally small sandstone xenoliths are present. Similar to the alkaline and subalkaline MORB-like basalts, the transitional types also display some geochemical signatures that are characteristic of alkaline OIB and/or MORB (Fig. S). However, the pronounced negative Nb anomalies (Fig. SA) strongly argue for a subduction-related setting. Although most geochemical evidence favours an

Fig. 5. (Opposite. ) (A) MORB normalized trace-element diagrams for volcanic rock olistoliths. See Fig. 2 for location and

stratigraphical position of different olistolith types of mafic volcanic rocks. (1) alkaline metabasalts, (2) subalkaline MORE­ like metabasalts of olistolith located at Kalvegjerdsneset (higher stratigraphical position compared to subalkaline MORB-like metabasalts of olistolith located at Botnaneset), (3) subalkaline MORB-like metabasalts of olistolith located at Botnaneset, (4) transitional type metabasalts, (5) !AT metabasalts, and ( 6) acidic volcanic rocks. The order of the elements and the values of the normalizing constants are from Pearce ( 1983) , except for Cr, which is from Pearce (1980) . (B) Discriminant plots for olistoliths of ( 1 ) volcanic rocks, and (2) and (3) mafic volcanic rocks. See Fig. 2 for location and stratigraphical position of different olistolith types of mafic volcanic rocks. Discriminant plot (8 . 1 ) is from Winchester & Floyd ( 1977) , whereas discriminant plots (B .2) and (8.3) are from Pearce ( 1980).

2S2

R . Ravnas and H. Fumes

original back-arc setting, some element abundances and element ratios (e.g. high AI, low Ni, low FeO/ MgO) are more characteristic of calc-alkaline basalts and island arc tholeiites. Basalts associated with incipient back-arc rifting or initial stages of back-arc spreading are thus thought to be the best analogues of these m etavolcanics (Bruhn et at., 1978; Saunders etal., 1979; Weaver et at., 1979; Tarney eta/., 1982). Island arc basalts (IAT) comprise scoriaceous, porphyritic aa flows with phenocrysts of saussuritis­ ized plagioclase and minor FeTi-oxides. The MORB normalized, multi-element variation diagrams (Fig. SA) show typical IAT character with negative Nb anomalies and relative depletion of Ti, Zr, Y and Cr, thus testifying the subduction-related (i.e. island arc or active continental margin) affinity of these metabasalts (see also Fig. SB). Acidic volcanic rocks comprise massive, por­ phyritic rhyolite and rhyolitic ignimbrite, with an inferred co-ignimbrite ash layer. Phenocrysts or crystaloclasts are quartz (commonly blue), zoned, twinned and 'chessboard' sodic plagioclase, sanidine, myrmekite, quartz/alkali-feldspar spherulites and inferred glass shards. In addition there are pumice and lithic fragments present in the ignimbrite. Geo­ chemical analyses of the ignimbrite matrix, rhyolitic core of individual ignimbrite flows , ash-layer and massive rhyolite, suggest that these acidic volcanic rocks are co-magmatic (Fig. S; Ravnas, 1991). The geochemical and isotope composition of these silicic rocks favour a subduction-related original tec­ tonic setting (Tarney et a/., 1977; Saunders et a/., 1979; Pearce et a/., 1984; Faure, 1986; Wilson, 1989). The pronounced depletion of Nb, P and Ti (Fig. SA) indicate derivation from a depleted IAT source (Pearce et a/., 1984; Holm, 198S), whereas positive anomalies of Zr and Y are more typical of silicic rocks from tholeiitic associations (Pearce et a/., 1984, and references therein). The isotope com­ position, which is close to that of MORB (Table S), excludes magma generation by partial melting of continental crust and suggests little influence of melted sediments in the magma source region (Faure, 1986). The acidic volcanic rock olistoliths are, accordingly, interpreted to represent island arc/ active continental margin volcanics erupted at a considerable distance from the trench, i. e. behind the arc with respect to the trench. Alternatively, these volcanics may have been derived from a slightly contaminated MORB source, representing oceanic back-arc, rift-related acidic volcanics (Gill et a/., 1984).

As noted above, the mineralogical and geo­ chemical similarities between the granitic and rhyolitic conglomerate debris and the olistoliths of acidic volcanic rocks suggest a common protolith. The higher K, Rb, Zr, Nb, Y, and Th content of the rhyolitic rocks and the higher plagioclase content of the granitic rocks may accordingly reflect fractional crystallization, and particularly plagioclase fraction­ ation (Taylor et a/., 1968; Saunders et al., 1979). Notably, the different volcanic-rock olistolith types are present at certain stratigraphical levels. Olistoliths of IAT lava-flow basalts and acidic volcanic rocks are found at a lower stratigraphical level, although generally above 'mega-olistoliths' of terrigenous sediments, whereas transitional and subalkaline MORB-like metabasalts of increasingly alkaline affinity are found at successively higher stratigraphical levels (Fig. 2). However, the olistolith consisting of alkaline metabasalt makes an important exception. Whereas volcanic-rock olistoliths gener­ ally are found in the upper part of the melange succession, the olistolith of alkaline metabasalt occurs in the lower part of the melange succession in an interval containing olistoliths of rhyolite, bedded chert, pebbly mudstones and shallow-marine sediments. Most major and trace elements of the mafic vol­ canic rock olistoliths seem to display 'gradational' variations from alkaline through subalkaline MORE­ like to transitional types, and finally to !AT-type basalts, particularly demonstrated by their Fe, Mg, K, Ti, Nb , Zr and Y content. All metabasalt types have high AI and similar Rb, Ba and Sr contents (Fig. SA & Table 3). Moreover, all volcanic rock olistoliths, including the alkaline metabasalts, show evidence of a primary subduction-related setting. These points suggest that the protoliths of the mafic volcanic rock olistoliths may be genetically related and represent parts of a continuous volcanic series. Below it is argued that the volcanic rock olistoliths represent parts of a volcanic succession in which the alkaline and the IAT types represent the recorded end-members. Olistoliths of sedimentary rocks

Geochemical analyses have been performed on samples from all types of sedimentary-rock olisto­ liths. However, only the lithic-volcanic arenites from the olistoliths of shallow-marine sediments and deep­ marine sandy turbidites will be discussed in some detail here.

Geochemical data for determining provenance The lithic-volcanic arenites consist of angular to subrounded quartz (blue or colourless), plagioclase (saussuritisized, zoned, twinned or 'chessboard'), microline, sanidine, lithic-volcanic grains, lithic­ sedimentary grains, minor iron oxides, ilmenite, calcite (locally dolomite), apatite, epidote and accessory sphene and zircon. The lithic-volcanic grains represent massive or foliated metabasalts, m eta-andesites, acidic volcanics and minor myrme­ kite. The lithic-sedimentary grains represent chert or quartz schist and minor inferred intrabasinal sand­ stones and mudstones. Interbedded pebbly sand­ stones and polymictic conglomerates show a similar gravel admixture to the channel-fill conglomerates. Bhatia ( 1 983) and Bhatia & Crook ( 1986) sug­ gested a simplified plate-tectonic classification of continental m argin and oceanic basins based on the nature of the crust from which the sediments were derived. Four types of plate tectonic setting were recognized: oceanic island arcs, continental island arcs, active continental margins and passive margins. The geochemical signatures of the lithic-volcanic arenites (Table 6) clearly suggest derivation from island arcs, either continental or oceanic (Fig. 6). According to the classification schemes of Bhatia ( 1983) and Bhatia & Crook ( 1986) most m ajor elements favour a continental ist'and arc as prov­ enance, whereas most trace elements and ratios argue in favour of an oceanic arc. The 'oceanic island arc' setting includes sedimentary basins adjacent to intra-oceanic island arcs or island arcs partly formed on thin continental crust, whereas the 'continental island arc' type includes sedimentary basins adjacent to island arcs formed on well­ developed continental crust or on thin continental margins (Bhatia, 1983 ) . A comprehensive data set on the geochemistry of detritus from an intra-oceanic arc was provided by Hiscott & Gill ( 1992) . The lithic-volcanic arenites of the sedimentary rock olistoliths of the Kalvag Melange differ from the volcaniclastic deposits discussed by Hiscott & Gill ( 1 992) by their higher content of Nb, and, although less pronounced, also Zr, La and Th, which is here interpreted as the signature of the addition of alkaline and/or continental source rocks (see below). A m agmatic arc (Dickinson, 1985 ; Valloni, 1985) is also suggested as the source for these arenites by the modal analyses, although such analysis is in this case hampered by post-depositional alteration effects, e.g. metamorphic recrystallization and subsequent partial obliteration of primary grain boundaries. Further constraints on the provenance are

253

provided by the framework grains and the gravel population of the interbedded pebbly sandstones and conglomerates. The type of framework grains and the polymodal distribution of the gravel popu­ lation of most debris types strongly suggest that volcanic and sedimentary rocks now present as vol­ canic rock olistoliths and extrabasinal conglomerate debris of the channel-fill deposits formed a m ajor source for these lithic-volcanic arenites. The geo­ chemical composition of the lithic-volcanic arenites in fact can be explained by mixing of detritus derived by erosion of such rocks alone (Fig. 7). Moreover, the deviations in element and element ratio ranges compared with those suggested for basins of island arc settings (Bhatia, 1983; Bhatia & Crook, 1986; see Table 6) should be expected according to this interpretation (Fig. 7). Alternatively the apparent andesitic composition of the lithic-volcanic arenites (Fig 7A), may suggest that andesitic igneous rocks originally formed a larger volumetric proportion of the provenance region. Notably, there appears to be no systematic stratigraphical variations in mineral­ ogical and/or geochemical composition, suggesting that the different source rock types were all present in the provenance region(s) throughout the life-span of the depositional basin. Groundmass deposits

The m atrix of the olistostromal groundmass deposits is geochemically similar to the m udstones in the olistoliths of shallow-marine deposits. The m atrix is, however, enriched in K, Fe, Mg, V, Cr, and to a lesser degree Zr and Nb, and depleted in Si, Ca and Na compared with the lithic-volcanic arenites. These effects are attributed to a larger volumetric pro­ portion of original m ud-sized p articles (clays and m icas) in the groundmass deposits (Roser & Korsch, 1986; Argast & Donnelly, 1987). This supports the interpretation that the finer grained, more distal mem bers of the shallow-marine deposits provided most of the m aterial of the melange' s groundmass (Ravnas, 1 99 1 ) . The enrichment of fines in the olisto­ stromal matrix can be ascribed to remoulding, during which most of the sandier deposits remained as clasts/olistoliths, or, perhaps more likely, to incor­ poration of shelf and slope mud into the ambient mud slurries or debris-flows from which the olisto­ stromal 'groundmass' was deposited. The 'matrix' of the channel-fill conglomerates and the gully- and channel-fill turb idites is, on the other hand, mineralogically and geochemically similar to

N Ul .,.

Table 6. Representative geochemical analyses of sedimentary-rock olistoliths Sample no.

location

Si02

Olistolith type

Ti02

Al203 Fe203"

MnO

MgO

CaO

Na20

1<20

P205

LOI

Total

v

Cr

Co

Ni

Cu

Zn

Rb

Sr

y

Zr

Nb

Ce

Nd

Ba

La

Th

109

144

30

34

36

91

87

144

35

242

13

89

45

361

37

10

1 04

154

25

35

10

79

52

166

28

203

11

49

46

458

24

151

15

42

27

78

0

220

41

203

10

24

50

27

36

13

135

27

39

23

90

65

211

29

228

13

56

48

305

29

19

12

91

R8-1 1 0

Annevika

Shallow marine sed.

63,41

0,84

14,92

6,99

0,09

3,81

4,16

2,90

1 ,84

0,11

1 ,73

1 00,81

R8-124

Annevika

Shallow marine sed.

65,97

0,87

1 1 ,93

6,19

0,09

4,62

6,09

3,13

1,13

0,12

0

1 00,84

R8-1 1 5

Annevika

Shallow marine sed.

54,64

0,91

9,60

5,21

0,21

3,80

18,09

2,16

0,04

0,17

6,65

101 ,47

6,58

0,08

4,83

5,17

2,83

1 ,53

0,12

0,82

1 00,96

5,12

10,17

1 ,59

2,36

R8-1 1 9

Annevika

Shallow marine sed.

64,55

0,88

R8-1 1 7

Annevika

Shallow marine sed.

R8-122

Annevika

Shallow marine

R8-127

Annevika

R8-129 R9-10

13,56

)1

94

105

3

56,85

0,80

12,63

5,81

1 00,87

96

137

20

46

103

345

29

209

13

54

42

328

38

11

64,39

0,91

1 3,07

6,41

0,08

4,67

4,97

2,83

1 .88

0,13

0,80

100,14

1 04

143

26

36

22

85

81

202

32

282

14

70

39

373

39

6

65,76

0,96

1 1 ,83

7,54

0,10

4,86

4,57

2,35

1 ,70

0,12

0,73

100,51

134

193

31

37

20

79

84

199

32

211

7

50

27

,244

29

8

Annevika

sed. Shallow marine sed. Shallow marine sed.

67,43

0,89

12,07

6,55

0,09

4,03

3,92

3,44

1 , 34

0,13

0,75

100,64

107

153

26

35

25

78

67

200

29

262

12

84

44

259

44

10

Varenesvika

Shallow marine sed.

64,63

0,88

14,27

6,88

0,07

3,76

2,10

3,24

2,35

0,12

1 ,50

99,82

116

140

30

30

21

88

109

117

31

252

14

83

33

376

39

16

19

93

107

117

36

11

5,28

0,11

0,14

R9-11

Varenesvika

Shallow marine sed.

63,40

0,96

14,65

7,01

0,08

4,01

2,80

3,51

2,23

0,13

1 ,48

1 00,27

114

136

30

30

300

15

106

51

439

45

R8-213

Varenesvika

Shallow marine sed.

64,42

0,86

12,95

6,16

0,09

3,93

6,57

2,62

1 ,84

0,13

0,90

1 00,48

101

112

23

29

23

83

85

138

34

242

13

75

50

369

48

16

R8-21 1

Varenesvika

Shallow marine sed.

63,71

0,94

1 5,26

6,98

0,08

3,88

2,25

3,03

2,42

0,15

1 ,64

1 00,33

117

139

31

34

20

95

118

111

38

296

17

101

50

439

41

15

R9-7

Varenesvika

Shallow marine sed.

68,01

0,86

12,90

6,73

0,08

3,67

2,17

1,60

1 00,23

117

145

29

29

22

81

79

108

31

252

14

86

46

214

35

8

R9-4

Varenesvika

Shallow marine sed.

67,63

0,76

1 3,02

6,54

0,09

3,78

2,71

2,95

1 ,43

0,10

1 .46

1 00,48

112

143

27

33

33

83

70

127

27

184

8

80

43

255

41

9

4,15

4,75

0,50

0,10

0,97

1 00,51

132

185

31

35

18

71

20

179

25

126

7

42

30

123

12

13

99,44

147

40

135

28

2,46

1 ,63

0,11

R8-31

Langeneset

Deep marine turb.

68,45

0,86

1 1 ,13

7,95

0,13

3,85

R8-32

Langeneset

Deep marine turb.

64,95

0,96

1 1 ,93

7,95

0,13

4,15

4,75

2,82

0,12

0,88

201

34

39

22

75

5

38

20

168

15

12

R8-33

langeneset

Deep marine turb.

68,06

0,71

12,76

6,32

0,07

3,54

2,13

2,93

2,11

0,13

0,81

99,57

113

145

28

35

19

76

97

125

27

170

9

40

31

301

12

16

7,04

0,09

3,81

2,40

3,59

2,28

0,14

1 ,03

99,99

121

158

31

39

22

93

1 02

121

38

326

15

76

38

385

35

18

7,27

0,08

3,83

2,44

4,29

2,25

0,14

1 ,21

99,04

126

159

31

37

28

94

99

122

36

303

16

72

45

455

31

4

0,07

3,44

1 00,08

112

147

26

29

21

74

92

129

24

160

8

38

34

299

20

11

111

R8-34

Langeneset

Deep marine turb.

64,36

0,99

14,27

R9-61

langeneset

Deep marine turb.

62,04

1,00

14,49

Deep marine turb.

69,27

0,82

137

1 ,96

0,10

1 ,06

R9-63

langeneset

Deep marine turb.

70,57

0,80

1 1 ,79

6,18

0,08

3,33

2,66

2,69

1 , 54

0,11

0,92

1 00,68

150

26

33

18

74

80

127

24

215

9

48

283

28

7

R9-64

Langeneset

Deep marine turb.

69,69

0,86

1 1 ,94

6,04

0,09

3,10

2,18

5,10

1 ,46

0,12

1,31

101 ,90

107

156

25

31

28

77

62

110

30

298

12

72

51

274

37

7

R9-65

Langeneset

Deep marine turb.

67,37

0,86

13,31

6,57

0,08

3,52

2,00

3,38

1 ,99

0,12

1 ,19

1 00,39

114

161

29

33

20

83

89

117

29

256

14

83

39

322

30

R9-66

Langeneset

Deep marine turb.

70,07

0,85

1 1 ,27

6,29

0,10

3,69

3,46

2,91

0,99

0,11

0,86

1 00,60

118

169

27

34

21

71

50

157

28

1 92

11

72

47

289

31

R9-62

LOI,

Langeneset

loss on ignition;

F�O�',

0,69

total iron as

12,50

5,98

F�03 + 1 . 1 x FeO.

2,18

2,82

70

:;.;;, :;.;;, .,

"<: ;:s !':lo "'

.,

;:s "'-

� � ..., ;:s

�

(b)

(a)

1 ,4 1,2

� N'

0

i=

,'

0,8

.... ... _

0,6

I

0

/,- ,_'- -'--- - - - - "1

0,4 "

"

0,2

60

_

,_ ....

,

/D "/

/

"

,

0

,

"

:_ _

c

... ... ...

""

,

/

"

_

..

-

--

.. ... ..

....

Do • •- - - - - - .oD

'\ I

-'

- -

.,. ..... - --

0,3

A;n��. c::.". ' J ,' . . .... D� a -

,/ "

-- -'- - -: B ,.,..

0

,"

-

�

•'

/

N

Q

� 0 N

A

0,2

/

,' C

) j, .. ..;

0,1

/"

/ /� "" D ./

, , ..._ _ ..

.... ... -

"' .,..

:- - '

...0,

r"

•

"'ib",,- ' ,' B ',, ....( - -...� ..... �·

,'-'\

<

I

/,.

/

/" -.... ....

./

6 -.-

. .

/

I

�·t

�oo,_-·.., � c;, o o oaD

,"

� 3 \:!.

g2 ..,

I I I I I I I

D

s+

/

�

, ,'

(c)

I I I I I

I

I

�7 :

/ c

('-

<

' ..... _ _ /,"

I I I I I I r-- -

-

....

•• ',

�-:.. ... • / B : - -� --·------,

-- --

\

',,

/

--� �

!(

0

o: _,.,.

-''),'-•'/ \',�

,, ,�

9::-. oa 0" A ' o• ....... -•- - o- - - " .

C'l "' 0

g..

"'

4

2

10 8 6 Fe203 MgO (%) +

12

14

16

0

+---�--�---+--��-�

0

2

8 10 6 Fe203 MgO (%)

4

+

(e)

(d)

14

16

0 +---�--�---+--� 0 2 10 12 14 16 4 8 6 Fe203 MgO +

!} s

C)> ...,

Th

»' � � v

50

12

;:! r;· �

�

� s· �·

40

"<:::

(3

'<:

E

15: 30

.. .....

, ....... ..,...

/./

�

0 !"' I 0 5

I

10

I

15

I

20

I

25

I

30

( Co

.....

_

__

__

A

__ _

_

1-

-

/ ·r... _, ... .. ""'!. ', ,' __ D '\ 0. • "'� 0 \ • B \'-'... o · , .o. . ... ' \ / ',_ __

20 10

c /:

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R. Ravnas and H. Fumes

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the lithic-volcanic arenites contained in olistoliths of shallow-marine sediments and of deep-marine turbidites. The channel- and gully-fill de posits are accordingly interpreted to represent resedimented, originally shallow-marine sands and grave ls.

PROVENANCE AND TECTONIC SETTING OF THE KALVAG MELANGE

On the basis of geochemical d ata, most volcanic­ and sedimentary-rock olistoliths unequivocally favour a subduction-related setting for their original depositional environment. The great thickness of

Geochemical data for determining provenance

the melange succession (> 2500 m) and its facies assemblage and their distribution, all suggest depo­ sition in the lower part or at the base of a fault­ controlled, steep submarine slope that was subjected to persistent mass-failure processes (Ravnas, 199 1 ) . The type, abundance and geochemical signature of olistoliths of volcanic rocks (basaltic, andesitic and acidic) and first-cycle , marine, volcaniclastic rocks indicates that the resedimentation zone was on the submarine slope of a volcanic arc. The stratigraphical distribution of the various types of olistolith lithology with the shallow-marine material in the lower part of the melange, and the deep-marine turbidites, cherts and volcanics in the upper part (Figs 2 & 3 ) , suggest that the melange was formed by progressive 'unroofing' of the volcanic terrane (Fig. S). This interpretation is further supported by the vertical changes in the geochemical affinity of the volcanic rock olistoliths. The general lack of pre-emplacement tectonic deformation of olistoliths argues against an origin of the melange as multiple slides tapping a terrain of tectonically juxtaposed blocks. The occurrence of clasts and olistoliths of sedimentary and volcanic rocks within the same debris-flow units suggests that the volcanic debris represents volcanic rocks that either were interbedded with the sedimentary deposits and/or constituted the substratum upon which the sediments were originally deposited. Progressively older deposits of the volcanic arc margin are inferred to have been uplifted by tectonic faulting (thrusting) and resedimented during sub­ sequent gravitational mass movement (Fig. SA). The resulting 'olistolith-composition stratigraphy' in the melange succession (Fig. 2) would thus be an inversion of the original stratigraphy of the volcanic/ clastic sedimentary apron of the volcanic arc. Accordingly, the olistoliths at the lowest strati­ graphical level in the melange succession are thought to represent the resedimented youngest portion of the volcanic-arc edifice, whereas those present at successively higher stratigraphical levels would represent progressively older portion of the arc margin. The apparently systematic geochemical variations of the mafic volcanic-rock olistoliths from alkaline through subalkaline MORB-like to transitional types, and finally to IAT-type basalts (Fig. 5 & Table 3) are interpreted to reflect (parts of) the gradual magmatic evolution of the melange's provenance region. The IAT-type basalts, occurring at a lower stratigraphical level than the transitional and subalkaline MORB-like

257

basalts (Fig. 2) , are interpreted to represent the latest stage(s) of this inferred continuous magmatic evolution. Even though the single olistolith of potassic, alkaline basalt is found at a lower strati­ graphical level, the alkaline volcanics are thought to represent the oldest record of this assumed volcanic rock series (Ravnas, 199 1 ) . The magmatic evolution suggested b y the volcanic­ rock olistoliths can best be explained by a gradual development from continental intraplate, rift-related potassic magmatism, possibly associated with con­ tinental flood-basalts, through incipient stages of rifting to initial back-arc spreading, and finally to island arc calc-alkaline and tholeiitic magmatism. The progressively increasing influence of subduction­ related processes in the magma source region during the final stages of the recorded magmatic evolution is demonstrated by the more pronounced negative Nb anomalies. Magmatic development is therefore related to an area undergoing extension and rifting of continental crust behind an active continental margin. The rifting is thought to finally have separated a volcanic arc, detached from the mainland by a back-arc basin or marginal sea (see discussion below) . A comparable system would be the present Bransfield Strait, which is a narrow basin separating the South Shetland Islands from the Antarctic Peninsula, and which is thought to have formed by recent back-arc extension behind the South Shetland volcanic arc (Weaver et al. , 1979; Garrett & Storey, 19S7 ) . The formation and evolution of the Middle Jurassic to Lower Cre­ taceous 'rocas verdes' of southern Chile (Bruhn et al., 197S; Saunders et al., 1979) may constitute another analogue. Shortly after the formation of the back-arc basin, a thick pile of first-cycle, lithic­ volcanic terrigenous clastic materials and possible intercalated volcanics were deposited in inferred fault-controlled basin(s) (Fig. SA; Ravnas, 199 1 ) . The most favourable setting for the Kalvag Melange is thought to be along the volcanic arc margin of the back-arc basin. The rhyolite and rhyolitic ignimbrite olistoliths are interpreted to have been derived from acidic volcanics associated either with the incipient rifting to initial back-arc spreading stages, or alternatively to reflect the inferred final island arc magmatism. Based on the assumption that the granitic pebbles and boulders represent the plutonic equivalents of these acidic volcanics, the presence of such material as conglomerate debris suggests that the volcanic arc itself was deeply eroded.

258

R. Ravnas and H . Furnes

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Fig. 8. (A) Suggested depositional model for the Kalvag Melange and geological evolution of its provenance and tectonic

setting. 1, Deposition of the back-arc basin, volcanic arc margin along which the Kalvag Melange originated, 2, Initial destabilization and gravitational resedimentation of the volcanic arc apron. 3-4, Deposition of the lower (3) and upper (4) part of the melange. See text for further discussion. (B) Evolutionary model for the Lower Ordovician ophiolite complexes and island arc sequences, and the Upper Ordovician to Lower Silurian rift-related sequences. 1 - 2 , Initiation of subduction and development of immature island arc systems. 3, Build-up of mature island arcs that became colonized by shallow­ marine, Toquima- Table Head (Laurentian) faunas. 4, Accretion of the island arc system to the Laurentian continental margin. 5-6, Rifting of the active continental margin and formation of a late Ordovician to early Silurian marginal basin inhabited by Hirnantian and Holorhynchus (linked Laurentian and Baltic) faunas. From Pedersen et al. (1992).

Despite some slight variations in geochemical composition , the IAT- and MORB-type metabasaltic debris, in addition to the calcareous material, may have been derived from intercalated mafic volcanics of the proposed volcanic arc apron, representing eroded IAT- and transitional-type basalts, respect­ ively (compare multi-element variation diagrams of

Figs 4A & 5B). However, metabasalts of boninitic affinity have not been observed as olistoliths in the melange. Boninitic rocks within the Scandinavian Caledonides have so far been reported only from the Lower Ordovician ophiolite complexes/island arc sequences (e.g. see Pedersen et at., 1988; Pedersen & Fumes, 199 1 ) , which thus provide the only likely

Geochemical data for determining provenance

source for this debris type among the terranes exposed along the orogen today. Moreover, ongoing research has revealed the presence of large accumu­ lations of gabbroic rocks of boninitic affinity associ­ ated with one of these complexes (Fumes & Pedersen, 1994) . Characteristically, the Lower Ordovician ophiolite complexes/island arc sequences show a long-lived magmatic development (c. 500-470 Ma) , ranging from typical MORB to IAT, boninites, calc-alkaline and alkaline metabasalts and differentiates (Pedersen et al., 1988; Pedersen & Fumes, 199 1 ) . It is note­ worthy that the relatively flat MORB normalized pattern displayed by the MORB metabasaltic conglom­ erate debris (Fig. 6a) is typical for the metabasalts from the axis sequence of the Lower Ordovician ophiolite complexes, thereby suggesting an alter­ native source for this debris type . Based on mineral­ ogical and geochemical data the trachyandesite/ dacite debris are also assumed to have been derived from the Lower Ordovician island arc sequences. The sedimentary cover to these ophiolite complexes, consisting of interbedded metamorphosed volcanics, phyllites and chert, may thus represent a probable source for the quartzite and possibly also the sub­ arkosic conglomerate debris (Ravnas, 199 1 ) . The back-arc basin within which the volcanic arc apron originally formed received sediments from two different source rock associations or provenance regions: the newly formed volcanic arc and an ?older ophiolite complex/island arc sequence. As no systematic vertical variations in mineralogical and/ or geochemical composition have been observed , either in the sedimentary-rock olistoliths or in the olistostromal 'groundmass', sediments were ap­ parently shed continuously from both source rock associations into the original receiving basin(s). This suggests that sediments were derived either (i) con­ tinuously from two areally distinct provenances; (ii) from a volcanic arc formed by arc splitting of the Lower Ordovician ophiolite complexes/island arc sequences; or (iii) from a deeply eroded volcanic arc separated from and formed by extension of an active continental margin onto which the Lower Ordovician ophiolite complexes/island arc sequences had already been emplaced. Based on the isotope character of the volcanic-rock olistoliths combined with the present knowledge of the development of marginal basins within the outboard terranes of the Scandinavian Caledonides, the latter is preferred (see discussion below) . The accreted or obducted ophiolite complexes/island arc sequences are inter-

259

preted to have formed part of the crust upon which the volcanic arc was founded, thus suggesting that the volcanic arc was built on a sliver of continental crust.

REGIONAL CORRELATIONS AND INTERPRETATIONS

The Scandinavian Caledonides, being dominated by thrust-emplaced tectonic units, is made up of accreted suspect and exotic terranes (sensu Coney et al., 1980), representing respectively the outer part of the continent margin (suspect terranes) and 'oceanic' or island arc/ophiolite complexes, meta­ morphic complexes and continental lithosphere con­ sidered to have been derived from outboard of the early Palaeozoic continent Baltica (exotic terranes) (Stephens & Gee, 1985 , 1989; Roberts , 1988). The development of marginal basins and associated magmatism within the outboard terranes of the Scandinavian Caledonides are summarized in a number of papers (e.g. Pedersen & Fumes, 1991 ; Pedersen et al., 1992). Two generations of ophiolite complexes and spreading related mafic magmatism are now established within the Caledonian­ Appalachian orogenic belt; these are of early Ordovician and late Ordovician to early Silurian age, respectively (e.g. see Pedersen & Fumes, 199 1 ; Pedersen e t al., 1992, and references therein) . The Lower Ordovician ophiolite complexes and island arc sequences of the Scandinavian Caledonides are thought to reflect the build-up of an island-arc- marginal-basin system within the Early Palaeozoic Iapetus Ocean (Pedersen & Fumes, 199 1 ; Pedersen e t al., 1992, Pedersen & Dunning, 1995). Pedersen & Dunning ( 1995) suggest a continuous development from an immature island arc system comparable to the present Palau- Kyushu ridge of the western Pacific into a mature island arc similar to the present Sunda- Banda arc. The Upper Ordovician-Lower Silurian ophiolite and associated igneous complexes are related to rifting of an active continental margin and the development of small, fault-controlled marginal basins (Tucker et al., 1990; Pedersen & Fumes, 199 1 ; Pedersen et al., 199 1 , 1992) . Of these Upper Ordovician t o Lower Silurian marginal basins, only that represented by the Solund- Stavfjord Ophiolite Complex developed into an oceanic basin (Fumes et al., 1 990) . The relationship between the Lower Ordovician ophiolite complexes/island arc sequences and the

260

R . Ravnas and H. Fumes

Upper Ordovician to Lower Silurian magmatic complexes is poorly known. However, a line of evidence suggests that the Upper Ordovician to Lower Silurian marginal basins formed by extension along a continental margin onto which Lower Ordovician ophiolite complexes/island arc sequences had already been accreted or emplaced (Fig. 8B) (Pedersen & Fumes, 199 1 ; Pedersen et a!., 199 1 , 1992; and references therein). Based on faunal and isotopic data, the Lower Ordovician ophiolite complexes/island arc sequences are interpreted to have been accreted onto the Laurentian margin of the Iapetus Ocean in the early to mid-Ordovician (Bruton & Bockelie , 1980; Pedersen et a!., 1988, 199 1 , 1992; Pedersen & Fumes, 199 1 ) . A major problem in the reconstruction of the origin and evolution of the Kalvag Melange is the lack of age constraints and rather poorly defined tectono-stratigraphical relationship to other early Palaeozoic marginal basin complexes/sequences of the Scandinavian Caledonides. The geochemical characteristics of volcanic-rock olistoliths from the Kalvag Melange compare well with various types of volcanic rocks from the Lower Ordovician ophiolite complexes/island arc sequences reported by Fumes et a! . ( 1986) and Sivertsen ( 1992) . These volcanics are now regarded as representing the late, evolved stages of the build-up of a mature island arc (Sivertsen, 1992; Pedersen & Dunning, 1995 ) . However, the combined geochemical and isotope signature of the volcanic-rock olistoliths, and par­ ticularly that from the rhyolitic ignimbrite olistolith , do not compare favourably with any known magmatic rocks from the Lower Ordovician com­ plexes. Moreover, based on geochemical data and the relatively undeformed nature of the alkaline and subalkaline MORB-like volcanic-rock olistoliths, these olistolith types are not thought to have been derived from a hypothetical oceanic island or seamount that shed material into the trench along the subduction zone (e.g. see Ogawa, 1985; Pautot et a!., 1987) above which the Lower Ordovician ophiolite com­ plexes/island arc sequences formed. The suggested setting and evolution of the back­ arc, depositional basin of the Kalvag Melange resembles closely that proposed for the late Ordo­ vician to early Silurian marginal basins, e.g. the Solund-Stavfjord Ophiolite Complex and the Sulitjelma Gabbro (Fumes et al., 1990; Pedersen & Fumes, 199 1 ; Pedersen et a!., 199 1 , 1992 ) . A cor­ relation of the proposed volcanic arc apron along which the Kalvag Melange originated with the Upper

Ordovician to Lower Silurian rift-related igneous complexes and their cover, is further strengthened by the close spatial relationships with the Solund ­ Stavfjord Ophiolite Complex and the suggested tectono-stratigraphical position of the Kalv�lg Melange (Fig. 1) (Fumes et al., 1990; Hartz et a!., 1 994) . On the basis of isotope data obtained from the sedimentary cover (single zircon dates) , the marginal basin represented by the Solund-Stavfjord Ophiolite Complex also received sediments from a composite terrane, which included the Lower Ordo­ vician ophiolite complexes/island arc sequences (Pedersen & Dunning, 1993 ) . Moreover, the inferred rift-related alkaline and subalkaline MORB-like vol­ canic rocks present as olistoliths in the melange are geochemically similar to metabasaltic dykes in the Hs;;yvik Group of the middle tectonic unit on Atls;;y further south (Ravnas, 1991 ) . Although the origin of these dykes is still uncertain (Fumes eta!., 1990) , they may possibly be related to the same rifting event inferred from the volcanic-rock olistoliths of the Kalvag Melange above . The lower part of the melange is interpreted to reflect the initial destabilization of the volcanic arc slope and gravitational failure of the sedimentary wedge, whereas the upper part reflects repeated mass-failure processes that finally lead to resedimen­ tation of the deeper levels of the original arc apron (Fig. SA) . The spatial distribution of facies assem­ blages and olistolith types , together with the large thickness of the melange succession , thus suggest major fault-related uplift of the volcanic arc margin. Tectonic uplift and steepening of the basin margin may be attributed to at least two mechanisms, which may not be mutually exclusive. The Andaman Sea has been advocated as a modern analogue for the formation and evolution of the Solund- Stavfjord Ophiolite Complex (Fumes et a!., 1990). Tectonic faulting may, accordingly, be related to transpression due to large-scale strike-slip faulting along the island arc, similar to the transcurrent movements along the present Andaman-Nicobar Ridge (Curray et a!., 1979; Hla Maung, 1987) . Alternatively, the tectonic uplift may be attributed to thrusting during the closing stages of the back-arc basin history. In the latter case the Kalvag Melange may be coeval with the Sunnfjord Melange, which is interpreted to have been formed during obduction of the Solund ·­ Stavfjord Ophiolite Complex, i . e . during the closing stage of this Late Ordovician to Early Silurian marginal basin (Berg, 1988; Andersen et a!., 1990; Alsaker & Fumes, 1994, Osmundsen & Andersen,

Geochemical data for determining provenance

1994). In either case, continued uplift led to oversteepening of the basin margin , gravitational instability and deposition of the Kalvag Melange as a series of successive gravitational resedimentation events which tapped progressively deeper levels of the volcanic arc apron. Final emplacement of the melange onto the continental margin probably occurred during the closure of the Iapetus Ocean as a result of the continent- continent collision of the Laurentian and Baltic shields in the late Silurian (Bryhni & Sturt, 1985; Ziegler, 1985) .

AC KNOWLEDGEME N T S

Financial support for this study has been provided through grants from the Norwegian Research Council for Science and Humanities (NAVF) , project code D .4 1 .3 1 . 147 , and the University of Bergen. We are grateful to Drs W. Nemec and R . B . Pedersen for their encouragement, helpful criticism and discussions during the course of this study. We thank Drs J. Winchester, R.M. Hiscott and A . G . Plint for thorough reviews o f a n earlier draft of this manuscript, and J. Ellingsen for help in preparing the illustrations. This work represents publication no. 155 in the International Lithosphere Project (ILP) .

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Geochemical data for determining provenance

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Spec. Pubis int. Ass. Sediment. (1995) 22, 265-281

Differential subsidence and preservation potential of shallow-water Tertiary sequences, northern Gulf Coast Basin, USA MARC B . EDWARDS 5430

D umfr e i s, H o uston, Texas77096-4020, USA

ABSTRACT

Growth faulting, which accompanied shelf-edge progradation and filling of the northern Gulf Coast Basin, resulted in partitioning of the basin margin into fault blocks with contrasting subsidence rates. Study of correlative sections in juxtaposed fault blocks reveals that contrasting subsidence rates can result in strongly differing facies patterns in neighbouring areas. This complicates the task of predicting sandstone reservoir occurrence and properties. Shallow-water clastic sections from Eocene to Miocene in age were investigated using extensive well-log observations, supplemented with micropalaeontology and seismic profiles. All depositional environments involve an ongoing, complex interplay between sedimentation and erosion at different time- and physical scales. In certain settings, a greater subsidence rate causes the preservation of certain facies that would otherwise have been eroded at lower subsidence rates by processes inherent to the environment. The critical subsidence rate that separates preservation from non-preservation is termed the preservation potential threshold for a particular depositional facies. Examples are provided for progradational mouth-bar facies in a deltaic setting (Wilcox), and storm­ deposited shoreface-shelf muds in a prograding shoreline setting (Frio). Where rates of subsidence are even greater, the growth fault may produce a topographical scarp at the surface, which will influence the disposition of depositional environments. Here, the concept of preservation thresholds is not adequate to account for the observed facies changes. Rather, the presence of the surficial scarp as the surface manifestation of the subsurface fault causes the preferential development and preservation of channel activity in the topographic lows, and progradational environ­ ments with channel bypass in the topographical highs . An example is provided for a series of prograding stacked deltas (Miocene). These concepts may help to focus attention on the role of subsidence in constraining the appearance of the sedimentary record .

INTRODUCTION

In growth-faulted regions, such as the northern Gulf Coast Basin of Texas and Louisiana, contempor­ aneous faulting structurally offset stratigraphical surfaces shortly after their formation. Through time, the sediment on the upthrown block (footwall) sub­ sided at a lower rate than sediment on the down­ thrown block (hangingwall). This setting provides the opportunity to study the relationship between the preserved stratigraphical record and changing subsidence rates, while other important variables, such as eustatic sea-level , sediment supply and depositional environment, remain comparatively unchanged.

Exploring in such regions, It IS common to drill into fault blocks in which strata of a particular age have not been penetrated previously. Typically, little is known about how the downthrown section will differ from the comparatively well-known equivalent section on the upthrown block. This paper describes several examples of characteristic changes that take place across growth faults , and attempts to explain the observed changes. The most obvious manifestation of the effect of growth faulting on sedimentation is an increase in thickness of a genetic unit from the upthrown to the downthrown blocks. Expansion ratios (downthrown

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

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thickness divided by upthrown thickness; Thorsen, 1963) as great as 10 have been documented. I n addition t o thickness, other properties that vary across the faults include net sandstone, percentage sandstone and log facies development. The way in which these properties change across growth faults will be described in the examples below. The interpretation of three stratigraphical units, the Palaeocene -Eocene Wilcox Group, the Oligo­ cene Frio Formation, and the Lower Miocene (no formal nomenclature) of the Texas-Louisiana Gulf Coast all illustrate contrasting relationships between growth-faulting and sedimentation patterns. One important relationship concerns the relative rates of subsidence and sediment supply in shallow-water clastic depositional systems, in which sea-level is the major constraint on base level (Wheeler, 1964). If sediment supply is much greater than subsidence, then significant relief is unable to develop where the fault trace emerges at the depositional surface. How­ ever, where differential subsidence rates are suf­ ficiently high, significant topographical relief can be created, which can then influence sedimentation patterns. Sedimentologists and stratigraphers commonly invoke mechanisms such as changes in sea-level, subsidence rate and sediment supply to explain the distribution of sedimentary facies and the character­ istics of vertical profiles. In ancient sediments it is generally impossible to validate these interpretations independently or to isolate their respective effects. However, growth faults that rise to the depositional surface cause contrasting subsidence rates in juxta­ posed areas, while other conditions are relatively

TEXAS

unchanged. Hence it is possible to identify those changes that most likely relate to subsidence rate. Sequence stratigraphy has heightened interest in the temporal significance of certain types of strati-· graphical surfaces. Peculiarly, emphasis has been given to surfaces that are admittedly not isochronous (e .g. sequence boundaries; see Posamentier & Weimer, 1993) as opposed to surfaces that are (e .g. flooding surfaces; see Galloway, 1989). The require-· ment that sequence boundaries everywhere separate older from younger sediments adds to the demands placed on the stratigrapher to resolve the age of a stratigraphical section. Thus, the lessons learned while studying the effects of growth faulting can be applied to the problems of attempting to distinguish surfaces formed by normal environmental processes, such as channel migration (termed 'source diastems'; Swift et al., 1991), from those formed by imposition of external relative sea-level controls, such as 'incised valleys' (which are components of sequence boundaries).

STUDY AREA AND SCOPE

Examples from diverse geographical locations andl geological ages (Fig. 1) have been chosen to illus­ trate the principles set forth in this paper: the Palaeocene-Eocene Upper Wilcox Group of South Texas, the Oligocene Frio Formation of South Texas, and the Lower Miocene of southwest Louisiana. The data set consists primarily of well logs at a scale of 1 in. =100ft, supplemented by seismic and micropalaeontological data. With cali-

LOUISIANA

San Patricio County Live Oak County

GULF OF MEXICO

100 Miles 100 Km

Fig. 1. Index map to the location of the four examples presented in this paper : 1, Upper Wilcox in Zapata County; 2, Upper Wilcox in Live Oak County, 3, Frio in Nueces and San Patricio Counties; 4, Lower Miocene in Vermilion Parish. General location of regional sandy shelf-edges are indicated in stippled patterns for Wilcox, Middle Frio and Lower Miocene trends. (From fig. 2, Winker & Edwards, 1983.)

