A sedimentological evaluation of the devonian-carboniferous boundary stratotype in southern Ireland

.S'edi~nenzam (;c,)logy, 34 (1983) 97 l lN

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l{Ise~ier Scientific Publishing Company, Amsterdam

Printed in The Netherlands

A S E D I M E N T O L O G I C A L EVALUATION OF THE D E V O N I A N - C A R B O N I F E R O U S B O U N D A R Y STRATOTYPE IN S O U T H E R N IRELAND

PIERS R.R. G A R D I N E R and IV()R A.J. MacCARTHY

(;eolo~,,ieal ,S'urcer of Ireland, 14 Hume Street, Dublin (lrelund) Department ,,t Geology, tSm:er~i(r ('olleee, Cork (Ireland) i Received April 20. 1981 : accepted November 8, 1982)

•~.BSTI~,A(" 1 Gardmer, P.R.R. and MacCarthy, I.A.J.. 1983. A sedimentological evaluation of the l)e~onian (arb, mil'erous boundary stratotype in southern Ireland. Sediment. Geol.. 34:~7 118. The stratotype section for the base of the British Dmantian (Devonian Carboniferous houndarxl is currently located at the interface of two marine formations at the Old tlead of Kinsale. This corresponds with the boundary of the LN and VI miospore subzone floral assemblages. The ~ucccssion immedialelx below the interface records an upward change from inner shelf into transgressive backbarrier >edimcnts, due t~, the initiation and progressive influence of a storm-dominated migrating microtidal barrier complcx. The formation junction is considered to reflect a non-sequence arising from subsequenl erosive shorcface retreat and destruction of the barrier superstructure, which was followed by apparenth sudden deepening and the deposition of "pro-delta" facies. Comparisons with other sections at this stratigraphical level in southern Ireland indicates that barrier construction with landward migration into backbarrier bays occurred during a destructive phase of deltaic development in latest Devonian times. ~\ rapid transgression associated with a cessation of sand supply then took place, resulting in variable re~orking of the underlying sequences and discontinuous shoreline retreat. This wits followed by the deposition of ~torm influenced calcareous shelf and inshore bay sediments over laterally diverse deha platform ~uccessions. With continued deepening a 'pro-delta" environment was established over the area m the earliest Carboniferous (VI miospore subzone), which thus records the relati,~elv unusual preserx:mon of barrier and backbarrier sequences in a regionally transgressive situation.

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unraxelling the sedimentology of the facies variations. This has essentially arisen due to the impetus of Prof. J.F.M. de Raaf and his research school, who highlighted the importance of detailed process orientated analyses in d e t e r m i n i n g local e n v i r o n m e n tal interpretations (De Raaf, 1970: De Raaf and Boersma, 1971: Kuijpers, 1971. 1972, 1975: Leflef, 1973a, b: Van Gelder, 1974: De Raaf et al., 1977). By integrating these results with other studies it has recently been shown thai the transgression was spasmodic in nature, and that the marine succession in this area reflects fixe major transgressive pulses ( G a r d i n e r and MacCarthy, 1980). In bulk terms the intervening sequences consist of a fine-grained depositional phase follo,~,,ed by an upward coarsening and shallowing sequence. The third of these transgressive pulses occurs at the base of a widespread, 6 60 m thick, "pro-delta' deposit I ( a s t l e Slate Member, Kinsale Formation). This overlies the laterally variable t'acics of a deltaic succession (Coomhola F o r m a t i o n ) deposited in a semi-enclosed setting IGardiner. 1975: G a r d i n e r and MacCarthy. 1981). Deposition of the ('o,mHlola F o r m a t i o n ( 4 0 - 8 0 0 m thick) was structurally influenced (MacCarthy and Gardiner, 1980), with the result that although there is evidence of delta progradation in the lower part where fluvially influenced sequences overlie low-energy subtidal deposits IVan Gelder, 1974), the upper part reflects complex facies variations arising from smaller scale barrier and shoreline fluctuations. The j u n c t i o n between these two deltaic formations at the Old Head of Kinsale (Fig. 1) was considered bv Nax.lor t 1966, p. 244) to be a " n a t u r a l transition", and taken by George et al. (1976) as the D e v o n i a n Carboniferous b o u n d a r y stratotype. However, both the regional ( G a r d i n e r and MacCarthy, 1981) and adjacent local relationships ( G r a h a m , 1975: Sleeman et al., 1978: Van Gelder and Clayton, 1978) suggest the possibility of a depositional break at the interface in this particular section. This was therefore evaluated in detail. The results are presented below, together with a brief comparison v
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Fig. t. Geological sketch map of southwest Ireland showing the position of the localities mentioned in the text. Those circled have yielded I.N and VI subzone miospore assemblages. All sections contain the lunction of the Coomhola Formation with the overlying Castle Slate Member (Kinsale Formation). l+ithological data sources as follows: 1, 2, 5 = authors, unpubl.: .,' - Naylor et al. 119771:4 Na,,Ior <19751: ,5= Reilly and Graham (1976); ,5. 7-Graham (1975): ,~', v Graham and Reillx ([9761: 11)= Naylor ~1966), Kuijpers (1972), this paper: 1 1 - N a y l o r and Higgs (19801, Mac('arthv. unpuhl.: 12 = Naylor 11969), MacCarthy, unpubl.; 1 3 - 2 3 - MacCarthy et al. (1978). Palynologicaldata sources: ] VanVeen(1981):2=VanderZwan(1980):3=Nayloretal. l'19771; 4= Naylor (19751: v , Keegan 11977): 9=Van Gelder and Clayton (1978): 10=George et al. (1976): 1 1 - Navlor and ttiggs (1980): 12 = Clayton et al. (1974): 1 3 - 1 6 Sleemanet al. (19781: IX, ]O, 22.23 Van der Zwan and Van Veen 19781.

