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

Barremian (/) ::::l a::O

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Berriasian

Tithonian

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Oxfordian

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Callovian

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Evaporite

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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

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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 l