267

D ff i erent a i l subsd i ence bration from whole cores, it is possible to make fairly reliable inferences about lithology from the electric logs (Fig. 2). In this paper, diagrams show only the spontaneous potential (SP) curves, but all logs were examined and interpreted by integrating the SP with the induction logs. The Upper Wilcox of South Texas (areas 1 and 2, Fig. 1) was deposited in a series of shelf-edge delta lobes as part of the Rosita delta system (Edwards, 1980, 1981). A variety of sand geometries reflecting a varied depositional environment were identified. Strike continuous, upward-coarsening sand bodies, appear to represent wave-dominated shorelines, such as strandplains, whereas sections with scattered, blocky and fining upward sandstones suggest depo­ sition along wave-influenced deltas with pronounced mouth-bar development. The Frio of South Texas includes an interdeltaic embayment dominated by stacked and prograding barrier bar and strandplain sandstones that have pronounced strike continuity (Boyd & Dyer, 1964; Galloway et al., 1982a, b). Sandstones pinch out into lagoonal and continental mudstones and siltstones up-dip, and grade down-dip into storm-deposited interbedded siltstones and mudstones of the shelf. The Lower Miocene of southwest Louisiana was deposited as a series of stacked delta lobes at the initiation of the major sediment influx during the Miocene (Curtis, 1970). Intense contemporaneous structural activity focused sediment into a series of

HC suppression of SP

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Increasing silt content upwards

INDUCTION

Sample electric log showing response of SP and induction resistivity logs to lithology. In this paper, only the SP curves are shown, but in all cases both sets of curves were utilized in the stratigraphical studies upon which this paper is based. (Modified from Edwards, 1984.)

structural basins (Sloane, 1971), probably largely controlled by subsurface salt movement. Detailed mapping indicates that deltas were supplied by large distributary channels and incised valleys. Numerous subsurface studies in this region have shown that the major sand-bearing stratigraphical units on the contemporaneous shelf and upper slope can be divided into regressive packages, or cycles, that can be correlated from tens to hundreds of miles along strike (e .g. Curtis & Picou, 1978). These packages are bounded by transgressive shales that are associated with marine flooding surfaces or maxi­ mum flooding surfaces. Recent studies tying planktic foraminiferal assemblages to a chronostratigraphical framework suggest that the cycles have durations of about 100-200 ka (Edwards, 1990; Mitchum & van Wagoner, 1991). A recent attempt to relate Lower Miocene oxygen-isotope cycles to well-log cycles suggested a dominant periodicity of 100 ka (Ye et al., 1993). In the examples that follow, most of the correlation markers shown are thought to bound cycles of this order (see Figs 5, 9, 10, 11 & 14).

EFFECTS OF GROWTH FAULTING ON SEDIMENTATION

A considerable literature covers many aspects of the setting and effects of growth faulting on sedimen­ tation . Curtis & Picou (1978) placed the major growth-faulted trends of the Gulf Coast Basin into an offlapping delta model. Winker & Edwards (1983) examined the delta model in a shelf margin setting and pointed out some of the ways in which this setting differs from a platform setting. We can envision growth faults as part of the extensional head region of a large gravity driven slope failure that also includes a contractional toe region with folds and thrusts in deep water. Study of salt tec­ tonics in the past few years has focused on the role of salt in terms of both lateral and vertical flowage. Growth faults in salt-dominated areas can form in response to: the evacuation of deep salt into extrud­ ing diapirs, the evacuation of salt sheets, or combi­ nations of salt-driven and slope-driven gravity systems (VendeviUe & Jackson, 1992). The most obvious effect of growth faulting on sedimentation is the change in thickness (Thorsen, 1963). Aside from such thickness changes, the character of a stratigraphical unit can appear unchanged, or can show significant changes across a growth fault. In either case, it is important to attempt

268

M.B. Edwards

to predict sandstone properties for the purpose of hydrocarbon exploration and production. The fol­ lowing sections present examples of various depo­ sitional responses across growth faults.

2900m

Jonnell gas

1

Ramos

1--

9500'

J onnell gas

--11

Yzaguirre

Expansion only

Stratigraphical sections with large expansion ratios usually display significant stratigraphical changes across a growth fault. A noteworthy exception occurs in the Upper Wilcox in Zapata Country. A fault with an expansion ratio of greater than three, shows no obvious changes in percentage sandstone or log facies (Fig. 3). These sandstones were illustrated in stratigraphical and structural sections (see Edwards, 1981, Figs 10 & 11). Comparison between the upthrown and down­ thrown blocks is facilitated by changing the vertical scale of one of the logs, in order to make correlation markers subparallel (Fig. 3). This procedure of 'double datuming' attempts to reconstruct the strati­ graphy as if all of the locations had been subject to the same subsidence rate. In this and the following examples, no attempt has been made to decompact the sections in order to compensate for greater compaction of shale relative to sandstone. Con­ trasting amounts of compaction could influence present-day thickness where lithology changes sig­ nificantly across the fault. However, these sections are overpressured, which has resulted in compara­ tively small amounts of shale compaction. Inspection of the two logs (Fig. 3) indicates that percentage sandstone and log facies are largely unaf­ fected by the fault, except for the presence of a high­ frequency signal in the middle part of the down-dip well. In this example, the proportion of net sand in the two wells is the same as the expansion ratio . Facies changes and channel erosion

Extensive mapping in the South Texas Upper Wilcox (Fig. 1) reveals the presence of stacked delta lobes with mappable distributary channels (Edwards, 1980, 1981; Winker & Edwards, 1983). A set of eight wells arranged in a dip section (Fig. 4) has been selected to show thickness and log facies changes (Fig. 5). In an unfaulted area, with a uniformly increasing subsidence rate down-dip (flexure), a gradual change in log facies would be expected, with a change from distributary channels up-dip, through mouth bar to distal mouth bar down-dip. However, in this example

"

.><

£ E ro

"

.><

£ E "'

(") � 0 0

c: 0

·�

Q) !/) a.

:0

a. ::J

"'

�

0

t

0 ro

(") c: 0

'£

Q) !/) a.

:0 c: ::

0 0

Fig. 3. Stratigraphical dip section across a large growth fault in Zapata County, South Texas. The vertical scales of stratigraphically equivalent section both upthrown and downthrown to the fault have been adjusted to compensate for differential subsidence. Note the similarity of the SP logs, despite the expansion ratio of 3.45. Similar 'double­ datumed' sections are also shown in Figs 5, 1 1 & 13. (Modified from Edwards, 1984.)

(Fig. 5), a striking change in log facies occurs at the up-dip growth fault (between wells 2 and 3), with more subtle changes at the other faults. The effects of growth faults on stratigraphical preservation can be observed by comparing and contrasting adjacent

Dfi ferent a i l subsd i ence Texas

Louisiana

INDEX MAP

I

5ooo· 1500 m

I

Live Oak County

4. Index map showing the location of wells depicted in Fig. 5, Live Oak County, Texas (see Fig. 1 for context).

Fig.

subsidence-normalized wells that are in the same versus different fault blocks. For example, in the Luling regressive cycle (Fig. 5B), wells 1 and 2 are very similar, as are wells 3 and 4. However, wells 2 and 3 are clearly different. The preserved components of each progradational deltaic cycle (Fig. 5) include mouth bar and distal mouth bar deposits, typically arranged in upward­ coarsening successions; distributary channel deposits represented by blocky sandstones and upward-fining successions; and coastal plain and interdistributary bay deposits, which consist of mixed sandstones and shales in a variety of patterns, usually thinner than the overall regressive succession . Subsidence creates space for the preservation of facies beneath erosion surfaces. In this example, major fluvial erosion is represented by the base of the blocky and upward-fining sandstones. Subsid­ ence allows the deposition and burial of pro­ gradational facies beneath the depth of erosion attained by channels. This relationship becomes critical in up-dip areas, where the decreasing thick­ ness of the space available for preservation of progradational facies approaches the depth of dis­ tributary channels. Up-dip of this point, the Slick and Luling genetic regressive successions are charac­ terized by amalgamated sandy channel-fill deposits. Subsidence rate also affects regional facies trends. High subsidence rates favour stacking of delta lobes

269

with 'foreshortening' of dipwise facies gradients (the distance along dip from proximal to progressively distal facies) rather than down-dip translation of facies belts by continued progradation into the basin. A schematic facies preservation diagram (Fig. 6) enables prediction of the depositional sequences that can be juxtaposed across a growth fault in a prograding delta. At the top (Fig. 6A) is a subsidence rate graph . The duration of the progradational phase is suggested to have lasted about 1000 yr, although it could have been much shorter or longer. The growth­ fault curve (solid line) shows subsidence rates that reflect the thickness changes observed due to growth faulting and roll-over (up-dip thickening on the downthrown block toward the fault). The vertical steps show the locations of individual growth faults. Shown for reference is the flexure curve (dashed line), which assumes a linear increase in subsid­ ence rate down-dip. Arrows labelled U and D are explained below. Figure 6B illustrates some of the facies relation­ ships and inferred environments of deposition. After transgression and abandonment of the previous delta lobe, the cycle begins by progradation of the delta front, with superimposed higher frequency trans­ gressive -regressive cycles. The progressive basin­ ward shift of environments down-dip has been shown by the down-dip termination of channel erosion surfaces and mouth bars. Coastal-plain facies are depicted as blocky sandstones, although they are in reality complex intercalations of various channel and bay deposits. There is a total of 5 0-200ft (15 -60m) of subsidence at the up-dip end, whereas at the down-dip end there is 25 0-2000ft (75 -600 m) of subsidence . It is assumed that distributary chan­ nels are 5 0-200ft deep regardless of location, although it is likely that they would decrease in depth down-dip due to channel bifurcation and dis­ charge through crevasses. The subsidence curves can be used to estimate preservation of vertical sections on either side of a growth fault. At a growth fault, the regionally aver­ aged subsidence rate does not occur. Instead, much higher subsidence rates occur on the downthrown block, and much lower rates occur on the upthrown block. On the upthrown block (see Fig. 6A), the lower subsidence rates resulted in a vertical section that resembles that developed up-dip. This is shown by the arrow U, which is projected to the left until it intersects the flexure curve, where the appropriate lower subsidence rate would have occurred in a non­ growth-faulted setting (depicted by flexure curve in

M.B. Edwards

270

Downdip

Updip 3

2

4

5

6

8

7

- -100m - -500'

91m

- -1000'

\F

-

-1500'

A - -500m u L T -2000' ' -;-5 614m

�

19 579m

A Paleostructure B Subsidence normalized 3

2

4

PD

5

7

6

8

PD F L T

Fig. 5.

Thickness and facies changes in two stacked delta complexes, the older Luling (A and B) and the younger Slick (C and D) units in the Upper Wilcox of South Texas (see Figs 1 & 4 for locations) . For both deltas the upper panels (A and C) show a palaeostructural cross-section with the top to the delta complex as the upper datum, and growth faults that were active during delta formation. The lower panels (B and D) are 'double-datumed' sections in which vertical scales were adjusted to normalize for differential subsidence. Generalized deltaic facies are identified using electric log characteristics. CH, channel; MB, mouth bar; PD, prodelta. No horizontal scale.

Fig. 6A and facies distribution in Fig. 6B). At the intersection point, the arrow is then extended down into the facies diagram, in order to determine the vertical section that corresponds to that subsidence rate. A similar procedure can be carried out for the downthrown block by projecting the downthrown subsidence rate to the right (arrow D, Fig. 6A) until it intersects the flexure curve, and then extending it

down into the facies diagram below. This procedure predicts that the vertical sections on either side of the growth fault would show considerable differences in facies profiles that could not be explained solely as a function of changing palaeogeographical lo­ cation. The thickness of the upthrown and down­ thrown profiles could then be expanded or contracted to restore their present-day relative thicknesses. However, this method may predict excessive lateral

Different a i l subsd i ence

271 Downdip

Updip 2

3

4

5

6

7

8

- -100' - -200' - -300' - -400' - -500' - -600' C Paleostructure

D Subsidence normalized 2

4

5

6

7

8

PO

Fig. 5.

:><: u ::::;

(f)

(Continued.)

facies changes because the distance between the two schematic wells (below arrows U and D, Fig. 6) will be greater than the actual distance between two wells on either side of a growth fault. Facies changes and storm/wave erosion

The South Texas Frio Formation (Fig. 7) illustrates shorelines that prograde as a line source (i . e . wave­ dominated deltas, barrier islands and strandplains), rather than point sources (such as river-dominated deltas). It is based on complex stratigraphical and structural relationships determined from approxi­ mately 600 well-logs (Weise et a!., 1981; Bebout et a!., 1982; Edwards, 1986), coupled with micro­ palaeontological studies that indicate neritic to coastal environments throughout the section, with overall shallowing upwards (Martin, 1969, 1970). The sandy, wave-dominated Frio shoreline was bordered by muddy coastal plains up-dip, muddy slope deposits down-dip, and by major deltaic depo­ centres along strike (Galloway et a!., 1982a, b). These palaeogeographical relationships strongly

suggest that sand was supplied to the Frio shoreface by wave-driven alongshore transport in the fore­ shore/upper shoreface, rather than by prograding deltas (Martin, 1969). A typical well-log in the Frio shows an upward transition from shelf mudstone to strandplain sand­ stone over a depth range of about 5 000 ft. Subdivision of the section into component cycles reveals that the thickness of the cycles increases markedly with greater depth, from less than 5 0 ft to almost 5 00ft (Fig. 8). This suggests that structurally deeper sedi­ ments were deposited at much greater subsidence rates than shallower Frio sediments. Rigorous cor­ relation of hundreds of well-logs in the area allows the tracing of these varied cycles across growth faults. Palaeostructural cross-sections (Figs 9 & 10) show the position of growth faults and their effect on thickness and log facies. The analysis of log facies changes is presented in a set of four well-logs, each in different fault blocks (Fig. 11). The log correlations look questionable at their original depth-scales, but when normalized for differential subsidence, the correlation of the indi-

M.B. Edwards

272 A Subsidence rates (thickness/arbitrary unit of time)

0' 1-��;===�-----�� ��---- ------------------------�_0' ----1�� Upthrown . _1 U - -Growth faults \ = ---I= � ----J / Downthrown . _ F lexure � / .= rr:- I= D 1= 50'= 250'r 200' 15-60m \"-0II0"e 2000' u D 75-600m -

-

-

�

-

---� --

B Facies Updip

Downdip

a

o 0

E

0

� � b � � ....

Fig. 6. Schematic diagram of facies preservation relationships in a deltaic environment characterized by significant channel erosion. (A) Two subsidence rate relationships are shown, a straight dashed line indicating uniformly increasing subsidence rate with distance as in a flexure, and a complex solid line indicating abrupt increases in subsidence rate down-dip across growth faults and decreasing subsidence rate up-dip within some fault blocks due to roll-over. (B) Facies preservation trends and characteristic SP logs. Proximal to distal facies trends have been exaggerated. The figure can be used as a kind of nomograph to estimate facies changes across growth faults in this setting. See text for additional explanation. (Modified from Edwards, 1984.)

vidual cycles is much more obvious. Figure 11 illus­ trates the change in log facies from thick sand bodies up-dip, to thin interbedded sandstones (siltstones) and mudstones down-dip. The palaeostructural sec­ tions emphasize that the facies changes take place abruptly across the growth fault. However, the cor­ relatability of the individual cycles across the fault suggests that significant sea-floor relief was not main­ tained at the location of the fault. Core study of the down-dip interbedded facies (Berg & Powell, 1976) suggested that the sandstones were deposited by turbidity currents. Studies of the modern Texas shelf, however, suggested that shelf sands are deposited from storm-generated geo­ strophic flows that produce turbidite-like beds (Morton, 1981). Transport is oblique to the shore­ line, but the resulting thin beds have enormous strike continuity (Snedden & Nummedal, 1991). The continuity is consistent with the excellent gas

production that can be achieved despite the thin, ratty character of the sands. The large amount of interbedded mudstone was presumably derived from the adjacent deltaic depocentres. A schematic diagram (Fig. 12) suggests how abrupt facies changes can develop across growth faults, without the necessity of an abrupt change in depo­ sitional environment. The model proposed in Fig. 12 relies upon differential subsidence to produce a con­ trast in the sediments preserved, even though the depositional processes on both sides of the fault are virtually identical. Unfortunately, the available data do not allow many aspects of the model to be resolved until additional core data are obtained . The dynamics of shoreface -shelf systems have been discussed by many authors (e .g. Swift & Thorne, 1991), however, a considerable number of issues remain contentious. For the present dis­ cussion, it suffices to emphasize the key points. The

Dff i erent ai l subsd i ence Louisiana Texas

N

t

Nueces

I

10 Miles 10 Km

Kleberg

Fig. 7.

Index map showing location of cross-sections A-A' and B-B' in Figs 9 & 10, and general location of Frio investigation around Corpus Christi Bay.

precise roles of fair-weather versus storm conditions in transporting sediment on the upper shoreface is often unclear and may vary from one shoreline to another. The Frio shoreline prograded steadily through time due to the high sediment supply, and significant wave transport along shore and storm transport to the inner shelf. Logs and whole cores indicate a progression from relatively coarse-grained homogeneous sandstones up-dip, through interbedded thick sandstones and shales, to thinly interbedded 'ratty' siltstones and

273

shales down-dip. It seems likely that the coarse­ grained shelly sandstones up-dip were deposited in a foreshore and upper shoreface setting and are com­ posed of wave-reworked and amalgamated storm beds. The ratty interbedded section down-dip was deposited on the lowermost shoreface to shelf and consists of interbedded storm-deposited sands and muds. The abrupt facies change at the fault remains to be explained. The important distinction is that mud layers are preferentially preserved on the more rapidly subsiding down-dip block. The mud could have been deposited from suspension either by turbid nearshore, semi-permanent currents, or by waning flow following a storm event. In either case, it is important that the mud was not eroded subsequently, either by storms or fair-weather wave activity. The Frio shelf system can be regarded in terms of profiles of equilibrium and changing base levels. Areas of greater subsidence rate permitted less sedi­ ment to be reincorporated by storm currents and bypassed on to the shelf, and hence more sediment was allowed to accumulate. If eustasy and sediment supply are held constant, then the high subsidence rates on the downthrown block effectively raise the local base level (increase accommodation ) , increasing the probability that storm-deposited muds would be preserved. At some critical subsidence rate, sand beds stop being amalgamated, and instead are separated by thin mud layers. At this point, the reservoir characteristics of the sand body change drastically. The most rapidly subsiding areas immediately down-dip of the growth fault would probably have

A

8

7000' -

�

8000'-

�

J 0 Fig. 8.

Example of Frio log from Corpus Christi Bay showing vertical changes in cycle thickness, suggesting increasing subsidence rate with depth. Small portions of the log have been enlarged to illustrate cycle styles and thickness. The numbers next to logs A, B and C refer to the thickness of the log segments illustrated.

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3000m 10 000' -

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M. B. Edwards

274

NW A

32

31

30 GULf 2 ST TR

47 2

-1> (8560')

GULf ST TR

53

-1> (8678')

33

CITIES SER !c SUNRAY 1 ST TR 52 1 ST TR 51

-1> (8670')

-1> (8740')

34 1

ATlANTIC ST TR 81

-1> (8885')

38

37

35 CHERRYVILLE 1 ST TR 81

6

-1> (8983')

ATlANTIC ST 470

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(10360')

1

TENNECO w IA TANG I

SE

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(10038')

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OAlUM HORIZON 10

/

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INDEX MAP

E

I

oo oo I{) PALEOSTRUCTURAL DIP SECTION IN FRIO FORMATION, NUECES CO., TX. DAlUM: LOCAL MARKER IN FRIO V. E.

=

32 X

Fig. 9.

Frio palaeostructural section A-A' (see Fig. 6 for location and Fig. 1 for setting) showing an overall regressive stratigraphical interval bounded by marine flooding surfaces expanding dramatically across two growth faults, with concomitant facies changes. (Modified from Edwards, 1986.)

served as an effective sediment trap, especially as the local onshore-directed subsidence gradient was opposed to the obliquely offshore-directed sediment­ transport gradient. Storm-deposited sands typically pinch out down-dip, leading to the development of combination traps in off-structure positions. Facies changes at faults that develop topographical relief

The balance between sediment supply and subsid-

ence occasionally results in a significant topographi­ cal scarp being developed where the fault intersects the sea-floor. This is illustrated by the Lower Miocene of southwestern Louisiana (Fig. 1), where a large depositional basin formed as a result of the sudden and rapid removal of subsurface salt, and then stabilized when the salt body was fully evacu­ ated (Edwards, unpublished). Thin stratigraphical units in this area were mapped using almost 2000 well-logs, micropalaeontological data, and seismic data. The maps show the devel--

275

Different ial subsd i ence 39

40

41

BRITlSH-AMERICAN

CITlES SERVlCE

HAMON

#1

12

13

i} (8549')

(8288')

(8185')

ST TR

ST TR

f/1

(9400')

10 #1

ST TR

i} (9121')

42 15 #3

ST

44

3

TENNECO

RENWAR I< PHILLIPS

43

HAMON HAMON

8#2

ST

8#2

ST TR

5

(9029') (9021') (9220')

#1

HOGG

f/1

(8905')

sE B'

HORIZON 20

B

E

1

*-......

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INDEX MAP

2

*

\ 39

N

1

40

*-0

\ 41

*

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10,000' 3000

m

42 43

PALEOSTRUCTURAL DIP SECTION

44

*·-¢-.0

'----

NUECES AND SAN PATRICIO COS., TX

3

0

DATUM: LOCAL MARKER IN FRIO

B'

V. E.

=

32 X

Fig. 10. Frio palaeostructural section B-B' (see Fig. 6 for location and Fig. 1 for setting) showing dramatic expansion of two sandy regressive cycles, each bounded by marine flooding surfaces, across a growth fault, with concomitant facies changes. Note excellent roll-over structure with expansion and improved development of sandstones up-dip to the northwest into the growth fault. (Modified from Edwards, 1986.)

opment of channel sandstone bodies that traverse the up-dip stable shelf and trend toward the down­ dip basin, where large quantities of sand were depo­ sited on the downthrown side of the fault (Fig. 13). The pattern is repeated in several successive units (Fig. 14). The coincidence of the structural boundary and the facies boundary indicates that the basin was a topographical as well as a structural low, which not only collected sediment but also attracted distribu­ tary channels and incised valleys from the adjacent highs (Fig. 13B). Similar relationships between struc­ tures, relief and facies have been recognized in

other settings (e.g. Hopkins, 1987; Leeder & Alexander, 1987). The up-dip areas (Fig. 13B) appear to be com­ prised largely of shallow-water mudstones and local­ ized mouth-bar sandstones, which were deposited while sea-level was relatively high . The development of narrow incised valleys allowed the preservation of these fine-grained deposits, and contrasts with the Wilcox example, described above, in which there was extensive scouring at the bases of distributary channels. The apparently stable channels suggests a relative fall in base level, due either to eustasy or to local structural uplift beneath the upthrown block

276

M.B. Edwards

Updip

Down dip 2

3

4

Down dip

Updip

Datum

2

4

3

Datum Top

Top J

J

Paleostructure

2000' A

610m

Base J B

230'

405'

500'

2000'

Fig. 11. (A) Frio schematic palaeostructural section of one stratigraphical unit ('J') with well-log segments selected from four wells. (B) Frio stratigraphical section with vertical scales adjusted to normalize for differential subsidence across growth faults. The correlations suggest that thin digitated SP facies up-dip have closely time-equivalent ratty/serrated SP facies down-dip. Figures 9 & 10 demonstrate that the facies changes occur at growth faults.

due to movement of deep salt. The down-dip basin­ fill attests to the huge sediment volumes that were bypassed through the relatively small valleys.

PRESERVATION POTENTIAL THRESHOLDS

The above examples raise the question as to why some of those faults that are not associated with depositional topography show dramatic facies changes, whereas others do not. It is customary to express the significance of a growth fault in terms of its growth ratio, as this can be measured directly in the subsurface data. However, this ratio gives no information about the absolute rates of subsidence on either side of the fault. It seems likely that a major factor controlling facies development is the absolute subsidence rate. Along a continuum of subsidence rates, will be rates that separate domains that are associated with the preservation or non-preservation of a particular facies component, or sedimentary feature (Fig. 15). These specific subsidence rates can be referred to as preservation potential thresholds for a particular sedimentary feature in a particular depositional

environment. As noted above, absolute subsidence rate has to be evaluated in the context of the other variables that affect base level: eustatic fluctuations, and the amount and calibre of sediment supply (e.g. Swift & Thorne, 1991). For example, in the deltaic setting discussed above, distributary channels are a significant source of erosion. With sufficiently low subsidence rates, migrating channels will remove all or most of the progradational facies, resulting in amalgamated multistory channel sand bodies. When the threshold subsidence rate is exceeded, progra­ dational facies will be preserved . Another example is the balance of erosion and deposition on the shoreface . A threshold value of subsidence, in the context of the other controlling variables, separates regimes in which mud layers will be eroded from those in which they are preserved. Analogous thres­ holds could be postulated for other environments and sedimentary features.

CONCLUSIONS

1 The interplay between erosional and depositional processes in sedimentary environments is controlled to a large extent by subsidence rates. The latter

277

Dff i erent ial subsd i ence

l

FORESHORE Bay/

Barrier/

lagoon

strandplain

SHOREFACE

Storm erosion

SHELF

_

A Stable setting

B Unstable setting

Prograding wave-dominated shoreline

Fig. 12.

Schematic cross-sections showing wave-dominated shoreline progradation in (A) stable and (B) unstable settings. Characteristic well-logs (SP only) are shown. (A) This section indicates the presence of wave grading shaping the foreshore and upper shoreface, and storm events resulting in erosion of the upper shoreface and deposition of a storm couplet on the shelf. (B) The enhanced subsidence rates in the unstable setting allows for the preservation of the muddy portion of the storm couplet from erosion, whether by wave grading or a subsequent storm event. The enhanced preservation seems to have resulted from greater subsidence rates, which modified local base levels. (Modified from Edwards, 1984.) A

Topography

B

healed by sedimentation

Topography controls sedimentation

Fig. 13.

Schematic block diagram contrasting: (A) normal sedimentation across growth fault without development of topographical relief, versus (B) growth fault with associated topographical scarp and bypassing of coarse sediment across up-dip block to rapidly subsiding down-dip block. Based on a study of the Lower Miocene. (Modified from Edwards & Tuttle, 1993.)

Amalgamated channel complex

Expansion with drastic facies changes across growth fault Expansion without facies changes across growth fault

M. B. Edwards

278 - DowndipHawthorne 1 Delcambre E 135-04E-001 (12281')

A

Texaco lle Blanc B 135-04E-Q49 (12036')

B

- Downdip143'/44m

137'/42m

Jllt lllf Jllt�JIIr �II).: (IIi ill}.} ill[

Fault/channel relationships

1 Fault buried; no channels in area

,04

168'/51m

108'/33m

153'/47m

132'/40m

316'/96m

2 Fault active but channels not closely related to fault traces

283'/86m

3 Fault active with tight control on channel location. Channel updip and downdip

743'/226m

4 Fault active with tight control on channel location. Channel in downdip well only

780'/238m

5 Fault active with tight control on channel location. Channel in downdip well only

411'/125m

-1700'/520m

111r� �111

6 Fault active with tight control on channel location. Channel in downdip well only

14. (A) Palaeostructural cross-section showing change of Lower Miocene section across a prominent growth fault that had topographical relief developed during part of its history. The section was subdivided into six units for mapping and correlation, of which the deepest unit has no clear lower boundary in this well. (B) The six units are shown with vertical scales adjusted to normalize for differential subsidence across the fault. For each layer, the original thicknesses are shown, and the expansion ratio is shown between the logs. Fig.

279

Dff i erent ial subsd i ence

Examples of different fault/subsidence relationships 0)(1) (1) c·- :J ro � V>0 "' Vl "' "' (.) � .0 c (.) "' "' .!: "0

Fault A

Fault

1

Fault 2

Fault 4

Fault 3

'iii

D

.0 ::J Vl

Preservation threshold: storm-deposited mud

-

Preservation threshold: progradational facies

-

Channel inc;sion threshold:

Preservation of storm-deposited mud Preservation of progradational facies Updip bypass due to channel incision

u

D -----------

------- -

---

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D - ---

-------

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- - - - - --- -

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Presence of features

:I I I I I I I I I I I •

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Fig. 15.

Representation of preservation potential thresholds in growth-fault settings. The vertical axis shows increasing absolute subsidence rate, but is schematic with no scale. Most of the stratigraphical units depicted in this study are considered to be 'fourth order', and hence have a duration of the order of approximately 100 ka. Preservation thresholds are depicted for progradational facies at a lower rate than that for storm-deposited mud in the transitional lower shoreface to shelf environment. Five different fault situations indicate how various combinations of depositional environment and subsidence rates determine whether particular facies will or will not be preserved across specific growth faults. Note that expansion ratio is related to relative rather than absolute subsidence rate. If blocks on either side of a growth fault are on the same side of the preservation threshold, then facies patterns will be similar with regard to the particular process threshold, although thickness will be different. Settings that represent examples described in this paper are fault A, Lower Miocene of southwestern Louisiana; fault 1, amalgamated sandstones on both sides of fault common in up-dip areas; fault 2, Wilcox in Live Oak County, Texas; fault 3, Frio in South Texas; fault 4, thick sections with similar facies on both sides of the growth fault, common in down-dip areas.

affects preservation potential via thresholds that separate domains of facies preservation versus erosion. 2 Subsidence rates, through their effect on preser­ vation potential, influence facies composition, including geometry, bedding characteristics, fabric, palaeontology and seismic response. The effect of subsidence rates is readily demonstrated in the Gulf Coast Basin due to the presence of growth faulting. In tectonically stable basins, the effect of subsidence rate on facies composition may not be determined readily. 3 Where subsidence rates were sufficient to create topographical scarps at the depositional surface, the effect of faulting on sedimentary facies is greater, as there is then a feedback effect on the location of

major facies, and not just the preservation potential of sensitive component facies. 4 The documented effects of preservation thresholds on reservoir architecture indicate that it would be useful to be able to predict preservation patterns on unexplored fault blocks in growth-faulted basins. However, the present study indicates that expansion ratios are generally inadequate to make such predic­ tions: absolute subsidence rates are required, but are often difficult to obtain or predict.

ACKNOWLEDGEMENTS

The Wilcox and Frio work was initiated while I was with the Bureau of Economic Geology of the

M. B. Edwards

280

University of Texas. I thank my colleagues there for their generous support, especially Don Bebout, Bonnie Weise and Bill Galloway. Competent assist­ ance was provided by Rick Schatzinger, Jim Lockley, Down Downey, Susan Hallam, Steve Mann and Doug Wilson. Contract funding which led to some of these results was provided by the US Department of Energy and the Gas Research Institute. This early and much of my subsequent work on Gulf Coast onshore stratigraphy has been supported by many oil and gas companies, for which I express my appreciation. The Miocene work benefited from the micropalaeontological expertise of J . Loyd Tuttle and seismic data from Geophysical Pursuit, Inc. Yvonne Bowlin ably assisted in all parts of this study. This manuscript has been improved by the reviews of William R. Dupre, William C. Ross and the Editor Guy Plint. In addition, many colleagues , too numerous to mention, have discussed the concepts discussed herein, offering anything from constructive criticism to disbelief. Finally, I thank Harold Reading for picking me up at Gatwick after my first trans-Atlantic journey, and for attempting to teach me to drive on the correct side of the road.

REFERENCES BEBOUT,

D.G., WEISE, B.R., GREGORY, A.R. & EDWARDS, M.B. ( 1982) Wilcox sandstone reservoirs in the deep subsurface along the Texas Gulf Coast. Tex. Univ. Bur. econ. Geol. Rep. Invest., 117, 125 pp. BERG, R . R . & PowELL, R . R. (1976) Density-flow origin for Frio reservoir sandstones, Nine Mile Point Field, Aransas County, Texas. Trans. Gulf Coast Assoc. Geol. Soc. , 26, 310-319. BoYD, D.R. & DYER, B.F. ( 1964) Frio barrier bar system of south Texas. Trans. Gulf Coast Assoc. Geol. Soc., 14, 309-321. CuRTIS, D.M. ( 1970) Miocene deltaic sedimentation, Louisiana Gulf Coast. In : Deltaic Sedimentation ­ Modern and Ancient (Eds Morgan, J.P. & Shaver, R.H.), Spec . Pub!. Soc. econ. Paleontol. Mineral, Tulsa, 15, 293-308. CuRTIS, D .M. & Picou , E.B., JR. (1978) Gulf Coast Cenozoic; model for application of stratigraphic concepts to exploration on passive margins. Trans Gulf Coast Assoc. Geol. Soc. , 28, 103- 120. EDWARDS, M.B. (1980) The Live Oak delta complex : an unstable shelf-edge delta in the deep Wilcox trend of South Texas. Trans Gulf Coast Assoc. Geol. Soc., 30, 71-79. EDWARDS, M.B. (1981) Upper Wilcox Rosita delta system of South Texas: growth-faulted shelf-edge deltas. Bull. Am. Assoc. petrol. Geol. , 65, 54-73.

EDWARDS,

M.B. (1984) Stratigraphic and structural analysis of growth-faulted regions using well logs: a workshop. Houston, Texas. Unpublished lecture notes and problems. EDWARDS, M.B. (1986) Sedimentary effects of differential subsidence in Frio shoreface-shelf sediments, Gulf Coast Tertiary. Houston geol. Soc. Bull., 28, 10-14. EDWARDS, M . B . (1990) Stratigraphic analysis and reservoir prediction in the Eocene Yegua and Cook Mountain Formations of Texas and Louisiana. In: Sequence Stra­ tigraphy as an Exploration Tool (Ed. Armentrout, J .M.), pp. 151-164. Gulf Coast Section, Soc. econ. Paleontol. Mineral. 1 1th Ann. Res. Conf. EDWARDS, M.B . & TuTTLE, J.L. (1993) Regional Sequence Stratigraphy and Exploration Potential of the Lower Miocene of Southwest L ouisiana. Proprietary industry study (unpublished) 121 p. GALLOWAY, W . E . (1989) Genetic stratigraphic sequences in basin analysis I: architecture and genesis of flooding;­ surface-bounded depositional units. Bull. Am. Assoc. petrol. Geol. , 73, 125-142. GALLOWAY , W.E., HoBDAY , D . K. & MAGARA, K. (1982a) Frio Formation of the Texas Gulf Coast Basin - depo­ sitional systems, structural framework, and hydrocarbon origin, migration, distribution, and exploration potential. Tex. Univ. Bur. econ, Geol. Rep. Invest., 122, 78 p. GALLOWAY, W.E., HOBDAY, D . K. & MAGARA, K., (1982b) Frio Formation of the Texas Gulf Coast Plain: depo­ sitional systems, structural framework, and hydrocar­ bon distribution. Bull. Am. Assoc. Petrol. Geol. , 6�>, 649-688. HoPKINS, J.C. (1987) Contemporaneous subsidence and fluvial channel sedimentation: Upper Mannville C Pool, Berry Field, Lower Cretaceous of Alberta. Bull. Am. Ass. petrol. Geol. , 71, 334-345. LEEDER, M.R. & ALEXANDER, J. ( 1987) The origin and tectonic significance of asynunetrical meander-belts. Sedimentology, 34, 217-226 . MARTIN , G.B. (1969) The subsurface Frio of South Texas: stratigraphy and depositional environments as related to the occurrence of hydrocarbons. Trans. Gulf Coast Assoc. Geol. Soc. , 19, 489-499. MARTIN , G.B. (1970) Depositional history: key to explo­ ration. Oil Gas J. , January 12, 98-106. MiTCHUM, R.M. & VAN WAGONER, J.C. (1991) High­ frequency sequences, and their stacking patterns: sequence-stratigraphic evidence of high-frequency eustatic cycles. Sediment. Geol. , 70, 131-160. MoRTON, R . A . (1981) Formation of storm deposits by wind-forced currents in the Gulf of Mexico and the North sea. In: Holocene Marine Sedimentation in the North Sea Basin (Ed. Nio, S.D.), Spec. Pubis int. Ass. Sediment., No. 5, pp. 385-396. Blackwell Scientific Publications, Oxford. PoSAMENTIER, H . W . & WEIMER, P. ( 1993) Siliclastic sequence stratigraphy and petroleum geology - where to from here? Bull. Am. Assoc. petrol. Geol. , 77 , 731-742. SLOANE, B.J. (1971) Recent developments in the Miocene Planulina gas trend of south Louisiana. Trans Gulf Coast Assoc. Geol. Soc. , 21, 199-210. SNEDDEN , J.W. & NU MMEDAL, D. (1991) Origin and geometry of storm-deposited sand beds in modern sedi­ ments of the Texas continental shelf. In: Shelf Sand and

Dff i erent ial subsd i ence Sandstone Bodies: Geometry, Facies and Sequence Stra­ tigraphy (Eds Swift, D.J.P., Oertel, G.F., Tillman, R . W. & Thorne, J.A.), Spec. Pubis int. Ass. Sediment., No. 14, pp. 283-308. Blackwell Scientific Publications, Oxford. SwtFT, D.J.P. , PHILLIPS , S. & THORN E , J .A. (1991) Sedi­ mentation on continental margins, IV: lithofacies and depositional systems. In: Shelf Sand and Sandstone Bodies: Geometry, Facies and Sequence Stratigraphy (Eds Swift, D.J.P. , Oertel, G.F., Tillman, R.W. & Thorne, J.A.), Spec. Pubis int. Ass. Sediment., No. 14, pp. 89- 152. Blackwell Scientific Publications, Oxford. SwtFT, D.J.P. & THORNE , J.A . (1991) Sedimentation on continental margins, I: a general model for shelf sedimentation. In: Shelf Sand and Sandstone Bodies: Geometry, Facies and Sequence Stratigraphy (Eds Swift, D.J.P., Oertel, G.F., Tillman, R.W . & Thorne, J.A.), Spec. Pubis int. Ass. Sediment . , No. 14, pp. 3-3 1 . Blackwell Scientific Publications, Oxford. THORSE N , C. E. (1963) Age of growth faulting in southeast Louisiana . Trans. Gulf Coast Assoc. Geol. Soc., 13, 103- 1 10. VENDEVILLE, B.C. & JACKSON , M.P.A. (1992) The rise and fall of diapirs during thin-skinned extension. Tex. Univ.

281

Bur. econ. Geol. Rep. Invest. , 209, 60 p. B.R., EDWARDS, M.B . , GREGORY , A.R . , HAMLIN, H.S. , JIRIK, L . A. & MORTON , R.A. (1981) Geologic Studies of Geopressured and Hydropressured Zones in Texas: Test-well Site Selection, Final Rep ort. Texas University Bureau of Economic Geology, unpublished contract report prepared for Gas Research Institute, 308 pp. WHEELER, H.E . ( 1964) Baselevel, lithosphere surface, and time-stratigraphy. Geol. Soc. A mer. Bull., 75, 599-6 10. WINKER, C. D. & EDWARDS, M.B. (1983) Unstable progra­ dational clastic shelf margins. In: The Shelfbreak: Critical Jnte1Jace on Continental Margins (Eds Stanley, D.J. & Moore, G.T.), Spec. Publ. Soc. econ. Paleontol. Min­ eral . , Tulsa, 33, 139-157. YE, Q . , MAn-HEWS, R.K., GALLOWAY, W. E . , FROHLICH, C. & GAN , S. (1993) High-frequency glacioeustatic cyclicity in the Early Miocene and its influence on coastal and shelf depositional systems, NW Gulf of Mexico Basin. In: Rates of Geologic Processes: Tectonics, Sedimen­ tation, Eustasy and Climate (Eds Armentrout, J.M., Bloch, R. & Olson, H.C.), pp . 287-298. Gulf Coast Section, Soc. econ. Paleontol. Mineral. 14th Ann. Res. Conf.

WEISE,

Sequence and Seismic Stratigraphy in Facies Analysis

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

Spec. Pubis int. Ass. Sediment. (1995) 22, 285-303

Seismic-stratigraphical analysis of large-scale ridge-trough sedimentary structures in the Late Miocene to Early Pliocene of the central North Sea J O E CAR T WRI G HT Department of Geology, Imperial College of Science, Technology and Medicine, Prince Consort Road, London SW7 2BP, UK

ABSTRACT

This paper describes the geometry, seismic-stratigraphical characteristics and palaeogeographical setting of a system of ridges and troughs measuring approximately 5-20km long, by 1-3km wide by 100-300 m high, developed in the mudstone-dominated Upper Miocene and Lower Pliocene of the central North Sea. The ridge-trough system occupied a base-of-slope to basin floor setting centred on UK Quadrant 22 at the time of deposition. The ridge crests are oriented NNW-SSE and are almost straight in plan view. The troughs are characterized by incision at the base, followed by drape, and subsequent onlap fill. The ridges and troughs have a distinctly aggradational character. Two possible mechanisms are proposed to explain the development of these structures. The geometry and scale of the ridge-trough structures is similar to that described for mud waves formed on deep-marine sediment drifts by the action of oceanic bottom currents. By analogy, it is possible that the ridge-trough structures could have formed in response to bottom currents flowing along the axis of the central North Sea, linking the Norwegian-Greenland Sea with the Bay of Biscay. This mechanism requires a marine connection in the southern North Sea, and direct evidence for this possible connection has been removed by erosion during the Pleistocene. An alternative mechanism is that the ridge-trough system formed in response to downslope-directed currents of unspecified type developed on the pro­ delta slope. This mechanism is appealing in that it explains the orientation of the ridges and troughs (orthogonal to the nearest delta front) , and, most significantly, explains the significant incision observed along the troughs. The proximity of the ridge-trough structures to a rapidly prograding Neogene delta (the Skaggerak Delta) suggests a close link between pro-delta processes and the formation of these structures. The necessary confinement of the current system may have been provided by slope gullies, but conclusive evidence of this on the regional seismic data is lacking at present.

INTRODUCTION

described from the Cenozoic of the North Sea. They are on average 5-20 km long, and have a fairly regular spacing of 1-3 km. They have not been described in any previous publications on the North Sea Basin, and directly analogous structures have not, apparently, been described from any other basin. The lack of attention given to these extraordinary structures can probably be attributed to the fact that they are contained in an interval that has no economic potential from a petroleum perspective.

This paper describes a set of large-scale sedimentary structures that are developed in Upper Miocene to Lower Pliocene mudstones in the central North Sea. The structures are a series of ridges and troughs that are distributed throughout an elliptical area centred on UK Quadrant 22 (Fig. 1). The ridge-trough structures formed in a physiographic base-of-slope­ basin-floor setting, although the maximum water depths were relatively shallow, between 200 and 3 00m. The ridge-trough structures are by far the largest type of discrete sedimentary structure so far

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

285

286

J.

Cartwright

spacing of 25 km was used to map the distribution of the ridge-trough system (Fig. 1). The area of interest is located in the axis of the Cenozoic Basin, close to the region of maximum total Cenozoic: subsidence and sediment accumulation (Ziegler,, 1982; Joy, 1993).

The quality and quantity of petrophysical, litho­ logical and biostratigraphical data available for the Neogene interval in the North Sea is greatly inferior to that available for the more prospective Palaeogene and older formations. This contrast in the well data­ base explains why there have been so few studies on the Neogene of the North Sea, and contrasts markedly with the vast literature available on the Palaeogene (see Lovell (1990) for a comprehensive bibliography). The main aims of this paper are to describe the seismic facies characteristics of the ridge-trough structures with particular emphasis on their internal reflection configurations and external morphology, and to place these structures in a palaeogeographical context. The paper concludes with some preliminary ideas on possible depositional models.