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THE OLD HEAD OF KINSALE Strati,graphv The full sequence was described by Naylor (1966), Kuijpers (1972) subsequently providing further sedimentological details of the Devonian succession. The latter consists of the locally named Old Head Sandstone Formation (Coomhola Formation of Gardiner and Horne, 1972~ 1976). The uppermost 50 m of this Formation comprises parallel laminated or cross-stratified sandstones, sometimes scour based, interleaved with silt or sand lensed mudstones that comprise over 30% of the succession. Crinoid ossicles are common, and bioturbation is locally strongly de,,eloped (Naylor, 1966). However, because of access difficulties, no further details can be obtained except from the topmost 15 m on the west coast, where the contact with overlying dark grey claystones of the Castle Slate Member can be reached. George et al. (1976) provided a summary log of this part of the section, for which Navlor (1966) and Kuijpers (1972) gave no specific information. The results of a detailed examination of this section are presented in Fig. 2. f'acies anuhwis Navlor (1966) suggested that the Old Head Sandstone Formation represented deposition on a rapidly subsiding delta front platform, while Kuijpers 11972) favoured a lagoonal or interdistributary bay setting alternating with high-energy tidal environments for the upper part. The succeeding 60 m-thick Castle Slate Member was thought to have accumulated in a pro-delta setting (Naylor, 1966). A more refined depositional interpretation is however possible for the west coast interface section, with six successive facies being recognised within the accessibility limits (Fig. 2). Facies A. This is at least 15 m thick, although only the top 3.6 m is accessible. It is composed of sharp or erosively based greyish sandstones (60%,) interleaved with darker grey heterolithic units (Reineck and Wunderlich, 1968). The sandstones range in thickness from 5 to 110 cm, with the thinner units sometimes showing parallel lamination. Units thicker than 15 cm commonly show large-scale cross-stratification, with uniformly directed low angle sets up to 25 cm thick. Some contain siltstone or mudstone clasts arrayed along foresets or at the base of units. The heterolithic units are typically clay dominant, with delicate silt or sand laminae or discontinuous ripple forms less than 2 cm thick. The latter show opposed laminae, draping and pinch-and-swell, features characteristic of wave generation (Boersma~ 1970: De Raaf et al., 1977). The sparse palaeocurrent data indicate a transport direction towards the north-northeast, with wave-ripple crests trending east-northeast. Bioturbation is rare in the exposed section, although locally well developed in the lower part of the facies (Kuijpers, 1972). Brachiopod debris was noted in one sandstone.