Cenozoic Basin evolution

The Cenozoic North Sea Basin is widely regarded as a type example of a post-rift thermal sag basin (Dewey, 1982; White, 1989). Rifting ceased at the end of the Jurassic along most of the con­ stituent grabens of the North Sea Rift (Ziegler, 1982; Bertram & Milton, 1989; Cartwright, 1991). After a period of infilling of remnant depositional topography in the Early Cretaceous and early part of the Late Cretaceous, the rift axis and rift shoulders began a phase of regional sag-type subsidence that persists to the present-day (Joy, 1993). Uplift of the hinterlands bordering the North Sea Basin com-· menced at approximately the same time as regional subsidence (Watson, 1985; Rundberg, 1989), and

GEOLOGICAL SETTING

A grid of regional two-dimensional seismic data covering the central North Sea, with an average line

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Ridge structures, central North Sea

sediment supply changed from dominantly carbonate material during the Late Cretaceous, to dominantly clastic material throughout the Palaeogene and Neogene (Ziegler, 1982; Lovell, 1990). Basin-margin deltaic systems were established along certain sectors of the basin at different times, presumably in response to changes in the location of the main centres of marginal uplift. Uplift of the Scottish Massif in the Late Cretaceous to Early Cenozoic (Watson, 1985) is generally considered to pre-date the main phase of uplift of Fennoscandia, which probably dates from the Oligo-Miocene (Holtedahl, 1960; Spjeldnaes, 1975). One result of this diachroneity in uplift of the eastern and western marginal areas is that the pattern of infill of the Cenozoic is asymmetric. The main progradational clinoforms built out from the western margin for the Palaeocene to Oligocene interval, with sediment mainly supplied from the Scottish Massif and the Shetland Platform (Parker, 1975; Rochow, 1981; Lovell, 1990). A major reorganization of the clastic supply into the basin occurred at approximately the beginning of the Middle Miocene, and progradation switched to a dominantly easterly source. Over large areas of the Norwegian, Danish, and German sectors of the North Sea, the Middle Miocene to Late

287

Pleistocene interval is dominated by a large deltaic system, which prograded from the eastern basin margin in Denmark, crossing the Danish Sector en route to the central North Sea Basin (Fig. 2). This large deltaic complex is referred to in this paper as the Skaggerak Delta. This name was chosen because the main axis of sediment supply has been mapped along a NE-SW course running through the present­ day Skaggerak (Morgan, 1992). The scale and influence of the Skaggerak Delta on the Neogene fill of the basin is illustrated in two schematic sections based on interpret­ ations of regional seismic lines calibrated with the released well-data (Fig. 2). The Skaggerak Delta is represented on these two sections by large, westward-building sigmoidal and tangential­ oblique clinoforms above the regional middle Miocene marker (MM). The magnitude of the clastic input from the east is particularly apparent in the way that the clinoforms cross the basin axis and downlap or onlap the western margin of the basin. These two sections also illustrate the asymmetric pattern of basin-fill in the Cenozoic, with contrasting westerly and easterly sources of sediment supply in the Palaeogene and Neogene, respectively. The Late Miocene to Early Pliocene ridge-trough

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the configurations of the post-rift succession. The Late Miocene to Early Pliocene ridge-trough system is located at the base of slope of large-scale prograding clinoforms associated with the Skaggerak Delta. MM, middle Miocene marker; uM, upper Miocene; eP, lower Pliocene; Ku, upper Cretaceous.

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288

Cartwright

system occupies a position in the basin axis in the two sections illustrated in Fig. 2. The ridge-trough system is associated with the toesets of major west­ ward prograding clinoforms of the Skaggerak Delta. Progradation and aggradation of the clinoforms is matched in part by aggradation and southwesterly migration of the ridge-trough system.

EXTERNAL MORPHOLOGY AND INTERNAL CONFIGURATION OF THE RIDGE-TROUGH SYSTEM

Database and methodology

The descriptive parts of this paper are based largely on interpretations of reflection seismic data. Litho­ logical data used in this study are restricted to cut­ tings descriptions compiled in completion logs for over 100 released well-data in the central North Sea, and a limited amount of petrophysical data. Unfor­ tunately, the biostratigraphical control of many of these released well-data is very poor, in part due to lack of sampling and in part due to drilling methods (Gradstein et at., 1992). In addition, the lack of velocity information means that accurate well-to­ seismic calibration has not been possible. Conse­ quently, detailed lithostratigraphical correlations have not been successful. The analysis is therefore based on a seismic-stratigraphical approach, which

although limited by the vertical seismic resolution of the order of 1 0-15 m in the Neogene interval, is none the less sufficient to give some preliminary indications of the processes responsible for the devel­ opment of the ridge-trough structures. Detailed interpretation and mapping of the ridge­ trough structures was carried out on a study area of approximately 400 km2 in Quadrant 22, of the UK Sector of the North Sea. This study area is located in the central portion of the larger region in which the ridge-trough structures are developed (Fig. 1). The area was selected for detailed study because well­ and seismic data were readily available. Mapping was based on a seismic grid with an average line spacing of 1 km, which was sufficient to allow indi­ vidual ridges and troughs to be correlated without difficulty. The seismic data is a late 1980s vintage, and the quality and resolution of the Late Miocene to Early Pliocene interval are excellent. External morphology

The external morphology of the ridge-trough struc­ tures is illustrated with reference to a representative seismic section from the study area (Fig. 3). The section is oriented NE-SW, approximately orthog­ onal to the long axes of the ridge-trough structures in the study area. The base of the main ridge-trough system in the study area is picked at a low-frequency reflection that appears to mark the upper limit of a

Fig. 3. Seismic profile from the study area showing the cross-sectional geometry of the ridge-trough system. Note the diachroneity evident in the abandonment of active ridges from right to left. Also note the asymmetry of the post­ abandonment infill of the troughs. Section oriented SW-NE. Datum at top of section is 600ms TWTT (two-way travel time).

289

Ridge structures, central North Sea

precursor set of ridges and troughs (reflection A on Fig. 3). The upper boundary is picked wherever a clear onlap can be seen in a trough (e.g. reflection B on Fig. 3). The ridges have an external form that is best described as mounded (Mitchum et at., 1977). The intervening troughs have a channel-like profile. The ridge-trough structures have crest-to-crest spacing of between 1 and 3 km, and vary from almost perfectly symmetric to slightly asymmetric in form. The crest-to-trough amplitude of the structures defined using the upper and lower bounding reflec­ tions varies between 1 00 and 200m. The ridges have broad crests with fairly steep flanks, with dips of up to 7° (uncorrected for compaction). Throughout the study area there is clear evidence that the system of ridges and troughs is diachronous from northeast to southwest. This diachroneity can be seen, for example, in Fig. 3 in the contrasting topography and external form of the reflections comprising the ridge-trough system across the section. The relief on the shallow, high amplitude reflections that can be seen draping the ridges on the northeastern side of the section (e.g. reflection C) increases to the southwestern side of the section. Reflection C changes from a post-abandonment drape of the ridge in the northeast, into an internal reflection of the ridge-trough structure in the southwest. This indicates a younging of the active growth of the ridges to the southwest.

Internal reflection configurations

The typical internal configurations of the ridges and troughs are illustrated in Fig. 4, a seismic section parallel to that shown in Fig. 3, but reproduced at an enlarged scale to enhance the resolution. The ridges are composed of low- to moderate-amplitude reflections with moderate continuity. The lowermost reflections are concordant with the base, but there is increasing convexity exhibited by successively shallower reflections, such that the maximum relief between ridges and adjacent troughs is at the upper bounding reflection (e.g. at reflection A). Numerous reflection terminations can be observed within the ridges illustrated in Fig. 4, particularly towards the margins of individual ridges. Some of these reflection terminations may be indicative of genuine stratal truncation, pinch-out or downlap. However, considering the probable distortive effect on ray-paths of the troughs, and the possibility of multiples being generated in the overlying high­ amplitude reflections, it is possible that many of the reflection terminations are artefacts produced by interference or tuning. Genuine downlap is observed only rarely within the ridges, and consequently the most common diagnostic feature used to define a constructional sedimentary mound (Mitchum, 1985) cannot be confidently applied to the ridges illustrated in Figs 3 & 4. 1:

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290

J.

Cartwright

The troughs imaged on the section in Fig. 4 are characterized by an almost concordant set of high­ amplitude reflections whose lower contact is difficult to define. The overall configuration of the troughs is almost equivalent to an upside-down version of the ridges. No obvious incision can be seen at the base of the lower reflections defining the troughs, and direct correlation between reflections from the troughs to adjacent ridges is hampered by changes in reflection character and continuity. However, at least some of the low-amplitude internal reflections of the ridges pass directly into the concave, high­ amplitude reflections of the troughs. It is possible that the contrast in reflection amplitude between reflections in the troughs and ridges is due to lateral variations in lithofacies. For example, some form of lag deposit in the troughs could explain the higher amplitudes. However, it is also possible that the variation is due to tuning and the concave shape of the reflectors. At the upper levels of the ridge­ trough system, the relief across the troughs is gradually diminished by the development of onlap fill (e.g. Fig. 4, point X). Thus, although the troughs seen in Fig. 4 have the general morphology sugges­ tive of some type of channel, there is little conclusive evidence that the troughs on this section formed by an erosive process. A clearer indication of a possible erosional origin of the troughs can be seen on another NE - SW

oriented seismic section (Fig. 5), located several kilometres to the northwest of Fig. 4. Figure 5 shows an example of a trough that has a clearly recognizable basal reflection, with a classical channel profile (trough A). Erosional truncation can be seen on both margins of the trough, and there are sufficient genuine reflection terminations to define the position of the basal erosion surface. This basal surface (reflection X) passes laterally into the neighbouring ridges, and can be correlated with the erosional bases of the two neighbouring troughs (points B and C). Above the erosional base in trough A, the overlying reflection is concordant with reflection X, and above this there is a clear onlap of the trough margin by three subhorizontal reflections. The evol­ ution implied by these reflection configurations is: (i) incision, (ii) drape, and (iii) onlap fill. Taking the truncation and onlap relationships together, trough A can only reasonably be interpreted as an erosional feature. A similar pattern of erosion, drape and onlap can be seen on the other two troughs in Fig. 5. The variability in the detailed configurations of the three troughs seen on the section in Fig. 5 is typical of that observed throughout the study area. Incision is invariably observed along at least part of the axial extent of every trough. However, the degree of incision, and the clarity with which the incision is expressed, varies widely, both from trough to trough, and along any individual trough. The

Fig. 5. Seismic section oriented at right angles to long axes of ridge-trough structures, showing contrast in reflection configurations of three troughs, and evidence of erosional truncation at their bases . Datum at top of section is 600ms TWTI.

Ridge structures, central North Sea

troughs in Fig. 4 are typical of those in which basal incision is difficult to define from the cross-sectional reflection configurations, whereas trough A in Fig. 5 is a type example of a trough whose seismic expression most clearly conforms to that of an ideal incised profile. Figure 6 is a comparison of two representative seismic sections located along the axes (NW-SE) of a ridge and trough pair. The internal configurations of the ridge and trough are comparable, with a consistent parallel form with respect to the lower and upper bounding reflections. There is no evidence of any significant thickness variation along the axes of either the· ridges or the troughs, and no evi­ dence of internal downlap, progradation, onlap or erosional truncation in the along-axis direction of either the ridge or the trough. Interpretation

The reflection configurations of the ridge-trough system observed in cross-section are summarized in a schematic line drawing in Fig. 7. One of the most

291

puzzling aspects of the ridge- trough system illus­ trated in this figure is that there is a combination of features that are diagnostic of both constructive and destructive process elements. Included for com­ parison in Fig. 7 are two idealized line drawings through a definitive constructional mound and an incised channel separating interchannel ridges based on criteria outlined in Mitchum et al. ( 1977) and Mitchum (1985). The incision observed at the base of at least a portion of every trough clearly points to the activity of an erosive current regime for at least part of the time represented by the growth of the ridge-trough structures. In contrast, the increasingly convex­ upwards configurations of the reflections within the ridges is interpreted as evidence that the ridge struc­ ture is at least partly constructional. The common development of drape configurations in both ridges and troughs indicates that the constructional and erosive components of the system alternated with a dominantly pelagic or hemi-pelagic mode of deposition. Applying the criteria outlined by Mitchum (1985)

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..Ill

1km

I

Fig. 6. Comparison of two longitudinal seismic sections along a ridge (A) and adjacent trough (B). Note the absence of erosional truncation, onlap, downlap or progradation in the internal reflections of either the ridge or the trough.

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-

0

EROSIONAL CHANNEL

1 km

--- ---

BASE MIDDLE MIOCENE MARKER

INTER-MOUND TROUGH (NON-EROSIONAL)

Fig. 7. Summary diagram showing the main stratigraphical features of the ridge-trough system observed from the study

area. Idealized sections across an erosional channel and a pair of constructional mounds shown beneath main figure for comparison. Note that the ridge-trough structures have affinities with both these contrasting idealized cases. Note also that with low-resolution seismic data it may be difficult to distinguish the two cases, particularly If multiples obscure the true stratal configurations at the ridge margins.

for the recognition of different types of construc­ tional mounds, it is difficult to make a conclusive argument that the ridges are primary constructional features. Neither is it reasonable to interpret the ridges as purely remnant relief produced by erosion along the troughs. Without the benefit of detailed lithological calibration of the seismic data it is impossible to determine the extent to which erosive and depositional processes alternated or dominated throughout the evolution of the ridge-trough system.

MAPPING THE RIDGE-TROUGH SYSTEM

Mapping of the ridge-trough system was under­ taken in order to determine the spatial relationship between ridges and troughs and to determine their orientation relative to the major physiographic features of the basin. Detailed mapping of indi­ vidual ridges and troughs was restricted to the small study area (Fig. 1) because it was only there that

the relatively closely-spaced seismic grid ( 1-2 km) allowed correlation of individual ridges and troughs. Structural contour mapping of specific reflections within the Late Mioc�ne to Early Pliocene interval was attempted initially, but correlating horizons across the grid proved to be too difficult, mainly due to loss of reflection continuity in the troughs. Consequently it was decided to map the individual trough axes because these are narrower than the ridges and are therefore less likely to be miscorrelated. The map of trough axes in the study area is presented in Fig. 8. The most important features of this map are all well constrained by the seismic grid. These are: 1 the orientation of the trough axes is confined to a narrow range from 320°N to 340°N; 2 the troughs are linear to gently sinuous, and die out abruptly in both directions along their axes; 3 there is no interconnection between individual troughs - they neither bifurcate nor anastomose in either direction; 4 the length of those troughs that die out along

293

Ridge structures, central North Sea

··············

Seismic lines

-�- Axis of Trough • --·•

Termination of Trough

-
(decreasing relief in direction shown)

0

N

+

2

Fig. 8. Map of the trough axes in the study area. Also shown is the seismic grid used for the mapping. Note the

discontinuous array of trough axes in both an upslope and downslope direction (regional palaeoslope to the northwest).

their axes in both a northwesterly and southeasterly direction within the confines of the mapped area varies from 2 to 10 km - one trough is over 20 km long, and continues outside the limits of the seismic grid; 5 the trough axes are fairly consistently spaced over the area, with the intertrough spacing varying between 1 and 3 km. One of the most surprising aspects of Fig. 8 is the lack of any interconnection between adjacent troughs. The incision observed in the troughs (e.g.

Fig. 5) clearly indicates that currents strong enough to erode muddy sediments flowed along the trough axes. Because it has not been possible to correlate the erosional bases of individual troughs across the study area, the extent to which erosional activity in the troughs was synchronous is not known. For example, did the basal incision along the troughs take place at approximately the same time through­ out a narrow zone, a broad zone, or was it confined to a single trough at any one time? The overall diachroneity in the growth of the

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system from northeast to southwest evident on the seismic data (e.g. Fig. 3) and on the regional sections (Fig. 2) suggests that only a few of the trough axes indicated on Fig. 8 were active at any one time. This interpretation is based on the assumption that, because the depositional unit containing the ridge­ trough structures is diachronous as a whole, then it is probable that the time of initial incision of troughs was also diachronous. If this is correct, then Fig. 8 should not be viewed as a palaeogeographical reconstruction, but as an overall progression estab­ lished during several million years of deposition and erosion. Another important feature of the organization of the troughs is the observation that the troughs die out in both a northwesterly and southeasterly direction. Of the 20 troughs partly or wholly mapped within the study area, 13 die out along their axes at one end of the trough within the mapped area, and five die out at both ends along the trough axis (Fig. 8). A series of line drawing cross-sections constructed at intervals along one of the troughs

I I

� I

:

�

l

(A-B of Fig. 8) with terminations in both directions along-axis are presented in Fig. 9. This diagram illustrates the subtle changes in internal configuration of the trough, with variable depth of incision and a general decrease in relief towards both ends of the trough. At the limit of the trough, only a vague indication of an erosive feature can be discerned, and beyond this point, the trace of the trough axis projects directly into a ridge structure.

PALAEOGEOGRAPHICAL SETTING OF THE RIDGE-TROUGH STRUCTURES

A palaeogeographical reconstruction of the Early Pliocene North Sea Basin is presented in Fig. 10. The reconstruction is based on a basin-wide cor­ relation of a seismic horizon marking the end of the period in which the ridge-trough system formed . This horizon is dated as approximately 4 Ma, i.e . Early Pliocene. The depositional shoreline break

UNE1

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Fig. 9. Line drawings of seismic sections oriented SW-NE along the trough A-B (location in Fig. 8) , showing the change in reflection configurations of the trough along its axis. Note that the trough dies out towards the north-northwest and merges into a neighbouring ridge structure.

Ridge structures, central North Sea

was defined at the break of slope between topsets and foresets for clinoforms along the progradational sections of the basin margin (Posamentier & Vail, 1988). Progradational shorelines were differentiated from retrogradational or static shorelines by defining the extent of marine onlap for the Late Miocene to Early Pliocene interval. Major depocentres were

295

identified by mapping regions dominated by pro­ gradational clinoforms. The palaeobathymetry of the basin was estimated by measuring the relief on clinoforms using topset surfaces as mean sea-level projection datums. Figure 10 shows that the North Sea Basin was a narrow marine gulf or seaway in the Early Pliocene,

Actively prograding shoreline at 4ma

\9ma

\

Passively onlapped shoreline at 4ma

I

Position of depositional shoreline break for Skaggerak Delta at various times Progradation route of Skaggerak Delta

Bathymetric contour in metres I

�istribution of Ridge-Trough System �bandoned at 4ma

Other major distributary systems

Active at 4ma

t �·

. ,, ,: ,,

i :

i :

...... ................\� ........

Fig. 10. Palaeogeographical reconstruction for the Late Miocene to Early Pliocene. The North Sea is shown as a semi­

enclosed marine seaway. Marine connection to the Norwegian and Greenland Sea was open via a narrow strain between two deltas. The marine connection to the south is uncertain (see text). The position of the depositional shoreline break of the Skaggerak Delta is shown for various time intervals, and illustrates the rapid advance of this major depositional system.

296

J.

Cartwright

with a NNE-SSW trending elliptical central basin narrowing to the north and south. The bathymetry estimated from clinoform relief indicates that maxi­ mum water depths of 300-400 m were developed in a basin-plain region centred on Quadrant 22 of the UK Sector (ct. Figs 1 & 10). Palaeobathymetric constraints from the limited sampling of microfauna in wells in the central North Sea are imprecise, but are broadly consistent with the magnitude of water depth interpreted from the clinoform geometry. Water depths derived from an analysis of subsidence of the central North Sea are in close agreement with the values presented here (Joy, 1993). The region of active ridge-trough structures is shown to be confined to the southern end of the basin plain in the Early Pliocene. The full extent of the ridge-trough structures (active and abandoned) is also indicated in Fig. 10, in order to show the overall southwestward migration of the ridges and troughs from the Late Miocene to the Early Pliocene. The most probable marine connection into the North Sea was at the northern end of the basin, with the Norwegian-Greenland Sea (Ziegler, 1982). Whether there was an additional connection to the Bay of Biscay to the south is not clear. Neogene strata are absent throughout much of the south­ western part of the North Sea (south of this study area), and a major erosional unconformity is developed near the base of the Late Pleistocene (Cameron et al., 1987). Hence it has not been possible to extend the palaeogeographical recon­ struction into the southern North Sea and clarify the nature of any possible marine connection. Four main depocentres prograded into the North Sea in the Early Pliocene. The courses of distributary systems supplying these depocentres are sche­ matically illustrated in Fig. 10, to give some indi­ cation of probable source areas, although it should be noted that there is little direct evidence concerning the exact positions of the distributary systems. In the extreme north of the basin, two large deltaic systems prograded towards each other simultaneously, narrowing the marine connection to the Norwegian­ Greenland Sea. In the south (south of 57°30'), clino­ forms of the Skaggerak Delta dominated the infilling of the basin. The rapid westward progradation of the Skaggerak Delta is indicated on Fig. 10 by the mapped positions of the depositional shoreline break through the Late Miocene to Early Pliocene interval. A relatively minor depocentre was confined to an embayment in the otherwise passively onlapped southwestern margin of the basin.

In summary, the ridge-trough system in the Early Pliocene occupied a base-of-slope-basin-fioo:r position, in water depths of 300-400 m, along the southern portion of a narrow marine basin that was being rapidly infilled by a major delta prograding east to west and either narrowing or closing any possible marine connection to the south.

DISCUSSION

Any explanation of the genesis of the ridge-trough system needs to take account of the following key observations. 1 The ridge-trough system developed during a short interval of time (c. 1-2 Myr) spanning the Late Miocene to Early Pliocene. No remotely similar features have been documented from comparable base-of-slope depositional systems in the North Sea at any other period in the Cenozoic. 2 The dimensions of the ridges are large in com­ parison to most constructional bedforms. They are of the order of 10 km long, 1-3 km wide, and 100-200 m thick. 3 The orientation of the long axes of the ridges and troughs range from 320°N to 340°N. 4 Ridges and troughs exhibit a variety of construe-· tiona) and erosional characteristics, from stratal thickening towards the ridge axes (constructional), to truncation of ridge flanks and incision along trough axes (erosional). 5 There is no evidence of lateral accretion of indi­ vidual ridges. The internal reflection configurations indicate that ridges grew by aggradation: abandon-­ ment was followed by a migration of active growth to new ridges to the southwest (of the inactive ridges). 6 The troughs do not interconnect and, where mapped to their full extent, die out along both axial directions. 7 The zone of ridge growth migrated from northeast to southwest with time. Depositional models Mudwaves formed by the action of bottom currents

Large-scale, constructional muddy waveforms formed by the long-term action of bottom currents have been described only from slope and abyssal­ plain settings in oceanic basins (see Pickering et al. (1989) for a comprehensive review), but there is no

Ridge structures, central North Sea

reason why similar structures should not be found in locations other than oceanic basins, provided that the necessary current regime was operative for a sufficient period. The dimensions and internal stratification of the ridge-trough system are com­ pared with a compilation of line drawings of mud­ waves deposited by bottom currents in various oceanic localities (Fig. 11). The relief of these mud­ wave examples is of the order of 10-20 m, i.e. considerably less than the c. 100-200 m relief of the ridges in the central North Sea. However, the shape, internal reflection configurations and spacing of some of the mudwaves are comparable with the ridge­ trough structures. Better grounds for comparison can be made with large-scale bedforms associated with a sediment drift in the Northeast Atlantic, described by Kidd & Hill (1986). The mudwaves developed on the Feni and Gardar Drifts were studied by these authors using a combination of sparker and GLORIA sidescan sonar data. The resulting map of the mudwaves (Fig. 12) shows a pattern with some similarity in scale and organization to that exhibited by the ridges and troughs from the central North Sea study area, although there is less regularity to the mudwave organization than to the ridge-trough system (Fig. 8). An interesting feature of the map shown in Fig. 12 is that the mudwave crests close at both ends of their long axes. This is in contrast to the ridge­ trough system where the mapping showed that it is the troughs that terminate in both directions along­ axis. It is not clear whether this attribute is diagnostic of mode of origin. In common with the ridges in the North Sea, the mudwaves on the Feni and Gardar Drifts show no evidence for lateral migration on an individual basis. The long axes of the mudwaves are thought to be oriented obliquely to the bottom­ current flow direction. The limited basis for comparison means that it is hard to make an unequivocal case that the ridge­ trough system in the central North Sea was the result of bottom-current activity. The similarity with the published descriptions of mudwaves associated with sediment drifts is sufficient to at least consider bottom currents as a possible driving mechanism for the development of the ridge-trough system. However, although a bottom-current mechanism may offer a reasonable explanation for both the distribution of the ridge-trough system in the basin axis· and the long-term migration of the zone of active ridge-trough structures towards the south­ west (i.e. along the likeliest route for currents linking

297

the two oceanic basins), it is difficult to reconcile this mechanism with the often significant incision observed along many of the troughs. In particular, it is difficult to see how bottom currents could have produced the incision of tens of metres of fine­ grained sediments along the troughs in a north­ northwest orientation (i.e. transverse to the long axis of the basin and hence the likely flow route for bottom currents), and on a longer term lead to a general southwesterly migration of the entire ridge­ trough system. In conclusion, the interpretation of the ridges as some form of bottom-current deposit cannot be discounted, particularly in view of the uncertainties regarding a southerly marine connection, but the observation of incision along the troughs argues against this depositional model. Pro-delta current regime

An alternative model for the development of the ridge-trough system is suggested by two observations: 1 the approximately orthogonal relationship be­ tween the axes of the ridges and troughs (NNW­ SSE to NW-SE) and the strike of the nearest depositional shoreline break (NE-SW); 2 the skewed distribution of the ridge-trough struc­ tures to that sector of the basin plain immediately adjacent to the toe of the Skaggerak Delta, which was by far the most significant Mio-Pliocene pro­ gradational system in the North Sea. The evidence of erosional truncation and incision in the troughs (Fig. 5) implies that currents flowing along the trough axes were of sufficient strength to result in the cumulative erosion of tens of metres of muddy sediment. If the troughs were foci for strong currents flowing along the trough axes, it is reason­ able to infer that their origin was directly upslope, either at the delta front or on the upper reaches of the pro-delta. Because the nearest delta front to the ridge-trough system was along the southeastern margin of the basin, it is reasonable to argue that the ridge-trough system developed in response to a current regime associated with the rapidly prograding Skaggerak Delta. If some form of pro-delta current regime was responsible for the development of the ridge-trough system, several questions need to be addressed. 1 What was the specific nature and flow regime of the current system? 2 How and where did the currents originate?

298

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0

1km

Fig. 1 1 . Line drawings of high-resolution shallow seismic images of mud waves from the Central and Southern Atlantic

showing typical dimensions and internal configurations. Taken from Pickering eta/. (1989). Line drawing of ridge-trough structures displayed at the bottom for comparison.

299

Ridge structures, central North Sea

N

1

0

KILOMETRES

5

r-------7'"--.:--� so·N

Fig. 12. Detailed bathymetric map

over a field of sediment waves located on the Feni Drift, Rockall Trough, North Atlantic. Note that the sediment wave crests (stippled) are discontinuous. Taken from Kidd & Hill (1986).

so•

30'

How did these currents result in the stratal patterns observed on the seismic data? 4 In the absence of any evidence linking similar ridge-trough structures to the other deltas along the basin margin, why was this current system uniquely developed in connection with the Skaggerak Delta? 5 Why was this current regime operative only during

3

the Late Miocene to Early Pliocene, given that the Skaggerak Delta was first established in the Early Eocene, and continued prograding until the Early Pleistocene? 6 Why did the locus of current activity migrate towards the southwest from the Late Miocene to the Early Pliocene? The postulated downslope current system must

300

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have been organized in linear, confined flow paths with lateral dimensions of 1-2 km in order to produce the regular, isolated and linear troughs mapped in the study area. To explain the lack of interconnection between the troughs and the obser­ vation that the troughs die out at both ends of their axes (Figs 8 & 9), it is necessary to invoke downslope currents that were either of a finite extent in the flow direction, or currents whose strength varied along the flow direction. Both of these conditions would be expected to apply to a downslope pro-delta current system. Any downslope currents would reach their maximum velocity at the base-of-slope, and would gradually lose velocity along the horizontal basin floor. The finite longitudinal extent of the incision observed along the troughs could, therefore, be interpreted as a result of the current exceeding a critical erosive velocity threshold for the muddy sediments as the currents accelerated downslope in a northwesterly direction. Then, as the power of the current decreased as flow proceeded along the basin floor, the degree of incision would also decrease until the trough was no longer recognizable as such (Fig. 9). One of the problems with this pro-delta current model is the large distance from the likely point of origin of the currents at the head of. the slope, to the toe region where the ridge-trough structures formed. A question that cannot be resolved as yet is how sufficient flow confinement would have been maintained from the delta front, where the currents may have originated, for example, as some form of turbid underflow (Bates, 1953), over a distance of tens of kilometres to the northwest at the base of slope, where incision commenced. The extent to which the Coriolis effect would have deflected downslope flowing currents originating at or close to the delta front is also a problem with this depositional model. If a modern system, such as the Huanghe (Wright et al., 1990), represents a reason­ able analogue, then along-slope deflection should be expected (Hill, 1984). It is also difficult to envisage a mechanism whereby the downslope flow would maintain its confined flow pattern once it began to decelerate at the lower break of slope. One method of providing the necessary flow confinement would be if there was a pro-delta channel or gully system linking the distributary channels at the delta front with the troughs at the base of slope (Fig. 13A). Linear gully systems have been observed and mapped on the present-day Scotian slope of eastern Canada (Piper & Sparkes, 1987). The scale and

spacing of these slope gullies is closely comparable to those exhibited by the troughs in the central North Sea, but the erosional relief of the Scotian gullies is only of the order of 10 m. It is possible, however, that even shallow erosional gullies could provide the necessary confinement and guide the flow paths from the upper reaches of the pro-delta to the toe region. Minor erosional features have been observed on the pro-deltaic foresets of the Skaggerak Delta on regional seismic data (Morgan, 1992), but the wide spacing of the seismic grid available at this time makes any direct correlation of these features with the ridges and troughs in Quadrant 22 extremely tenuous. A direct relationship between the distributary channels on the delta front and confined axes of flow at the base of slope is appealing in that it may explain the overall southwesterly migration of the ridge-trough system from the Late Miocene to the Early Pliocene. If downslope currents originated at the outlets of the major distributary channels on the delta front, possibly during one of the glacially forced relative lowstands of sea-level (Haq et al., 1987), then the locus of trough development would prob­ ably have been confined to the sector of the basin floor directly downslope from the major channels. If for any reason the channels avulsed at the delta front, then the locus of active ridge-trough growth would be expected to migrate accordingly (Fig. 13B) . A depositional model linking the evolution of the Skaggerak Delta directly to the development of the ridge-trough system also provides a satisfactory explanation for the aggradational nature of the stratal patterns interpreted on the seismic data (Fig. 7). The Late Miocene to Early Pliocene is known to have been a period of rapid basin sub­ sidence (Kooi & Cloetingh, 1989; Joy, 1993), and this led to increased aggradation of the basin-margin deltas (Morgan, 1992). The aggradational character­ istics observed for the troughs and ridges can be viewed as a response to the aggradation of the deltaic depositional system as a whole, including the foreset and bottomset depositional units, and any downslope current system originating at the delta front. The pro-delta current mechanism can thus explain many of the observations documented in this paper, but a major unresolved question is why the ridge-­ trough structures were restricted to the Late Miocene to Early Pliocene in particular? What was so unique about this period that either a pro-deltaic downslope current system or a bottom-current system should

301

Ridge structures, central North Sea

NE

A Pro-delta

"'

-

Topset

} aggradation Trough

' Bottomset

500m

�

C u r rent flow path

ll----.,--

'�

SW

--,

0

10

20km

Approximate scale

B

Palaeogeography time t1

Palaeogeography time t2

N

'

Avulsion of major distributary c h a n nel

Progradation of depositional shoreline break

Fig. 13. (A) Evolutionary reconstruction of the development of the ridge-trough system based on a pro-delta current

regime. Flow confinement from the delta front is provided by a system of slope gullies. Maximum flow velocity is reached at the base of slope, and the trough incision begins at this position. Deceleration along the basin floor results in the elimination of the troughs in a north-northwest direction. (B) Migration of the ridge-trough system is achieved by avulsion and channel switching at the delta front.

have been operative for this specific period, and not throughout the entire Cenozoic? The major climatic change that took place with the onset of the glacial­ interglacial cycles in the Late Miocene to Early

Pliocene (Lidmar-Bergstrom, 1982) may be part of the solution. The resolution of this question will require considerable additional research, not least of which would be a comprehensive study of the

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morphological evolution and palaeogeographical setting of the Skaggerak Delta throughout the Cenozoic. The impetus for additional work on the commer­ cially unprospective Neogene interval of the North Sea Basin may be provided by the enigmatic linear sandbodies of the Eocene in the North Sea. The troughs described in this paper may be examples of large sedimentary features that resulted from a process involving the transport and confinement of reservoir-prone units from the delta-top into deep­ water environments. If the troughs were a locus for sand transport and deposition, then the resultant reservoir geometry would be a linear body compar­ able in scale and thickness to reservoir sands devel­ oped on the Eocene slope and basin floor.

CONCLUSIONS

1 The ridge-trough structures of the central North Sea are unique features in the context of the Cenozoic stratigraphy of the North Sea, with apparently no direct analogues in the published literature. 2 The linearity of the structures and the regularity of their spacing suggest that they either formed in a current regime that was highly oblique or even transverse to their long axes, or in one that was parallel to their long axes. 3 Two contrasting current regimes are regarded as the basis for a genetic model: (i) bottom currents flowing along the NNE-SSW basin axis, or (ii) pro­ deltaic currents flowing downslope in a NW-SE direction from the delta front of the actively pro­ grading Skaggerak Delta. 4 Although both current regimes can provide satis­ factory explanations of the dimensions and mor­ phology of the ridge-trough structures, the often pronounced incision observed in the troughs is dif­ ficult to reconcile with bottom currents flowing obliquely across the long axes of the troughs. The incision is compatible with currents flowing along the trough axes. Hence the preferred model is the pro-delta current regime. 5 A weakness in the pro-delta current model is the lack of a conclusive explanation as to how the necessary confinement in the current system was initiated and maintained down the slope and across the basin floor so as to produce the linearity observed for the troughs. Slope gullies observed on foresets of the Skaggerak Delta may have provided the

necessary confinement, but much additional map­ ping is required to confirm this suggestion.

ACKNOWLEDGEMENTS ·

I would like to thank Fina Exploration Ltd for financial support and Fina and partners for access to well and seismic data. Thanks also to colleagues at Fina for valuable discussions, namely Rod Laver, Dave Fassom, Mick Cope and Joe Staffurth. The manuscript was greatly improved as a result of the many suggestions made by the reviewers, Chris Baldwin and Phil Hill, and the Haroldesque editorial skill of Guy Plint. Thanks also to Howard Johnson and Dick Selley for suggestions for improving the manuscript. Finally, my sincerest thanks to Harold for putting me on the right track over the last 15 years (derailments are my responsibility !). Tony Brown and Jubrail produced the figures.

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K.T . , HISCOTr, R.N. & H EIN , F.J. ( 1989) Deep Marine Environments, Clastic Sedimentation and Tectonics . Unwin & Hyman, London, 416 pp. PIPER, D.J.W. & SPARKES, R. ( 1987) Proglacial sediment instability features on the Scotian slope at 63 degrees W. Mar. Geol. , 76, 1 5 - 3 1 . PoSAMENTIER, H.W. & VAI L , P.R. ( 1988) Eustatic controls

PICKERING,

on clastic deposition I I : sequence and system tracts models. In: Sea-level Changes: an Integrated Approach (Eds Wilgus, C.A. , Hastings, B.A. , Kendall, C. G.St C. , Posamentier, H.W. , Ross, C.A. & Van Wagoner, J.C.), Spec. Publ. Soc . econ. Paleontol. Mineral. , Tulsa, 42, 125 - 154. RocHow, K.A. ( 198 1) Seismic stratigraphy of the North Sea 'Palaeocene' deposits. In: Petroleum Geology of the Continental Shelfof North- West Europe (Eds Illing, L.V. & Hobson, G. D.) , pp. 255-266. Institute of Petroleum,

London. Y. ( 1989) Tertiary sedimentary history and basin evolution of the Norwegian North Sea: an integrated approach . Dr. Ing. thesis, University of Trondheim. SPJELDNAES, N. ( 1975) Palaeogeography and facies dis­ RuNDBERG,

tribution in the Tertiary of Denmark and surrounding areas. Norges Geol. Unders. , 316, 289- 3 1 1 . WATSON , 1 . ( 1985) Northern Scotland as an Atlantic-North Sea divide. J. geol. Soc. London , 142, 22 1 -243. WHITE, N.J. ( 1989) Nature of lithospheric extension in the North Sea. Geology, 17, 1 1 1 - 1 14. WRJGHT, L . D. , WISEMAN, W .J . , YAN G , Z.S. , BORHOLD, B . D . , KELLER, G.H. , PRJOR, D.B. eta!. ( 1990) Processes of marine dispersal and deposition of suspended silts off the modern Huanghe (Yellow River). Continent. Shelf Res. , 10, 1 - 40 .

ZIEGLER, P. A. ( 1982) Geological A tlas of Western and Central Europe . Shell International Petroleum BY, The Hague, 130 pp.

Spec. Pubis int. Ass. Sediment. (1995) 22, 305-327

Millstone Grit cyclicity revisited, II: sequence stratigraphy and sedimentary responses to changes of relative sea-level O L E J . M A R T I N S E N* J OH N D . C O L L I N S ONt and B R I A N K. H O L DS W O RTH:j: *Norsk Hydro Research Centre, N-5020 Bergen, Norway; tCollinson Jones Consulting, Westgate House, Hill's Lane, Shrewsbury SYI 1 QU, UK; and

:j:

Department of Geology, Keele University, Keele, Staffordshire ST5 5BG, UK

ABSTRACT The Namurian (Upper Carboniferous) Millstone Grit cyclothems of northern England were formerly interpreted as the results of deltaic progradational episodes formed during lowstands of glacial-eustatic sea-level, punctuated by widespread fossil-bearing marine bands formed during sea-level highstands. New research, based on detailed fieldwork in the Askrigg B lock and Craven Basin area, has led to a reinterpretation of Millstone Grit cyclothem stratigraphy. Although the broad framework of a cyclothem bounded by goniatite-bearing marine bands is retained, a more complex internal archi tecture is recognized . The tectonic framework of northern England in the Carboniferous evolved from an active extensional phase in the Dinantian to a thermal sag phase in the Namurian. As a response, two different marine­ band-bounded sequence types developed in the Askrigg Block and Craven Basin area: (i) a 'block-edge' sequence, formed when there was significant differential bathymetry between earlier active fault blocks and adjacent deep basins, and (ii) a ramp sequence, formed when clastic fill and decaying differential subsidence had reduced the bathymetric contrast between the Askrigg Block and the Craven Basin to a gently inclined ramp. The main differences between these two sequence types are: (i) the formation and preservation of fine-grained highstand strata, which is a particular featu re in the 'block-edge' sequence; (ii) formation of a falling-stage systems tract, which occurred only in the ramp sequence; (iii) the formation of a lowstand turbiditic wedge in the 'block-edge' sequence; and (iv) the greater thickness of the 'block-edge' sequence. These differences can be attributed to varying basin bathymetry and available accommodation as a response to the evolving tectonic and palaeogeographical situation.

INTRODUCTION

1992; Maynard, 1992). This area is particularly favourable for an appraisal of the consequences of sequence stratigraphical analysis particularly because of the highly refined Namurian biostratigraphical framework (Fig. 2; Holdsworth & Collinson, 1988). In particular, based on a detailed data set from the Askrigg Block and Craven Basin area in northern England (Martinsen, 1990), we discuss: (i) the goniatite-bearing marine bands and their relation to sea-level curves; (ii) the architectural arrangement of Millstone Grit cyclothems; (iii) the sedimentary responses to relative sea-level change; and (iv) the

Recent advances in stratigraphical methods, par­ ticularly in sequence stratigraphy (e.g. Embry & Podruski, 1988; Van Wagoner et a/., 1988; Galloway, 1989; Cross & Gardner, 1991), have caused new light to be shed upon the interrelationships of stra­ tigraphy and sedimentology (ct. Walker (1990, 1992) for a critical discussion). One basinal succession for which new interpretations based on sequence stra­ tigraphical analysis are being presented and debated is the classic Namurian (Upper Carboniferous) Millstone Grit of northern England (Fig. 1; cf. Martinsen, 1990, 1993; Read, 1991; Collinson et al.,

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

305

O.J. Martinsen, J.D. Collinson and B. K. Holdsworth

306

ALSTON BLOCK BLF STAINMORE BASIN

N

(Cross-section in Fig. 6a) ASKRIGG BLOCK

0

4, 6

2

(Cross-se . Ctton Lancaster CRAVEN BASIN

in F'

tg. 6b) Cross-section in Figs 6c and 8)

Fig. 1. Location map of northern England and simplified structural map of the study area. For a detailed structural map of northern England, see Gawthorpe eta/. (1989). The location of cross­ sections in Figs 6 & 8 are shown, and the location of localities mentioned in the text are indicated by numbers: 1, Lad Gill, Stonesdale; 2, Kettlesbeck, near Settle; 3, Cononley Beck, Airedale; 4, Backstean Gill, Nidderdale; 5 , Horn Crag Q uarry, Silsden; 6, general area of several localities on Askrigg Block used in Fig. 8. For grid references, see Figs 6 & 8.

PF, Pennine Fault; BLF, Bu tterknowle-Lunedale Fault.

existence of a Namurian major-scale cyclicity. The aim is to extend existing stratigraphical models for the transition from the Askrigg Block into the Craven Basin (Martinsen, 1993) and to provide the impetus for a thorough re-evaluation of Millstone Grit cyclicity.

NAMURIAN GEOLOGICAL FRAMEWORK

Stratigraphy

The concept of sedimentary cyclicity is not new in Carboniferous geology and dates back to Phillips (1836), who recognized that the Carboniferous Yoredale cyclothems of northern England consist of repeated lithological successions and that there are lateral variations within these. For the Namurian, the detailed stratigraphy is dependent largely on the detailed goniatite bio­ stratigraphy established by the work of Bisat (1914, 1924) and Hudson (1934, 1939, 1941, 1944). Because of inadequate exposures, complex structure, and local, inconsistent stratigraphical naming of coarser grained lithologies, the goniatite-bearing marine bands are vital for stratigraphical analysis. The later work of Ramsbottom and co-workers (Ramsbottom et al., 1962, 1978; Ramsbottom, 1 977, 1979) has been instrumental in consolidating and expanding Bisat's and Hudson's work. Holdsworth & Collinson (1988) presented a goniatite-based biostratigraphical scheme. It presents a refined subdivision of Namurian stratigraphy and will be used in the following. Other

schemes include those of Riley (1985) and Riley et al. (1987). There are about 60 marine bands in the Namurian succession, most containing their own goniatite species (see Fig. 2; Holdsworth & Collinson, 1988). The average duration of a Millstone Grit cycle is about 180 000 yr, although significant variability probably existed (Maynard & Leeder, 1992). Most marine bands have an interbasinal extent (Ramsbottom et al., 1978). Some marine bands are particularly widespread and extend over much of Europe, in fact, a few can be traced from western Ireland to the Urals (Wagner et al., 1979; Martinsen, 1990). Other essential work for establishing a reliable Namurian stratigraphical framework was carried out in numerous mapping projects organized by the British Geological Survey. Perhaps the most sig­ nificant of these is the classic Rossendale Anticline Memoir (Wright et al., 1927), where the authors not only presented a reliable stratigraphical scheme based on goniatites, but also showed a sophisticated understanding of how a particular Millstone Grit cyclothem operated in terms of depositional setting and relative sea-level change. Structure and regional depositional patterns

The central and northern England area (the Central Province) was in early Carboniferous time domi­ nated by a pronounced differential bathymet1ry related to carbonate build-ups on tilted fault blocks (Gutteridge, 1989), some of which were granite­ cored (Miller & Grayson, 1982; Fig. 3A). The

Millstone Grit cyclicity MINOR CYCLE H2c(ii) z

<X: f=

a: 0 "�

<X:

::J

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H. smithi-undulatum

H2a(i)

H. smithi/Hd. proteus

H1b(iv)

H. aff. beyrichianum

H. beyrichianum H. beyrichianum

Q

H1a(iv)

Homoceras sp. nov.

H1a(iii)

I.

subglobosum

H1a(ii)

I.

subglobosum

H1a(i)

I.

subglobosum

E2c2(iii)

N.

nuculum

E2c2(ii)

N.

nuculum

E2c2(i)

N.

nuculum

E2c1

N.

stellarum

E2b3

Ct. nititoides

E2b2(ii)

F.

E2b2(i)

Ct. nitidus

E2b1(iii)

C. subplicatum Grp.

E2b1(ii)

Ct. edalensis/Ct. bisati

E2b1(i)

C. subplicatum Grp.

E2a(iii)

E.

yatesae

E2a(ii)

E.

ferrimontanum

E2a(i)

C. cowlingense

E1c

C. malhamense

UJ

0 :I: u

Q. Q. ::J c co ·c: ::J

E

C, Cravenoceras; Ct, Cravenoceratoides; E, Eumorphoceras; F, Fayettevillae; H, Homoceras; Ht, Homoceratoides; N, Nuculoceras; V, Valites.

. H2a(ii)

eostriolatus

H1b(i)

Q;

Fig. 2. Excerpt of the goniatite band stratigraphy established for the Namurian. Based on Holdsworth & Collinson (1988).

V.

H. undulatum

<X: a:

co

co

H2c(i) H2b

H. aff. beyrichianum

u

"-

Ht. prereticulatus

H1b(iii)

-2

co z ..... 0 t

GONIATITE SPECIES

H1b(ii)

z

(j)

307

z

<X: 0

a: UJ C!l (f) z a:

<X:

Pendleian (part)

shallow-water limestone platforms were separated from adjacent deeper water basins by steep sub­ marine slopes (Fig. 3A; Collinson, 1988; Gawthorpe et al., 1989). Most authors agree that this configur­ ation resulted from N -S extension north of the northwards migrating Variscan front (Leeder, 1982, 1988; Leeder & McMahon, 1988). Active extension probably ceased in early Namurian time, causing a change from mainly fault-controlled subsidence to thermal sag, which progressively decreased during the Namurian (Fig. 3B; e.g. Leeder & McMahon, 1988). The area discussed in this paper, the Askrigg Block to Craven Basin transition, lies at the northern margin of the central Province, and the two components are separated by the North Craven

holmesi

Fault (Fig. 1). The Askrigg Block is cored by the Wensleydale Granite (Fig. 3A), which contributed significantly to maintaining the Askrigg Block as a structural high throughout Carboniferous time. The Askrigg Block was tilted to the north during the main extensional phase, and a half-graben basin, the Stainmore Basin, was formed along its northern edge. This basin is bounded to the north by the Butterknowle- Lunedale Fault, which also marks the southern end of the Alston Block. The Stainmore Basin ceased to subside differentially by early Namurian time (based on uniform thickness data of overlying strata across basin), and throughout the rest of the Namurian, had no significant impact on depositional patterns (Martinsen, 1990). Therefore, for the cyclothems investigated in the present paper,

308

0.1. Martinsen, J.D. Collinson and B . K. Holdsworth

A END-DINANTIAN BLOCK-AND-BASIN BATHYMETRY CENTRAL PROVINCE N-S CROSS-SECTION

N 30km

Rossendale Block

25km

s

30km

Wensleydale Granite

Only relative vertical scale TRANSITION OF SOUTHERN END OF ASKRIGG BLOCK

B N

S

s

N

E1c-E2a(iii); pronouced 'block-edge' with turbidite-fronted deltas

E2b2(i)-Kinderscoutian: gently-inclined ramp with shallow-water deltas

Inactive Craven Fault system Only relative vertical scale

Ca. 40km

Fig_ 3. (A) North-south cross-section across the Central Province to show the end-Dinantian block-and-basin physiography. Modified from Collin son (1988). (B) Cartoon illustrating the change of the southern end of the Askrigg Block from a pronounced 'block-edge' to a ramp in the Namurian. Based on Martinsen (1993).

the Stainmore area is considered part of the Askrigg Block. Although the Namurian basin-filling sediments of the Central Province were partially deposited in deep water and therefore are expected to be largely marine, it is important to realize that the basin complex formed in an intracratonic setting far removed from the nearest ocean. In fact, the contemporaneous Mid-European Ocean probably lay as far south as the present position of Portugal (Anderton et at., 1979; Ziegler, 1986, 1988), possibly connected to the Central Province through the Variscan foredeep in central Europe. The Central Province may also have had a connection to the Canadian Maritime Provinces (Ziegler, 1988). This intracratonic setting and isolation from the ocean is important for explaining marine bands and Millstone Grit cycle salinity variations (Collinson, 1988; Holdsworth & Collinson, 1988). The initial Namurian depositional systems pro­ graded from the north and northeast into the Central Province. At the southern end of the Askrigg Block, across the Craven Faults at the transition to the Craven Basin, a pronounced

bathymetric break occurred, which was a first-order control on early Namurian depositional patterns. To the south, turbidite sedimentation was import­ ant in the earliest Namurian cycles (Pendleian and earliest Arnsbergian; Baines, 1977; Collinson, 1988; Martinsen, 1993). With time (from latest Arnsbergian), this bathymetric break transformed into a gently inclined, southerly dipping ramp (Fig. 3B), resulting from decreased differential sub­ sidence rates due to transition to thermal sag and progressive fill of the Craven Basin. Consequently, the potential for significant turbidite sedimentation immediately south of the Craven Faults declined, and shallow-water deltaic sediments dominated deposition in the later Namurian cycles of this area (cf. Martinsen, 1993). Further south in the Central Province, in the Pennine Basin, basinal shales dominated deposition until the mid-Namurian (Kinderscoutian). These shales were succeeded first by turbidite-fronted deltas and then by shallow-water deltas (Walker, 1966; Collinson, 1969, 1988). Although this tran­ sition occurred at a much later stage compared with the northern margin of the Craven Basin, it was

Millstone Grit cyclicity

perhaps not a result of ongoing, prolonged differ­ ential subsidence in the south. Rather, because the clastic supply was mainly from the north, the feeding systems did not manage to prograde into the area until the middle part of the Namurian. Basin-filling was therefore delayed in the Pennine Basin compared with the Craven Basin, probably as a result of the clastic supply being absorbed by depocentres beneath the southern North Sea. In the Staffordshire Basin, a slightly different situation existed. There, protoquartzitic turbidites were supplied from the southern Midland land mass in several minor cycles from as early as E1b (Holdsworth, 1963; Trewin & Holdsworth, 1973). In the Widmerpool Gulf to the east, the same situation seems to have taken place (Aitkenhead, 1977). Although sand was supplied to these areas at this early stage, the supply was insufficient to fill the available accommodation space, and complete basin filling did not occur until the Marsdenian R2b(ii) Ashover/Roaches Grit was supplied from the north (Jones, 1980; Collinson, 1988).

GONIATITE-BEARING MARINE BANDS AND RELATIONSHIP TO SEA-LEVEL

The marine bands, which allow such a refined bio­ stratigraphical breakdown of the Namurian, are virtually the only fossiliferous sediments in Namurian basinal successions. Fossiliferous intervals range from a few centimetres up to around 2 m in thickness and generally constitute only a small proportion of the total Namurian succession (Collinson, 1988; Holdsworth & Collinson, 1988). Sediments between marine bands are virtually devoid of body fossils, irrespective of whether the sediments are fine­ grained, deep-basinal mudstones, thick sequences of turbidite sandstones or progradational deltaic suc­ cessions (Fig. 4). Where the intervening sediments are deltaic progradational successions, the marine bands occur closely above an abandonment surface or just below the base of the next upward-coarsening pro-delta unit, leaving a clearly asymmetric cyclo­ them motif (Fig. 4). These observations, coupled with the widespread nature of many marine bands (see above) suggests an external control on the occurrence of goniatite-bearing marine bands. Most Namurian researchers have believed this mechanism was glacial-eustasy, a theory which is increasingly supported by growing evidence for widespread

309

glaciations in Gondwana during the Carboniferous (Caputo & Crowell , 1985; Veevers & Powell, 1987). Although goniatites constitute the stratigraphically most useful faunal phase of marine bands, it is not uncommon for various other faunal elements to occur in distinct faunal phases stacked one above the other. Based on detailed analysis of a cored condensed basinal succession at Ashover, Derbyshire, an ideal succession of faunal phases has been constructed (Ramsbottom et al., 1962; Holdsworth & Collinson, 1988; Fig. 4). In the Ashover boreholes the total succession from base Namurian to the Marsdenian is of basinal mudstones with no coarse clastics. The Ashover succession was deposited on part of the Derbyshire Massif, a mudstone-draped Dinantian carbonate platform, isolated from clastic supply by a partial moat of deep-water gulfs. The isolation from clastic supply means that there is virtually no change of lithofacies between marine bands and intervening barren mud­ stones. In such circumstances it is very difficult to conclude that marine bands are purely reflections of sediment starvation and condensation. The presence or absence and the changes of fauna in a succession like Ashover almost certainly result from changes in the overlying water column, in particular its salinity. Of the faunal elements present in the sequence, only the thick-shelled goniatites are thought to reflect fully marine salinity (Fig. 4). The other phases are thought to record progressively less saline conditions. The implication is that the marine bands do not reflect faunal concentration but represent the arrival of fauna that had been absent from the basin during intervening times. From the succession of faunal phases, the arrival of a fully marine fauna appears to be associated with pro­ gressively increasing salinity (Fig. 4)·. Because many of the faunal elements are nektonic or pelagic, it is not clear whether the salinity changes applied to the whole water column or whether they affected only the photic zone, where most of the fauna swam. Whatever the extent of desalination, it is apparent that the time of maximum salinity in the basin is most likely to coincide with or occur very slightly after the peak of the eustatic sea-level curve. If one considers the wider palaeogeographical context of the Namurian basins of the British Isles (see above), it is clear that they occupied an enclosed intra­ continental setting, removed by considerable dis­ tance and topographical complexity from the nearest ocean. Connections between the ocean and the British basins were therefore probably long and

O.J. Martinsen, J.D. Collinson and B . K. Holdsworth

310

Eustatic curve F-R

E1c minor cycle in Craven Basin (Ca. 800 m)

E2a(i)

E 1 c minor cycle in Staffordshire Basin (Ca. 25-90 m)

6

Delta top sands

Delta slope heterolithics

Ideal faunal/salinity cycle (based on Ashover cycles)

6 5

E2a(i) Barren shale

Turbidite sands and ironstones

E_j_

Barren s ale Calcareous siltstone 6

E1c

SIMPLE, MINOR CYCLE

Goniatite AID 4 Spat 3 Lingula 2 Planolites 1 Fish BARREN 1 Fish 2 Planolites 3 Lingula 4 Spat 5 AID 6 Goniatite 5

4

3

L

AEN

T Ideal

faunal/salinity oscillation

1

Fig. 4. Minor cycle stratigraphy showing variations of the basic cyclicity, from the ideal faunal/salinity cycle erected from Ashover borehole data (cf. Ramsbottom et al. , 1962; Holdsworth & Collinson, 1988), to thicker, sand-prone cycles from the Staffordshire Basin and the Craven Basin. The probable relation to the eustatic curve is also shown. (Modified from Holdsworth & Collinson, 1988.)

A, Anthracoceras; D, Dimorphoceras (thin-shelled goniati tes) .

tortuous. That, coupled with the fact that the British basins were supplied by very large volumes of fresh water, which delivered an abundant supply of clastic sediment, makes it easy to envisage how fluctuations in eustatic sea-level would have induced changes in the salinity of basinal waters. During lowstands, the cross-sectional areas of the connecting channels would be much reduced, especially if they were of low relief. Their main function would then be to accommodate the discharge of fresh water from the basins to the ocean. During highstands, expanded cross-sectional areas would have allowed oceanic waters to enter the basins and thereby increase salinity. Salinity changes are recorded from the Red Sea as a result of Pleistocene/Holocene changes in sea-level, although in this case lowstand led to hyper­ salinity because of the climate and Jack of river discharge (Thunell et al., 1988 ). The fact that sal­ inities appear to have oscillated throughout the Namurian within rather constant limits suggests that the straits which connected the basins with the ocean were rather stable and not subjected to major tectonic change or sedimentary choking. Partial analogues at the present day, such as the Skagerak­ Kattegat, the Hudson Straits and the Straits of

Hormuz are all relatively sediment free. Simple but conservative calculations based on likely river discharges (approximately 1 x 10 4m3 s-1, based on Kinderscout Grit channels; Collinson, 1969 ) and basin water volumes (around 1 x 10 13 m3, based on an assumed areal extent of the basins of 5 x 10 4km2 and a depth of 200 m) suggest that river water could have replaced basinal water in a matter of 30-35 yrs if the process was 100 % efficient. Given the average duration of a minor cycle was 180 000 yrs, desalination seems likely even at much lower efficiencies. In order for basinal waters to have become fully saline after a period of lowstand, and thereby allowing the deposition of a marine band, it is clear that some threshold of eustatic level must have been exceeded and that sea-level must have remained above the threshold for a period of time. The marine bands, therefore, on this basis, are thought to record periods of time at and around eustatic maxima. Given that there was probably some lag in the response of salinity to eustatic level, it is likely that the interval of marine band deposition may be skewed slightly later about the eustatic peak (Fig. 4). From their positions close to the bases of pro-

Millstone Grit cyclicity

gradational deltaic cyclothems it is likely that marine bands are associated with periods of relative sedi­ ment starvation and of rapid deepening. If marine bands are considered only in this context then it could be suggested that they record the time of maximum rate of rise in relative sea-level, i.e. around the R inflection point of the Exxon sequence stratigraphical model (Jervey, 1988; Posamentier et at., 1988). However, we believe this to be unlikely. If we are correct in believing that full marine salinity was necessary to support a fauna including thick-shelled goniatites, a view supported by the common occur­ rence of Radiolaria in this faunal phase and (as far as is known) no other, then, given the ocean-distant setting of the Millstone Grit basins, it seems likely that such salinities could have been achieved only close to the eustatic maxima. This would normally be expected to significantly post-date the R inflection point. It could even be inferred that, in some cases, parts of the curve between the R inflection point and the eustatic maximum may be represented by the inferred lower salinity phases of fauna that occur within marine bands below the thick-shelled goniatites. In other cases, there is compelling evi­ dence to place the response to the R inflection point clearly below the marine band. In the E1c- E2a cycles in Staffordshire (Trewin & Holdsworth, 1973), units of 'transitional shale', unfossiliferous and lacking carbonate, invariably separate a sideritic protoquartzitic turbidite unit from calcitic marine siltstones below and goniatite-bearing marine bands above (Fig. 4). Such shales can readily be interpreted as recording periods of rapid adjustment of environ­ ment when conditions favoured neither marine fauna or mineralogy nor influx of delta-derived clastics. The transitional shales below a marine band probably closely approximate the R infection point, whereas the overlying transitional shales mark the earliest stages of eustatic sea-level fall. In cyclic deltaic sequences, maximum conden­ sation, predicted as a response to the R inflection point, is most likely recorded by heavily bioturbated flooding surfaces or coal seams, which commonly terminate sandstone sequences. Such surfaces commonly are separated from the overlying marine band by a thin interval of mudstone. This mudstone is therefore thought to be the record between the R inflection point and the sea-level maximum of the marine band. Whilst these modifications regarding marine band

311

occurrence in relation to eustatic oscillation differ only slightly from those of conventional sequence stratigraphical theory, they illustrate ways in which the principles of sequence stratigraphy must be adapted in settings other than those in which they were formulated initially.

M ILLSTONE GRIT CYCLOTHEM ARCHITECTURE: RESPONSE TO RELATIVE SEA-LEVEL CHANGE

Introduction

Since Wright et al. (192 7) published the classic Rossendale Anticline memoir, the established view has been that Namurian cyclothems represent more or less continuous progradational episodes formed during falling or low relative sea-level, punctuated by marine transgressions, highstands of sea-level and deposition of goniatite-bearing marine bands (Fig. 5). In the following, two Millstone Grit cyclo­ thems are discussed in order to illustrate what has been interpreted as a different and more complex internal architecture. The two cyclothems were selected based on outcrop quality, and because they show different lateral facies and thickness develop­ ments. It is focused on the cyclothem development from the Askrigg Block and into the northern part of the Craven Basin. The depositional history of each cyclothem is reviewed briefly (for a full dis­ cussion of this aspect for the lower Namurian (Pendleian -Chokierian) cyclothems, see Martinsen, 1990, 1993). In the following, we take a less rigid view concern­ ing the use of the word 'sequence' than some existing definitions. In the American Geological Institute definition, for example (Bates & Jackson, 1987), 'sequence' is constrained to a depositional unit bounded by unconformities. However, as there is in existence more than one sequence stratigraphical model (Embry & Podruski, 1988; Van Wagoner et a/., 1988; Galloway, 1989; Cross & Gardner, 1991), and the models differ in what type of strati­ graphical surface is chosen as a boundary, we use the term 'sequence' for the stratigraphical unit that is practical to use in a given basin. The most practical sequence type is chosen based on an objective evaluation of the nature and extent of various strati­ graphical surfaces. The most easily defined, most extensive surfaces, based on detailed facies analysis,

0.1. Martins�n, J.D. Collinson and B . K. Holdsworth

312

,_ Fluvio-deltaic distributary

Goniatite band Barren mudstones

channel systems

�

Goniatite band

Slope channels feeding turbidites

Continuous progradation with successive clinoforms downlapping base-of-cycle goniatite band

Base-of-slope turbidites related to

Adjacent goniatite bands

gravitational instability or underflows

only separated by thin shale

during progradation

in basinal areas

Coarsening-up siltstone m

t#

Goniatite band

Fig. 5. The traditional Mill stone Grit cycle interpretation as a result of d elta progradation (turbidite-fronted delta example). See text for discussion.

are chosen as sequence boundaries. To avoid conflict with allostratigraphy, the same type of surface is picked throughout the succession studied. Goniatite-bearing marine bands are considered to represent the most significant stratigraphical levels in the Namurian of the Askrigg Block to Craven Basin succession, because of their lateral extent and distinctive faunal characteristics (Martinsen, 1993). They represent maximum transgression between successive progradational episodes and compare with the genetic stratigraphical sequence boundaries of Galloway (1989). Marine bands have a thickness and are not surfaces in a strict sense. However, when we describe them in the following, we see their mid-point as a 'surface', representing peak marine­ band conditions. Detailed facies analysis and regional understand­ ing of marine bands suggest that they represent the units across which the most significant palaeogeo­ graphical changes occurred in the Namurian suc­ cession. The marine bands are therefore chosen as sequence boundaries. Within a marine-band­ bounded sequence, however, the idea is to establish a detailed stratigraphical architecture, using both other stratigraphically significant surfaces (e.g. lowstand erosion surfaces or unconformities) and systems tracts.

The E2a(i) cyclothem Boundaries and thickness

The Cravenoceras cowlingense goniatite band forms the base of the E2a(i) cyclothem (Fig. 2), and can be found across most of the Askrigg Block, as well as in the western and easternmost parts of the Craven Basin (cf. Martinsen, 1990). The top of the cyclothem is marked by the base of the Eumorphoceras bisul­ catum ferrimontanum band, which forms the base of the overlying E2a(ii) cyclothem. This band is found throughout most of the Craven Basin, and on the northwestern side of the Askrigg Block. On the Askrigg Block, the cyclothem is approxi­ mately 70 m thick in the northwest, whereas it thins to around 40 m in the southeast (Fig. 6A-C). In the Craven Basin it is highly variable in thickness. In the northeast of the basin it is apparently less than 1 m thick due to condensation (further discussion below), whereas in the Lancaster Fells in the west, the E2a(i) cyclothem approaches 150 rn in thickness (Fig. 6A; cf. also Moseley, 1954, 1956). Depositional model and tectonism

Detailed studies (cf. Martinsen, 1993) suggest the following depositional model for the E2a(i) cyclo-

Millstone Grit cyclicity

them (Fig. 7). A phase of mudstone deposition occurred on the Askrigg Block. A zone of barren mudstones were succeeded by a phase of normal marine conditions represented by a fossiliferous interval with the Cravenoceras cowlingense fauna. A slight coarsening upward trend within the mud­ stones overlying the fossiliferous level suggests progradation of a clastic shoreline from the north. A drop of relative sea-level probably followed, and a rapid basinward shift of facies is interpreted to have taken place, causing establishment of the shoreline at the southwestern margin of the Askrigg Block. From this shoreline, sediment was supplied down into the Craven Basin as turbidity currents and suspended silts, leading to the establishment of a slope apron (Fig. 7). During the ensuing rise of relative sea-level, fluvial aggradation took place within the incised palaeo­ valleys (Upper Howgate Edge Grit) and eventually led to the deposition of a fluvial sheet sandstone (the Tan Hill Grit; Figs 6A & 7; Martinsen, 1993). Deposition corresponded with continued pro­ gradation and fill of the basin, with the development of a deltaic plain in the northwestern parts of the Craven Basin. The deltaic plain was subsequently flooded and transgression took place. In the initial phase of the transgression, shallow-water conditions persisted on the northwestern parts of the Askrigg Block, allowing for sporadic, high-frequency pro­ gradations of small delta lobes (Figs 6A & 7). Finally, the relative sea-level rise overwhelmed sediment supply, causing transgression and establish­ ment of offshore shelf conditions where mud­ stone sedimentation dominated. The conditions were shallow enough for occasional storms to rework benthic fauna into a shell bed, the Lad Gill Limestone. Fluvial channels were confined largely to the western parts of the block, leaving the eastern part relatively starved of sediment (Fig. 7). The cause for this was probably increased subsidence along the Dent Line, a long-lived lineament bounding the Askrigg Block in the west (cf. Underhill et al. , 1988). The Dent Line probably experienced minor reactivation during the early Namurian (Martinsen, 1993). The lack of sediment supply to the eastern parts of the block caused sediment starvation in the eastern parts of the Craven Basin (Figs 6C & 7), because most sediment was supplied to the basin across the western part of the Askrigg Block. This continued sediment supply to the western parts of the Craven Basin caused the slope apron or turbiditic

313

wedge to develop there (see above), because there was still a significant bathymetric contrast between the block and the basin. E2a(i) cyclothem architecture and sequence stratigraphy

The Cravenoceras cowlingense marine band itself is interpreted to represent the eustatic maximum or shortly thereafter (discussion above; Fig. 6A- C), whereas the thin, barren mudstones below the marine band probably represent maximum trans­ gression (maximum flooding). The fine-grained deposits above the marine band (eustatic maximum), and directly below the major fluvial complexes, but separated from these by a major erosion surface across which a major basin­ ward shift of facies occurs, are interpreted as late highstand deposits (Fig. 6A & C). The mudstones are interpreted to conform to the upper part of the falling limb on the eustatic curve, and on the relative sea-level curve prior to initiation of the fall, but at a time when the rate of rise had decreased to the point where clastic systems could prograde back into the area from the north. The erosion surface underneath the Upper Howgate Edge Grit fluvial complex (Fig. 6A) is interpreted as a lowstand erosion surface. Based on the interpretation that a prominent basinward shift of facies takes place across the surface, it is also interpreted as an Exxon-type sequence boundary. The surface is unlikely to correlate with the base of the turbiditic slope sediments in the Craven Basin. This is because incision probably evolved with time, and the final position of the incision surface most likely corresponds to some cryptic level within the upper parts of the turbidite slope succession (Fig. 6A). Therefore, the downlap surface below the turbidite deposits in the Craven Basin is more likely to record initiation of sea-level fall rather than cor­ relate directly with the final position of the lowstand erosion surface below the Upper Howgate Edge Grit on the Askrigg Block. The stratigraphical level corresponding to the ultimate lowstand of relative sea-level is considered as the boundary between a lower and upper part of the lowstand systems tract (Fig. 6A). Architecturally, the unit below this surface is a wedge in the Craven Basin, and the correlative of this lowstand wedge on the Askrigg Block is the lowstand erosion surface. When relative sea-level started to rise, and the incised valleys started to fill, turbidite sedimentation

314

O.J. Martinsen, J.D. Collinson and B . K. Holdsworth

NW Askrigg Block Lad Gill (SE 888046)

Sea-level curves (Lad Gill section)

Ca. 50km

-

3 -

j, 1 1

NW Craven Basin Kettlesbeck (SD 7 43632)

Wave-influenced shallow-water deltaic sediments Braided, fluvial channel complexes and overbank fines

-- -

IFS

Upper LST

---

E

"' c. �

Q)

-

�

g_

Lower LST

R

F Relative

c"O 0 Q) ·- "' � 0 "' c.

EX �

F

Q)

��

Eustatic

.,-o -"' co.o "0-

Distal part of highstand delta complex

E

:J .0 Q) 0 a:

Lower sequence boundary (Cravenoceras cowfingense marine band)

ms

A

Sea-level curves Relative

NW Craven Basin +----- 4okm Kettlesbeck (SD 743632)

-----•

NE Craven Basin Cononley Beck (SD 986468)

Eustatic River mouth bar

E

"' c.

Upper sequence boundary and superimposed initial flooding surface

�

Q)

� .2-

--o c Q) 0 "' 0 �c.

·-

co X E., 0�

u..-o Q) "' -.0 "' "

E

:J .0 Q) 0 a:

Upper LST --Lowstand, turbiditic slope wedge

--

Lower LST

Cravenoceras cowfingense

--

10

-

_

�

Edge Marine Band (E2a (ii))

0

m

m

fs

Lower sequence boundary (base E2a(i)) and conformable lowstand surface

marine band (E2a(i)) F ..

R ..

F ..

R ..

B

315

Millstone Grit cyclicity

Sea-level curves Relative F

R

,._____..

Eustatic F

SE Askrigg Block +----- 25km Backstean Gill (SE 058723)

.,.

____

NE Craven Basin Cononley Beck (SD 986468)

R

,._____..

40 Upper sequence boundary

Lower sequence boundary

c

Cravenoceras cowlingense marine band

All surfaces merge into condensed section in basin

(base E2a(i) cycle) on Askrigg Block

Fig. 6. (Opposite. ) (A) Stratigraphical correlation and proposed sea-level curves for the cross-section of the E2a ( i ) sequence from the northwestern part of the Askrigg Block into the northwestern part of the Craven Basin. (Opposite.) (B) Like (A), but cross-section from the northwestern part of the Craven Basin to the northeastern part of the basin. (Above. ) (C) Like (A) , but correlation from the southeastern part of the A skrigg Block into the northeastern part of the Craven Basin. For section locations, see Fig. 1.

MFP, Mirk Fell palaeosol ( Elc) ; UHEG, Upper Howgate Edge Grit; THG, Tan Hill Grit, LG Lst, Lad Gill Limestone; MFS, maximum fl ooding surface; LES, lowstand erosion surface; IFS, initial flooding surface; SB, sequence boundary. Systems tracts; HST, highstand; LST, lowstand; TST, transgressive.

may have gone on for some time. The time for cessation of turbidite sedimentation within the E2a(i) cycle may have been a result of elimination of the differential bathymetry (and accommodation space) at the southern margin of the Askrigg Block. How­ ever, during times of rising sea-level, rivers are more likely to aggrade and trap sediment, particularly in their downstream reaches (Schumm et al., 1987; Schumm, 1993), which in turn may reduce the poten­ tial for voluminous turbidity current generation. Several factors complicate the correlation of major stratigraphical surfaces from the basinal part of a succession upstream to a significant distance into fluvial systems. Firstly, there is likely to be a lag in the time the relative sea-level behaviour is recorded in upstream reaches. For instance, it is known from field data as well as flume experiments that upstream reaches may incise while downstream reaches aggrade (Shepherd, 1979; Schumm, 1993). This suggests that sequence boundaries commonly may have very complex histories and are not strictly correlative over large distances in upstream­ downstream profiles.

Secondly, the effects of a particular sea-level change may reach only a surprisingly short dis­ tance upstream from the coast. Saucier (1981; cf. review in Schumm, 1993) suggested that the effects of the approximately 120-m sea-level lowering during the late Wisconsin lowstand extended only 370 km upstream in the Mississippi river valley, in contrast to the 1530 km estimated by Fisk (1944). Also Blum (1994) showed similar ·effects for the Colorado River of Texas. Both points are important for the E2a(i) cyclo­ them. If there was a lag time between the initial sea­ level rise and the fluvial response by aggradation, the fluvial deposits on the northwestern part of the Askrigg Block may not be coeval, but perhaps be later than those in the basin or at the edge of the Askrigg Block, 60 km to the south. This is similar to back-filling during relative sea-level rise, where deposition starts at sea-level and with time progresses back up the alluvial tract. Such a relationship cannot be verified from the existing data set. An even more refined chronostratigraphy is needed to work out such details.

316

O.J. Martinsen, 1. D. Collinson and B. K. Holdsworth

A

B

Condensation in NE part of basin (lateral setting, no supply from block)

c

D

Bay deltas on SE block, condensation in NW part of basin

Distributary channels in basin, braided fluvial style on block leaving sheet sandstone (Tan Hill Grit)

� t N

I

Rapid coastal transgression, but minor, shallow-water deltas briefly interrupted retrogradation Salinity increase and faunal reintroduction

Start of prolonged period of slow deposition and condensation in basin and on the block

..,.

..a �

9

Fig. 7. Palaeogeographical summary diagram of the E2a(i) cyclothem related to different stages of sea-level. A, Eustatic maximum; B, lowstand; C, rising sea-level; D, rapidly rising sea-level.

It is fairly certain that the effects of the E2a(i) sea­ level fall extended as far upstream as the north­ western part of the Askrigg Block (provided the equivalent shoreline during the sea-level minimum lay at the southern end of the Askrigg Block), because the facies interpretation clearly indicates that a sea-level fall and basinward shift of facies occurred (Fig. 6A; Martinsen, 1993). On the southeastern part of the Askrigg Block, the abrupt transition from offshore mudstones to burrowed and rooted, thin crevasse splay deposits probably records the lowstand phase (Figs 6C & 7). Overbank deposition was probably limited during the early parts of the lowstand period when most sediment was transferred directly to the block-edge. During later aggradation, more sediment may have been supplied to floodplain or interchannel areas, but the proposed tectonic tilting of the Askrigg Block to the west would have prevented substantial volumes from reaching the southeastern parts of the block (Fig. 7). On the Askrigg Block, the incised Upper Howgate

Edge Grit lies directly on offshore mudstones. The sheet sandstone occurs above (the Tan Hill Grit), within the part of the cyclothem where relative sea­ level probably was rising steadily and significant: alluvial aggradation occurred (Fig. 6A). During very rapid rates of relative sea-level fall, the only possi­ bility for the fluvial system is to incise, particularly when the substrate is fine-grained sediments with pronounced cohesive stability and little possibility to compensate by channel-belt widening and increased bed roughness (Schumm, 1993). The resulting fluvial morphology of the Upper Howgate Edge Grit is therefore likely to be different from the braided, amalgamated model of Shanley & McCabe (1993). This may have been the case for the Upper Howgate Edge Grit of the E2a(i) cyclothem. Only when relative sea-level had risen further, perhaps to where the channels were no longer confined by an incised valley did the braided, sheet-like morphology of the Tan Hill Grit develop (Fig. 7). The thin deltaic coarsening upward succession on top of the Tan Hill Grit in the northwestern part of

Millstone Grit cyclicity

the Askrigg Block suggests that a flooding surface occurs between the two units (Fig. 6A). This is the first indication of inundation and transgression within the E2a(i) cyclothem. It also suggests that the boundary between a lowstand and transgressive systems tract should be put at this surface (see above). Moreover, the mudstones and goniatite­ bearing Lad Gill Limestone on top of the deltaic succession indicate further transgression and delta abandonment. These data support a transgressive systems tract interpretation of the 8-m deltaic suc­ cession, perhaps including the 4-5 m of mudstones below the Lad Gill Limestone. In the basin, there is no evidence for coarser grained, transgressive sediments. Instead, the thin deltaic succession is inferred to thin basinwards and merge with the mudstones on top of the distributary channel complex (Fig. 6A). In effect, the trans­ gressive systems tract in the basin is represented by a condensed section, which probably merges with the condensed shales at the base of the overlying E2a(ii) cyclothem (Fig. 6A). In the southeastern parts of the block and in the eastern parts of the Craven Basin, there is little evidence for a transgressive phase (Fig. 6C). Only, the intense burrowing on top of the crevasse sand­ stones indicate delta abandonment and trans­ gression. The burrowed level is sharply overlain by dark, offshore mudstones (Fig. 6C). The E2a(ii) marine band, the E. bisulcatum ferrimontanum band caps the E2a(i) cyclothem. Like its E2a(i) counterpart, the C. co wlingense band, it is interpreted to represent the peak of the eustatic curve. Dark mudstones underlie the marine band. It is considered likely that maximum transgression occurred at some point during deposition of these barren mudstones. The E2c2 cycl othem Boundaries and thickness

One of three upper Arnsbergian cyclothems with at the base (i.e. E2c2(i, ii or iu); cf. Fig. 2) is described as a counterpart to the E2a(i) cyclothem. There is uncertainty concerning the exact age of both the interpreted basal marine band and the cyclothem itself, because all the three nuculum bands are not developed everywhere. On the Askrigg Block, no N. nuculum bands are found , probably as a result of occlusion or erosion (see below). In the northeastern part of the Craven

Nuculoceras nuculum

317

Basin, only one or two bands are developed at any one location (e.g. Wilson, 1977), leaving a great deal of ambiguity as to which exact cyclothem the described rocks belong. For the present purpose, it is described as 'the E2c2 cyclothem'. Based on facies analysis (Martinsen, 1990), map­ ping (Stephens et al. , 1953), and local stratigraphical studies (Wilson, 1960, 1977), there is a high degree of confidence in correlating the described cyclothem between the southeastern part of the Askrigg Block and the northeastern parts of the Craven Basin, over a N-S distance of approximately 25 km. Therefore, the exact age and correlation of the described cyclo­ them with the minor cycle chart (Fig. 2), is of no importance for the present purpose. The E2c2 cyclothem varies considerably in thick­ ness. It is 5 -36 m in the southeastern corner of the block, around 34 m in the northeastern part of the Craven Basin, only 12 m in the north-central parts of the basin, and around 40 m thick in the northwestern parts of the basin (Fig. 8). On the Askrigg Block, the cyclothem base is interpreted to be the erosive base of the Lower Follifoot Grit, whereas in the basin, the base of the cyclothem is one of the three N. nuculum bands (Fig. 8). Martinsen (1993) interpreted the underlying E2cl cyclothem to be entirely eroded on the block by the Lower Follifoot Grit, which also erodes into the E2b3 cyclothem (Fig. 8). The cyclothem top is interpreted to coincide with the base of a multi­ cyclothem condensed interval, which encompasses at least nine minor cycles (E2c2(i or ii)- H1b(iii)), and which can be traced across the entire Craven­ Askrigg region (Fig. 8; ct. Martinsen, 1990). Depositional model and tectonism

In the southeast part of the Askrigg Block, the Lower Follifoot Grit is a trough cross-stratified sandstone, which has been interpreted to represent deposition in low-sinuosity, braided river channels (Martinsen, 1993; Figs 8 & 9). Geometrically, it has two modes of occurrence. For the most part, it is a tabular sheet sandstone recognizable over several tens of square kilometres (Fig. 8; Wilson, 1960). However, locally, two well-defined storeys of sand­ stone are observed. The lower one is very lenticular, forming ribbon-shaped bodies up to 15 m thick and hundreds of metres wide (Fig. 8). These sandstone bodies are clearly incised into the underlying depo­ sits. The upper storey is much more extensive lateraLly, forming a pronounced sheet. The Lower

318

0.1. Martinsen, J.D. Collinson and B . K. Holdsworth

SE Askrigg Block ---- 20-25krn ---+--• NE Craven Basin S

N Great Whernside West Scrafton Beldin G i l l S l i pstone Crag Thrope Edge SE 0074 SE 082824 SE 095799 SE 1 38821 SE 1 06752

Sea-level curves Relative

Eustatic

!Horn Crag Quarry, Silsden F R F R SE 053480 ---- ----

Scar House Beds (E2b3) 35m Mu lti-cycle condensed section (contains maxi m u m flooding s u rface and top seq u e n ce boundary of this cyclothem)

Lowstan d erosion su rface (coincides with sequence boundary)

Highstand systems tract

0

Fig, 8, Cross-section showing stratigraphical correlation and proposed sea-level curves of the studied E2c2 cyclothem from the southeastern part of the Askrigg Block to the northeastern part of the Craven Basin. See Fig. 1 for section location.

FSST, falling stage systems tract; LST, lowstand systems tract; MFS, maximum flooding surface; LES, lowstand erosion surface; FRS, forced regression surface.

Follifoot Grit is abruptly overlain by a multi­ cyclothem condensed section, which contains gonia­ tites of the lsohomoceras subglobosuml Homoceras beyrichianum marine bands of H1a- H1b age (Fig. 2; Wilson, 1960). In the basinal section above Silsden in Airedale (Fig. 8), to the south, offshore mudstones occur directly above the Nuculoceras nuculum marine band at the base of the cyclothem (Stephens et al. , 1953). These mudstones are only a few metres thick, and are then abruptly overlain by a 26-m-thick, entirely sandy succession (Fig. 8). This sandstone succession has been interpreted as friction-dominated and wave-influenced delta-mouth-bar deposits (Fig. 9; Martinsen, 1990). The mouth bar sandstone is growth-faulted to the depth of the base of the sand­ stone, suggesting that the fault soles out within the underlying mudstones. The mouth bar sands are overlain by an erosively based, very coarse-grained pebbly sandstone, which is entirely trough cross­ stratified (Fig. 8). The sandstone is interpreted as a braided fluvial complex, and is eventually overlain by fossiliferous mudstones of the overlying multi­ cyclothem condensed section (cf. Stephens et al. , 1953; Wilson, 1977).

E2c2 cyclothem architecture and sequence stratigraphy

The Nuculoceras nuculum goniatite band probably represents maximum eustatic sea-level (discussion above). Consequently, the dark mudstones that overlie the goniatite band in the basin are interpreted to represent late highstand deposits (Fig. 8), laid down when the nearest shoreline was still far to the north. In the basin , the abrupt contact between mud­ stones and the delta sandstones, and the close juxtaposition of the goniatite band and the delta sandstones (only separated by a few metres; Fig. 8), suggest that a relative drop of sea-level took place, represented by the lower boundary of the sandstone. One interpretation of the lower boundary is as a lowstand surface of erosion or Exxon-type sequence boundary formed during the most rapid rate of sea­ level fall. The delta sandstones above the lowstand surface would further be interpreted as lowstand sediments, suggesting that the entire delta is a lowstand delta. An alternative interpretation is that the sharp transition from the mudstone to the deltaic sandstone was produced by accretionary forced regression during relative fall of sea-level, making the delta

Millstone Grit cyclicity

A

319

B

Falling-stage (forced regression) delta with growth faults c N

N

r

coastal transgression

Start of prolonged period of slow deposition and condensation in basin and on the block

Salinity increase and faunal reintroduction

Fig. 9. Palaeogeographical summary diagram of the studied E2c2 cyclothem related to different stages of sea-level .

A, E ustatic maximum; B, falling-stage; C, lowest/ slowly rising sea-level; D , rapidly rising sea-level.

sandstones part of a falling stage systems tract (Fig. 8; e.g. Plint, 1988; Hunt & Tucker, 1992; Posamentier et al. , 1992). This interpretation is supported by the fact that the braided, fluvial sand­ stone overlying the deltaic rocks is so much coarser than the deltaic sandstones, shows entirely different palaeocurrents and therefore does not seem to be related to the deltaic sandstones (i.e. it is not their distributary channel). The erosion surface below the fluvial sandstones could represent the relative sea­ level lowstand and thus equate with an Exxon-type sequence boundary (Fig. 8). On the Askrigg Block, the erosion surface below the lenticular channels of the Lower Follifoot Grit probably marks an extensive erosion surface related to falling and low relative sea-level (Fig. 8). Martinsen (1993) suggested that extensive incision took place on the block at this time, and the under­ lying E2c l cyclothem was largely eroded. The incision phase may correlate with the deposition of

the inferred falling-stage delta in the basin (Fig. 8), but perhaps more likely, it represents the incision related to relative lowstand and correlates with the erosion surface between the deltaic sandstones and the very coarse fluvial sandstones in the basin. This interpretation implies that erosion below the Lower Follifoot Grit on the Askrigg Block involved both initial formation during the falling stage and further development and modification during the relative lowstand. If this interpretation is correct, the fluvial sand­ stones on the block and in the basin were deposited during a relative rise of sea-level (Fig. 8). Questions as to whether the shoreline prograded during filling of the incised fluvial channels, or if enough sediments were trapped on the alluvial plain to enhance trans­ gression at the contemporaneous shoreline cannot be answered due to inadequate exposure. In the latter case, the fluvial interval should be termed part of a transgressive systems tract. However, the lack

320

O.J. Martinsen, J. D. Collinson and B. K. Holdsworth

of evidence for transgressive deposits, and analogy with the E2a(i) cyclothem suggests that it is more likely that progradation continued up until or very close to cyclothem abandonment. In that case, the fluvial rocks should be termed part of the lowstand systems tract (Fig. 8). It is likely that sea-level fall across a slowly sub­ siding area like the Askrigg Block would enhance incision and stabilize single channels or 'valleys', particularly if incision was rapid and occurred in a fine-grained substratum (see Schumm, 1993). Only during sea-level rise and consequent increasing accommodation would the incised channels be 'overflooded' and fluvial deposition take place out­ side the incised channels. If grain size, discharge and gradient were favourable for a braided fluvial style to develop at this time, the resulting fluvial sediments would probably be sheet-like. Provided the fluvial rocks of the E2c2 cyclothem form part of a lowstand systems tract, no trans­ gressive systems tract seems to have developed above, because there is an abrupt transition to marine mudstones with goniatites, both in the Craven Basin and on the Askrigg Block (cf. Wilson , 1960, 1977). This transition suggests a rapid rise of relative sea-level, flooding and cyclothem abandon­ ment with no time for development of a 'late' progradational interval, like that of the E2a(i) cyclo­ them, at least in the study area. Therefore, in summary, within the E2c2 cyclo­ them, falling-stage sediments are interpreted to predominate in the basin , whereas on the block, lowstand systems tract sediments dominate above a complex lower erosional surface. Highstand sediments were probably eroded on the block, whereas in the basin they are restricted to a thin interval immediately overlying the Nuculoceras nuculum marine band (Fig. 8). Transgressive sys­ tems tract sediments are condensed, and probably occur as a thin zone between the fluvial sediments and the overlying multi-cyclothem, goniatite-bearing condensed section (Fig. 8).

post-extensional subsidence but also of decreasing overall subsidence (cf. Leeder & McMahon , 1988). One similarity between the two cyclothems is the proposed prolonged absence of clastic supply from the time of maximum transgression (thin shales below marine bands), until maximum eustatic highstand (marine bands; see Figs 6 & 8). This encompasses one-quarter of a eustatic curve, and, depending on the rate of subsidence, probably corresponds to the middle part of the rising, extended limb on a relative sea-level curve. Therefore, it seems that clastic supply was absent for a prolonged period compared with conventional sequence strati­ graphical models (cf. Van Wagoner et al. , 1990). The decreasing overall subsidence is important for explaining why late highstand deposits are not pre­ served on the Askrigg Block in the E2c2 cyclothem in contrast to the E2a(i) cyclothem within which highstand deposits are relatively thick. A lower rate of subsidence at this time probably enhanced and amplified the fall of relative sea-level on the Askrigg Block causing deeper incision and erosion. In addition, a lower subsidence rate during the relative highstand generated less new space for sediment accumulation, which gave less accommodation and a thinner highstand succession to start with. Another contrast is the interpreted development of a lowstand, turbiditic wedge during E2a(i) time whereas a falling stage, forced regression delta evolved during E2c2 time. These depositional patterns follow closely the shelf-edge and ramp models of van Wagoner et al. (1988). However, in addition , in the higher subsidence situation of the E2a(i), the response time of the sedimentary system to falling relative sea-level was longer. This longer response time may have slowed regression during E2a(i) time. The opposite was probably the case for the E2c2 delta. A similar conclusion was reached by Holdsworth & Collinson (1988, p. 138), where the response time of sand deposition to lowered sea­ level was related to bathymetry, the latter being a combined effect of compaction of previously deposited sediments, eustasy and subsidence.