102 The heterolithic units are thought to have been deposited in a low-energy wave-dominated environment. The delicate alternations of coarse and fine sediment, allied with a lack of current-ripple forms, suggests that both are suspension deposits, with the coarser levels due to fallout after storm stirring (Reineck and Singh, 1972: Kumar and Sanders, 1976). Subsequent wave action resulted in ripple bedforms. If the wave rippling took place under storm conditions, the waterdepth would have been in the range of 7-30 m according to the threshold curves of Levell (1980). In contrast, the thicker sandstone units clearly represent higher-energy traction current incursions, and probably reflect minor migrating sand sheets. The presence of intraformational clasts, allied with the sharp or erosively based nature of the sandstones~ suggests that they are the products of high-energy storm-surge conditions (cf. Bhattacharyya et al., 1980). In view of the apparent orthogonal relationship between the wave-ripple crests and palaeocurrent direction, such currents may have been wind induced, since significant sand transport can occur by this agency (Boggs, 1974; Swift et al., 1978). These depositional conditions point to a storm influenced shallow coastal or shelf setting (Reineck and Singh, 1973). In view of the interleaved association of the heterolithic units with tidal deposits lower in the Formation (Kuijpers, 1972)~ which suggests that these were lateral time equivalents resulting from migrating facies belts, an inner shelf setting seems likely for Facies A. This interpretation is supported by the occurrence of shelly debris and the crinoidal fragments recorded by Naylor (1966). Facies B. This 8.4 m thick facies is, in contrast, strongly bioturbated and lacks well-developed wave-generated features, The dominant rock type (90%) consists of claystones (0.2--2 cm thick) delicately interbedded with laminated coarser levels. Sparse weakly rippled or lenticular sandstone levels may be wave generated~ but internal structures are obscured by bioturbation. These heterogeneous units show a vertical sequential arrangement, occurring typically as CU sequences 0.6-2 m thick which in one case is capped by laterally impersistent sandstone. However, a thinner C U F U and an FU sequence also occurs. The sand-dominant levels are generally intensely bioturbated and contain thin (mm-size) laterally extensive clay or silt laminae (Fig. 3b). The remainder of the facies (10%) is composed of weakly laminated claystones 8-20 cm thick, No fauna was observed. The facies reflects relatively slow deposition from suspension in a low-energy setting where there was abundant biogenic activity. Various studies (e.g. Hayes, 1967; Gadow and Reineck, 1969; K u m a r and Sanders, 1976) have documented how storms can spasmodically introduce sand into low-energy shelf and nearshore environments, and a similar origin for the coarser levels is envisaged to explain the association of lithologies in this facies. In view of the limited amount of any wave activity, the environment was either protected or below effective wave base. De Raaf et al. (1977) have described very similar facies and considered that the sequential character was due to incipient shoal formation. It is however equally likely that the

Fig. 3. a. I)evonian-Carboniferous stratotype section at the Old Head of Kinsale showing the boundar\ contac! (open arrow) between the Old Head Sandstone (Coomhola) and Kinsale Formations. Note the irregular erosive basal contact of Facies C sandstones and lateral persistence of interbedded sandstones and clavstones in Facies D. Facies E sandstones are abruptly overlain by clavstones (Facies F, (astle Slate Member). Interval shown corresponds approximately to 11-17 m in Fig. 2. Solid arrow shows tile direction of stratigraphical younging, h. Lower part of bulk coarsening up unit near the top of Facies B ~howing interlaminated sandstone, siltstone and claystone. Note relatively intense bioturbation, gentle lensing of some of the coarser units and delicate sand laminae within sihstone and claystone unit~,, Taken at the 10.80 11.15-m interval of Fig. 2. c. Alternations of laminated claystone levels with thinly laminaled sandstones in the lower part of Facies D. Note the disruption of laminae by vertical burrows and the wave ripple forms capping the central sandy unit, 13.08 13.34-m interval of Fig. 2.