Sequence model Choice of sequence stratigraphical model Introduction

The E2a(i) and E2c2 cyclothems illustrate different stratigraphical architectures as responses to con­ tJ·asting physiographic relationships between the Askrigg Block and the Craven Basin. The contrast evolved as a result not only of decreasing differential,

Local sequence stratigraphical models should fulfil the same requirements as facies models (Reading, 1986; Walker, 1992). Thus, a sequence strati-· graphical model for the E2a(i) and E2c2 cyclothems will have to serve several purposes: (i) act as a norm so that the model can be compared with sequence

Millstone Grit cyclicity

models from other areas; (ii) the model should be rigorous enough to be a framework for future studies; (iii) it should aid prediction of sequence stratigraphical architecture in adjacent areas; and (iv) it should be a basis for interpretation of Namurian sequence stratigraphy. The models described here are considered to fulfil these requirements, and they are modified versions of the general models of Galloway (1989) and Exxon (e.g. Jervey, 1988). Several local factors prevent the general models from being applied as templates in the present situation, and modified versions are required. First, erosion surfaces at the bases of the incised fluvial complexes are laterally discontinuous and cannot be correlated with confidence between individual outcrops (e.g. through palaeosols), nor in a proximal- distal direction (cf. Martinsen, 1993). Secondly, goniatite-bearing marine bands are the most laterally continuous stratigraphical horizons (see above), and are interpreted to be the levels across which the most important palaeogeographical reorganizations and stratigraphical breaks occur (cf. Figs 7 & 9). The erosion surfaces are recognized as lowstand surfaces of erosion and compare with Exxon-type sequence boundaries, but they are not considered as the most prominent breaks in the succession, nor are they laterally continuous enough to be reliable for correlation (Martinsen, 1993). For example, the lowstand erosion surfaces are not re­ cognizable in the Craven Basin (Figs 6A-C & 8). Thus, the goniatite bands are the most sensible stratigraphical levels to pick as sequence boundaries. Thirdly, since lowstand wedges of sediment occur both in the E2a(i) and E2c2 cyclothem, the Galloway (1989) model cannot be applied in a straightforward way because such units are not emphasized in that model. However, Xue & Galloway (1993) described low­ stand prograding complexes from both shelf-edge margins and ramp margins in a study from the Songliao Basin in China, thus extending the original Galloway (1989) model. The Xue & Galloway (1993) model is therefore a reasonable comparison for the models presented herein for the E2a(i) and E2c2 cyclothems, with the added complexity of a forced regressive unit/falling stage systems tract for the E2c2 cyclothem. Cyclothem asymmetry, glacio-eustasy and spatial development of systems tracts

In the E2a(i) and the E2c2 cyclothems, as well as in

321

most other Millstone Grit cyclothems deposited near the basin margin (Wright et al. , 1927; Stephens et al. , 1953; Wilson, 1960; Collinson, 1988; Holdsworth & Collinson, 1988; Martinsen, 1993), there is a pronounced lithological asymmetry (see Figs 6 & 8; cf. also modelling by Collier et al. , 1990). The asymmetry developed in the thicker Millstone Grit cyclothems contrasts with the symmetry of the basinal cycles at Ashover (Rams bottom et al. , 1962; Holdsworth & Collinson, 1988; Fig. 4). The Ashover faunal cycles were probably controlled by salinity (see above; Holdsworth & Collinson, 1988). Although salinity probably fluctuated symmetrically with time (Fig. 4; discussion above), the supply of coarser sediments did not. Consequently, rather than salinity, sand deposition must have been related to other mechanisms, such as supply variations and subsidence. Therefore, relative sea-level rather than eustasy controlled the succession in the Askrigg Block and Craven Basin area. Because of the added influence from subsidence, relative sea-level curves are characterized by longer periods of rise (e.g. Fig. 6A). The sedimentary expression and character of the succession at Ashover would, therefore, be expected to be different from the Askrigg Block to Craven Basin succession, even if sediment input was similar. Given the inherent nature of a glacial world, with major glaciations in Gondwana (Caputo & Crowell, 1985 ; Veevers & Powell, 1987), glacigenic sea-level falls tend to be much slower than corresponding sea­ level rises due to slow ice-cap build-ups and rapid melt-downs (e.g. Williams, 1988). As pointed out by Read (1991), although periods of eustatic sea-level fall were relatively short, the rate at which sea-level fell must have been high enough to overcome sub­ sidence rate and produce a relative sea-level fall. During these periods, sandy systems filled the basins. The rate of relative sea-level rise must have been extremely rapid, particularly in basinal areas, where it was amplified by high subsidence rates. Consequently, in low-lying coastal plains, which probably formed an important part of the Namurian systems (cf. Holdsworth & Colbnson, 1988), the rate of coastal transgression must have been high, with little possibility for the sedimentary systems to leave significant sandy deposits behind. Also, the generally observed lack of significant wave or tidal action decreased the possibility for significant reworking of the previously deposited sediments. These relationships suggest that the asymmetric lithological response to relative sea-level change was

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0 .J. Martinsen, J. D. Collinson and B. K. Holdsworth

inversely related to the time needed to deposit the facies. The thick, coarsening upward successions in the basins were probably deposited in a relatively short time, whereas the thin condensed mudstones and marine bands record most of the time rep­ resented by each Millstone Grit cyclothem. This asymmetry contributed to the areal distri­ bution of systems tracts for both the 'block-edge' and ramp situation. The rapid transgression resulted in a very thin transgressive systems tract, with, at the most, only a thin shallow-water delta on top as a last sedimentary response to the accelerating sea-level rise. The transgressive systems tract developed a recognizable thickness only on the Askrigg Block. The highstand systems tract of the two cyclothems, which is mudstone-dominated in the study area, is probably better developed to the north (discuss­ ion above), closer to the palaeoshoreline. The ex­ pectedly slower rate of glacio-eustatic sea-level fall resulted in a slower rate of relative sea-level fall, and a well-developed lowstand systems tract in the Craven Basin. However, even in an ideal situation, with no along-strike displacement and variability of systems tracts, no one vertical section is expected to show a complete development of a marine-band­ bounded sequence. The systems tracts are displaced either down-dip or up-dip with relation to each other (Figs 6A & 8), and their complete presence in any one vertical section is unlikely. With the added complexity of a strike-variable development of systems tracts (see Fig. 6A -C; compare cross­ sections), a common situation along any shoreline (Martinsen, 1994), it is self-evident that only very few two-dimensional cross-sections would show a complete development of systems tracts.

MULTI-CYCLOTHEM CONDENSED SECTIONS AND MAJOR SCALE CYCLICITY

Background

Although the basic cyclothem pattern ('minor cycles' of Holdsworth & Collinson, 1988) of the Millstone Grit has been established since Wright et a!. 's (1927) classic Rossendale Anticline Memoir, the issue of a cyclicity at a larger scale has been more actively debated recently (cf. also Read (1991) for a dis­ cussion of cyclicity at several orders in the Namurian of the Midland Valley of Scotland). In 1977, Ramsbottom argued that Millstone Grit cyclothems could be grouped into larger scale cycles, termed

'mesothems'. The mesothem theory has been subject to severe criticism and its validity questioned (see Holdsworth & Collinson (1988) and Martinsen (1990) for a thorough discussion). It is proposed that major cyclicity exists, but in a different form than mesothems. Nam urian maj or-scale cyclicity

Within the Namurian E1c- H2c succession of the Craven-Askrigg region, three regionally correlative stratigraphical intervals can be defined. In each are several minor cycles, which are condensed, represent non-deposition and/or are entirely fine-grained (Fig. 10). Interval 1 extends from cyclothem E2b1(i) to cycle E2b2(ii) and contains five minor cycles, interval 2 extends from cyclothem E2c2(ii) to cyclothem H1b(iii) and contains nine minor cycles, and interval 3 extends from cyclothem H2a(i) into the Kinderscoutian (R1) and contains five minor cycles to the base of the Kinderscoutian. Common to these three, multi-cyclothem con­ densed intervals, is that they are sand-free and only small thicknesses of fine-grained sediments were deposited. Only locally within one section can evi­ dence for subaerial exposure be documented, and erosion has not been proven. Therefore, these fine­ grained intervals are mainly of subaqueous origin. There are ambiguities to some of the correlations, particularly on the Askrigg Block where few goniatite bands occur. Nevertheless, the general framework seems to hold, not least because the intervening cycles are dominated by sandstones, which them­ selves can be correlated regionally both because of adjacent goniatite faunas, classic mapping and their distinctive depositional environments (cf. Martinsen, 1990). Based on these multi-cyclothem condensed mud­ stone intervals, it is suggested that three major cycles of deposition can be defined within the E1c­ H2c succession (Fig. 10). Major cyclothem 1 (MC1) comprises the E1c- E2a(iii) minor cycles, bounded below by the sub-E1c unconformity on the Askrigg Block (ct. Dunham & Wilson, 1985) and the base of the Cravenoceras malhamense goniatite band in the basin, and above by the base of the E2b1(i) Cravenoceras subplicatum group marine band. This band forms the base of the lowest multi-cyclothem mudstone interval. All the cycles in MC1 contain sandstones. Major cyclothem 2 (MC2) encompasses minor cycles E2b1(i)- E2c2, and is bounded below by the base of the Cravenoceras subplicatum group

Millstone Grit cyclicity

323 H l bliv)

SE Askrigg Block .----- 1 5-20km --------. NE Craven Basin

Base E2c2(i?)­ unconformity MC2

Base E 1 c­ unconformity

\

E1c

Craven Fault System

� t

\ 1 50

\

MC1

om

• Multi-cycle mudstones III:t!Il Dominantly sandstones D Other mudstones and siltstones Minor cycles

Major cycles

Fig. 10. Proposed division of the Namurian Elc-H2b stratigraphy into major cyclothems. See text for discussion.

marine band, and above by the base of the second Nuculoceras nuculum marine band at the base of the E2c2(ii) cyclothem, which also forms the base of the second multi-cyclothem condensed interval. Whereas the lower five cycles in MC2 are sandstone­ free, the upper three contain sandstones (see below). Major cyclothem 3 (MC3) contains ten minor cycles comprising E2c2(ii) to H1b(iv). The lower boundary is the base of the second Nuculoceras nuculum marine band at the base of the E2c2(ii) cyclothem, whereas the top boundary is defined at the base of the Homoceras smithi/ Hudsonoceras proteus goniatite band at the base of the third multi­ cyclothem condensed interval. Only the uppermost cyclothem, H1b(iv), contains sandstones, whereas the underlying nine make up the second multi­ cyclothem condensed interval.

Maj or c ycl othem architecture and genesis

In only one major cyclothem (MC2) is there indi­ cation of a systematic organization of the contained minor cycles. The E2b3, E2c1 and the one sandy E2c2 cyclothem show an increasingly progradational character (cf. Martinsen, 1993). In addition, Martinsen (1993) argued that the lowstand erosion surface below the E2c2 Lower Follifoot Grit on the Askrigg Block was, in addition to the unconformity below the E1c Grassington Grit, the most prominent within the E1- H2 part of the Millstone Grit in this area. Therefore, it is probable that the sequential evolution of the MC2 cycles reflects progressively lower sea-level with time. Such a pattern is compar­ able to a progradational sequence-set architecture, as described by Mitchum & Van Wagoner (1990), and may reflect longer term sea-level changes than

324

0.1. Martinsen, J.D. Collinson and B . K. Holdsworth

those at minor cycle scale. However, none of the other proposed major cycles show a systematic stacking pattern. Although there is no direct reason to argue against sea-level changes as the main control on the major cycles, in most of the multi-cyclothem condensed sections, the basic cyclicity is still present, with faunal bands being interbedded with barren mud­ stones (e.g. Wilson & Thompson, 1959; Arthurton et al. , 1988). This suggests that lowstands were still reducing salinities below the critical marine-band threshold. In addition , there is evidence for subaerial exposure in the E2b1(i)- E2b2(ii) condensed section in the Colsterdale Marine Beds on the eastern side of the Askrigg Block (Wilson & Thompson, 1959). Therefore, it appears that there are no major changes to the sedimentation and cyclicity patterns beside the decreased thickness and the lack of sand. Possibly, one would expect a change of the cyclicity patterns, particularly in the condensed sections (the assumption then being that these would represent a particularly high sea-level), if major-scale sea-level changes caused the major cyclicity. Major changes of sediment supply are more likely to have caused the major cyclicity. Within the E2c2(ii)- H1b(iii) cyclothems, which form a con­ densed section, particularly in the eastern part of the Craven-Askrigg area, the mudstones thicken dramatically towards the west, include turbidite sandstones and reach 165 m in thickness (Wilson et al. , 1989). This relationship suggests that a major switch of sediment supply to the west occurred within the E2c2(ii)- H1b(iii) time interval, and probably explains the condensation in the Craven -Askrigg area. It is suggested also that the other multi­ cyclothem condensed sections were formed by large­ scale switches of sediment supply. To the east, over the southern North Sea, large Namurian depocentres are known to have been present. However, the stratigraphical subdivision in this area is yet not refined enough to test the hypothesis that those areas were alternative sites of sediment accumu­ lation . Therefore, a sequence stratigraphical scheme at major cyclothem scale is not considered to be possible at present.

CONCLUSIONS 1 To illustrate the lateral and temporal develop­ ment of Namurian Millstone Grit cyclothems , the Arnsbergian E2a(i) and E2c2 cyclothems in the

Askrigg Block and Craven Basin area have been re-evaluated in terms of architectural and strati­ graphical response to relative sea-level change. 2 Goniatite-bearing marine bands are thought to represent the highest eustatic sea-level evidenced by salinity data. This interpretation contrasts with the conventional sequence stratigraphical models where faunal concentrated bands are considered to represent maximum transgression and flooding. Instead, thin , condensed shale below the marine bands probably represent maximum rate of rise of sea-level. 3 The E2a(i) cyclothem developed when there was a significant bathymetric contrast between the Askrigg Block and the Craven Basin. Prolonged condensation took place, and clastic deposition did not occur until the late part of the highstand of sea­ level. Subsequent lowstand sediments are developed both on the Askrigg Block as fluvial systems and in the basin as a turbiditic slope apron. The trans­ gressive systems tract is very thin due to rapid trans­ gression and condensation in the basin. 4 The E2c2 cyclothem developed at a time when there was a gently inclined ramp basinward of the Askrigg Block. Prolonged condensation and/or non­ deposition took place until sea-level started to fall and a forced regression delta formed. This delta was incised during the subsequent sea-level lowstand , and fluvial deposition took place o n the Askrigg Block. No underlying highstand or overlying trans­ gressive systems tract sediments are preserved or were deposited. 5 Major cycles of deposition are interpreted to exist in the Namurian of the Askrigg Block and Craven Basin area. The cycles are bounded by condensed, multi-cyclothem mudstone levels. Except in one case, the major cycles show no systematic architec­ ture, suggesting they were not related to major changes of relative sea-level. Instead, they prob­ ably represent major avulsive shifts of sediment depocentres.

ACKNOWLEDGEMENTS

We wish to thank numerous colleagues who have inspired and guided our thinking about the Millstone Grit both in the UK and in Ireland, particularly Harold Reading, Trevor Elliott, Roger Walker and Colin Jones. Ole Martinsen wishes to thank Total Norway for sponsoring 3 years of doctoral research, Norsk Hydro for sponsorship of 3 years of post-

Millstone Grit cyclicity doctoral

research ,

and

Frank

Heggernes

and

Karianne Krag Kjos for providing new ideas about the northwest quarter of the Askrigg Block. The referees Charles Bristow, Robert

L.

Gawthorpe and

Peter J. McCabe are thanked for reviewing the manuscript.

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Millstone Grit cyclicity

R.G. ( 1979) River channel responses to bedrock lithology and stream capture, Sandy Creek Drainage, central Texas. In: Adjustments of the Fluvial System (Eds Rhodes, D . D . & Williams, G . P . ) , pp. 255-276. Kendall/Hunt, Dubuque, Ia. STEPHENS, J.V., MITCHELL, G . H . & EDWARDS, W. ( 1953) Geology of the country between Bradford and Skipton. Mem. geol. Surv. G. B. , Eng. Wales, Explan. Sheet, 69, 180 pp. THUNELL, R . C . , LOCKE, S . M . & WILLIAMS, D . F. ( 1 988) Glacio-eustatic sea-level control on Red Sea salinity. Nature, 334, 601 -604. TREWIN , N.H. & HoLDSWORTH, B . K . ( 1 973) Sedimentation in the lower Namurian rocks of the North Staffordshire Basin. Proc. Yorks. geol. Soc. , 39, 3 7 1 - 408. UNDERHILL, J . R . , GAYER, R . A . , WOODCOCK, N . H . , DoNNELLY, R . , JOLLEY , E . J . & STIMPSON, I . G . ( 1 988) The Dent Fault System, northern England reinterpreted as a major oblique-slip fault zone. J. geol. Soc. London, 145, 303 - 3 1 6 . VAN WAGONER, J.C. , POSAMENTIER, H.W., MITCHU M , R.M. JR. , VAIL, P.R., SARG , J.F., LouTIT, T.S., et al. ( 1 988) An overview of the fundamentals of sequence stra­ tigraphy and key definitions. In: Sea-Level Changes: an Integrated Approach (Eds Wilgus, C . K . , Hastings, B . S . , Kendall, C.G .StC . , Posamentier, H . W . , Ross, C . A . & Van Wagoner, J.C.), Spec. Pubis Soc. econ. Paleontol. Mineral . , Tulsa, 42, 39-45. VAN WAGONER, J . C . , MITCHU M , R.M. JR . , CAMPION , K . M . & RAHMANIAN , V . D . ( 1 990) Siliciclastic sequence stra­ tigraphy in well-logs, cores and outcrops. Am. Assoc. petrol. Geol. Meth. Explor. Ser. , 7, 55 pp. VEEVERS, J.J. & POWELL, C.McA. ( 1 987) Late Paleozoic glaciations in Gondwanaland reflected in transgressive­ regressive depositional sequences in Euramerica. Geol. Soc. Am. Bull. , 98, 475-487. WAGNER, R . H . , HIGGINS, A.C. & MEYEN, S . V . (eds) (1979) The Carboniferous of the USSR. Occas. Pubis. Yorks. geol. Soc . , 4 , 247. WALKER, R . G . ( 1966) Shale Grit and Grindslow Shales: transition from turbidites to shallow-water sediments in the Upper Carboniferous of northern England. SHEPHERD,

J. sediment. Petrol. , 36,

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R.G. ( 1990) Facies modeling and sequence stra­ tigraphy. J. sediment. Petrol. , 60, 777-786. WALKER, R . G . ( 1 992) Facies, facies models and modern stratigraphic concepts. In: Facies Models - Response to Sea-Level Change (Eds Walker, R . G . & James, N . P . ) , p p . 1 - 14 . Geological Association o f Canada, St John's. WILLIAMS, D .F. ( 1 988) Evidence for and against sea-level changes from the stable isotopic record of the Cenozoic. In: Sea-Level Changes: an Integrated Approach (Eds Wilgus, C.K . , Hastings, B . S . , Kendall , C.G .StC . , Posamentier, H.W. , Ross, C . A . & Van Wagoner, J . C . ) , Spec. Pubis Soc. econ. Paleontol. Mineral. , Tulsa, 42, WALKER,

31 -36.

A.A. ( 1960) The Millstone Grit Series of Colsterdale and neighbourhood, Yorkshire. Proc. Yorks. geol. Soc. , 32, 429-452. WILSON , A.A. ( 1977) The Namurian rocks of the Fewston area. Trans. Leeds geol. Assoc. , 9, 1 -42. WILSON , A . A . & THOMPSON, A.T. ( 1 959) Marine bands of Arnsbergian age (Namurian) in the south-eastern portion of the Askrigg Block, Yorkshire. Proc. Yorks. geol. Soc. , 32, 45-67. WILSON, A . A . , BRANDON , A . & JOHNSON, E . W . ( 1 989) Geological Context of the Wyresdale Tunnel Methane Explosion . Geological Survey of Great Britain, Technical Report, WA/89/30, HMSO, London. WRIGHT, W . B . , SHERLOCK, R.L. , WRAY , D . A . , LLOYD, w. & ToNKS, L.H. (1927) The geology of the Rossendale Anticline. Mem. geol. Surv. G. B . , Eng. Wales, 182 pp. XuE, L. & GALLOWAY , W.E. (1993) Genetic sequence stratigraphic framework , depositional style, and hydro­ carbon occurrence of the Upper Cretaceous QYN formations in the Songliao lacustrine basin, northeastern China. Bull. Am. Assoc. petrol. Geol. , 76 , 1792- 1808. ZIEGLER, P . A . ( 1 986) Geodynamic model for Palaeozoic crustal consolidation of western and central Europe. Tectonophysics, 126, 303-328. ZIEGLER, P .A. ( 1 988) Evolution of the Arctic-North Atlantic and the Western Tethys. Mem. Am. Assoc. petrol. Geol. , 43, 196 pp. WILSON ,

Facies Analysis in Reservoir Sedimentology

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

Spec. Pubis int. Ass. Sediment. (1995) 22, 331-354

Productive Middle East clastic oil and gas reservoirs: their depositional settings and origins of their hydrocarbons Z I A D R. B E Y D O U N American University of Beirut, Lebanon

ABSTRACT The Middle East region contains 67% of the world's reported recoverable oil reserves and 30% of its reported recoverable gas reserves. The greater bulk of these are housed in carbonate reservoirs of Late Permian-Miocene age, with sandstone reservoirs of that age range being minor and restricted to specific areas of the region. The predominance of carbonates since the latest Permian is related to the plate tectonic position of the region, located principally in low latitudes and warm climatic belts. By contrast, from the Infra-Cambrian to the Early Permian, the Middle East region was located principally in middle to high southern latitudes in cool-wet to cold-dry climatic belts, when clastic sedimentation predominated and carbonates and evaporites were deposited mainly during brief excursions into lower latitudes and/or during short episodes of aridity in higher latitudes. Only fairly recently has the petroleum potential of the Palaeozoic clastic reservoirs begun to be properly appreciated after discover­ ies in Oman demonstrated ample sourcing from extensive organic-rich Infra-Cambrian to Early Cambrian sediments, with both vertical and lateral migration into Cambro-Ordovician and Permo­ Carboniferous sandstone reservoirs. More recently still, important discoveries in Central Arabia have confirmed sourcing of Permo-Carboniferous sandstone reservoirs from regionally deposited organic-rich Lower Silurian shales related to transgressive global sea-level rise. Either of these sources, where overmature, have most probably also provided most of the region's vast Permian Khuff carbonate reservoir gas reserves. The depositional environments and relevant petrophysical properties of the productive sandstone reservoirs of the region are reviewed by stratigraphical position and age and the source rocks charging these are briefly discussed. It is concluded that tremendous petroleum potential exists in the Middle East for substantial new Palaeozoic sourced and reservoired hydrocarbons. However, unlike the extensive lateral facies continuity of the post-Palaeozoic source-reservoir-seal relationships and the giant traps of the very wide Arabian shelf region (resulting in almost predictable accumulation trends), the Palaeozoic clastic reservoirs and seals generally exhibit rapid lateral and vertical facies changes coupled with the generally smaller sizes of local fault-block traps and/or low-amplitude fold traps. Consequently, the distribution of hydrocarbon accumulations in Palaeozoic sequences is less easily predicted, and in some areas is critically dependent upon seal distributions. Provided source and 'kitchen' distributions can be better understood and migration paths to traps better constrained, major additional reserves in Palaeozoic sandstone reservoirs will be established. Additional discoveries in post-Palaeozoic sandstone reservoirs are likely to be mainly restricted to the identification of hidden rift plays similar to those of Yemen and the Euphrates area of Syria.

I N T RO D U C T IO N west Iran-southeast Turkey, lying, respectively, southwest and south of the Zagros and Bitlis (Taurus) sutures, together with the Levant states and including the Sinai Peninsula but not Egypt proper (Fig. 1).

The geological boundaries of the Middle East region, as discussed in this survey, are those of the Arabian Plate and the adjacent Levant-Sinai fragment to its northwest (Fig. 1). Geographically, this encompasses the countries of the Arabian Peninsula and south-

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

331

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PreCambrian

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Approximate extent of currently hydrocarbon-productive basins with predominance of sandstone reservoirs and their ages P.M. C. (see legend for key)

N.B. Small lener (e.g. pl denotes only minor production or potential.

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Fig. I. Geological sketch map of the Middle East.

Middle East clastic oil and gas reservoirs

The Middle East region, as outlined above (includ­ ing the Gulf of Suez rift basin), at the start of 1992 contained a reported 665.6 billion (B) barrels of recoverable oil and 1325.8 trillion (T) cubic feet (tcf) of recoverable gas (88.51 B tonnes and 3 37.54 T m respectively) (Oil and Gas Journal, 30 December 1991). This is equivalent to 67.2% of the World's reported recoverable oil reserves and 30.3% of the reported recoverable gas reserves. In 1991, the region accounted for 28.5% of the World's annual oil production (Oil and Gas Journal. , 30 December 1991); this figure, however, is determined more by economic and engineering constraints than by limitations on reservoir potential and productive capacity. Oil initially was discovered in the region in 1908, at Masjid-i-Suleiman in southwest Iran in the Oligo-Miocene carbonate reservoir of the Asmari Formation, and that field and others went on pro­ duction shortly afterwards. Commercial oil was found subsequently in Iraq, at Kirkuk in 1927 in a similar carbonate reservoir. Oil initially was found in Egypt at Gemsa field, located onshore on the western side of the Gulf of Suez in 1909, again in a carbonate reservoir of Miocene age. This was fol­ lowed by the discovery of the first clastic production in the Gulf of Suez at Hurghada field near to Gemsa in 1915, from Cretaceous 'Nubia' sandstone reser­ voirs as well as from a Miocene carbonate reservoir. On the eastern side of the Arabian Peninsula, the next discovery was made in Bahrain in 1932 in the Cretaceous sandstone and carbonate reservoirs of the Awali field, and in 1938 in eastern Saudi Arabia at Dammam in Upper Jurassic carbonate reservoirs and also in Kuwait at Burgan in Cretaceous sand­ stone reservoirs. World War II interrupted further exploration activity, but shortly after it was over a remarkably rapidly expanding exploration pro­ gramme commenced in many different parts of the region and continues to the present; discoveries have been made in every single country of the region except Lebanon, where the protracted phase of civil strife commencing in the early -mid- 1970s has cur­ tailed all exploration activities. The last country to join the producers club was Yemen in 1984. The magnitude of the many discoveries and devel­ opments varies considerably, from the 'giant' (0.5 -4.5 B barrels estimated ultimate recoverable (EUR) reserves of oil or oil equivalent) and 'super­ giant' (> 5 B barrels) fields of Iran, Iraq, Kuwait, Saudi Arabia, Qatar, the UAE and Oman, to the major fields of Syria and Yemen and the Gulf of Suez region, to the relatively modest fields of Bahrain

333

and southeast Turkey and to the minor production of Jordan and Israel (Beydoun, 1988). Cumulative Middle East oil production from all reservoirs since the first Iranian discovery has totalled about 175.3 B barrels, making the ultimate recoverable oil reserves for the region around 841 B barrels of oil (Oil and Gas Journal, 30 December 1991). Cumulative gas production figures are not discussed here because all gas production in the region is at present for local usage (desalination, power generation, fertilizers, feedstock, light industry) or for reservoir pressure maintenance by reinjection; some gas is still being flared during oil and gas separation, pipeline oper­ ations and refining. No gas as yet is being exported, although plans to export are well advanced in several countries of the region and exploration for non­ associated gas to establish additional reserves has been stepped up. Most of these immense reserves are housed in reservoirs predominantly of Permian to Oligo­ Miocene age, of which 80% are carbonate; the bulk of these reserves (98.7% of the oil and 98.8% of the gas) are located along the prolific northeast Arabian shelf. This region had formed the Arabian passive Tethyan margin of Gondwana (and subsequently of Afro-Arabia) from the initial cratonization of the Arabian Shield in the late Proterozoic until the collision and suturing with Eurasia. Initial collision commenced in the medial Tertiary and continued to the present, overprinting the Zagros foreland basin on the outer edge of this margin; the northeast margin extends from northwest Iraq through south­ west Iran and the eastern Arabia-Gulf region to northern central Oman (Fig. 1). The current hydrocarbon-producing reservoirs of the region and those being developed for production are shown by country in Table 1. The clastic reser­ voirs are shown in Fig. 2 and the principal areas of clastic reservoir production are indicated on Fig. 1. Apart from the relatively long established major production areas of the Gulf of Suez ('Nubia' Palaeozoic-Mesozoic, and Miocene, constituting Egypt's most prolific hydrocarbon province) and the Basra -Kuwait-northeast Saudi Arabia-Bahrain Cretaceous reservoir areas near the head of the Arabian-Persian Gulf, the remainder of the import­ ant producing areas constitute new additions to the list, with the Oman Palaeozoic discoveries, many dating from the mid- 1970s, the Syria and Yemen Mesozoic discoveries, from the mid-1980s, and the Central (Saudi) Arabian Upper Palaeozoic discover­ ies, from the late 1980s. Other minor clastic reservoir

Z. R.

334

Beydoun

Table 1. Petroleum producing formations of the Middle East. Expanded and updated from Beydoun (1991), reproduced by permission of AAPG Periods, Epochs, Ages

(Modified fonn Dunnington MIQ. PLIOCENE

1967)

GULF OF

YEMEN

SUEZ

ISRAEL/

PA L ES T INE

JORDAN

S.E. TURKEY

SYRIA

MIQ.PLIOCENE M · U MIOCENE L-M MIOCENE OGLIQ.

OLIGO. MIOCENE

U • M EOCENE L.EOCENE·PALAEDCENE

EOCENE/ PALAEOCENE

Habshiya : 0. (Gulf of Aden)

Belayim: 0 Kareem: 0 Rudeis: 0 Nukhul: 0

JOJTibe: O.G. Chilou: O.G. Jaddala : 0. Sinan : 0.

MAESTRICHTIAN U.CAMPANIAN L.CAMPAN-/SANTDNIAN CONIACIAN

u c

Sudr: 0

Garzan: 0

Matulla: 0

Beloka : 0. Karabogaz : 0.

Shiranish: O.G. Amah: O.G. ''Massive"

R E T A

TURONIAN M

c

APnAN L

BARREMIAN HAUnERIVIAN

VALANGINIAN BERRIASIAN

J

u

Karababa : 0.

PORTLANDIANM. KIMMERIDGIAN

Limestone : 0.

+ Derdere: o. (Mardin)

"Judea" : 0.

Nubia "A"

Qishn Clastics: 0

("Malha"): 0 Heletz: 0

Rutbah : 0.

Naifa: O.G. Alit: O.G.

u L. KIMMERIDGIAN

R A

s s I c

Ajlun Gp: 0. (Hummar & & Na'ur)

Raha: 0

ALBIAN

E

0 u s

CENOMANIAN

OXFORDIAN/CALLOVIAN M L

BATHONIAN BAJOCIAN

Amran : G.

Nir Am: 0

Kohlan: 0

Brur: O.G (Zohar: G)

LIASSIC L. LIASSIC

Butmah : 0. Camurlu: G--+Kurachine : O.G.

TRIASSIC

(Mulussa) : G.O. PERMIAN PERMQ.CARBONIFEROUS

p A L A

Dubaidib:

G

Katin : G.O.

CAMBRQ.ORDOVICIAN

I

CAMBRIAN I INFRA· CAMBRIAN

c 0 =Oil

Markadah : G.

Nubia "C"

CARBONIFEROUS SILURIAN-DEVONIAN

E 0 z 0

G

=

Gas, as predominant production

G after 0

=

Gas alone in some fields

Condensate in many of the non-Permian Gasfields

hydrocarbon production in the region occurs in Turkey (Palaeozoic), Jordan (Palaeozoic) and Israel (Cretaceous). The identified or inferred source rocks that charge these Middle East reservoirs, both car­ bonate and clastic, are indicated in Fig. 2 and those that have a known relevance to sourcing the clastic reservoirs are discussed briefly in the text, together with the environmental characteristics and some relevant petrophysical properties of the reservoirs. In a survey and review of this scale, no attempt has been made to describe the clastic sediments

concerned in the interpretive detail reached by some of the other contributors or, indeed, that which might be considered adequate by the dedicatee of this volume. Nevertheless, the review is offered with the intention of drawing out the important distinctions necessary in viewing the hydrocarbon­ productive clastic sediments of the prolific Middle East region in terms of age, plate tectonics, palaeolatitude, climate and broader depositional environments. Numerous oil and gas field names are mentioned

335

Middle East clastic oil and gas reservoirs

Table 1. (Continued. )

N. IRAQ

S.W. IRAN

OFFSHORE IRAN

CENTRAL & S. IRAQ

KUWAIT & DIVIDED ZONE

SAUDI ARABIA

QATARBAHRAIN + OFFSHOR E

UNITED ARAB EMIRATES

OMAN

Fars: G. "Main" O.G. limestone :

Asmari: O.G. (Kalho<)

AvanaiJaddala /Aaliji: 0.

Ahwaz/Asmar i

"Ghar": 0.

G.O. Jahnum: 0.

Shiranishl "Pilsner" :0.

Gurpi : 0.

Belayim/Rudois : G.O. (Rod Sea basin)

Ghar: 0

Pabdeh:

Dammam : O.G. U. Radhuma : 0.

G.

Hartha : 0.

llam : 0. Magwa: 0. (Mishri� Komotan/ Dokan I 0. "Qamchuqa" :

L. Qamchuqa and 0. Garagu

Mishrif: 0.

Mishrif: 0.

Nahr Umr : 0.

Rumaila : 0. Ahmadi: 0. Wara : 0. Mauddud: 0. Burgan : 0.

Servak : 0. Ahmadi : 0. Wara: 0. Mauddud: 0. Burgan: 0.

Wara: 0. Mauddud: 0. Burgan: 0.

Dariyan: G.

Shuaiba: 0. Biyadh :

Zubair: 0.

Zubair : 0.

(Zubair) :

Yamama : 0.

Minagish :0. (Ratawi)

Yamama : 0. Sulaiy: 0.

0. Fahliyan : G.

Fahliyan : 0. (Yamama)

0. Bubwaib: 0.

Hith: 0. Arab A�: 0. Arab D : 0 Jubaila : 0. Hanila : 0. Tuwaiq: 0.

Arab : 0. Arab D : 0.

Dhruma: 0.

Najmah : 0.

Mishril: 0. A hmadi/ Khatiyah: 0 Wara: 0 Maud d ud: 0 Nahr Umr: 0. Shuaiba : 0 Kharaib: 0

G. Mauddud : 0. Thamama (Shuaiba) : 0.

o.

Thamama (Shuaibo) : 0.

Thamama: 0.

Arab A�: G.O. Arab D I Darb: 0.

Arab I Darb. 0.

Araej: 0.

Araej : 0.

Marra!: 0.

Natih:

Mafraq: 0

Marrat: 0.

Kurachino : 0. Kangan: G. Dalan : G.

Dalan: G.

' Khuff: G.

Khuff: G. Unayza: O,G.

Sara: 0 Sajir: 0

Khuff: G.

Khuff: G.

Pre-Khuff :

G.

Khuff: O,G. Haushi : 0. (G harif, AI Khlata) Haima: O,G.

Haima: 0, G (Haradh, Amin, M ahawis/Andam) Ara: 0. Buah: G.

in the text but not given precise location on Fig. 1 because of the scale involved. A better overview of field locations (with the exception of those in the Levant-Sinai fragment) can be obtained from the enclosure map in Beydoun ( 199 1), which was also issued with Bull. Am. Assoc. petrol. Geol., 75(7) for July 1991.

P A L A EO G EO G R APH I C A L A N D P A L A EO C L I M A T I C S Y NOP S I S From the latest Proterozoic to the latest Palaeozoic

(Permian) the Arabian Plate, including the Levant­ Sinai fragment, was attached to the African Plate, which together formed part of Gondwana and was located in the southern hemisphere; it was part of the long and very wide shelf adjacent to the 'northern' margin of Gondwana and bordered the Palaeo-Tethys ocean. This wide shelf expanse was covered intermittently by shallow epeiric seas along­ side lowlands consisting of the cratonized Arabo­ Nubian and other shields; the areal extent of these shelf seas changed in response to succeeding trans­ gressions and regressions. Following plate tectonic reconstructions by Scotese et at. ( 1979) and Scotese

336

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& McKerrow (1990), the Arabian Plate (within the borders defined above) from latest Precambrian to Early Silurian had its 'long axis' (Fig. 1 inset) oriented) approximately NE-SW, with its present northern extremity in the southwest (Fig. 3). Palaeolatitudinal positions during this time changed through the following steps (Fig. 3):

Latest Precambrian Late Cambrian Middle- Late Ordovician Late Ordovician Early Silurian

between 7°S (Oman) and 2SOS (Sinai) between 30°S (Yemen) and 55°S (Sinai) between 28°S (Oman) and 52°S (Sinai) between 35°S (Oman) and 58°S (Sinai) between 29°S (Oman) and 50°S (Sinai)

The 'long axis' orientation changed with the advent of the Devonian and continued to shift orientation for the remainder of the Palaeozoic, as tabulated below, with the position of the present northern extremity ( PNE) indicated for each new orientation. (Fig. 3): Devonian Early Carboniferous Permian

orientation E-W, PNE west orientation NW-SE, PNE northwest orientation N-S, PNE north

Palaeolatitudinal positions during this time span changed as follows (Fig. 3): Early Devonian Early Carboniferous

Late early Permian

Late Permian

between 22°S (Zagros suture) and 39°S (Yemen) between 14°S (southeast Turkey/northern Iraq) and 37°S (Yemen) between 28°S (southeast Turkey/western Iraq) and 47°S (Yemen) between l2°S (southeast Turkey/northern Iraq) and 36°S (Yemen)

This sequence of changes would have placed the Middle East region throughout most of the Palaeo­ zoic mainly in temperate middle to high latitudes, where a humid continental climate with large con­ trasts in seasonal temperatures would have prevailed and where clastic sedimentation would have pre­ dominated. The region drifted into the edge of the warm and dry (arid) climate belt towards the end of the Palaeozoic. Only sporadic migration into or

337

influence of warmer climatic belts occurred before the Late Permian, allowing warm water carbonate shelf sedimentation, with or without accompanying aridity and deposition of evaporites, for relatively limited spans of time. For example, during latest Proterozoic to Early-Middle Cambrian time, evap­ orites and carbonates were deposited in the Gulf region and Oman (latitudes 5° -15°S) (Huqf and Hormuz Groups) and carbonates (Middle Cambrian Burj Formation) in the Levant (25°S); during the Devonian and Early Carboniferous, Jauf Formation limestones were laid down in Saudi Arabia (latitude 22°-20°S) and Markada and Harur Formations car­ bonates in Syria and Iraq (latitudes 15°-28°S). Finally, in the Late Permian, the Khuff cyclic car­ bonates and evaporites, which were laid down over most of eastern Arabia (S0-30°S) during a rise in sea-level following the end of Gondwanan Permo­ Carboniferous glaciation heralded the Mesozoic predominance of low-latitude carbonates (Fig. 3). During the Late Ordovician, the present part of central and western Arabia formed the northernmost extent of the Ordovician (Saharan) glaciation, lying at about 50°S (Zarqa and Sarah Formations of Saudi Arabia), and during Late Carboniferous to Early Permian, the present southern part of Arabia (Oman-Yemen and southwest Saudi Arabia­ Ethiopia) again underwent an episode ofglaciation (the AI Khlata Formation of the Haushi Group of Oman and its correlatives in Yemen and Ethiopia). However, considering that this glaciated strip of Arabia (and adjacent Afghanistan-India to its south) was located in middle latitudes during that phase (average 42°-45°S) the glaciation there was probably of uplifted (upwarped) areas prior to Permo-Triassic rifting of this part of Gondwana. During the Mesozoic, the Arabian Plate region moved to tropical latitudes, and its orientation ('long axis') was similar to that of the present. In the Early Triassic it was still wholly in the southern hemisphere (0°-25°S), but from the Early Jurassic onwards it straddled the equator, extending to no more than 20°S but generally averaging about equal areas north and south of the equator (Beydoun, 1991). In the Cenozoic the move northwards accelerated and from the late Palaeocene onwards it lay totally in the northern hemisphere, its northern tip not extending northwards beyond about 35°N. In consequence, warm-climate carbonate sedimentation has been predominant from the Late Permian onwards, accompanied from time to time by episodes of aridity; the deposition of associated clastics was

Z. R.

338

Beydoun

3o• s

30° s

so• s

c

so•

D

o•

o•

NUB - Nubia

IR - Iran

AFG - Afghanistan TUR - Turkey

IND - India

A.