104

sequences here have resulted from temporal changes in storm action or sediment availability. Such depositional conditions occur in a variety of marginal and open marine environments. In the absence of any further diagnostic features no specific environmental allocation is possible. Facies C. This 1.1 m-thick erosively based unit consists of coarse-grained greycoloured sandstones, containing low-angle cross-stratified sets 20-60 cm high. These are overlain by a 35 cm-thick parallel-laminated coarse orange-coloured sandstone (Fig. 3a). The latter contains dispersed siltstone and sandstone clasts and scattered crinoidal fragments. Within the context of the low-energy setting of the underlying facies, such scouring and subsequent traction current deposition could have been caused by rip currents on the upper shoreface (Shepard et al., 1941; Vos, 1976), by washover action (Reinson, 1979), or by shoreface storm processes (Kumar and Sanders, 1976). The former appears unlikely, since there is no evidence of associated middle and upper shoreface deposits. A pure storm genesis is also thought improbable in the absence of any waning flow evidence and lack of suspension deposition (Stubblefield et al., 1975; Kumar and Sanders, 1976), although comparable internal stratification has been noted in shelf storm sands (Clifton, 1973: J.R. Boersma, pers. commun., 1981) and attributed to storm-wave surge action (Clifton, 1973). A washover origin is favoured in view of the abrupt incursion of this facies and the presence of low-angle and horizontal lamination, which are characteristic of such deposits (Schwartz, 1975). Facies D. This 1.7 m-thick sequence consists of parallel-sided greyish sandstones or siltstones interbedded with dark grey structureless claystones (Fig. 3c), individual beds being laterally uniform for at least 30 m. The claystones comprise some 16% of the facies and are 1-2.5 cm thick. The coarser units range in thickness from I to 21 cm, with the thicker beds containing internal laterally persistent clay laminae, rare mud intraclasts and ripple bedforms. The latter typically show opposed foresets, drape structures and curved lower boundaries, features indicative of wave generation (Boersma, 1970), and in some cases have wave rippled upper surfaces with wavelengths of less than 10 cm. The facies consists of two CU sequences, 0.7 and 1 m thick, each of which is characterised by a progressive upward reduction in the thickness and number of clay intervals. The top of the lower CU sequence is capped by a thin wave-rippled sandstone. Bioturbation is often present at the claystone/sandstone interfaces and in thin sandstone levels. Although the facies resembles Facies B, it differs in having far less claystone content, notably less bioturbation, and evidences a significant amount of wave-built cross-lamination. The facies reflects periods of low-energy mud accumulation interrupted by influxes of coarser sediment deposited from suspension under the influence of wave action. The organic reworking at the clay-sand interface suggests that the coarser influxes were often preceded by a time lapse, and hence not genetically related to the claystone units. A significant feature is the presence of laterally persistent clay

laminae within the sandstones. This indicates that the thicker sandstones merely represent packets deposited under conditions of increased availabilit\ ~l' sand supply, which was still, however, episodic in nature. The resemblance of this facies to Facies B suggests that the sand influxes here have also resulted from storm action. with the higher level of wave energy indicative of a shallower depositional ~elting. Since the facies overlies an inferred washover unit, it is considered to haxc been deposited in a sheltered shallow backbarrier setting where quiet-water mud ~.lCCtllllUlation was the norm. During storms sand was either washed over in suspension or blown in from backbarrier flats (cf. Barwis, 1976), while wave reworking took placc during the waning stages of storms or in rough-weather conditions. Comparable facies have been described by Davies et al. (1971) from lagoonal sequences, v~ith the coarser interbeds in this case consisting of barrier-derived material. The ( L sequences are thought to reflect progressive shallowing (of. Vos, 1977) rather than migrating shoals in view of the lateral uniformity of the facies, Facies k'. This can only be reached with difficulty at the top of a small ca\e, where it is heavily weathered. Details are therefore limited, it consists of an 1N era-thick parallel-sided yellowish-grey coarse-grained sandstone with sharp upper and lower boundary contacts. Internal structures consist of poorly expressed parallel or very low angle cross-lamination within which discontinuous silt or clay streaks arc Iocall,¢ visible, especially in the upper few centimetres of the unit. No f,:lUlla or bioturbation was noted. Deposition of the facies occurred under relatively high energy traction currcnt conditions. Comparable parallel and low-angle laminae are typical of beach foreshores formed during swash and backswash (Clifton, 1969: Clifton et al., 1971). Similar fealures are also evinced by washover fan deposits (Schwartz, 1975: ftobda,, and Jackson, 1979). Both possibilities seem plausible for this facies. Hov, evcr, laminated sands deposited during the destructive phase of the North Santee l)cha, South Carolina (Stephens et al., 1976), also resemble this facies, and it is possible that Facies E may represent a comparable reworkmg phase. Some re~orking is certainly suggested by the discontinuous finer laminae in the upper part ~1" the facies. Facies F. This consists of dark grey claystones with a sharp basal contact. Fine sandstone laminae occur in the lowest 0.7 m (Fig. 2), above which arc uniform claystones with impersistent sandy streaks and sparse phosphatic nodules (Fig. 3al. The coarser laminae are interpreted as suspension deposits laid down in a low energy mud environment below wave base. A storm-sand genesis appears likely for thcse coarser incursions, although Naylor (1966) recognised small-scale sill-based graded units which may be of turbidite origin. There is little vertical lithological variation in the lower part of this 60 m-thick member, which has been interpreted as reflecting a pro-delta (Naylor, 1966) or deep-water setting (Van Gelder and Clayton, 1978). Although comparable facies are known from bay centres (cf. Kanes, 1970}, the thickness and relative lithological uniformity favour a pro-delta type environment.

106

The local depositional model With the exception of Facies C, Facies A - E show no obvious boundary discontinuities and have a common though variable motif of storm and wave influence, features which are inferred to reflect progressive genetically linked environmental changes. On this basis, Facies B is considered to have been deposited in a protected shallow-water embayment rather than below normal wave base, with the sequence recording the following environments: (1) Facies A, inner marine shelf with deposition above wave base; (2) Facies B, shallow sheltered bay: (3) Facies C, backbarrier washover; (4) Facies D, shallowing protected backbarrier bay or lagoon; (5) Facies E, beach or backbarrier washover. In part reworked (?); and (6) Facies F, 'pro-delta'. Upward shallowing on this scale can be caused either by prograding fluvially influenced linear or deltaic shorelines, or by migrating barrier systems. The sequence A E does not however conform with characteristic prograding shoreline sequences (cf. Reineck and Singh, 1973; Harms et al., 1975), and in view of a progressive barrier influence and absence of fluvial evidence the depositional site is considered to have been associated with b a r r i e r / b a r systems rather than a land-backed shoreline. In assessing possible barrier models critical elements are the type of barrier and its degree of lateral stability (Heward, 1975). The presence of repeated storm washover sediments allied with wave action is indicative of microtidal barriers (Hayes, 1975), which are typically long linear features with widely spaced inlets. The vertical facies changes reflect laterally shifting environments and hence some degree of barrier movement, which was shoreward directed in view of its encroaching influence into a protected embayment. A transgressive barrier model therefore seems appropriate for this part of the sequence. This is envisaged as having been initiated in a storm dominated marine shelf setting with a low or insignificant tidal range, with the Old Head of Kinsale located on the landward side of subsequent barrier formation. The barrier was of microtidal type, which apparently migrated laterally into a backbarrier bay or lagoon with consequent shallowing in the studied section. On the assumption that wave trends (Fig. 2) paralleled the slope strike, the barrier had an approximate east-west orientation. The succession thus records the influence of barrier initiation and the lower part of its transgressive backbarrier infill. In this context Facies E is considered to be a washover deposit rather than a remnant foreshore. In marked contrast, Facies F shows an abrupt and permanent cessation in coarse clastic supply, allied with deposition in a stable low-energy (pro-delta) setting, This radical change implies a sudden deepening, which should result in the virtually full preservation of the underlying barrier sequence (Kraft, 1971). The absence of a capping barrier superstructure above Facies E therefore merits discussion. There are two likely possibilities: (1) In place "drowning" before further barrier migration took place, which would

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have resulted in a shift of the barrier shoreline to a more inshore position (Sanders and Kumar, 1975). In this case the backbarrier sequence should be capped bx shoreface or inner shelf deposits (Hew'ard, 1975: Rampino and Sanders, 1~81). (2) Barrier shoreface erosion and retreat. Depending on the rate of relativc ri~c of ~,ea level and degree of sediment supply, this would cause partial or total dcstvuctio~-J of the harrier sequence (Swift, 1968: Swift et al.. 1971). The absence of shoreface or inner-shelf deposits above Facies E renders (he drowning mechanism alone unlikely. Evidence from Facies E is inconclusixe, but suggests some reworking prior to 'pro-delta' deposition. On this basis some degree of erosive shoreline retreat with barrier destruction is thought to have taken placc, v~ith the interface between Facies E and F marking a non-sequence. Since the reducti,m of sediment rates and shoreface erosion are dependent variables, both probabl 5 had an influence on the non-preservation of the barrier superstructure. It is, howcxer, unlikely that there was continuous shoreface retreat, since reworking during the passage of the breaker zone would have destroyed most if not all of the backbarricr sediments as well (Swift, 1968: Rampino and Sanders, 1981). It is therefore suggested that the Old Head of Kinsale section reflects barrier formation followed bx a period of shoreface erosion when sand supply was insufficient for barrier n~)urishment. Accelerated deepening subsequently occurred, causing a landward.jump of the surf zone over the backbarrier bay ("stepwise retreat": Sanders and Kumar. 1075: Rampino and Sanders, 1981) with resulting preservation of the backbarricr ~cq uence. This model differs from the "end member" transgressive harrier model of Reinson (1979) in the absence of recognised tidal flat, flood-tidal delta and marsh sequences. The absence of the first two facies may reflect an insignificant tidal range and too much wave action (Reineck, 1972), or merely subsequent erosion: the studied section is too limited for meaningful deductions to be made. The prescr~cd part of the backbarrier sequence was presumably located too far offshore to bc associated ,aith marshes. Two aspects concerning the proposed model cannot be resolved on a local ~calc. Firstly, the origin of the barrier. This could have developed by' the emergence of offshore bars or by longshore growth of a spit (Hoyt, 1967), both of which would give the same initial shelf/embayment separation. Secondly, whether the barrier itself was migrating or merely widening. The stationary barrier of Padre Ishmd, Texas, for example, is currently increasing its width by migration into the adiacent lagoon (Dickinson et al., 1972), with the resulting development of a ""transgressive" backbarrier sequence on its landward side. ( ' O M P A R A T I V I : ~ S E C T I O N S IN S O U T H W E S T