LATE CAMBRIAN

MAD - Madagascar

B.

MIDDLE-LATE ORDOVICIAN (LLANDEILIAN - CARADOCIAN)

SOM - Somalia

c.

LATE ORDOVICIAN (ASHGILLIAN)

D.

EARLY SILURIAN (LLANDOVERIAN)

E.

LATE EARLY DEVONIAN (EMSIAN)

F.

EARLY CARBONIFEROUS (VISEAN)

G

LATE EARLY PERMIAN (ARTINSKIAN)

H.

LATE PERMIAN (KAZANIAN)

Fig. 3. Palaeozoic plate tectonic reconstructions. Arabia and surroundings. Modified with permission from Scotese eta/.

(1979); Scotese

&

McKerrow (1990).

Middle East clastic oil and gas reservoirs

principally in the form of transgressive shales or pelagic marls, some of which form major source rocks and/or seals. Episodes of coarse clastic depo­ sition were restricted to either arid conditions in specific climatic rain-shadow belts (Triassic) or to major fluvio-deltaic regimes draining uplifted regions and confined to specific time slices and areas. Examples are the Lower and Middle Cretaceous systems of the southern Iraq-Kuwait and north­ eastern Saudi Arabian region, the Gulf of Suez area and the Levant; the Oligo-Miocene systems of the same Kuwait-Basra area; the clastic deposits associated with initiation of rifting and subsequent marginal infill (Marib-Jauf in Yemen, Gulf of Suez): the regressive/basal transgressive coarse clastic deposition following regional uplift (end Jurassic­ early Cretaceous in various sectors): or finally, the Mio-Pliocene molasse associated with the rising Zagros-Taurus orogens of southwest Iran, Iraq and southeast Turkey. It is these varied palaeolatitudinal settings of the Middle East Arabian Plate region that have localized the development of clastic (reservoir) facies from the Permian onwards but permitted their extensive regional development prior to the Late Permian. Post-Palaeozoic deposition of carbonates over the wide Arabian shelf region gave rise to very areally extensive reservoir facies, periodically alternating with widespread source-seal units, thus enabling the generation and trapping of the very large hydro­ carbon reserves that are located principally in the northeastern part of the region, in large to 'giant' structural traps with predictably identifiable 'kitchen' areas and migration pathways. By contrast, the gen­ erally quite rapid lateral facies changes of the predominantly clastic Palaeozoic sequences and particularly the relative paucity of thick efficient (unbreached) seals, precludes short migration path­ ways from source to reservoir into trap, and often permits considerable vertical migration until an effective (regional) seal is encountered. Moreover, the distribution of Palaeozoic source sediments has been dependent largely on deposition in tectonically controlled basins or during major transgressive episodes, which subsequent tectonic events have obscured or stripped off. So far, only two major regional source facies have been identified: the Infra­ Cambrian to Lower Cambrian Huqf Group of Oman and the Gulf region and its correlatives, which are now generally deeply buried beneath later sediments and are mainly overmature, except in Oman, and the Lower Silurian Qusaiba/Safiq shales and their

339

correlatives; other potential Palaeozoic source rocks are of local significance only. Because Palaeozoic formations, and especially the pre-Devonian ones, are stratigraphically much older and, generally, deeper as exploration targets in a region fortunate in its profusion of prolific Tertiary and Mesozoic carbonate-reservoired oil accumu­ lations and in Late Permian carbonate-reservoired gas accumulations, the widespread search for pre­ Permian targets had to await changes in global economic and energy demands before being initiated. The first such exploration was carried out in Oman, where the Palaeozoic sediments are not deeply buried in the southeastern part of the country, and later in other parts of the region lacking younger reserves ( Jordan), before the latest successful ven­ tures in Saudi Arabia. Based on these recent very major discoveries in central Arabia in Upper Car­ boniferous to Lower Permian sandstone reservoirs and the identification of the source rock for these substantial accumulations as being the Lower Silurian shales (see next section), the potential for finding considerable new Palaeozoic reserves else­ where in the Middle East region is particularly promising. Potentially, overall the Middle East reserves could well double in the coming two to three decades of intensive exploration of the stratigraphically deeper levels, provided due understanding of the palaeogeographical/subcrop distribution of regional source sediments, source 'kitchens' and maturation levels are arrived at to direct the search (Beydoun, 1991).

P A L A EO ZO I C C L A STI C R E S E RVO I R S At present, the most important Palaeozoic clastic reservoir production in the Middle East region is that in central and southern Oman (Fig. 1), which first commenced regular production in 1975. Heavy oil of 19°-22° API gravity had been discovered in southern Oman (Dhofar Province) as early as 1956, by the Cities Service Company at Marmul, but it was never developed because of the heavy and viscous nature of the oil and the low oil prices prevail­ ing prior to 1973, which made the venture non­ commercial; the age of the reservoir had not then been established precisely nor had the significance of the oil sourcing been appreciated or properly understood. The sharp rise in oil prices in 1973 coincided with the early phase of serious exploration for older

340

Z.R.

deeper reservoirs in central and southern Oman (additional to the Cretaceous reservoirs of central Oman, which had been on production since the mid1960s), and this gave added impetus for accelerated exploration and development following the first dis­ coveries of light oil in Palaeozoic sandstone reser­ voirs. First commercial production began in 1975, from the fields in the Ghaba basin, and there are now at least 33 fields producing from Palaeozoic reservoirs in southern and central Oman, which had cumulatively produced some 890 million barrels of oil until the end of 1991. There are several billion barrels of recoverable oil in these reservoirs. Clastic reservoirs in Oman There are two major clastic reservoir groups in Oman. The older group comprises continental and shallow-marine sandstones of the Haima Group, which range in age from Early Cambrian for the oldest (Amin Formation) to the Late Cambrian to Early Ordovician (Andam Formation); the second group comprises glaciogenic sandstones and shallow­ marine and fluvial-channel sandstones of the Haushi Group, which ranges in age from Late Carboniferous (AI Khlata Formation) to Early Permian (Gharif Formation) (de Ia Grandville, 1982; Hughes Clarke, 1988; Sykes & Abu Risheh, 1989). Haima Group

The oldest oil-bearing Haima Group clastic sedi­ ments in Oman are those of the Karim and Haradh Formations (the Lower Haima of South Oman; Hughes Clarke, 1988). The Haradh is of greatest significance as a reservoir and comprises chert-rich cross-stratified sands with abundant synsedimentary deformation features. It probably represents a braided alluvial apron shed from highlands to the southeast, with the abundant chert being derived from the underlying formations of the Huqf Group (Heward, 1989). The Amin Formation reservoir is predominantly an intercalation of aeolian dune, sabkha and water­ lain sands. It is remarkable for the generally small size of the dunes that deposited it and reflects a period of arid continental conditions of widespread extent. Conglomerates derived from a deformation front to the west of Oman interfinger with these desert deposits. The Amin Formation probably cor­ relates, at least in part, with the widespread Saq (sandstone) Formation of Saudi Arabia.

Beydoun

The Mahwis Formation of South Oman comprises waterlain sheet-flood sands and conglomerates, probably again derived from the west (Heward, 1989). It passes northwards into the coastal andl storm-dominated shallow-marine sediments of the Andam Formation, which consists of fine sandstones, shales and thin ooidal and stromatolitic limestones in the middle part. The sediment package, of Late Cambrian to Early Ordovician age on the basis of its trilobite and other fauna, represents the first marine transgression in the Palaeozoic of Oman (Hughes Clarke, 1988). The Ghudun Formation reservoir (Haima Group), although an excellent reservoir, is only of minor regional significance, being of importance in only one South Oman field. Haushi Group

For the Haushi Group reservoirs, the lower or the AI Khlata Formation reservoir consists of a complex of clastic lithologies with considerable lateral and vertical changes in both facies and in thickness; this is particularly the case in South Oman, where these changes occur over very short distances (Levell et a/., 1988; Heward, 1990). Lithologies range from coarse-fine diamictites, conglomerates (particularly at the base), gravels, pebbly sands, sands, silts, silty shales and shales that are often varved. Boulders can reach up to 10 m in diameter and are of very varied composition, including granite, granophyre and volcanics, as well as Infra-Cambrian sedimentary rocks of the Huqf Group. The AI Khlata is a glacially generated sediment package and was deposited by a series of glacial advances and retreats from the southwest towards the northeast, as evidenced by striations and 'whale backs' in glacial pavements in outcrop (Levell et at. , 1988). The reservoirs are predominantly glaciofluvial and glaciolacustrine (Heward, 1990). The AI Khlata correlates with simi­ lar deposits present elsewhere in southern Arabia and Ethiopia (Beydoun, 1988; McClure, 1980; Levell et a/. , 1988). The overlying Gharif Formation reservoir of the Haushi Group consists of a series of fluvial deposits derived from the southeast. The arkosic character is due to recycling of granitic material from AI Khlata, presumably being stripped off from the uplift that finally rifted between Oman- Afghanistan-India to the southeast. The formation fines and becomes more shaley to the northwest; its deposits reflect an ameliorating climate, most being red-beds, locally

Middle East clastic oil and gas reservoirs

with calcrete and frequent occurrences of silicified wood in outcrop, with growth bands indicating a seasonal climate. Marine sands and limestones occur to the north and west and reflect the interplay between marginal-marine -coastal-plain deposits and continental red-bed facies (Hughes Clarke, 1988). The formation generally correlates with the upper Unayza Formation of Saudi Arabia discussed below. The considerable volumes of hydrocarbons in place, housed in the Palaeozoic clastic reservoirs of the Oman region, are demonstrated to have been abundantly sourced from Huqf Group source sedi­ ments at least 550 Ma in age (but also of older Upper Proterozoic age) straddling the Infra-Cambrian to Early Cambrian boundary and terminating with a widespread evaporite unit- mainly salt (Gorin et al., 1982; AI Marjeby & Nash, 1986; Grantham et al., 1988; Mattes & Conway Morris, 1990). A few fields in western Oman, however, adjacent to the eastern side of the Rub al Khali basin, have oil accumulations in Haushi Group clastic sediments that have been geochemically typed to Silurian (Safiq and equivalent) source rocks (Grantham et al., 1988; Beydoun, 1991). The intermittent uplift of the southeast Arabian plate margin since mid-Palaeozoic times, and par­ ticularly since the Silurian, when the region became relatively positive, is the principal reason why these very old source rocks in South Oman remain in the oil window and continue to generate oil; they lie, on the whole, at moderate depths and seemingly have not been more deeply buried. The relatively heavy nature of some of the oil is considered to be due to subsequent transformation, although some may be ascribed to generation from still relatively immature sediments at shallow depth. Hydrocarbon charge from these old source rocks into Palaeozoic clastic reservoirs occurs from beneath the terminal Ara Formation (Huqf Group) salt edge and takes place in conjunction with progressive subsurface salt removal and dissolution; this leads to the formation of residual cores composed of basin-fill bodies of the overlying post-salt Palaeozoic clastic deposits (AI Marjeby & Nash, 1986). Short- to medium-distance lateral migration into the Palaeozoic clastic deposits, and/or subsequent long-distance vertical migration within these clastic deposits, occurs until stopped by an effective seal in trapping geometry; these seals are predominantly shales or argillaceous sediments. In the clastic depositional facies patterns that charac­ terized the Palaeozoic of the region, shale seals

341

generally have limited lateral extent and are often thin and easily breached. Migration of Huqf sourced oil has, thus, often moved vertically up as high as the Carboniferous-Permian reservoirs of the Haushi Group, where sealing is provided by glaciogenic or shallow-marine-floodplain shales and mudstone deposits. Furthermore, where these have locally become breached or have proved inadequate, verti­ cal migration has continued in central Oman to charge the Mesozoic carbonate reservoirs of the largest Oman fields (Sykes & Abu Risheh, 1989). The subject and wider implications of Palaeozoic oils in Middle East exploration has been discussed and reviewed in more detail by Beydoun ( 1991). The AI Khlata reservoirs are determined by a number of facies belts whose distribution is partly influenced by local salt dissolution (of the underlying Huqf Group Ara Formation salt); where subsurface salt dissolution occurred during the glaciation episode, lakes were formed with generally shaley deposits. Along the margins of the lakes, glaciogenic deltas gave rise to mixed sequences of sands and shales (multiple reservoir-seal couplets) and consti­ tute the most prospective facies belt. Within the main South Oman salt basin, terrestrial deposition dominated, with much waste rock (diamictites) and erratic reservoir character (Levell et al., 1988). Deltaic deposits are occasionally incised by sand­ stone-filled channels and rapid lateral and vertical variations in lithology and stratigraphy make it diffi­ cult to estimate and indeed develop all recoverable oil reserves from the AI Khlata reservoir. The over­ lying transgressive Gharif Formation sands often provide a peripheral stratigraphical trap component reservoir to the structural trap formed by drape over Haima Group sediment pods left by salt withdrawals and dissolution (AI Marjeby & Nash, 1986). Porosities and permeabilities of the South Oman Palaeozoic clastic reservoirs are generally good and range from 15 to 35% and 100 mD to over 10 D in the friable and poorly cemented reservoir levels (de Ia Grandville, 1982; Levell et al., 1988). Clastic reservoirs in Saudi Arabia In central Saudi Arabia, several billion barrels of high gravity (c. 43° API), low-sulphur Palaeozoic oil (and gas) have been discovered recently in the region south of Riyadh (Fig. 1), principally in the Unayza Formation sandstone reservoirs of ?latest Carbon­ iferous to Early Permian age (AI Laboun, 1986; Husseini, 1992; McGillivray & Husseini, 1992), but

342

Z.R. Beydoun

also including some oil housed in Devonian and Ordovician sand reservoirs. Oil to potential source rock correlations have identified the principal source of these hydrocarbons to be the basal, organically rich shales of the Qusayba Member of the Lower Silurian Qalibah Formation (Abu-Ali et al., 1991; Husseini, 1992; McGillivray & Husseini, 1992; Mahmoud et at., 1992). The Qalibah Formation was previously unrecognized as a separate formation and was included in the middle part of the Tabuk Forma­ tion (Powers, 1968). The first significant Palaeozoic sandstone reservoir hydrocarbon discovery was made in 1979 in Qirdi (south of the Khurais field, then the westernmost established oilfield in Saudi Arabia) in an Upper Permian pre-Khuff (carbonate) Formation sandstone sequence and in the same year pre-Khuff gas was also tested in the 'supergiant' Ghawar field (Husseini, 1992). Other discoveries followed in deep tests on the Khurais trend south of that field, but at depths in excess of 4000 m. It was not until 1989 that the major discoveries in the previously unexplored areas were made, and at moderate depths. The Unayza Formation is a siliciclastic succession of fluvial, coastal-plain, deltaic and shallow-marine facies whose age may extend back to Late Carbon­ iferous (AI Laboun, 1986), but is principally of Early to Late Permian age; it appears broadly cor­ relatable with the Gharif Formation of Oman whose deposition is associated with the Late Carboniferous to Early Permian deglaciation episode that affected Gondwana, including the southern part of Arabia (see earlier). Distribution of the varied Unayza facies is controlled by palaeogeography, sea-level fluctuations and by local and regional structure (Ferguson & Chambers, 1991). The complex inter­ play of these factors coupled with erosion during sea-level lowstands, has resulted in a variety of reservoir types, varying through poorly sorted coarse silt-sand of tentative marine origin, accumulated in fault bound depressions; fine to pebbly cross-bedded sandstone of alluvial origin; a variety of paralic sandstone facies, including shoreface-foreshore, delta channel and coastal-plain channel sands (it is these paralic facies that have proved to be the most prolific reservoirs); the final facies is a transgressive lag of unpredictable distribution but consisting of clean fine- to very coarse-grained reservoir quality sandstone (Ferguson & Chambers, 1991). Porosities in these facies are highly variable but average 20%, although permeabilities of several darcies are not uncommon (Husseini, 1992).

Other reservoir sandstones in this region of Arabia include Ordovician glacially related sandstones and Devonian ?continental sandstones. The hydrocarbon charge to the trap from the Silurian Qusayba shales is controlled by juxtaposition of reservoir to source through faulting and/or fluid communication with the main Unayzah reservoir. Stratigraphical pinch out of the Unayzah reservoir against the basal Khuff seal unconformity may affect the trap form significantly. The Lower Silurian Qusayba shales (and the equivalent Safiq in Oman) are widespread across Arabia and reflect the first transgressive cycle related to global sea-level rise consequent on the deglaciation occurring at the end of the Late Ordo­ vician to Early Silurian 'Saharan' glaciation. Similar shales are widely distributed throughout the Middle East and North Africa, but subsequent post- Middle Devonian tectonism in the Middle East and uplift have removed them from elevated areas; conse­ quently, their subcrop palaeogeographical distri­ bution in the· region beneath the Late Permian Khuff seal needs to be better controlled. The Qusayba shales are, however, seemingly present everywhere in the basins of Saudi Arabia and at depths where they are sufficiently mature to generate oil and gas (Husseini, 1992); this includes those portions of the Arabian Peninsula adjacent to Jordan and southern Iraq (AI Laboun, 1986). Oil generation and up-dip migration from some of the Arabian depocentres commenced as long as 160 Ma (Abu-Ali et a!., 1991). Clastic reservoir in the Gulf Region In other areas of the Gulf region, gas had been discovered in pre-Khuff clastic reservoirs in Abu Dhabi and Bahrain and non-commercial gas found in onshore Qatar. These gas discoveries are as yet undeveloped although potentially producible should market forces warrant it; but although the gas is generally H2S free, appreciable amounts of associ·· ated nitrogen and C02 cast doubt on commerciality. The principal accumulations in these countries occur in Unyiza-equivalent clastic deposits, and the sequences consist of continental fluvio-deltaic sands, silts and shales with several reservoir sandstone intervals separated by shale bodies. The source of the gas is generally assumed to be the Silurian shales, which, however, have not been penetrated in any Abu Dhabi well (Ali & Sildawi, 1989; Loutfi & El Bishlawy, 1989), although they may have

Middle East clastic oil and gas reservoirs

been partly penetrated Chaube, 1987).

m

Bahrain (Samahiji &

Clastic reservoirs in Syria In Syria, a number of gas accumulations occur in anticlinal structures in the central part of the country (southern Palmyrid basin) (Fig. 1). These are cur­ rently undergoing delineation drilling in preparation for commercial production. The reservoir facies is sandstone, occurring in intervals separated by clay­ stones, thin shales and intercalated dolomites of the Upper Carboniferous Markada Formation; there are rapid vertical and lateral lithological variations. The depositional environments are probably shallow­ marine to ftuvio-deltaic, with provenance of clastic material from erosion of the adjacent uplifts (Rutbah High and ?Aleppo High) created by an 'Hercynian' tectonic event. Periodically reduced clastic supply allowed carbonate deposition, commensurate with palaeolatitudinal location in tropical waters. Log porosities in the clastic sediments are reported to range between 5% and 16% . The source of the gas (predominantly methane but with C2 and C3 fractions and varying small amounts of nitrogen) has been inferred to be from shales within the Markada itself; however, based on regional considerations, it seems more likely to be from unconformably underlying Silurian (Tanf Formation) shales in a similar manner to the Unayza accumulations of Central Arabia. Although not proven to underlie this area, these shales are extensively distributed in the region and have been encountered in a number of widely scattered wells (Lababidi & Hamdan, 1985 ; Beydoun, 1988, 1991). Clastic reservoirs in Jordan In adjacent Jordan, the Risha field, close to the Iraqi border (Fig. 1) , is producing gas from a clastic reservoir of Late Ordovician age; this is the Dubaidib Sandstone Formation of the Khreim Group lying directly beneath the organic-rich Lower Silurian (Mudawara Formation) shales, which appear to act both as source and seal for this reservoir (NRA Jordan, 1989). This area of northeast Jordan was an Early Palaeozoic basin that accumulated a thick succession of Palaeozoic sediments, several thousand metres of which (Early Cambrian-Silurian) are still preserved following inversion and erosion in the Late Palaeozoic. The Ordovician reservoir sequence here consists of shallow-marine cyclic sandstone

343

sequences with stringers and intercalations of shales, claystones and siltstones; lateral and vertical con­ tinuity varies rapidly, however, being the main reason why only three or four of the dozen or more wells drilled over the (?)imprecisely defined sub­ surface feature to date are producers (NRA Jordan, 1989). Good permeabilities, however, are developed locally, as evidenced by the reported frequent mud losses during drilling. Light but non-commercial oil has been found in the same reservoir in well Wadi Sirhan 4, some 250 km to the southwest (NRA Jordan, 1989). Clastic reservoirs in southeast Turkey In southeast Turkey, several small accumulations of gas and condensate have been discovered in Devonian age clastic sediments of the Katin Forma­ tion, below the southern front of the Taurus thrust belt (Fig. 1). The reservoir facies consists of clean shallow-marine sands, which are oil impregnated nearby at outcrop (Hazro sands, previously thought to be Permian in age; Beydoun, 1991). These reser­ voir sands are considered to have been sourced by the Ordovician-Silurian organic-rich Bedinan For­ mation shales (locally defined as the Dadas Forma­ tion) with TOCs up to 5% (Harput & Erturk, 1991); these also contain reservoir sand levels (Handof Member) that have contained strong but non­ commercial hydrocarbon shows in a number of wild­ cat wells in the southeast Turkey region. The cumulative production from these Palaeozoic reser­ voirs has been modest to date; however, geochemical work in the area by a number of operators confirms that the Bedinan Formation has been the source charging a number of oil-productive Cretaceous car­ bonate reservoirs in southeast Turkey in areas where the Middle Cretaceous often unconformably over­ lies the Palaeozoic (Mardin-Bozova High). It is thought likely that accumulations in older clastic Palaeozoic reservoirs, such as those found in the Barbes and Katin fields, could be discovered with better subsurface well control and seismic mapping. Clastic reservoirs in the Gulf of Suez In the petroliferous Gulf of Suez rift basin of Egypt, on the south western edge of the Levant-Sinai fragment of the Arabian Plate, the bulk of the oil reserves are housed in sandstone reservoirs, the majority being of Miocene age, but a number of fields also produce from 'Nubia' sandstone reservoirs

344

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of Late Palaeozoic and Cretaceous ages. The Palaeo­ zoic reservoirs belong to the Nubia 'C' unit, which is of ?Devonian-Carboniferous age (the Nubia 'B' consists of Carboniferous black shales, which are regarded as a likely source for some of the oil by some authors- Ayouty, 1990). The Nubia 'C' depositional facies in the Gulf of Suez region is principally shallow marine with deltaic and fluvial intercalations over much of the area but with alluvial deposits in the southern part; the facies became predominantly continental to shallow marine towards the end of Nubia 'C' time in the Late Carboniferous to Permian (Klitzch, 1990). About six or seven oil fields produce from the Palaeozoic sandstone reservoir, including the now depleted Hurghada field in the west, where the bulk of the oil was housed in the Nubia 'C' reservoir and which produced during its life span (1915-1969) 41 million barrels of oil. Other Palaeozoic reservoir fields include the Ras Gharib, July, Ramadan, Ras Badran and Sidky fields, where reservoir porosities and permeabilities are fair to good and contain oils having gravities that range between 24° and 40° API (Ayouty, 1990). The oil in the Palaeozoic and younger reservoirs of the Gulf of Suez is considered by many to be sourced mainly from organic-rich Early Miocene shales and marls (Rudais-Karim Formations) brought into the right juxtaposition of source to reservoir in tilted fault-block traps or stratigraphical pinch outs. Additionally, and perhaps alternatively, Late Cretaceous and Eo-Palaeocene sediments also contain ample organic carbon, enough to generate the oil in Cretaceous and Neogene reservoirs (Ayouty, 1990), and the Carboniferous Nubia 'B' shales cannot be ignored as another possible source to be considered in the calculation of total oil charge volumetrics.

M E S O Z O I C C L AS T I C R E S E R V O I R S Commercial oil production from Mesozoic sandstone reservoirs was first established in the Middle East region on the western side of the Gulf of Suez, in what are now depleted fields (Hurghada, 1915; Abu Durba, 1918; etc.) in 'Nubian' sandstones of Cre­ taceous age (now known as the Nubia 'A') (Ayouty, 1990). Sustained and continuing commercial oil production from sandstone reservoirs in the Middle East region, however, was first made in Bahrain in 1932 with the discovery of the Awali field.

Beydoun

Clastic reservoirs in the Bahrain- Neutral Zone- Kuwait - Basra region The 'Bahrain zone' in the Awali field consists of several productive Middle Cretaceous sandstone-­ siltstone (Wara, Nahr Umr) and carbonate (Ahmadi, Mauddud) reservoirs. Subsequently, the discovery and development of large oil reserves from equiv­ alent age sandstone reservoirs occurred in Kuwait in 1938 (Wara and Burgan Formations), in Lower Cre ­ taceous sandstone reservoirs in the Basra area o:f southern Iraq at Zubair in 1949 (Zubair Formation) and in Albian age sandstone reservoirs (Safaniya/ Khafji Members of the Wasia Formation) in the northeastern part of Saudi Arabia and the adjacent part of the Saudi-Kuwaiti Divided (Neutral) Zone in 1951 (among others, the Safaniya-Khafji field ­ the largest offshore oil field in the world). The oil (and gas) fields of the southern Iraq ­ Kuwait-Divided (Neutral) Zone-northeast Saudi Arabia onshore and offshore region, were discovered and developed since the late 1940s and earliest 1950s. They produce (or are being prepared to produce) from these Middle and Lower Cretaceous sandstone reservoirs, and although only 23 in number, they constitute (both in production and recoverable reserves) by far the most important clastic reservoir oil province of the whole Middle East region . The recoverable hydrocarbon volumetrics of this province alone involve a sustainable production of several million barrels of oil per day and many tens of billions of barrels of remaining recoverable reserves (Beydoun, 1988, 1991), well in excess of the oil totals for the entire North Sea Basin ! The sandstone reservoirs housing these vast reserves represent the deposits of a huge fluvio-· deltaic-littoral-marginal marine depositional system, which commenced in early Cretaceous (Hauterivian-Aptian) time, when this area of the Middle East lay athwart the humid equatorial belt , 2°-3° to north and south of the equator, and ter-­ minated in the Albian to early Cenomanian, when the area drifted to between so and l0°N (Dercourt et al., 1986; Beydoun, 1991). This fluvial system (a virtual 'Nile sized' river complex) transported immense quantities of clastic detritus northeastwards from the Arabian -Nubian shield highlands, and from the exposed Palaeozoic clastic sediments over­ lapping the shield, on to the shallow and very wide Arabian passive margin shelf. The delta complex prograded during sea-level lowstands but retreated during transgressive high sea-level stands when shelf

Middle East clastic oil and gas reservoirs

carbonates or fine clastic sediments were deposited between the coarser lowstand sandstones. A distribu­ tary channel system also extended the delta complex to the Baghdad area in Central Iraq (Jawad Ali & Aziz, 1993). The repetitive nature of the Zubair Formation deltaic cycle (Hauterivian-Aptian and correspond­ ing to the subsurface Biyadh Formation of Saudi Arabia), which is one of the major productive reser­ voirs of the area, recently has been documented qualitatively (Khaiwka, 1989): the cycle includes pro-delta, delta front and delta plain, with marine shelf facies concluding the cycle. Subaqueous levee, distributary channels , distributary mouths, distal bars, freshwater lake, mud-fiat, marsh, subarea) levee and lobe abandonment facies have all been recognized and have a strong repetitive vertical pattern. Rapid facies changes provide the lithological con­ trols on the variable hydrocarbon distribution in the lower regional (Zubair) clastic reservoir; it provides the major production in southern Iraq but is a relatively minor reservoir in Kuwait and the north­ east Saudi Arabian fields. Potentially important new oil production from the Zubair reservoir was reported as being established recently in the East Baghdad oilfield of central Iraq (Beydoun, 1988). The Zubair Formation is overlain by the trans­ gressive carbonate cycle of the Shuaiba Forma­ tion, which concludes with a terminal reefal build-up phase (Dunnington et al., 1959; Powers, 1968). The Albian (Nahr Umr and Burgan Formations) clastic sequences of the Basra-Kuwait area and their equivalents in northwest Saudi Arabia and the Neutral Zone (Khafji and overlying Safaniya Members of the Wasia Formation) initiate a new prograding fiuvio-deltaic-littoral-shallow-marine depositional cycle following the Shuaiba (transgress­ ive) carbonate cycle, similar to that of the preceding early Cretaceous Zubair cycle. This Nahr Umr to Burgan progradational episode grades up into the widespread shallow-shelf carbonate sedimentation of the Mauddud Formation, which terminates the Albian. The early Cenomanian Wara clastic depositional episode that followed the Mauddud was of more limited duration and areal extent, covering Kuwait and eastern Arabia to Bahrain, and representing a brief terrigenous interlude in an otherwise marine Middle Cretaceous depositional phase. The Wara consists of a sequence of sandstones with interbeds

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of siltstone and shale and exhibits rapid lithological changes both laterally and vertically (Powers, 1968). In its lower part, in the Kuwait region, it includes a large lenticular sand body enclosed by grey, lignitic shales (Owen & Nasr, 1958; Dunnington et al . , 1959) which indicate a return t o fiuvio-deltaic depo­ sitional conditions. This sandstone body interfingers with littoral-shallow-marine facies to the southeast (Bahrain) and with deeper marine pro-delta facies to the north and northeast (southern Iraq); sub­ aerial delta deposits, tidal-fiat and coastal-plain deposits dominate to the west, with the submarine portion of the delta being located east of the present coastline. The Wara reservoir contains significant volumes of oil in Bahrain, in the Divided (Neutral) Zone and in northeast Saudi Arabia and southern Kuwait, where clean permeable sands are common. The Wara clastic cycle is terminated by a transgressive marine shale facies of the Ahmadi Formation, which forms the regional seal to the Wara reservoir (Brennan, 1990; Samahiji & Chaube, 1987). In the Basra area the Wara is not distinguishable as a sandstone sequence, the coeval interval being a facies of black silty shales and siltstones with a thin oil­ saturated sand (Dunnington et al., 1959) . Reservoir porosities and permeabilities for these different Lower and Middle Cretaceous sandstone reservoirs vary and data are generally sparse, but good reservoir properties are commonly present. (OAPEC, 1985 ; Brennan, 1990, 1991) . In southern Iraq, a petroliferous belt of Albian Nahr Umr clastic sediments coincides with a marginal-marine depo­ sitional area (Ibrahim, 1983) but production from this reservoir is subordinate to that from the Zubair reservoir. The hydrocarbons in the Lower and Middle Cre­ taceous clastic reservoirs of the 'head of the Gulf ' region of the Middle East are believed sourced from the Kazhdumi. Formation organic-rich sediments in the Kazhdumi basin immediately to the east (Brennan, 1990) and/or from a 'deeper' Lower Cre­ taceous source ( Ibrahim, 1981) within the Gotnia depositional basin of Murris (1980) (Sulaiy and Chia Gara/Sarmord and Garau source) with vertical migration and charging from the Zubair into the Nahr Umr reservoir of southern Iraq ( Ibrahim, 1981). This question of sourcing of the various Middle East reservoirs and the regional distribution of source rocks has been reviewed and discussed in more detail by Beydoun (1988, 1991) and by Beydoun et al. (1992).

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Clastic reservoirs in Yemen Two new and potentially major hydrocarbon pro­ ductive Mesozoic clastic reservoir regions have been established in the Middle East since 1984. These are in the Marib -Jauf rift basin and the Wadi Hadhramaut basin of Yemen in southwest Arabia. The Marib-Jauf basin now has about seven oil­ fields on production, or being developed for pro­ duction, where the oil and gas/condensate reserves are housed in sandstone reservoirs of latest Jurassic age. The discovery field was Alit, which is also the biggest. In 1991, this basin's oil production averaged about 190 000 barrels per day and recoverable reserves were estimated at about 3-4 B barrels (Oil and Gas Journal, 30 December 1991); the gas (with condensate) reserves are estimated to run into sev­ eral trillion cubic feet and means for developing these reserves are under consideration. The Upper Jurassic subsurface succession within this rift basin differs from that outcropping nearby in that a thick synrift sequence of sediments was laid down ; this NW-SE trending basin started to subside rapidly in the post-Oxfordian, with in excess of 3500 m of sediments being deposited in about 6 Ma (Maycock, 1987). The Callovian-Oxfordian Amran (Shuqra) platform carbonates here are overlain by a thick and predominantly shaley sequence followed by carbon­ ate turbidites rapidly deposited following graben inception. This sequence, informally referred to as the 'Lam Formation', is correlative with the organic­ rich Madbi Formation of the outcrop area to the southeast (Beydoun, 1964; Beydoun & Greenwood, 1968) and is the principal generative source for the Marib-Jauf hydrocarbon accumulations ; the 'Lam' is overlain by the 'Alif Formation', which is the main reservoir and consists of a shallow-marine-deltaic­ fluviatile succession of clean sands that are of Tithonian age (Huurdeman et al . , 1989) and are succeeded by sandstones and shales of braided flu­ viatile and deltaic systems, in turn followed by the Sabatain (Main Salt) evaporite sequence, also of Tithonian age (Maycock, 1987; Huurdeman et al., 1989) ; these form an excellent regional hydro­ carbon seal. Excellent reservoir porosities and per­ meabilities have been recorded ranging from 16% to 18% and 500- 1200 mD, with the oil being of 39°-41 ° API gravity (Huurdeman et al . , 1989). Some 300km to the east of the Marib-Jauf rift basin (and about 225 km east-northeast of its Late Jurassic carbonate-reservoired Shabwa hydrocarbon basin extension) recent exploration has revealed

Beydoun

other Late Jurassic rifts on an E-W trend (Mills, 1992), with a deep lying pre-rift sediment package underlying a Kimmeridgian-Tithonian synrift se­ quence, which includes an analogue of the organic­ rich Madbi source sequence; a series of important discoveries have been made in this area, the first in December 1990; estimated recoverable reserves to date are conservatively estimated at about 250 million barrels. The succession in this Hadhramaut-Jeza rift basin area is that described by Beydoun ( 1964, 1966) and Beydoun & Greenwood ( 1968) for the typical non­ evaporitic Jurassic Amran Group of Hadhramaut overlying the Liassic to Middle Jurassic clastic Kohlan Sandstone Formation ; this consists of pre-rift Shuqra Formation platform carbonates and synrift Madbi and Naifa (Kimmeridgian ­ Tithonian-Berriasian) organic-rich shales and marls and argillaceous limestones, respectively. No evap­ orites equivalent to the Sabatain Formation are developed here and hence there is no efficient seal on top of the Jurassic, as in the adjacent Marib -­ Jauf and Shabwa rift basins. Exploration wells have shown some unexpected subsurface differences in the post-rift Cretaceous stratigraphy from the out­ crop sequence described by Beydoun & Greenwood ( 1968). An early Cretaceous platform carbon­ ate unit occurs beneath the Lower Cretaceous (Barremian-Aptian) Qishn Formation consisting of carbonates and alternating shales unconformably resting on the terminal Jurassic Naifa Formation and separating it from the Qishn ; this sequence has been informally named the Saar Formation of Hauterivian to early Barremian age (Mills, 1 992) and appears to have been laid down in an E-W shallow marine embayment. Additionally, the Qishn itself exhibits a substantial development of clastics derived from bordering and still positive uplifts to the marine embayment, below its main Barremian-Aptian carbonate unit, which are not developed in the main reference surface section for the formation within the Tawilah Group (Beydoun & Greenwood, 1968) located near Mukalla along the coast about 150 km due south of the first discovery. These clastic sediments in the area of recent discoveries (Hadhramaut-Jeza rift) are not only well developed but have good reservoir qualities. In this area, the hydrocarbons were generated in the organic-rich and well-developed Madbi source rocks in the adjacent graben(s) and migrated into the suitably juxtaposed older Kohlan (Lower-Middle Jurassic) sandstone reservoir and also, in the absence of an

Middle East clastic oil and gas reservoirs

efficient seal (Tithonian evaporites of the Sabatain Formation being absent), through the overlying suc­ cession into the Qishn sands beneath the thick and tight main Qishn carbonate that forms the regional seal (Mills, 1992). Good quality oil of 36° API gravity is found in the Kohlan reservoir and 30° API oil in the Qishn clastic reservoir(s) and substantial flow rates were tested. Individual zones in the Qishn clastic deposits have flowed in excess of 4875 barrels per day of 29°-33° API gravity with little dissolved gas and low sulphur (Mills, 1992).

Clastic reservoirs in Syria In the Euphrates graben basin of eastern Syria, some 25 oilfields discovered since 1984 are now on production and collectively were producing about 330000 barrels per day of 34°-38° API gravity low sulphur oil in early January of 1992 ( Middle East Economic Survey, 35: 16, 20 January 1992). Recov­ erable reserves are reported to be well over 1 B barrels. The reservoir is the so-called Rutbah Formation sandstone of Early Cretaceous age (Beydoun, , 1988). The Rutbah type section was defined originally across the border in Iraq, south of the Euphrates River, where a 23-m section of vari­ coloured, white and ferruginous unfossiliferous sand­ stone is exposed, with the basal part probably being of continental origin and the uppermost part marine; the age was given as Late Triassic to early Ceno­ manian based on the relationship with the dated overlying and underlying formations (Dunnington et a/., 1959). The correlation of the Rutbah with the Nahr Umr of southern Iraq as a single diachronous sandstone deposit was not regarded favourably (see, however, Ibrahim ( 1981) who recorrelates the two). The informal ' Rutbah Sandstone' of Syria was for­ malized as the Cherrife Shale Formation of Late Jurassic to Early Cretaceous age (Daniel, 1963a). The type section was the Cherrife- 1 exploration well in the Palmyrid basin, where the formation comprises predominantly red-brown fet:ruginous sandy shale and green pyritic shale with intercalations of quartiz­ itic and calcareous sandstone and streaks of lime­ stone. Elsewhere in Syria, the succession includes decomposed basalt and ash and is more varicoloured, more sandy and, in certain levels, clayey and lignitic. Daniel ( 1963a) considered the Cherrife Forma­ tion an extension of the Hathira Formation of Jordan and the 'Gres de Base' of Lebanon, which are Upper Jurassic to Lower Cretaceous and

347

basal Cretaceous respectively (Daniel, 1963b; Dubertret, 1963). The 'Rutbah' reservoir of the Euphrates graben is sourced by the Upper Cretaceous, organically rich Shiranish-Soukhne Formation(s) marls; these are phosphatic as a consequence of upwelling, and were widely deposited pelagic sediments that extend in outcrop from Iraq across the Levant and northern Arabia through North Africa all along the southern Tethyan margin of North Africa. Their organic maturity level is low except where deeply buried in rapidly subsiding basins, such as the Euphrates graben, whose inception coincided with the start of the pelagic sedimentation regime (Beydoun, 1991). Juxtaposition of maturing source sediment with older reservoir sandstone facies in tilted extensional regime fault blocks in the subsiding grabens resulted in both oil charging and sealing of these generally good 'Rutbah' sandstone reservoirs (Beydoun, 1988, 1991). The presence of volcanics associated with the Euphrates rifting episode and the initial high palaeo­ heat-flow have together created locally 'dirty' sands, causing production problems that have necessitated remedial pressurization measures in some fields. Clastic reservoirs in other Middle East countries Elsewhere within the Middle East region, minor oil production or potential production from Mesozoic sandstone reservoirs occur in Oman, offshore Qatar and in southern Israel-Palestine. In the Oman area, producible oil has been found in a Lower Jurassic sandstone reservoir in (several) fields of the Ghaba basin of central Oman (Alsharhan & Kendall, 1986). The main production of the Saih Raw! field comes from the Upper Palaeozoic Haushi Group sandstone reservoirs and, more recently, considerable volumes of gas were estabbshed in the Lower Palaeozoic Andam Formation sandstone reservoirs. In 1978, commercial oil was discovered in the 'Marrat' reservoir (Beydoun, 1988). Hughes Clarke (1988) redefined the 'Marrat' of Oman and renamed it the Mafraq Formation (only partly equivalent to the Marrat of Saudi Arabia), with the lower Mafraq being continental and locally possessing the charac­ teristics of a palaeosol, whereas the upper Mafraq is shallow marine with variable thickness but including important clean reservoir sand levels (e.g. at Saih Raw!) and is regarded as being essentially a basal clastic unit to the Middle-Upper Jurassic Dhruma Formation; thus the upper Mafraq is younger than the Marrat of Saudi Arabia. Reservoir charge is

348

Z.R. Beydoun

probably from a deeply buried Huqf Group source that also charges the underlying Haushi Group reservoirs and not from organically rich units within the so called 'Marrat', as suggested by Alsharhan & Kendall (1986), nor from the overlying Middle Jurassic Dhruma Formation shales that provide the seal (Beydoun, 1988). It is appropriate here to note that Beydoun (1988), in discussing the 'Marrat' accumulation at the Saih Raw! field in Oman (p. 149) , made the passing remark that deeper exploration in producing fields in Kuwait (p. 185) had established substantial new reserves in lower Jurassic Marrat (assumed as clastic sediments in consequence of the Saih Raw I discovery in Oman and the Saudi Arabian surface section for the Marrat) and other formations. In reality, the Marrat facies in Kuwait, as in the adjacent por­ tion of Saudi Arabia, is predominantly carbonate (Powers, 1968), changing from the clastic facies of the surface section to the west, and the new Kuwait Marrat reserves are housed in carbonate reservoir(s). The minor offshore Qatar Mesozoic clastic reser­ voir accumulation occurs in an areally limited struc­ tural closure over a portion of the northern flank of the 'supergiant' Permian Khuff carbonate gas accumulation of the areally huge but very gentle North (Dome) field. The reservoir is a thin Nahr Umr (Albian) fine-grained sandstone, locally cemented, and tight to the southeast of the closure but shaling out westwards (Wells, 1987). Thus, the thin and areally limited reservoir sand facies, shaley at the base and top, represents the distal clastic regional facies of the Nahr Umr at this distance from the main fluvio-deltaic depocentres of the Basra­ Kuwait area. The hydrocarbon source for this sand­ stone reservoir (and also for the overlying and underlying Cretaceous carbonate permeability pinch-out reservoir accumulations on this same north flank of the North field) is the Upper Jurassic Hanifa bituminous carbonates, as discussed by Murris (1980). The migration route was complex ; initially up the flank of the Qatar Arch under the Upper Jurassic Hith anhydrite seal whence most of the oil moved towards the then culmination in southern Iran, but some was retained in palaeostructures or stratigraphical traps. Later, during late Tertiary times, when the North (Dome) structure was formed, this oil remigrated southwards, where the Hith anhydrite seal had been breached allowing vertical access into the Cretaceous reservoirs of the area (Wells , 1987).