IRELAND

Known sections showing the junction of the Castle Slate Member with the underlying Coomhola Formation occur in an area of some 150 km by 40 km m southwest Ireland (Fig. 1).

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Description Published information, with the exception of Graham (1975), is of a sparse stratigraphic nature. However, preliminary observations by the present authors indicate that five general facies sequences can be distinguished at this level Isee Fig. 5): (1) Where there is a sharp boundary contact between thick sandstones and the Castle Slate Member. Here the uppermost sandstone unit of the Coomhola t:ornmtion is typically 3 6 m thick, shows parallel or low-angle cross-stratification and has ~hellv material or intraformational clasts contained within the upper part or on the upper surface. Where fragmental debris is absent organic reworking has been noted. The underlying 5-10 m of strata contain thin sandstones and variably bioturbated facies similar to those described above at the Old Head of Kinsale. Rootlet lexels have been observed at one locality (Graham, 1975). This facies sequence is restricted to the present south coast sections (Fig. 1: Ioc. 3, 4, 6 8 and 22). (2) Sequences with a sharp formation boundary contact, but with a relatively thin (3 60 era-thick) capping sandstone at the top of the Coomhola Formation. The ()ld Head of Kinsale section above Facies A (Fig. 2) is typical of this facies sequence, but there is lateral variability in sand content, sequential organisation and degree of bioturbation between sections, This facies sequence occurs at the south c o a s t and in the north Cork Harbour area (Fig. 1: loc. 5, 10 and 13-16). (3) This typically consists of thick low-angle cross-stratified sandstones, in some cases capped by a thin pebbly sandstone containing intraformational clasts, which are sharply or erosively succeeded by a 0.5 5 m thick unit of shaly hic~clastic limestones variably interbedded with claystones. These calcareous units often show a bulk FU character and pass transitionally upwards by increased clay content into thick claystones (Castle Slate Member). Allochems include disarticulated brachiopod and crinoidal debris, which are usually arrayed as thin parallel bedded units. This sequence occurs around south Cork Harbour and further east (Fig. 1: loc. 11, 12 and 17 21) and also at Ballycrovane Harbour in the west (Fig. I: loc. I). (4) This differs from facies sequence 3 only in the sub-carbonate part, which consists of greyish, lenticular, scour-based cross-stratified sandstone units (2 5 m-thick) interleaved with plant bearing 0.5-5 m-thick yellowish silty claystoncs that locally contain poorly developed pedogenic carbonate levels. This facies is capped b~ a 30 cm-thick laminated pebbly sandstone, which has a sharp contact wilh the overlying shaly limestones. This facies sequence is seen at Ardmore (Fig. 1: Ioc. 23). (5) Sequence with a transitional contact. These consist of mud-dominant bioturbated strata containing thinly bedded heterolithic sandstones and sandstone laminae which pass upwards by progressive loss of sandstone over a 12 m interval into claystones of the Castle Slate Member. This sequence occurs at North Bantr\ (Fig. 1: loc. 2). For the majority of the sections miospore data are available (Fig. 1). showing that

110 the boundary between the LN and VI miospore subzones of Clayton et al. (1974) is apparently at or close to the formation interface. In all dated sections of facies sequences 1 and 2 along the south coast the sample gap includes the boundary contact, and in several cases is only separated by the uppermost sandstone unit of the Coomhola Formation (George et al.. 1976: Naylor et al., 1977). The base of the Castle Slate Member is diachronous, becoming older when traced northeastwards from the Old Head of Kinsale (Fig. 4), but for the majority of these sections the sample gap is too wide to permit an accurate evaluation. This same northward diachronicity is also evident along the west coast, as the L N / V I miospore subzone boundary occurs at the formation interface at North Dunmanus Bay (Fig. 1. loc. 3: Naylor et al., 1977), but within the Castle Slate Member at Ballycrovane (Fig. 1, Ioc. 1: Van Veen, 1981) and North Bantry (Fig. 1. loc. 2; Van der Zwan, 1980).