Clastic reservoirs in the Levant- Sinai region The Levant-Sinai fragment constituted an integral part of the Arabian Plate until the Neogene. In this region two areas of Mesozoic clastic reservoir oil production occur. The first is of minor significance and is located in the coastal region of southern Israel-Palestine. Here the only sustained oil pro­ duction is from the Heletz-Brur-Kokhav field complex, where the predominant reservoir facies is a Lower Cretaceous sandstone with some carbon­ ate interbeds belonging to the Lower Cretaceous Kurnub Group. The discovery was made in 1955 on re-entering and deepening a well abandoned in Aptian sediments in 1948 by the Iraq Petroleum Co. subsidiary Petroleum Development (Palestine) because of security conditions. The trap is a faulted anticline with a major NNE-SSW boundary fault to the west and a strong stratigraphical trap component consisting of channel sands and pinch-outs. Facies range from marine in the west to fluvio-continental in the east. The source of the oil is not firmly established but appears to be 'basinal' Middle Jurassic Barnea Formation carbonates located off­ shore to the west, with TOC values of 0.5-2.6% of type II kerogen; other source candidates are laterally equivalent basinal Lower Cretaceous shales filling canyons cut in the Late Jurassic ( Gilboa et al., 1990). The principal reservoir intervals are sands, but a Lower Cretaceous dolomite also produces. The sandstone reservoirs vary in characteristics and environmental facies and include three main types : an offshore (marine) environment calcite-cemented sandstone; a tidal channel and/or lagoonal environ­ ment sand, calcite or dolomite cemented; a coastal aeolian sand, very fine-medium grained, well sorted, loosely packed, low cement content and with porosities reaching 32% (average 24%) and good permeabilities at times exceeding 2000mD; most of the recoverable oil is in the coastal environment facies and the oil gravity is 27-31° API. However, reserves are modest with 38 MM barrels in place and 19 MM barrels being the estimated recoverable; with cumulative production until 1988 standing at 16.5 MM barrels the field must be close to exhaustion (Gilboa et al., 1990). By far the most important Mesozoic clastic reser-­ voir oil region of the Levant-Sinai fragment of the Middle East is the Gulf of Suez. There, at least 10 fields produce oil from multiple sandstone reservoirs of Cretaceous age, including Lower Cretaceous Nubia 'A' sands, Cenomanian sands , Turonian sands

Middle East clastic oil and gas reservoirs

and lower Senonian sands (Ayouty, 1990). The depositional environments charactizing the different Cretaceous reservoirs vary. In Late Jurassic to pre­ Aptian time, the Gulf of Suez region was positive with erosion or non-deposition taking place. By the Aptian, the northern extremity was within a delta­ front domain and during the Albian, the entire region was receiving alluvial deposition, with a flood­ plain environment belt straddling the Gulf and extending across the adjacent Eastern Desert of Egypt and Sinai. During the Cenomanian, the Gulf of Suez, much of Sinai and the northern part of the Eastern Desert were located in an open-marine regime, the Cenomanian reservoir sands thus being mainly shallow marine in facies. This situation con­ tinued in the Turonian but with the southern Gulf of Suez area alternating between shallow marine and fluvial -deltaic. By the Coniacian, the alternat­ ing shallow-marine and fluvio-deltaic facies had expanded to include the Egyptian Red Sea and all of the Eastern Desert while the northern part of the Gulf of Suez became uplifted as a ridge, with open­ marine conditions to its east and north. In the Santonian, the whole of the Gulf of Suez region was within the open-marine-embayment domain bounded by positive uplifts to the north and south and southeast. In the late Campanian the Gulf of Suez region became part of the phosphatic 'flint' facies or 'shale with bone beds' facies covering much of Egypt and extending into the Sinai-Levant region and across North Africa (Said, 1990). Porosities and permeabilities in these reser­ voirs are variable but generally good to fair, and oil gravities vary greatly in the different reservoirs , from 20° to 39° API (Ayouty, 1990). These oils, as with those in the Upper Palaeozoic Nubia 'C' sandstone reservoir, are all thought to be sourced principally from Lower Miocene sediments with Eo-Palaeocene and Upper Cretaceous sediments being probable additional sources (Ayouty, 1990). However, as stated in the previous section, there is an alternative view postulating the Carboniferous Nubia 'B' black shales as the source, at least for the Palaeozoic and Cretaceous reservoirs. Sourcing from the younger units would require fault or uncon­ formity juxtaposition of source against reservoir and involve only short-distance migration.

CENOZOIC CLASTIC RESERVOIRS Only two areas o f the Middle East have established

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Cenozoic commercial oil production from sandstone reservoirs. These are the Kuwait-Basra-Ahwaz (Iran) area at the head of the Gulf, and the Gulf of Suez rift of Egypt. Everywhere else, Cenozoic pro­ duction is from carbonate reservoirs, with the principal prolific producing reservoirs being the Oligo-Miocene Asmari Formation limestone of southwest Iran and its correlative of north Iraq. Towards the end of the Oligocene, the present head of the Gulf area was situated at about 25°N, a time coinciding with a global sea-level lowstand. Much of Arabia was positive and uplifted, with the exception of a narrow seaway along the line of the present Zagros range (Beydoun et al., 1992). A major river rising in the highlands of Central Arabia flowed northeast into the shallow shelf seas of the northern Gulf area, prograding a large delta in the Kuwait-Basra sector (Ghar Sandstone Formation); a pattern not unlike that initiating the deposits of the Cretaceous Zubair-Burgan-Wara Formations. From this delta, distal reservoir-quality sand tongues extended into the southwestern part of Khuzestan Province in Iran, to form the Ahwaz Sandstone Member of the Asmari (carbonate) Formation (Setudehnia, 1972; Beydoun et al . , 1992). The Ahwaz Member is considered stratigraphically equivalent to the Ghar Formation of Kuwait and southeast Iraq (Basra area) and is of imprecise Miocene age (Dunnington et al., 1959). During deposition of the lower part of the Asmari carbonates in Khuzestan, sands were being flushed from the west into this essentially carbonate shelf area; this situation continued until finally the carbonate facies of the Asmari prevailed (Setudehnia, 1972). The Ahwaz sands die out eastwards, but in a number of oil fields in southwest Khuzestan, such as Ahwaz, Mansuri and Marun, where the sands are of near­ shore, littoral -shoreline and aeolian environments, they constitute a dominant or significant part of the Asmari reservoir and are clean, loosely cemented, have porosities of 20-30% and permeabilities in the darcy range (Beydoun et al., 1992). The oil produced is of 30-33° API gravity and field reserves are sub­ stantial, with exceptional flow rates due to the excel­ lent reservoir properties. Clastic reservoirs in the Kuwait -Basra-Ahwaz area In the Kuwait-Basra area a few small shallow fields produce from the Miocene Ghar sandstone reservoir, but the oil is heavy and sulphurous, ranging between 18° and 20° API gravity. The fields with producible

350

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Ghar oil are Zubair near Basra in Iraq, and Bahra and Ratga in Kuwait, and the reservoir facies is fluvio-deltaic; the heavy nature of the oil is either because of biodegradation or of early maturity of source, but the source itself remains debatable (Beydoun, 1988). Clastic reservoirs in the Gulf of Suez In the Gulf of Suez, the principal productive sand­ stone reservoirs of the region are of Miocene age and contain the bulk of the reserves in many fields. The main oil-bearing levels are sandstones (although occasionally some limestones too) in the Nukhul (oldest), Rudeis and Karim Formations of the Gharandal Group (Lower to Middle Miocene), representing the main Neogene clastic cycle, and in the overlying Middle Miocene Belayin Formation, which marks the beginning of the main Miocene evaporite cycle represented by the Ras Malaab Group (Ayouty, 1990; Hughes et al., 1992). A num­ ber of thin sand reservoir levels are also present in the overlying Middle -Upper Miocene South Gharib Formation and the succeeding mainly Upper Miocene Zeit Formation (the youngest of the Ras Malaab Group). The main oil-bearing reservoir units occur in the lower four formations (Ayouty, 1990). The main reservoirs of the Morgan, Belayim Land and Marine fields, and July, Shoab Ali and Zeit Bay fields are in the Rudeis-Karim Formations. Belayim Formation oil-producing reservoirs are found in Belayim Land and Marine fields, Morgan and Shoab Ali fields, with less important production coming from the same reservoirs in several other fields. The Nukhul sandstone reservoir is oil productive in the Rudeis, Sidri, Shoab Ali and GS-173 fields, but some production is also obtained from contempor­ aneous carbonate reservoirs deposited as reefs or platform carbonates laid down on submerged palaeohighs. The Nukhul Formation is the basal marine Miocene unit and reflects an irregula_r reljef in the basin floor, leading to a diversity of sediments represented by bioclastic limestones, sandstones, or shales with sabkha-type anhydrites present in the central and southern parts of the Gulf of Suez. The Rudeis consists of highly variable lithology, thickness and depositional environments, controlled by syn­ depositional fault-block topography and represented by marls, limestones or shales with minor sands, willch, however, dominate locally at the expense of the shales (Hughes et al., 1992). The overlying

Beydoun

Karim Formation is similarly varied in facies but broadly consists of a lower, characteristically evap­ oritic member and an upper, fine-coarse-grained clastic member; the fine clastic sediments are shales and marls, and the evaporite member also contains intercalated shales and marls. The evaporites are absent locally owing to lateral facies changes, so that the formation becomes mainly shale-marl in com­ position, with sand levels that in marginal areas become the dominant clastic deposits and hence good reservoirs; environments range from shelf­ edge through the sublittoral (Hughes et al., 1992). The Belayim is typically evaporite and shale, but in marginal areas and laterally it can become pre­ dominantly sand, with carbonates developing on submerged palaeorelief; it is mostly of a sublittoral environment, generally shallow to deeper inner sub­ littoral with localized lagoonal conditions (Hughes et al., 1992). The South Gharib Formation is typically a thick evaporite body with intercalations of shale and fine to very fine-grained, well-sorted sandstones. The Zeit Formation represents a shallow-marine regime with alternating restricted and open con­ ditions (Hughes et at., 1992). Porosities and permeabilities in these formations are variable but generally good to fair. The reservoirs are sourced by oils generated in the Lower Miocene Rudeis and Kareem Formations, as in the case of oils in Mesozoic and older reservoirs (Ayouty, 1990), although, as discussed earlier, additional input from older sources is not ruled out. The northernmost Red Sea has an identical Miocene succession to that of the Gulf of Suez (as also the remainder of the Red Sea with recognizable Gulf of Suez units; Hughes & Beydoun, 1992). So far, however, only one hydrocarbon field, Barqan, has been discovered in the northernmost sector of the Arabian side in Saudi Arabia, where light oil, gas and condensate were tested in four sand reser­ voirs, in Rudeis-Kareem and Belayim Formation equivalents, but the field remains undeveloped despite promising flow rates (Beydoun, 1988).

CONCLUSIONS The bulk o f oil and gas production i n the Middle East region comes from Permian to Miocene reser­ voirs, the overwhelming proportion of which are carbonates that were deposited when the region lay mainly in warm low latitudes where marine carbon­ ate sedimentation dominated; sandstone sedimen-

Middle East clastic oil and gas reservoirs

tation during that time was subordinate and was related to particular climatic or tectonic conditions. Arid or warmer, wet climatic episodes coupled with local or regional uplifts and/or sea-level lowstands resulted in substantial localized aeolian or fluvio­ deltaic deposition. Tectonic phases allowed detritus to inundate local, rapidly subsiding rift basins; clastic detritus from developing and rising collisonal ­ compressional mountain belts along the northern and eastern peripheries rapidly infilled marginal seaways. During the pre- Late Permian Palaeozoic, by contrast, the Middle East region was located in the southern hemisphere in middle to high temperate (humid-cool) latitudes with only occasional influ­ ence from warmer belts , hence the dominance of clastic sedimentation in general and of sandstone facies in particula�;. The distribution of favourable (hydrocarbon) reservoir facies by itself is insufficient for the dis­ covery of commercial petroleum fields unless widely distributed organic-rich source facies are present to generate the hydrocarbons. Retention of hydro­ carbons in producible reservoirs also requires impervious seal facies that are suitably formed geometrically (by tectonics or stratigraphy) into a trap juxtaposed to the source. Moreover, the dis­ covery of these accumulations commences with the most easily accessible , invariably the shallowest (gen­ erally also the youngest) reservoirs. In time, older and deeper accumulations are explored for as the shallower ones are exhausted or become insufficient to meet rising demand. In the highly petroleum-rich Middle East region, the need to explore deeper than the Mesozoic reservoirs was for many years neglected because of the huge reserves established in the Mesozoic and Tertiary (predominantly carbonate) reservoirs. It is only in recent years that attention has turned to the (deeper) Palaeozoic sequences, where sandstone reservoirs predominate and recent successes in Oman and in Central Arabia have opened up a vast new exploration domain for the whole region. Exploration drilling results in Oman and in Cen­ tral Arabia have shown that the major new reserves of hydrocarbons housed in Cambro-Ordovician and Permo-Carboniferous sandstone reservoirs are sourced from two main regionally distributed source­ rock sequences of Infra-Cambrian/Early Cambrian and Early Silurian age. The former are thought to be present over much of the Gulf region and over Central Arabia and the

351

Rub al Khali basin, but in the Gulf region they are now too deeply buried and have become over mature; but the hydrocarbons they had generated in earlier geological time have probably charged Lower Palaeozoic reservoirs lying up-dip and awaiting dis­ covery. Elsewhere , exploration for these Infra­ Cambrian/Early Cambrian sourced hydrocarbons outside Oman has been minimal and their potential remains to be assessed by drilling, especially in the Rub at Khali basin. The Lower Silurian source sediments were depo­ sited regionally in response to global sea-level rise and they have been geochemically shown to have sourced the sandstone-hosted oils of central Arabia and, locally, of western Oman. Where they lie deeply buried under the Gulf, the consensus of specialist opinion is that they are responsible for the vast gas reserves housed in the Upper Permian Khuff car­ bonated reservoirs of the region (Beydoun, 1991). Post-Silurian sandstone reservoirs , located up-dip from the kitchen areas where these Lower Silurian shales are preserved , are highly prospective in an arc around the Arabian Shield, from the Rub al Khali to southwest Iraq and Jordan passing by central and northwest Saudi Arabia; similar areas in southeast Turkey and Syria up-dip from similar Lower Silurian kitchen areas are equally prospective. Many Palaeo­ zoic basins became inverted in the Late Palaeozoic and some of the sedimentary section was stripped off, so that better control of the palaeogeographical subcrop distributions of these source rocks is first required before exploratory drilling successes can follow routinely. The successful search for additional Mesozoic or younger sandstone reservoirs in the region depends mainly on the identification of 'disguised' rift basins, where source rocks have been buried to maturation and thus charge suitably juxtaposed sandstone reser­ voirs, as in the cases of Yemen and the Euphrates graben in Syria. The Jordan and other grabens merit more exploration with these models in mind. Unlike the wide areal continuity exhibited by Mesozoic-Cenozoic carbonate reservoir facies and evaporitic regional seal facies, the Palaeozoic clastic deposits generally show rapid lateral and vertical changes in continuity of the reservoir-seal couplets (although exceptions have been demonstrated in Oman and elsewhere); hence major vertical migration of hydrocarbons , either to dissipation at the surface or until stopped by younger more effec­ tive seals and diverted to accumulate in integral traps , can be expected. This factor will, in general,

Z. R. Beydoun

352

affect the size of many Palaeozoic accumulations, because the traps at this level will be subject to fault control and thin seals can be breached often. Nevertheless, the total hydrocarbons generated by regionally distributed source rocks are believed to be enormous and the exploration potential rewards are believed to be huge potentially.

ACKNOWLEDGEMENTS The author acknowledges with gratitude the typing and draughting support provided by the Exploration Department of Marathon International Petroleum (GB) Ltd in London during the preparation of this paper, and particularly to Mrs Wendy Milligan for drawing the figures and Miss Susan Hambrook for typing the manuscript. To Dr M.W. Hughes Clarke, sincere thanks are expressed for his continuing role as the authors' active 'sounding board' and for critically reading and discussing the contents of the paper and providing many constructive and helpful ideas and suggestions. Thanks are also expressed to Dr R. Crossley for reviewing the paper and for providing a number of helpful comments and suggestions. Gratitude is also expressed to Dr A.P. Heward for providing new insight into the Early Palaeozoic of Oman which has greatly improved that section of the paper, and for his other helpful remarks and suggestions. REFERENCES Asu-Au, M.A., FRANZ, U.A., SHEN , J . , MONNIER, F., MAHMOUD, M . D . & CHAMBERS, T.M. (1991) Hydrocar­ bon generation and migration in the Palaeozoic sequence of Saudi Arabia. Proceedings, 7th SPE Middle East Oil Technical Conference, November 1991, Bahrain, SPE 21376, pp. 345-356. Au, A. R. & SILDAWI, S.J. (1989) Hydrocarbon potential of the Palaeozoic pre-Khuff clastics in Abu Dhabi, U.A.E. Proceedings, 6th SPE Middle East Oil Technical Conference, March 1988, Bahrain , SPE 18009, pp. 819832. ALSHARHAN , A.S. & KENDALL, e.G. St C. (1986) Pre­ cambrian to Jurassic rocks of Arabian Gulf and adjacent areas: their facies, depositional setting and hydrocarbon habitat. Bull. Am. Assoc. petrol. Geol. , 70, 977 -1002. AYOUTY, M.K. EL- (1990) Petroleum geology. In: The Geology of Egypt (Ed. Said, R.) pp. 567-722. Balkema, Rotterdam. BEYDOUN, Z.R. (1964) The stratigraphy and structure of the eastern Aden Protectorate. Overseas Geology and Mineral Resources Bulletin , Supplement 5, HMSO, London, 107 pp.

Z.R. (1966) Eastern Aden Protectorate and part of Dhufar. In: Geology of the Arabian Peninsula. U.S. Geol. Surv. Prof. Pap., 560-H, 49 pp. BEYDOUN, Z. R. (1988) The Middle East: Regional Geology and Petroleum Resources . Scientific Press, Beaconsfield, 292 pp. BEYDOUN, Z. R. (1991) Arabian Plate hydrocarbon geology and potential - a plate tectonic approach. Am. Assoc. petrol geol. Stud. Geol. , 33, 77 pp. BEYDOUN, Z . R. & GREENWOOD, J.E.G.W. (1968) Aden Protectorate and Dhufar. In: Lexique Stratigraphique International, Vol. III, Asie, Protectorat d'Aden et Dhufar (Ed. Dubertret, L.). CNRS Paris, fasc. 10b2, 128 pp. BEYDOUN, Z.R., HUGHES CLARKE, M.W. & STONELEY, R. (1992) Petroleum in the Zagros Basin: a late Tertiary Foreland Basin overprinted onto the outer edge of a vast hydrocarbon-rich Palaeozoic-Mesozoic passive-margin shelf. In: Foreland Basins and Foldbelts (Eds Macqueen, R.M. & Leckie, D.A.), Mem. Am. Assoc. petrol. Geol ., Tulsa, 55, 309-339. BRENNAN, P. (1990) Greater Burgan field. In: Treatise of Petroleum Geology, Atlas of Oil and Gas Fields; Struc­ tural Traps I, Tectonic Fold Traps (Compilers Beaumont, E.A. & Foster, N . H.), pp. 103-128. American Associ­ ation of Petroleum Geologists, Tulsa. BRENNAN, P. (1991) Raudhatain field - Kuwait, Arabian Basin. In: Treatise of Petroleum Geology Atlas of Oil and Gas Fields; Structural Traps V, ( Compilers Foster, N . H. & Beaumont, E.A.), pp. 187-210. American Association of Petroleum Geologists, Tulsa. DANIEL, E.J. (1963a) Syrie Interieure. In: Lexique Stra­ tigraphique International, Vol. III, Asie, Liban, Syrie, Jordanie (Ed. Dubertret, L.), fasc. 10 cl, pp. 157-291. CNRS, Paris. DANIEL, E.J. (1963b) Jordanie (Ensemble). In: Lexique Stratigraphique International, Vol. III, Asie, Liban, Syrie, Jordanie (Ed. Dubertret, L.) fasc. 10 cl , pp. 293-399. CNRS, Paris. DERCOURT, J . , ZoNENSHAJN, C.P., Rtcou, L.-E., KAZMtN, V.G., LE PICHO N , X., KNIPPER, A.L., et al. (1986) Geologic evolution of the Tethys belt from the Atlantic to the Pamirs since the Lias. Tectonophysics, 1221, 241-315. DuBERTRET, L. (1963) Liban, Syrie. In: Lexique Stra­ tigraphique International, Vol. III, Asie, Liban, Syrie, Jordanie (Ed. Dubertret, L.) fasc. 10 cl, pp. 7-155, CNRS, Paris. D u N N INGTON , H.V., VAN BELLEN, R.C., WETZEL, R. & MoRTO N , D.M. (1959) Iraq. In: Lexique Stratigraphic International, Vol. III, Asie, Iraq (Ed. Dubertret, L.). CNRS, Paris, fasc. lOa, 333 pp. FERGUSON , G.S. & CHAMBERS, T.M. (1991) Subsurface stratigraphy, depositional history and reservoir devel­ opment of the early to late Premian Unayzah Formation in central Saudi Arabia. Proceedings, 7th SPE Middle East Oil Technical Conference, November 1991, Bahrain, SPE 21394, pp. 487-496. GILBOA, Y. , FLIGELMAN , H. & DERJN, B. (1990) Heletz-­ Brur-Kokhav field - Israel. In: Treatise of Petroleum Geology, Atlas of Oil and Gas Fields, Structural Traps IV, Tectonic and non-Tectonic Fold Traps (Compilers Beaumont, E.A. Foster, N.H.) pp. 319-345. American

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ScOTESE,

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Spec.

Pubis int. Ass. Sediment. (1995) 22, 355-380

The evolution of Oligo-Miocene fluvial sand-body geometries and the effect on hydrocarbon trapping: Widuri field, west Java Sea

R AY Y O U N G ,* W. E. H A RM O NYt

and

T H OM A S B U D I Y E N T Ot

*Maxus Ecuador Inc. , PO Box 650496, Dallas, Texas 75265-0496, USA; and tMaxus Southeast Sumatra Inc., Five Pillars Office Park, ]alan Letjen MT Haryono 58, PO Box 2759, Jakarta, Indonesia

ABS T R A C T

Widuri field is a faulted anticline located in a Tertiary back-arc basin at the western end o f the Java Sea in the Indonesian archipelago. The discovery well was drilled in early 1988 to a depth of 3735 ft subsea (1139 m SS) and penetrated 170ft (52 m) of net oil pay in late Oligocene and Early Miocene sandstones. Eight delineation wells, 47 development wells and five sidetracks confirmed a large oil field over an area of approximately 7500 acres. The oil is trapped in six separate reservoirs over a vertical interval of 325ft (99 m). The lowermost reservoir is a thick, immature, fluvial sandstone with multistorey and multilateral stacking, high permeabilities and a sheet geometry; the uppermost reservoir is a series of thin, lower permeability, 2000 ft (610 m) wide, distributary channel sandstones associated with a tide-influenced delta. The intermediate reservoirs show a complete gradation in size, geometry and reservoir quality between these two extremes. This gradual change from an immature fluvial system, through progressively more mature fluvial systems into deltaic sediments at the top is the result of a gradual relative rise in sea-level. The systematic upwards decrease in sand-body width is accompanied by a change in the trapping mechanism, from structural in the two sheet-like reservoirs at the base to stratigraphical-structural in the upper four reservoirs, where the existing structural closure is enhanced significantly through each reservoir having a restricted lateral extent and a favourable orientation with respect to structure. However, the Widuri reservoirs share a common oil-water contact defined by the spill point of the stratigraphically highest of the sheet-like reservoirs. The reservoirs must therefore be in communication and, without a stratigraphical component to trapping in every one of the overlying four reservoir intervals, the field oil-water contact would be significantly higher and the field size very much smaller. The understanding of sand body geometry and trapping has led to efficient field development (only two of the original holes needed to be side-tracked owing to lack of pay) and will form the basis for future reservoir management. From an exploration viewpoint , the Widuri work has provided analogue models in the search for other stratigraphically trapped oil in the contract area .

I N T R O DU C T I O N

500 million barrels of oil in late 1989 but virtually all of this oil came from the Sunda basin. It was not until late 1987 that the first oil was discovered in the Asri basin and this discovery subsequently led to the establishment of five new fields on the northwest flank of the basin (Fig. 2). Widuri field is by far the largest of these accumu-

Widuri field is located in the Southeast Sumatra Production Sharing Contract (PSC) area approxi­ mately 100 miles ( 161 km) north of Jakarta in the west Java Sea (Fig. 1). The PSC comprises approxi­ mately 5300 square miles ( 13 725 km2) and consists of two principal areas, the Sunda and Asri basins. Cumulative oil production from the PSC surpassed

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0

355

35 6

R. Young, W.E. Harmony and T. Budiyento

SOUTH CHINA SEA Natuna

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0

500 Miles

Fig. 1. Location of the Pertamina-Maxus , Production Sharing contract area showing the relative position of the Sunda and Asri basins and the distribution of oil fields.

lations and was discovered in early 1988 when Widuri-1 was drilled to a total depth of 3735 ft subsea (1139m SS) and penetrated 1 70ft (52m) of net oil pay in multiple sandstones of late Oligocene and Early Miocene age. Eight delineation wells, 47 development wells and five sidetracks (as of mid1992) confirmed a giant oil field with over 500 million barrels of original oil in place spread over approxi­ mately 7500 acres. Recovery efficiency is expected t9 be excellent and the discovery was the largest in Indonesia since 1974. The oil is a high pour point, low sulphur crude with a gravity of 31 API. First oil was produced in December 1990 and there are currently five platforms (A-E, Fig. 3). To optimize reservoir management it was decided to complete wells in single zones only and this strategy meant that a detailed and predictive geological model was vital. The primary aim of this paper is to show the o

relationship between reservoir geometry and struc­ ture at Widuri field and to demonstrate the resulting hydrocarbon trapping configurations.

BA C K G R O U N D

G E O L O GY

The PSC area is situated in a Tertiary back-arc basin and the history of tectonics in the region is therefore essentially extensional but with some wrench modi­ fication (see also Hutchison, 198 6). Structurally, the Asri basin is an asymmetrical syncline or half-graben with a roughly N-S axial trend (Fig. 2) and a maximum sedimentary fill of approximately 15 OOO ft (4773 m) of Oligocene and younger sediments. The oil accumulations are trapped in upthrown closures related to NE-SW trending, down-to-the-northwest, normal faults on the gently sloping northwest flank of the basin. These faults probably originated in

Fluvial sand-body geometries

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Cl 50 ms 10 Kms 10 Miles Fig. 2. Asri basin time-structure map on the top Talang

Akar Formation.

Fig. 3. Depth structure map on the top 34-2 reservoir showing (1) that Widuri field is a low-relief, faulted closure, (2) the location of the five platforms and 62 wells in the Widuri area and (3) the location of seismic line 88-196 (Fig. 4).

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357

early Oligocene times during basin formation but were subsequently reactivated at several times throughout the Tertiary. Widuri field is a low relief, three-way-dip, faulted closure (Figs 3 & 4), which formed approximately 19Ma (Early Miocene) and has not been altered significantly since. The stratigraphy in the Widuri field area is shown in Fig. 5. Lower Cretaceous, metamorphic basement was penetrated only over the A-platform area, which is the structurally highest part of the field. Basement is onlapped and overlain by upper Oligocene fluvial sandstones, shales and coals of the Zelda Member, Talang Akar Formation. The Lower Miocene Gita Member is a fluvio-deltaic succession of sandstones and shales with thin interbedded limestones and coals. The boundary between the Zelda and Gita members is the top of the Oligocene succession and is an unconformity over most of the Asri basin. At Widuri field, however, the boundary is probably a discontinuity that is associated with an important field-wide coal marker - Coal 'A' (see Figs 4 & 5). The Talang Akar sandstones are currently the only commercial reservoirs in the Asri basin. The top of the Talang Akar Formation is an unconformity that is generally marked by a relatively high gamma-ray (uranium) reading (Fig. 6). The overlying Batu Raja Formation has yielded non­ commercial oil shows from thin argillaceous lime­ stones and sandstones, but the predominant lithology is mudstone which acts as a local seal. The regional

358

R. Young, W.E. Harmony and T. Budiyento

NNW

Widuri Field (approx. 3 miles)

SSE TWT 0.7

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Fig. 4. Seismic line 88-196 over Widuri field (see Fig. 3 for location).

seal for both the Sunda and Asri basin is claystone of the overlying Gumai Formation (Bushnell & Temansja, 1986). The remainder of the succession consists of Miocene, shallow-marine shales, lime­ stones and sandstones, which are overlain by pre­ dominantly alluvial sediments and volcanics of the Plio-Pleistocene. An adequate source rock for the oil accumulation has not yet been penetrated in the Asri basin. However, by analogy with the Sunda basin, it is believed that lower Oligocene lacustrine shales are present in the central part of the basin and that these shales generated oil from Early Miocene until at least Late Miocene time.

R E S E RV O I R

OVE RVI E W

Figure 6 illustrates the full Talang Akar succession in the Widuri A-1 well, one of the few wells in the field that penetrated each of the main reservoirs.

The oil is trapped in six different sandstone units (from base to top the 36-1, 35-2, 35-1, 34-2, 34-1 and 33-series sandstones) over a vertical interval of approximately 325 ft (99 m). The reser­ voirs almost certainly share a common oil-water contact at 3680ftSS (1122 m SS - see below), but there is no evidence for direct, sand-on-sand com­ munication between any two reservoirs, nor any prediction that this might occur. However, the reservoirs must either be in communication or have been in communication at some time in the past through faulting, fractures or seepage. So far, however, the production history indicates that each reservoir is acting as a separate sealed unit. Reservoir descriptions are based on over 650ft (198 m) of core from seven wells. The sandstones vary from coarse- to very fine-grained and both conglomerate and granule layers are common . Compositionally, the sandstones are mature, with 75% being quartz arenites and the remainder split between sublitharenites and subarkoses. There does

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not appear to be any significant variation in detrital composition between the different reservoir units or between the different sedimentary facies. The maturity of the framework mineralogy probably reflects the provenance (see Folk, 1968) and this is supported by the preponderance of metasedimentary basement in the Sunda and Asri basins (see also Ben-Avraham & Emery, 1973). Clay minerals dominate the non-framework mineralogy but usually only account for less than 5% of the bulk rock volume in the majority of the reservoir quality sandstones. Kaolinite predominates with subordinate illite and chlorite, and minor smectite and illite-smectite mixed-layer clay. All

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other mineral cements are extremely sparse and the sandstones are generally unconsolidated. Reservoir quality is usually excellent but the measurement of porosity and permeability from core data has inherent errors because the reservoirs are uncon­ solidated. However, log porosities average 29% (using a 15% cut-off) and permeabilities from production tests on initial completions range up to 30 darcies (D). A reconstruction of the palaeogeography of the area during Talang Akar times is shown in Fig. 7 (see also Beddoes, 1981). The basin margin is defined seismically by the onlap of the Talang Akar Formation on to basement (Fig. 7) and the sediment

360

R. Young, W.E. Harmony and T. Budiyento

Fig. 6. The type well, Widuri A-1.

transport directions were then defined using interval isopachs (see also Wight et a!., 1986) and individual channel isopachs. The Talang Akar clastic sediments in the Widuri area of the Asri basin were derived exclusively from a northerly direction.

RES ERVOIR

GEOMETRY AND

TRAP

A discussion o f the reservoir character, geometry and trap for each of the six main reservoir units follows. These descriptions and interpretations are based on a data set that includes a 250-m grid of very . good quality two-dimensional seismic data, a mostly full set of log data from approximately 60 wells, and a total of over 650ft (198 m) of full-hole core from seven wells. A detailed discussion of the sand-body geometries and reservoir model was the subject of a previous publication (Young et a!., 1991) and it is not the intent of the present article to duplicate this. However, reservoir geometry is so critical to the

hydrocarbon trapping that an updated and abbrevi­ ated version is included. 36-1 sandstone

The 36-1 sandstone is the lowermost of the Widuri reservoirs and is oil-bearing only over the structurally higher A- and D-platform areas. Few down-dip wells penetrated the entire 36-1 sandstone and therefore less information is available compared with other reservoirs. However, the reservoir was fully cored in the discovery well, Widuri-1, where it is 43 ft (13 m) thick and rests directly on top of metamorphic base­ ment (Fig. 8). The base of the sandstone is erosive and the top has a sharp contact with the overlying rootleted siltstone and coal. Grain size is coarse to very coarse and sedimentary structures are domi­ nated by trough cross-bedding in 1-2-ft (0.3-0.6 m ) sets. The sandstone does not show any overall grain­ size trend but at least five sedimentary units can be recognized based largely on grain size (Fig. 8). Each

Fluvial sand-body geometries

361

Java Sea

r.::l �

Sediment source areas/Sunda Shield emergent (Talang Akar FM absent) Sediment transport directions Asri Basin oil fields

Fig. 7. Local palaeogeography

during Talang Akar Formation times.

unit fines upwards slightly and they are interpreted to represent a multistorey fluvial system. Elsewhere in the field, sidewall cores indicate a more variable grain-size with a slightly finer average, and the thickness reaches a maximum of 185 ft (56 m) in Widuri E-1 (see Fig. 8). Log character is commonly blocky but occasionally, as in the case of Widuri-6 (Fig. 8), it is clear that the 36-1 is made up of several different sandstones. However, this distinction between individual sandstones cannot usually be traced with any confidence beyond adjacent wells and the thick, blocky nature of the 36-1 over most of the field is almost certainly the result of both vertical and lateral channel amalgamation. The 36-1 sandstone overlies and onlaps basement and the net sandstone isopach (Fig. 9A) defines a relatively thick, sheet sandstone, which is more extensive than the available well control in all areas except the extreme southwest of the field. The sand­ stone is interpreted to be a coarse grained, immature, multistorey and multilateral fluvial sheet-sandstone deposited in a high gradient, high bedload river. Reservoir quality is excellent; porosities average 28% and production test permeabilities range from 20 to 30D.

Figure 9B is a depth structure map on top of the 36-1 sandstone. The structural spill point is located approximately 5000ft ( 1525 m) north-northwest of the D platform at a depth of 3740ftSS (1140m SS) and the maximum structural closure is shaded. The oil- water contact for this reservoir is at 3680 ft SS ( 1122 m SS) and this plots inside the closing contour. The trapping mechanism for the 36-1 reservoir is, therefore, clearly structural but the structure is only partially filled. In other words, there is more struc­ tural closure than oil column. 35-2 sandstone

The 35-2 sandstone is separated from the 36-1 by an average of 20ft (6 m) of mudstone, siltstone and coal. The reservoir has been cored in three wells but long sections of core that are representative of the 35-2 sandstone field-wide are not available. The sandstone, however, is similar to the 36-1, but with some important differences. Grain size is more vari­ able and slightly finer (field-wide average is medium to coarse grained) but the sedimentary structures are still dominated by trough cross-beds in sets up to 2 ft (0.6 m) thick. Bed contacts are still abrupt and

362

R. Young, W. E. Harmony and T. Budiyento A

6000'

11 000' Widuri E-1 20

API

Widuri-7 120

20

API

5000'

10 000'

Widuri-4 120

20

120

API

A'

Widuri-1 20

API

Widuri-6 120

20

API

120

0

50

100

FT

150 Basement of weathered basement 200

GAMMA RAY

0

API

OEP TH (FT)

100

Core log 36-1 sandstone Widuri-1 Rootleted siltstone and coal

LOCATION MAP

Trough cross-bedded, pebbly, coarse to very coarse grained sandstone in at least five different sedimentary units

Metamorphic basement and weathered basement

Fig. 8. 36-1 reservoir: vertical and lateral sedimentary character.

evidence for multistorey and multilateral stacking of sand bodies is common (see Young et at., 1991); in one instance, an individual channel sandstone can be traced for up to 3000ft (915 m) along depositional

dip using log character correlation. The most sig­ nificant difference to the 36-1, however, is the shape of the sand body. The net sandstone isopach (Fig. lOA) defines a sandstone that is thinner and

363

Fluvial sand-body geometries

+

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A

Fig. 9. 36-1 reservoir. (A) Net

sandstone isopach with the oil­ water contact marked. (B) Depth structure map on the top 36-1 sandstone showing that the oil­ water contact is 60ft above structural closure (shaded). The structural spill point is marked by a triangle.

less extensive. Net sandstone thickness from well penetrations varies from 0 to 67ft (0-20m) and a sand body is defined (or predicted from thickness trends) that is elongated roughly N-S and has an average width of approximately 11 000 ft (3354 m) with a zero-edge to the west and an inferred one to the east. The abrupt change to a westwards depo­ sitional strike at the north end of the sand body is defined by well penetrations to the north and west outside of Widuri field.

Cl100' 5000 FT

a

The 35-2 sandstone is interpreted to be a multi­ storey and multilateral fluvial system that is less extensive, slightly finer grained, and therefore probably slightly more mature than the 36-1 system. However, the sandstone was still probably deposited in a relatively high gradient, high bedload river. Reservoir quality is excellent, with log porosity averaging 29% and test permeabilities ranging from 20-30D. Figure lOB is the depth structure map on top of

364

R. Young, W.E. Harmony and T. Budiyento

+

+

+

+

+

+

+

+

+

+

+

Cl 20' 5000 FT A

Cl100' 5000 FT B

the 35-2 sandstone. The spill point has -remained in approximately the same position but is now at 3680 ft SS ( 1122 m SS) and this is coincident with the oil-water contact for this reservoir. The trapping mechanism is therefore again clearly structural, but this time the structure is full to spill. 35-1 sandstone

The Widuri E-1 core (Fig. 11) is lacking the basal

Fig. 10. 35-2 reservoir. (A) Net sandstone isopach with the oil­ water contact marked . (B) Depth structure map on the top 35-2 sandstone showing that the structural closure (shaded) is coincident with the oil-water contact. The structural spill point is marked by a triangle.

6 ft (1. 8 m ) of the 35-1 sandstone but adequately demonstrates the sedimentary character of the reservoir. The 35- 1 is an erosively based sandstone, which is trough cross-bedded in 1-2-ft (0. 3-0. 6 m) sets, and which averages medium grain size. Intervals of coarse grain size are common towards the base of the sandstone and fine/very fine grain size common towards the top, but the sandstone does not fine upwards systematically and several distinct units with abrupt contacts are evident (see Fig. 11). This

Fluvial sand-body geometries

365

A

A' D-8

1800'

20 API 120

�

20

D-6 API

D-5

3000' 120

L.-.-----1

20

API 120

L.-.-----1

B-3

3150' 20

4250'

API 120

A-1

4400' 20

20 API 120

.______.

B-6

1------4

API

C-3

3900' 120

�

20

API 120

�

CORE LOG 35-1 sandstone Widuri E-1 LOCATION MAP

Rootleted shale and coal

3750

3760

Predominantly medium to coarse grained,. trough cross-bedded sandstone Abundant mudstone clasts and carbonaceous debris

Fig. 11. 35-1 reservoir : vertical and lateral sedimentary character.

multistorey character is readily recognizable in all other 35-1 cores and the units are separated occasionally by coals (see Widuri A-1, Fig. 11). The 35-1 sandstone is the most variable of the Widuri reservoirs (see cross-section A-A', Fig. 11). The log character is predominantly gamma-ray increasing upwards (fining upwards) with a gra­ dational top, but is also commonly blocky with a sharp top. The most striking feature of the sand­ stone, however, is the rapid lateral thickness changes within an overall range of 0-64 ft (0-19.5 m).

Fortunately, these thickness vanatwns could be mapped and predicted seismically prior to any development drilling. The 35-1 interval is extremely well constrained between two readily correlatable markers - Coal 'A' at the top and usually a double coal at the base (see cross-section A-A', Fig. 11). Acoustic impedance contrasts associated with these coal markers are responsible for two seismic reflec­ tion which envelope the 35-1 sandstone and can be traced field-wide and beyond. The isochron between these two seismic markers, when linked with

366

R. Young, W.E. Harmony and T. Budiyento

well penetrations, proved to be a very effective predictor of net sandstone distribution (see Young et al., 1991). The net sandstone isopach (Fig. 12A) demon­ strates an anastomosing pattern and the 35-1 sandstone is interpreted to be a multistorey and multilateral fluvial system, possibly representing anastomosing rivers. Individual channels can be mapped with more confidence and over a larger area than those of the 35-2 and 36-1 (e.g. the sand body

trending southwest from the D platform is a single­ storey channel), and sand body widths vary from approximately 3000ft (915 m) in the west to 6500ft (1982 m) in the east. Finer grain size, thinner and narrower sand bodies, and more order to the sequence suggest a more mature fluvial system than the 35-2. Reservoir quality is again excellent, with log porosity averaging 29% and test permeabilities ranging from 10 to 20 D. The depth structure map on top of the 35-1 sand-

+

+

Cl100' 5000 FT

a

Fig. 12. 35-1 reservoir. (A) Net sandstone isopach with the oil­ water contact marked. (B) Depth structure map on the top 35-1 sandstone showing that the oil­ water contact is 20ft below structural closure (shaded).