Interpretation The interpretation of the Toe and Galley Head sections (Fig. 1: loc. 6 and 7) by G r a h a m (1975) indicates that facies sequence 1 reflects transgressive barrier deposits abruptly overlain by mudstone deposited in a low energy environment. The evidence for physical or organic reworking of the uppermost sandstone is indicative of a non-sequence prior to 'pro-delta' deposition. Similar conclusions regarding the interface were reached by Graham (1975), and also by Van Gelder and Clayton (1978) for the West Seven Heads section (Fig. 1; loc. 9). By analogy with the Old Head of Kinsale section (this paper), facies sequence 2 is thought to indicate a storm-influenced backbarrier bay setting with the possibility of a depositional break in some sections before subsequent mud accumulation in a low-energy environment. The sand-dominant section at Marino Point (Fig. 4) however also shows similarities with modern sand fiats (cf. Yeo and Risk, 1981) and may reflect an inshore landbacked location rather than a backbarrier bay. The lower part of facies sequence 3 is interpreted as barrier sands, with the capping pebbly levels representing lag deposits resulting from reworking during transgression. The succeeding calcareous units are thought to reflect a low-energy shelf or open-bay setting affected by periodic storm action, which resulted in winnowing and reworking of bioclastic material into thin sheets on the mud substrate. Storm-induced currents may also have been responsible for net sediment transport. Progressive deepening is indicated by the upward transition into the pro-delta' facies. Facies sequence 4 is considered to intially record a fluvially influenced succession in which the finer-grained overbank deposits accreted at sufficiently low rates for pedogenesis to occur (Leeder, 1975). A coastal-plain setting has been proposed (Gardiner and MacCarthy, 1981). The capping pebbly sandstone is interpreted as a transgressive sand sheet in view of the environmental juxtaposition recorded by the overlying shallow water calcareous facies, which is similiar to those seen in facies '

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sequence 3. In contrast, facies sequence 5 shows a gradational transition from ba\ or lagoonal sediments (Jones, 1974) into pro-delta facies with no evidencc of a depositional hiatus. FIlE PROVIN('IAL MODEL

The general palaeogeography of southern Ireland prior to deposition of Ihc (astle Slate Member is shown in Fig. 5. This consists of a deltaic complex backed to thc north and northeast by extensive alluvial plains and passing southwards into an open marine shelf. The disposition of the facies sequences (Figs. 5 and 61 gives further insight into the depositional patterns of the delta margin at this stratigraphic level. Where facies sequence 1 occurs along the present south coast lhere is no evidence of diachronism between sections, and the area is therefore interpreted as marking a line of discontinuous bars or barriers. The lateral association of this sequence with facies sequences 2 and 3 indicates a complex of barriers migrating landwards into backbarrier bays. A more proximal coastal plain location is indicaled bv facies sequence 4, while a stable bay setting unaffected b,¢ barrier migration is shown by facies sequence 5. The inferred lateral facies relationships between Galley Head and Ardmore :ire shown in Fig. 6. Two contrasting transgressive phases are recognised. A destructive delta model is proposed for the first phase, during which peripheral barriers migrated shorewards into the backing shallow marine embayments or lagoons under