Fluvial sand-body geometries

stone (Fig. 12B) defines a structural spill point or closing contour at 3660ft SS (1116 m SS). The oil-water contact for this reservoir, however, is the same as for the 36-1 and 35-2 at 3680 ft SS ( 1122 m SS). There is therefore 20ft (6 m) more oil column than can be accounted for by structural closure. Returning to Fig. 12A, it is clear that the structural contour of the oil-water contact does not close to the northwest but that closure is effected by the sand body pinch-out to the northwest. There is therefore a stratigraphical component to trapping and a combination stratigraphical-structural trap results. 34-2 sandstone

The 34-2 sandstone rests directly on top of Coal 'A' and the sedimentary character is demonstrated in core from the B-8 well (Fig. 13). The base is erosive on top of Coal 'A' and the thin, very coarse grained, basal lag contains coal clasts. The overlying sand­ stone consists of predominantly medium- to coarse­ grained sandstone in the basal and middle section before passing into fine- and very fine-grained sand­ stones in the upper part, and finally into rootleted siltstone and coal at the top. Trough cross-bedding in 1 ft (0. 3 m) sets is the dominant sedimentary struc­ ture. Maximum penetrated thickness for the 34-2 sandstone is 54ft (16.5 m) and in the thicker sections the log character tends to be blocky, with a gamma­ ray increasing (fining upwards) signature at the very top (see Fig. 13). In the thinner sections, however, the log character is mostly gamma-ray increasing (fining upwards). Immediately above the top coal is the first indication of fully marine conditions in the Talang Akar Formation. Calcareous mudstone grades up into muddy limestone containing skeletal debris of bivalves, large foraminifera, solitary corals and echinoid fragments. This limestone can be recognized throughout most of the field. The 34-2 net sandstone isopach (Fig. 14A) out­ lines a sinuous sand body with an average width of 4000ft ( 1220m). The shape and character suggest that the sand body is fluvial but the exact type of fluvial system is less clear. Several distinct sedi­ mentary units can be recognized within the 34-2 at B-8 (Fig. 13), but these are not as discrete as those of previous sandstones. Similarly, evidence from other full-hole core, dense sidewall cores in most of the other wells, and log character analysis suggests a broad systematic fining within a single-storey system (see also cross-section A- A', Fig. 13). This implies

36 7

a relatively mature fluvial system but a meandering river model (Allen, 1965) is discounted on the lack of evidence for lateral accretion and on the geometry of the reservoir which is not, so much a sheet as a ribbon of sandstone that tapers towards both sides (see Fig. 14A). Similarly, equations relating sand body thickness to channel depth and channel belt/ sand body width (see Leeder, 1973; Collinson, 1978; Lorenz et al., 1985; Fielding & Crane, 198 7), although having a wide margin of error, predict a much wider sand body if the 34-2 were a high sinuosity channel deposit. It is tempting to use a Mahakam River (East Kalimantan, island of Borneo) type model (Allen et al., 1979; Young & Atkinson, 1993) as the analogy for the 34-2 reservoir. In this case, the river channel does not meander significantly but does have a meandering thalweg. On the inside of every thalweg meander loop is a non-emergent lateral bar which is attached to the channel side and progrades down­ stream. Echo-sounding and dredge sampling suggest that the Mahakam River immediately above the delta has a width, sinuosity, maximum sand thick­ ness, grain size and bedform distribution that is remarkably similar to the 34-2 sandstone. However, irrespective of the exact mechanism of channel fill, the 34-2 reservoir is interpreted to be a single storey, fluvial sandstone. Proximity to the sea is suggested by the maturity of the fluvial system and by the overlying transgressive limestone. However, evi­ dence for marine conditions within the sandstone is absent. Reservoir quality is excellent, with average log porosity of 29% and test permeabilities of 10-30D. The depth structure map on top of the 34-2 sand­ stone is shown in Fig. 14B and the structural spill point/closing contour is at a depth of 3630ft SS (1107 m SS). The oil-water contact for the reservoir is still 3680 ft SS (1122 m SS) and there is again a discrepancy (50ft or 15m) between oil column and structural closure. It is clear on Fig. 14A, however, that the trapping mechanism is the same as for the 35-1. The structural contour of the oil-water contact does not close to the northwest but the sand body pinches out in that direction. The trapping mechan­ ism again combines stratigraphy and structure. 34-1 sandstone

The base of the 34-1 sandstone lies just above the limestone at the top of the 34-2 sequence and all but the top few feet of the sandstone was cored in the

368 A 20

R. Young, W.E. Harmony and T. Budiyento

B-2 API

1600' 120

20

B-3 API

1900' 120

20

B-4 API

2200' 120

5300'

Widu ri- 7

20

API

120

A'

Widuri-5

20

API

120

Top l i m estone

Base coal 'A'

DEPTH LOCATION MAP

GAMMA RAY API100 0

(FT)

Core log 34-2 sandstone Widuri B-8 Muddy l i mestone and calcareous m u dstone Bivalve and fora m i n ifera fragments Root leted si ltstone and coal

Trough cross- bedded sandstone g rading from med i u m to coarse at the base to fine-very fine at the top

4490

Si MCg ,...,... Sh�Cl SA ND I

NO RECOVERY 4500

4510

4520

Coal 'A'

Fig. 13. 34-2 reservoir: vertical and lateral sedimentary character..

B-8 well (see Fig. 15). The sandstone is erosively based and consists of mostly fine- and very fine­ grained, very well-sorted sandstone, which broadly fines upwards giving a predominantly gamma-ray increasing upwards log signature. Cross-bedding is sporadically present throughout but, towards the top, parallel lamination, ripple cross-lamination and ripple/wavy bedding predominate. Carbonaceous

laminations are common, and abundant mudstone clasts in the lower to middle part of the sandstone result in increased gamma-ray readings (Fig. 15). Towards the top of the sandstone (4382. 5-4391f t MJ?) a heavily bioturbated and interbedded·­ interlaminated unit of mudstone, siltstone and sandstone contains predominantly Thalassinoides but also Teichichnus type burrows, which are indica-

Fluvial sand-body geometries

\ ....____

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369

+

+

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+

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+

+

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I

:

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( )

;--s··:

(

"-.!

+

+

+

Cl 20' 5000 FT

+

Fig. 14. 34-2 reservoir : (A) Net

sandstone isopach with the oil­ water contact marked. (B) Depth structure map on the top 34-2 reservoir showing that the oil­ water contact is 50ft below structural closure (shaded).

tive of marine influence. In Widuri-1 the 34-1 was also cored and clay drapes, a common fei!ture of tidal sediments, are abundant in the upper part of the sandstone, where marine acritarchs also have been found. The net sandstone isopach (Fig. 16A) defines a sand body with a maximum penetrated thickness of 49ft ( 15 m) and an average width of 2000 ft ( 610 m). The detailed shape of the contours in Fig. 16A is not

a

E

·.-----....

n '\

,/

\

CllOO'

5000 FT

B

the result of extrapolated well data. Prior to any development drilling it was recognized that an acoustic impedance contrast between the base of the sandstone and the underlying mudstone was respon­ sible for a seismic peak (see Blue seismic horizon, Fig. 6, for example) which, when mapped, had a restricted areal extent. Seismic modelling suggested that the sandstone thickness could be estimated from the amplitude of the reflection through the use

370 A 20

R. Young, W.E. Harmony and T. Budiyento

D-11 API

2150' 120 20

A-12 API

1600' 120 20

Widu ri-1 API

5000' 120 20

C-6 API

DEPTH

LOCATION MAP GAMMA RAY API

(FT)

3850' 20 120

B-8 API

A' 120

Core log 34-1 sandstone Widuri B-8

Biotur bated mudstone, si ltstone and sandstone

Thalassinoides/ Teichichnus type bu rrows

Fine and very fine grained sandstone

Calcareous mu dstone with shell fragments

Fig. 15. 34-1 reservoir: vertical and lateral sedimentary character.

of a tuning curve (Young et al., 1991), and the net sandstone isopach map was constructed based on reflection amplitude. This map proved to be very accurate and has changed little after development drilling. The shape and dimensions of the sandstone, together with the log and sedimentary character, suggest a single-storey channel which was reworked by tides in the upper part. Within Widuri field, the interval that is laterally equivalent to the channel sandstone consists of mudstone and siltstone with rare, thin sandstones (see cross-section A-A',

Fig. 15). When traced southwards away from the sediment source area, however, Young & Atkinson (1993) have shown that the entire productive interval at Widuri becomes more marine and is dominated mostly by 6-24 ft thick, stacked, coarsening upwards units, which are interpreted to be distributary mouth bars and various other marginal marine bars. Dia­ chronism of the whole interval is clear and the overall interpretation is a deltaic succession domi­ nated by tidal and fluvial processes (Young & Atkinson, 1993). The 34-1 is therefore interpreted to be a distribu-

Fluvial sand-body geometries

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Fig. 16. 34-1 reservoir. (A) Net

sandstone isopach with the field oil-water contact marked. (B) Depth structure map on the top 34-1 sandstone showing that the oil-water contact is 1 00ft below structural closure (shaded).

tary channel and the tidal reworking at the top is perhaps mostly a result of estuarine conditions after abandonment of the distributary. The apparent absence of delta-front bars may reflect the proximal position of the Widuri area with respect to the delta complex and/or the low preservation potential of tidal bars (oriented normal to the shoreline) as the delta prograded. The shape, dimensions and sedi­ mentary character of the 34-1 are very similar to

Cl100' 5000 FT

8

those of a modern Mahakam delta distributary channel (see Allen et at., 1979; Allen, 1985}. Reservoir quality is still very good with average log porosity of 29% and test permeabilities of 5-10 D. The depth structure map on top of the 34-1 sand­ stone (Fig. 16B) identifies the structural closing contour at 3580ftSS (1091 m SS). An oil-water contact for the 34-1 sandstone has not been penetrated in any well. However, a reservoir simu-

372

R. Young, W.E. Harmony and T. Budiyento

stratigraphical levels over a vertical interval of about: 75ft (23m) immediately above the 34-1 sandstone. None of the reservoirs has been cored, but Fig. 17 demonstrates the log character and limited lateral extent of the sandstones. Nine separate sandstones have been identified so far (Fig. 18A), although this number could increase or decrease with further drilling. The sandstones are fine to very fine grained with sharp bases and sharp or gradational tops and usually show a gamma-ray increasing upwards log signature. Penetrated net sandstone thickness does not exceed 15ft (4.5 m), with one exception, the 33-2, which is 26ft (8 m) in Widuri-6. Sand body widths average 2000 ft (610 m) but well penetrations of individual sandstones are too few to define geo-· metries accurately. However, at least some of the sandstones are thought to have shoestring geometries and occasionally, as in the case of the 33-2, reflection

lation study using the geological model described here and an oil-water contact at 3680 ft SS (1122 m SS) gave a very good match with the available production history. The oil-water contact is there­ fore assumed to be the same as all the underlying reservoirs. A discrepancy of 100ft (30.5 m) between the oil column and the structural closure can again be accounted for by a stratigraphical component to the trapping, where northwestwards closure is effected by pinchout of the reservoir. 33-series sandstones

There is less information on the 33-series sandstones than on any of the other reservoirs and they are the least understood and most difficult to define. The 33-series is a group of several different sandstones which, for the most part, occupy slighty different A 20

D- 1 API

3,800' 120

20

D-4 API

120

12,600'

C- 2 20

API

32 , 00' 120 20

API

0

Talang Akar formation

50 FT

150

120

Top

0

100

A'

C-1

FT

Top 33-1 sa ndst one i e �L_ �!_ .!

50

100

Top 34-1

150

san dst one inte rval

5000FT

t.----J

Fig. 17. Cross-section showing the electric-Jog character and lateral discontinuity of 33-series reservoirs.

3 73

Fluvial sand-body geometries

A

Cl100' 5000 FT

B Fig. 18. 33-series reservoirs. (A) Net sandstone isopach with the field oil-water contact marked. (B) Depth structure map

on a middle 33-series limestone marker showing that the field oil-water contact is 150ft below structural closure (shaded).

R. Young, W. E. Harmony and T. Budiyento

374

amplitude at the base of a sandstone can be used to map it. The sand bodies are laterally equivalent to finer grained sediment, which was cored in the C-2 well (Figs 1 7 & 19). The basal 32ft (9. 75m) of the core consists of carbonaceous mudstone with laminations of siltstone and very fine-grained sandstone. These laminations are millimetre-scale and sparse at the base but centimetre-scale and frequent towards the top. Both bioturbation and ripple cross-lamination increase upwards and fragments of fish bones, foraminifera and echinoids are scattered through­ out. The basal 27ft (8. 25 m) is a coarsening upwards unit which grades into a 5 ft ( 1.5 m) thick, fining upwards unit. The upward fining trend continues into highly bioturbated siltstone before a very sharp contact at 3899 ftMD marks the top of the unit. The overlying dark brown, carbonaceous mudstone contains scattered marine fauna. The coarsening­ fining couplet is thought to reflect prodelta to lower mouth-bar progradation and abandonment, whereas

DEPTH (FT)

GAMMA RAY

0

API

the overlying mudstone is clearly a shallow-marine deposit with minimal coarse clastic sediment input. However, only 16 ftMD (4. 9 mMD) above the top of the core is a 3 ft (1 m) thick coal. Using similar reasoning as for the 34-1 sandstone, the 33-series sandstones are interpreted to be single­ storey distributary channel deposits of a tide­ influenced delta. Reservoir quality is variable but log porosities average 27% and test permeabilities are of the order of 1-5 D. The depth structure map on a marker in the middle of the 33-series interval identifies a structural closing contour at 3530ft SS (1076 m SS, Fig. 18B). The oil-water contact for the 33-series sandstones is unknown and reservoir simulation studies do not help because production histories are too short. The oil-water contact is assumed to be at 3680 ft SS (1122m SS), the same as the other reservoirs, although multiple oil-water contacts are possible. However, irrespective of the position of the oil­ water contact, the oil column is again greater

100

3880

Widuri C-2

Fig. 19. Core log of part of the 33series interval in Widuri C-2 (see Fig. 17 for location of cored interval).

Fluvial sand-body geometries

375

Feet NW

0

1000

2000

3000

SE

Time Feet 1000 3100 1020 1040

3200

1060 3300 1080

3400

1100 1120 1140

3500

3600

1160 3700 1180 3800 1200 3900 1220 1240 4000 1260 1280 1300

4100

4200

1 320 1340 1360

4300

4400

Fig. 20. An uninterpreted seismic line from the three-dimensional volume over the Widuri area. The line is located just

south of Widuri field and shows an oblique view through one of the features that are interpreted to be incised valleys associated with the top Talang Akar horizon. See Fig. 6 for the stratigraphical location of the seismic markers.

than the structural closure and a stratigraphical component to the trapping, although less well defined, is again necessary. The interval between the top 33-1 sandstone and the top Talang Akar Formation consists of coal,

mudstone and siltstone with some thin, interbedded limestone. A low-energy, marginal marine setting is envisaged prior to the development of fully marine conditions in the overlying Batu Raja Formation.

376

R. Young, W.E. Harmony and T. Budiyento

Reservoir

33

Series

34-1

34-2

35- 1

35-2

36 -1

Average grain size

Multi storey/ m u ltilateral

fine to very fine

NO/NO

fine to very fine

medium

medium

medium to coarse

coarse

NO/NO

?NO/NO

YES/YES

YES/YES

YES/YES

®

*

Reasonable maximum thickness (ft)

715'-730'

40'

50'

60'**

65'

175'

@Average width

(ft)

72000'

2000'

4000'

6500'* *

1 1000'

Sheet

®I@ ratio

7 1 :67 to

Average porosity a n d Aerial distribution Environment of deposition (to scale) permeability range

darcies

1:50

29% 5-10

1:80

1: 108

1: 169

Sheet

�

27% 1-5

7 1 : 1 33

29% 10-30

darcies

29% 10-20

darcies

29% 20-30

darcies

28% 20-30

darcies

>CC UJ


<>: 1- w f-' ::::l Z

�

darcies

([: 1flj 0 �

"l

\�

Lithostratigraphy

g f-'� I�

cc w m

�u 0

::;; w ::;;

j:!

(3

z 0

� ::;;

UJ cc ::::>

�

::;; ::;;

cc 0 u.. cc

..: "" ..:

<.:J z

w 1Ul >Ul ...J

..: ...J

cc UJ m

..: >

::::> ...J u..

UJ cc ::::>

�

j:!

::;; UJ ::;; ..: 0 ...J UJ N

::;; :::;

__ ;

Fig. 21. Summary and comparison of sedimentary character, geometry and environment of deposition of the Widuri field

reservoirs. * Using 'reasonable' maximum thickness is an attempt to eliminate thickness anomalies; the parameter is not rigorously defined but nor is it arbitrary. * * Only the east channel system was considered ; the west channel system may be structurally controlled.

Discussion

The upper part of the productive interval at Widuri (the 34-2, 34-1 and 33-series sandstones) is inter­ preted as deltaic based largely on evidence outside of the Widuri area. Core data from approximately the same stratigraphical interval in the Sunda basin and the Ardjuna basin (approximately 70 miles to the southeast of Widuri in the adjacent PSC) indicate a predominance of tidal and fluvial processes at the shoreline (Young & Atkinson, 1993). In this setting, tidal reworking at the delta front fashioned delta­ front bars, oriented predominantly normal to the shoreline, which were thin (up to 25ft thick) compared with the channel deposits (approximately 15-40ft thick). The delta-front bars therefore had low preservation potential as the delta prograded (see e.g. Allen et al., 1979; Allen, 1985; Coleman &

Wright, 1975, their Fig. 16). The distributary channels of tidally influenced deltas are therefore not always juxtaposed against delta-front bars and, once abandoned, they become estuaries. These same characteristics (namely, channel features that erode into marine sediments and are estuarine in their upper parts), however, could be used to interpret the reservoirs (especially the 34-2 and 34-1) as incised valley fills, which formed as a response to periods of falling relative sea-level (see e.g. Van Wagoner et al., 1990). It is also reasonable to expect that relative sea-level changes caused by eustasy and/or tectonic movement had, at times, a major effect on deposition of the Talang Akar Formation over the 7.5Myr period. An example of this is at or near the top of the Talang Akar Formation, where a horizon characterized by down­ cutting is clearly seen on seismic records (Fig. 20).

Fluvial sand-body geometries

Reservoir

Aerial distribution ('to scale')

�

33

series

34-2

35-1

35-2

Fig. 22. Summary of Wid uri field

36- 1

Structural closing contour (ft SS)

Oil col u m n (ft)

3530

325

F i l l w.r.t. structure

Trap type

Q)

3680

q;� rf!! ��'

5'?

/f

(ft/SS )

3680 ?

�

34-1

owe

377

�1

3680

3580

3630

2 70

240

-o �

c Vl 0 0

- ·� - .c

.o c;; -o �

u rn :o · -

�u


·- ::0 LJ._ !:: "'

3680

3660

180

3680

3680

110

F i l l ed to spill

3680

3740

55

Part i a l ly fi l l ed

� 0. ::0 "' �

-

- "' Ul !::

"'

-

� 3 u 2 cil

trapping characteristics.

When mapped, this down-cutting feature has a channel form with a roughly N�S orientation, a width of approximately 1000-3000ft, and a thick­ ness or 'down-cut' of about 30-lOO ft. An interpret­ ation of recent three-dimensional seismic coverage over the northwest flank of the Asri basin has identified several more of these features, all associ­ ated with the same horizon. Unfortunately none of these features has yet been cored, but the event can be traced seismically all over the Asri basin and 60miles south into the western Sunda basin, where it correlates to a karst surface within limestone. The event therefore has regional extent and is interpreted to be the result of a lowstand of sea-level. There is strong seismic stratigraphical evidence that the low­ stand is a response to tectonic reactivation of the Asri basin. In contrast, within the productive interval at Widuri field there is no such seismic expression and no such evidence for regional lowstand surfaces of erosion (see Fig. 20). Until further evidence surfaces, therefore, the retrogressive stacking of the reservoirs at Widuri field is interpreted to be a response to essentially autocyclic switching of sediment supply within an overall transgressive systems tract. How­ ever, the primary aim of this paper is to demonstrate trapping configurations resulting from the interaction of sand body geometry and structure, and these are independent of the mechanism of stacking.

R E S E RV O I R

SUM M A RY

The more important aspects of the geometry and character of the Widuri reservoirs are documented in Fig. 21. The productive interval as a whole shows a gradual and systematic change from base to top. Grain size gradually decreases upwards from coarse grained in the 36-1 to fine grained - very fine grained in the 33-series reservoirs, and multistorey and multi­ lateral channel systems in the basal three reservoirs give way to single-channel systems in the upper three. The maximum sand body thickness gradually decreases upwards, as does the average width, and the thickness/width ratio shows a corresponding change. Porosities are similar throughout but test permeabilities show a significant decrease upwards. The change in sand body size, shape and character is the result of a maturing depositional system associated with a rise in sea-level between 25. 5 and 22Ma (see Fig. 5; Haq et a/., 1988). The immature fluvial system at the base (36-1 reservoir) is suc­ ceeded by progressively more mature fluvial systems before deltaic conditions prevail at the top. The change in character of the sand bodies therefore reflects the change in environment of deposition and this is supported by the available palaeontological data from the intervening shales. Rare occurrences of palynomorph taxa commonly associated with a brackish water setting have been found as low down

R. Young, W.E. Harmony and T. Budiyento

378

SE

NW

SE

Fig. 23. (A) Schematic diagram of the Widuri field trapping configurations. (B) A projection of what the Widuri trapping configuration would have been if the 34-2 reservoir had been a sheet sandstone.

as the 35-2 interval and increased frequencies in the overlying 35-1 and 34-2 intervals indicate increasing proximity to the sea.

TRAP

SUM M A R Y

The important aspects of the trap are summarized in Fig. 22. With the possible exception of the 33-series, the reservoirs share a common oil-water contact and therefore must either be in communication now or have been in communication some time after oil migration. The structural closing contour on top of the 36-1 sandstone is below the oil-water contact and the reservoir is only partially filled. The d osing contour on top of the 35-2 sandstone is coincident with the oil-water contact and is therefore full to spill. This is an extremely important point because it

is the spill point of the 35-2 reservoir that therefore defines the oil- water contact for the entire field. The lower two reservoirs are structural traps, but the upper four reservoirs are filled beyond structural closure and rely on a favourable orientation (northwestwards pinch-out) with respect to structure in order to trap oil by a combination of structure and stratigraphy. The trap model is drawn schematically in Fig. 23A but it is interesting to evaluate how the trapping would change if one of the upper four reservoirs, the 34-2 for instance, had been either a sheet sandstone or did not have a favourable orientation with respect to structure. In other words, what would have happened if the 34-2 did not have a stratigraphical component to trapping? This case is shown in Fig. 23B. The reservoirs are still in communication and so the oil would migrate upwards through the

Fluvial sand-body geometries

lower reservoirs into the 34-2, where it would be trapped with respect to a new field oil-water contact as defined by the 34-2 structural spill point (3630 ft SS or 110 7 m SS, see Fig. 22). The net effect is exaggerated in the schematic diagram of Fig. 23B, but a 50 ft (15 m) higher oil-water contact would significantly reduce the volume of oil in the field. An overall conclusion then is that the trapping at Widuri is intimately related to reservoir geometry and that the large size of the field is the result of a remarkable coincidence of events. A relative rise of sea-level helped cause a gradual change in the environment of deposition which, in turn, produced a systematic reduction in the size of sand bodies, the upper four of which each had to be oriented favour­ ably with respect to structure in order to trap oil stratigraphically.

APPLICATIONS

The geological model has contributed to the devel­ opment of Widuri field in two important ways: (i) improved drilling success and (ii) improved reservoir management. Early recognition of sand body geometries resulted in more efficient develop­ ment drilling, such that of the 47 development, infill and step-out wells drilled so far only two lacked economic pay, and both of these were side-tracked to successful locations. Similarly, reservoir manage­ ment has been optimized by having a detailed and predictive geological model, which forms the basis of reservoir simulation studies. This application will become increasingly important as the field matures, water production increases, and further infill drilling is required to optimize hydrocarbon recovery. From an exploration viewpoint, the work on Widuri field has provided an analogue model in the search for new fields in the PSC area, in which stratigraphy is an important means of trapping hydrocarbons. Widuri is by definition a large oil field but approximately 75% of its reserves have a strati­ graphical component to trapping. Any future dis­ covery of stratigraphically trapped oil could therefore potentially be large.

AC KNOWLEDGEMENTS

The authors wish to thank the management of Pertamina, Maxus and partners for permission to publish this paper. Special thanks are extended to

379

C.L. Mora and C.D. Atkinson for a critical review of the manuscript, and to J.F. Mitchell for his technical input in the early stages of the study. We also acknowledge the help and support of all the other members of the multi-disciplinary team that was responsible for developing Widuri field.

REFERENCES ALLE N ,

G.P. (1985) Delta sedimentation: modern and ancient examples in the Mahakam delta basin. Indonesia Petroleum Association, Field Trip Guide.

G.P. , LAURIER, D. & THOUVENIN , J. (1979) Etude sedimentologique du delta de Ia Mahakam. Compagnie Francaise Petrol. Notes Mem. , 15 , 156 pp. A LLEN , J.R.L. (1965) A review of the origin and character­ istics of recent alluvial sediments. Sedimentology , 5, 89-192. BEDDOES, L.R. (1981) Hydrocarbon plays in Tertiary basins of Southeast Asia. Energy, 6, 1141-1163 . BEN-A VRAHAM , Z. & E MERY, K.O. (1973) Structural framework of Sunda Shelf. Bull. Am. Ass. petrol. Geol. , 57, 2323-2366. B USHNELL, D . C. & TEMANSJA, A.D. (1986) A model for hydrocarbon accumulation in Sunda Basin, West Java Sea. Proceedings of the Indonesian Petroleum Associ­ ation, 15th Annual Convention, Vol. 1 , pp. 47-75. COLEMAN, J. & WRIG HT , L.D. (1975) Modern river deltas: variability of processes and sandbodies. In: Deltas (Ed. Broussard, M.L.), pp. 99-149. Houston Geological Society, Houston, TX. COLLINSON , J.D. (1978) Vertical sequence and sand body shape in alluvial sequences. In: Fluvial Sedimentology (Ed. Miall , A.D.), Mem. Can. Soc. petrol. Geol. , Calgary, 5, 577-588. F IELDING , C.R. & CRA N E , R.C. (1987) An application of statistical modelling to the prediction of hydrocarbon recovery factors in fluvial reservoir sequences. In: Recent Developments in Fluvial Sedimentology (Eds Ethridge , F . G. , Flores, R.M. & Harvey, M . D.), Spec. Publ. Soc. econ. Paleontol. Mineral. , Tulsa, 39, 321-327. F oLK, R.L. (1968) Petrology of Sedimentary Rocks. Hemphill's, Austin, TX, 170 pp . HAQ, B.U . , H ARDENBOL, J. & VAIL, P.R. (1988) Mesozoic and Cenozoic chronostratigraphy and eustatic cycles. In: Sea-level Changes: an Integrated Approach (Eds Wilgus , C. K. , Hastings, B.S. , Kendall, C.G.StC. , Posamentier, H . W. , Ross, C.A. & Van Wagoner , J.C.) , Spec. Pubis Soc. econ. Paleontol. Mineral. , Tulsa, 42, 71-108. H uTCHISO N , C.S. (1986) Tertiary basins of S.E. Asia ­ their disparate tectonic origins and eustatic strati­ graphical similarities. Bull. geol. Soc. Malay. , 19, 109-122. LEEDER, M.R. (1973) Fluviatile fining-upward cycles and the magnitude of palaeo channels. Geol. Mag. , 110, 265-276. LORENZ, J.C. , H EINZE, D. M. , C LARK, J .A. & S EARLS, C .A. (1985) Determination of widths of meander-belt sandstone reservoirs from vertical downhole data, Mesaverde Group , Piceance Creek Basin, Colorado. ALLEN ,

380

R.

Young, W.E. Harmony and T. Budiyento

69, 71 0-721. R.M . , CAMPION , K.M. & RAHMANIAN , V . D . (1990) Siliciclastic sequence stra­ tigraphy in well-logs, cores and outcrops: concepts for high resolution ·correlation of time and facies. Am. Assoc. petrol. Geol. Meth. Explor. Ser. , 7, 55 pp. WIGHT, A. , SUDARMONO, & lmron, A. (1986) Stratigraphic response to structural evolution in a tensional back-arc setting and its exploratory significance: Sunda basin, west Java Sea. Proceedings of the Indonesia Petroleum Association, 15th Annual Convention, Vol. 1 , pp. 77100. Bull. Am. Ass. petrol. Geol. ,

VAN WAGONER,

J. ,

MITCHU M ,

R. & ATKINSO N , C.D. (1993) A review of Talang Akar Formation (Oligo-Miocene ) reservoirs in the off­ shore areas of Southeast Sumatra and Northwest Java. In: Clastic Rocks and Reservoirs of Indonesia (Eds Atkinson , C.D. , Scott, J. & Young, R.). Indonesian Petroleum Association, Core Workshop Notes, October 11, pp. 177-21 0. YOUNG, R. , HARMONY, W . E . , GUNAWAN, J. & BUDIYENTO, T. (1991) Widuri field, offshore Southeast Sumatra : sandbody geometries and the reservoir model. Proceed­

YOUNG,

ings of the Indonesian

Petroleum Association,

Annual Convention, Vol.

1 , 385-417.

20th

Index

Abyssal plain,iron formations 149-51, 154 Acidic volcanic rocks,geochemistry 252 Ahwaz, Cenozoic clastic reservoirs 349-50 AI Khlata Formation 340,341 Allostratigraphy 72-3 Alluvial fan facies association, Tryfan Fach Member 172-4,184-5 petrography

178

Alluvial fans 203 Amin Formation 340 Anglesey 178 Animikie Group, iron formations 139-40, 142 Arabian Plate

335-9

Arenites lithic, geochemistry 247-8 lithic-volcanic,geochemistry 252-3 Askrigg Block 307-8,312-20 Asri basin 355-8, 377

Baby Formation,iron formations 143-6

85-94 Bideford Group 101-3, 110 Black shale facies,Bude Formation 126-7, 130 description 110-11 interpretation 120-2

Cherty iron formation Chesil Beach 91, 93

Black shales, geochemistry 248 Boss Point Formation 117 Bottom-current activity,ridge-

Computer maps,preparation Conglomerate debris, Kalvag Melange geochemical analyses 245 geochemistry 241-8 Conglomerates

trough system 296-7 Bovis Formation,Pororari Group 204,208-10 palaeocurrent pattern 212-13 Braided stream association,Tryfan Fach Member 166-9, 174-5 bounding surface with marine shelf association 177 petrography 178-9 Breccia 198-203 Bude Formation 101-5,130-1 bedtypes 108-9 facies analysis 110-26 description 110-20 interpretation 120-6 facies succession 106-10 description 126 interpretation

Burrowing, Woburn Sands 29,35-6

Basalts, geochemistry

Camaqua Basin, Brazil Cameros Basin 221 stratigraphy 248-

transitional 251-2 Basra, clastic reservoirs Cenozoic Mesozoic

349-50 344-5

Beachfaces bedforms and sedimentary structures characteristics

92-4 95

facies successions 94-6 gravelly 76 sediment transport 78-9 Beardmore Volcanic Associated Iron Formation 151-2 Beardmore-Geraldton Clastic Associated Iron Formation 146-9 Bedfordshire Strait 25 Bedforms,gravelly shorefaces

Clastic rocks, Upper Jurassic 232-3 Clasts 80-5 Coastal zone terminology 76-7

grain sizes 80-4 Guaritas sequence Ricinus 62

73

5-7,9-10

Senora de Brezales Formation 223-7 shoreface 75 Crackington Formation 104-5, 107-10 Craven Basin 308-9,312-18 Cryptobioturbation 35 Culm Basin 102-4 tectonic setting 105-7 tectonically juxtaposed successions 107-10

126-30

Back-arc basin 257 Bahrain, Mesozoic clastic reservoirs 344-5 Bars, gravelly shorefaces 85-6 alkaline 248 island arc 252 subalkaline MORB-Iike 51

139, 140

27-8,

3-4,6

Debris flows 186 breccia 203 Diorite 247 Dune migration, Guaritas sequence 10-11

220, 422

Canadian Shield, iron formations 139 Capel Curig Volcanic Formation 159, 183 Carbonate rocks,Senora de Brezales Formation 227-8 Cardium Alloformation,Alberta 48-50,73 allostratigraphy 49-50 comparison of allomembers stratigraphy and

Erosion channels 268-70 ridge-trough system storm/wave 271-4

297

surfaces 26-7 Estuaries 69-70 Estuarine environments, essential elements 41-2

65 Facies modelling

187-9

sedimentology 50-5 Cardium Formation 73, 76

Fan-deltas Feni Drift

gnivelly shorefaces 94 Carstone 38

Floodbasin facies association, Tryfan Fach Member 169-

80-4, 86-

Cenozoic clastic reservoirs 50

72, 174-7 349-

petrography 178-9 Frio Formation 267, 271-5

Channel erosion 268-70 Chert,geochemistry 247

Sedimentary Facies Analysis: A Tribute to the Research and Teaching of Harold G. Reading Edited by A. Guy Plint © 1995 The International Association of Sedimentologists ISBN: 978-0-865-42898-0 381

213, 216 297, 299

382

Index

Gabbro 247 Gardar Drift 297 Gas, Middle East 333 Gault Clay 22-3 Gharif Formation 340 Goniatite-bearing marine bands 309-11 Goniatites 111 Graphitic shear zones

149

Gravel deposition, Guaritas sequence 8-10 Gravel plunge step 90-1 Gravel wave ripples 85-6 Gravelly shorefaces 75-6 Gravels clast shape 84-5 grain size and shape 80-5 textural trends 80-4 wave-worked 90 Groundmass, olistostromal, Kalvag Melange 240-1 geochemistry 253, 256 Growth faulting 265-6 effects on sedimentation 26776 Guaritas depositional sequence, Brazil 3-5, 6 alluvial facies general features 5-11 interpretation of features 8-11 interpretation of lateral changes 11-14 lateral changes 11-14 pattern of textural dispersion 11 palaeocurrent pattern

11

Gulf of Suez, clastic reservoirs Cenozoic 350 Mesozoic 348-9 Palaeozoic 343-4 Gulf Region, Palaeozoic clastic reservoirs 342-3 Gunflint Formation, iron formation 139 Gwern Gof Tuff

159-60

Hadhramaut-Jeza rift basin 346-7 Haima Group, clastic reservoirs 340 Haradh Formation 340 Haushi Group, clastic reservoirs 340-1 Hawks Crag Breccia

199-201,

214-15 deposition 211 Hummocky cross-stratification, Bude Formation 129 Hydrocarbon trapping 378-9

Ignimbrite, geochemistry 252 Incised valley Ricinus 47-9, 66-71 Interbedded sandstone-shale facies, Bude Formation 126-7, 130 description 112-17 interpretation 122-5 Iron formations 137-9 case studies discussion objectives types

139-53 153-4 139

138

348

Jordan, Palaeozoic clastic reservoirs

241, 246-7 Middle East Clastic reservoirs

240-1

oil and gas reservoirs 333-9 palaeogeographical and palaeoclimatic synopsis

geological setting 238-40 profile through 242-3 provenance and tectonic setting 256-9 regional correlations and interpretations 259-61 Karst surfaces, Senora de Brezales Formation 228-9 Kuwait, clastic reservoirs Cenozoic 349-50 Mesozoic 344-5 Kyeburn Formation

335-9 source rocks Millstone Grit

336

cycle duration 306 cyclothem architecture and sea-level change 311-21 asymmetry 321-2 E2a(i) cyclothem architecture and sequence stratigraphy 313-17 boundaries and thickness 312 depositional model and

198

143

Laminar calcretes, Senora de Brezales Formation 228 Laminated beds, Bude Formation 112, 123 Late Jurassic unconformities 2334 Levant-Sinai region, Mesozoic clastic reservoirs 348-9 Lithic arenites, geochemistry 247-8 Lithic-volcanic arenites, geochemistry 252-6 Llewelyn Volcanic Group 15860, 183 180

tectonism 313 E2c2 cyclothem architecture and sequence stratigraphy 318-20 boundaries and thickness depositional model and tectonism 317-18

317

major cyclothems 322-3 architecture and genesis 323-4 major scale cyclicity

322-4

multi-cyclothem condensed sections 322-4 sequence model 320-2 choice 320-1 Morley Occurrence, iron formations

London- Brabant land mass

19, 20

Longshore currents, sediment transport 77-8 Lower Follifoot Grit 317-19 Lower Greensand Group, deposition 19-20 Lower Miocene, Louisiana 7, 274-6, 278

349-50

Mesozoic 344-9 Palaeozoic 339-44 geological boundaries 331-2

depositional model 257 geochemistry 241-56

volcanic source

stream association 177 petrography 179-80 Mesozoic clastic reservoirs 344-9 Metabasalts, Kalvag Melange

Cenozoic

343

Labrador Trough

Magma migration 185 Mahwis Formation 340 Marib-Jauf basin 346 Marine bands, goniatite-bearing 309-11 Marine shelf association, Tryfan Fach Member 161-6, 174 bounding surface with braided

Island arc basalts, geochemistry 252 Israel, Mesozoic clastic reservoirs

Kalvag Melange

Lower Ordovician ophiolite complexes/island arc sequences 259-60

266-

152-3 Muddy siltstone facies, Bude Formation 126-7, 129-30 description

111-13

interpretation Mudstones

122

Guaritas sequence Ricinus

8

60-2, 65, 66

Index deposition 70 distribution 63-5 Tryfan Fach Member 161-6 Mudwaves, ridge-trough system 296-7

383

pebbly sandstone and massive or graded sandstone 20310 geological setting 199 palaeocurrent pattern 212-13 tectonic implications 213-15 Pororari River 211

Namurian Basin, Central Province, 30-8 Namurian stratigraphical

Preservation potential thresholds 276 representation in growth-fault

framework 306 Neutral Zone, Mesozoic clastic reservoirs 344-5

settings 279 Pro-delta current system, ridge­ trough system 297-302

New Zealand 197-9 North Sea, ridge-trough system

Puysegur Formation

285-6 database and methodology 288 depositional models 296-302 external morphology 288-9 geological setting 286-8 internal reflection configurations 289-91 interpretation 291-2 mapping 292-4 palaeogeographical setting of structures 294-6

Oil, Middle East 333 Olistoliths, Kalvag Melange 240-1 sedimentary rocks geochemistry 252-3, 255 representative geochemical analyses 254 spatial distribution volcanic rock

244

geochemistry 248-52 representative geochemical analyses 249 Oman, clastic reservoirs Mesozoic 347-8 Palaeozoic 340-1 Outer shelf and slope, iron formation Outwash fans

143-6, 154

169

198

Pyroclastic material 158-9 syneruptive, aggradation 184-5

Qatar, Mesozoic clastic reservoirs 348 Quasi-planar lamination Qusayba shales 342

164

Ramp-fan system, iron formation 147-9, 154 Rhyolite, geochemistry 252 Rhyolitic braided-stream sandstones, Tryfan Fach Member 166-9 Ricinus, Alberta

47-9, 70-1

Allomember 54 comparison with other allomembers 65 cross-sections 71-2 deposition in estuary

212-13

Palaeozoic clastic reservoirs 339-44 Palestine, Mesozoic clastic reservoirs 348 Pennine Basin 308-9 Petroleum producing formations, Middle East 334-5 Planar lamination 164 Porcupine Basin 232-3 Pororari Group 199 depositional systems 211-12 facies assemblages 199-210

69-70

facies 57-62 lateral and vertical distributions 62-5 geometry 55-7 incised valley 56 sand body 56-7 incised shoreface 65-7

laminated mudstone 210 matrix- and clast-supported breccia 199-203

Sand volcano 120-1 Sand-body geometries, Widuri field 377 Sandstone-dominated facies, Bude Formation 129 description 117-20, 122-3 interpretation 124-6 Sandstones Bude Formation

112-17

graded silty 112-13, 123-5 Guaritas sequence 7-8, 10 irregularly laminated 117, 125-6 red, Senora de Brezales Formation 220-3 Ricinus 57-60, 69-70 parallel-laminated 59-60 structureless 57-9 Tryfan Fach Member conglomeratic alluvial fan 172-4 rhyolitic braided-stream 166-9 structure 179 thinly bedded floodbasin 169-72

Scandinavian Caledonides

259

Sea-level change and Millstone Grit cyclothem architecture 311-21 goniatite-bearing marine bands 309-11 Seaweed, sediment transport 78 76-9 Sedimentary rocks, olistoliths,

49

valley orientation 70 Ridge-trough system, North Sea 285-6 database and methodology 288 depositional models 296-302 external morphology 288-9 geological setting 286-8 internal reflection configurations 291-2

mapping 292-4 palaeogeographical setting of structures 294-6 Rip currents, sediment transport 78 River, trunk braided

Salinity, and marine bands 309-11 San Leonardo Formation 228-9

Sediment transport processes

interpretation 65-70 origin of valley 66-8 palaeogeography 68 structural problems 49

289-91 interpretation

347

Widuri field 360-75 Saudi Arabia, Palaeozoic clastic reservoirs 341-2

turbidite interpretation Palaeocurrents

Rutbah Formation

12-14

geochemistry 252-5 Sedimentary structures, gravelly shorefaces 85-94 Sedimentation and growth faulting 265-6, 267-76 Senora de Brezales Formation clast derivation 230-1 depositional history

230-1

depositional model 231-2 environmental synthesis 231 erosion 230-1 facies development, tectonic controls 229-30 late Jurassic unconformities, geometry 228-9 lithofacies 220-7 stratigraphy 220 Sequence stratigraphy

72, 266

234

384

Index

Shallow shelf, iron formation

139,

153-4 Shallow-marine sands 18 Shenley Limestone 38 Shorefaces bedforms and sedimentary structures 85-92 characteristics 95 facies successions 94-6 gravelly 75-6 sediment transport processes 77-8 Silica, iron formations 138-9 Skaggerak Delta 287, 297-300 Slaty iron formation 139, 140 Slumped beds, Bude Formation 116-17, 120-1, 125, 131 Slurried beds, Bude Formation 114-16, 118-19' 125 Snowdon graben 159 Solund-Stavfjord Ophiolite Complex

260-1

South Saskatchewan River 168 Staffordshire Basin 309 Stainmore Basin 307-8 Storm erosion 271-4 Submarine extrusive edifice, iron formation 152-3 Submarine lava plain, iron formation 151-2 Submarine rise, iron formation 146-9 Subsea fans 104-5, 106-7, 126,

associations

177-8

161-74

organization and distribution 174-8 relations 174-7 position of volcanic source 180 Tryfan Formation 159-61, 178 Turbidite system, Bude Formation 129-31 Turkey, Palaeozoic clastic reservoirs 343

Unayza Formation Unconformity 72 Underflows

164

Upper Howgate Edge Grit 16 Upper Jurassic clastic rocks Upper Wilcox Group

313232-3

266-70

Vertical-sequence analysis Volcanic rock olistoliths,

157-8

geochemistry 248-52 Volcanic terrains 158 bounding surface between syn­ and post-eruptive deposits 183-4

epiclastic interpretation 180-2 facies associations 161-74 modelling 187-9

357, 359,

376 palaeogeography 361 Tan Hill Grit 316-17 Terrace Bay Clastic Associated Iron Formation 149-51 Trace Fossils, gravelly shorefaces 91-2 Transgressiveestuarine-embayment system, depositional model 40-2 12-14 86-7 12-14

deposition 211-12 palaeocurrent pattern

212-13

Wave erosion 271-4 Wave motion and sediment transport 77 Widur field 355-6 background geology 356-8 palaeogeography 362 reservoir geometry and trap

360-77

overview 358-60 summary 377-8 stratigraphy 357, 359 trapping characteristics 378-9 Woburn Sands

377,

depositional models 20 facies 22-3 Heterolithic Sands 22-3 deposition 38-40

342

depositional systems

Talang Akar Formation

Troughs, nearshore Trunk braided river

extent of deposit facies

187

129, 130-1 Subsidence 269-70 rates 276, 279 Syria, clastic reservoirs Mesozoic 347 Palaeozoic 343

Tributary alluvial fan

Tryfan Fach Member 160, 186-9 depositional system 183-4 epiclastic interpretation 180-2 evidence for volcanism 178-80

description 28-9 interpretation 29-30 lithostratigraphical subdivision and relationships 20-4 Orange Sands 22-3 deposition 38-40 description 26-8 interpretation 28 palaeogeographic setting and depositional environments 19-20 palaeogeographical evolution 38-40 Red Sands 22-3 deposition 38, 40 description 35-6 interpretation 36-7 sedimentological characteristics 24-38 Silty Beds 22-3 deposition 40 description 34 interpretation 34-5

organization and distribution 174-8 geological framework 159-60 inter-eruptive sedimentation 187 sedimentation 185-6 terminology 158-9 Volcanic-associated iron formation 150-3, 154 Volcanism, stratigraphical record

Silver Sands 22-3 deposition 40 description 30-3 interpretation 33-4 stratigraphic framework

185-7 Volcano-tectonic uplift

Transition Series 22-3 description 37-8

183

interpretation Walther's Law 157, 187-9 Watson Formation, Pororari River 203-10

18-19

38

Yemen, Mesozoic clastic reservoirs 346-7