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storm and wave attack. This setting is comparable to the destructive phase recognised in a number of deltas (e.g. the Holocene Rh6ne Delta, Oomkens, 1970: the St. Bernard Delta, Coleman and Gagliano, 1964). The second phase was initiated by a sudden transgressive pulse associated with a cessation of siliciclastic supply. This resulted in variable reworking and truncation of a variety of earlier sequences by the landward advancing surf zone, with the preservation of backbarrier deposits indicative of a stepwise advance pattern during a rapid rise in sea level at this time (Rampino find Sanders, 1981). Relative sea-level stability then ensued, with the establishment of shelf or open-bay conditions over the area. From the biostratigraphical evidence it seems that the subsequently deposited shaly limestones passed laterally into mudstones further inshore in the Currabinny Marino Point area (Fig. 6). Comparable mudstone sequences cover the low- to moderate-energy shallow shelf off Suriname (Augustinus, 1980), which seems a possible modern analogue in this context. This interpretation explains the otherwise anomalous situation of the base of the Castle Slate Member being older in the palaeogeographicall_x nearer shore sections (Fig. 4), but also means that the previous categorisation of the Member as all being of pro-delta facies (MacCarthy and Gardiner, 1980) is an oversimplification. The patchy distribution of the shaly limestones may reflect local bathvmetric variations or areas of enhanced storm scouring, as on the southern Ncx~ England shelf (Twichell et al., 1981), with the possibility of a continued depositional break in limestone-absent sections such as the Old Head of Kinsale (Fig. 6). I:urther deepening took place at the base of the VI miospore (earliest Carboniferous), causing fin upward transition into a widespread low energy ('pro-delta') claystonc succession which subsequently overstepped the thin shelf limestones at Ardmore ( Fig. 6). On the basis of this provincial model, it seems more probable that the backbarrier sequence at the Old Head of Kinsale resulted from a migrating rather than a static barrier. Further detailed studies of the underlying successions along the present south coast are needed before the origin of the late Devonian barriers can be resolved. BIOSTRAT1GRAPHICALIMPLICATIONS The boundary between the LN and VI miospore subzones of Clayton et al. (1974) is currentlv recognised as the biostratigraphical base of the Carboniferous of the British Isles (George et al., 1976: Clayton et al.. 1978). The location of the two critical miospore samples which delimit this boundary in the stratotype section is shown in Fig. 2. The 2.25-m gap includes the inferred non-sequence. Since part of either or both sub-zones could therefore be missing and independent evidence concerning the distribution of other fossils across this stratotype boundary is lacking (George, 1978), this section would seem suspect for stratotype purposes, There are several sections where the L N / V I sub-zonal boundary occurs in the

114 Castle Slate Member with no apparent depositional break (Fig. 4). One of these may prove more suitable for a stratotype section. The section at Ballycrovane Harbour (Fig. 1, loc. 1) may be particularly attractive in this context, since conodont assemblages (Ducharme, 1968) have been obtained from limestones within the basal metre of the Castle Slate Member. These assemblages approximate in age to about the Devonian-Carboniferous boundary (Matthews and Naylor, 1973) and offer possibilities of biostratigraphical correlation between different fossil groups. CONCLUSIONS The Old Head of Kinsale section shows an upward transition from inner marine shelf to a shallowing backbarrier bay setting in the late Devonian. This change was due to the initiation and progressive influence of a storm dominated microtidal barrier, which was either transgressing or undergoing "in-situ" widening. Subsequent shoreface erosion was associated with diminished or absent detrital input and resulted in destruction of the barrier superstructure. Preservation of the backbarrier sediments is thought to reflect a relatively rapid rise in sea level at this time, which caused the shoreline to overstep the backbarrier bay in a landward direction. Claystone deposition took place in a low-energy 'pro-delta' environment in earliest Carboniferous times on top of the backbarrier succession, the interface marking a non-sequence. By reference to laterally equivalent sections in southern Ireland a provincial depositional model with two differing transgressive phases is recognised. The first was a destructive phase of delta development, during which there was storm and wave influenced transgressive barrier migration into backbarrier bays. A rapid rise in sea level associated with a cessation of coarse clastic supply marked the initiation of the second transgressive phase, with variable erosion and local evidence of discontinuous shoreface retreat. Storm-influenced muddy shelf or open-bay conditions were then established during a period of relative sea-level stability. This was followed by progressive deepening and further transgression in the earliest Carboniferous (VI miospore subzone), resulting in the widespread deposition of a 'pro-deha' clay blanket. The area thus records the relatively unusual preservation (Ryer, 1977: Reinson, 1979) of barrier and backbarrier sequences in a regionally transgressive situation. ACKNOWLEDGEMENTS It is with great pleasure that P.R.R.G. acknowledges the stimulating influence of the late Prof. J.F.M. de Raaf on his work in southern Ireland. Grateful thanks are also due to Dr. J.R. Boersma and the referees, whose perceptive comments resulted in significant manuscript improvements. We thank the continued forebearance of our respective wives during this study, and Nicola Gardiner and Anne Norberg for

115

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l.eflef, D.. 1973b. A change in the rock types associated with the approaching shoreline of the ()ld Red Continent, South of Cork, Ireland. Geol. Mijnbouw, 52: 335-350. Levell. B.K., 1980. Evidence for currents associated with waves m Late Precambrian shelf deposits l'ronl Finumark. North Norway. Sedimentology, 27: 153-166. MacCarthy, IA.J. and Gardiner, P.R.R., 1980. Facies changes in the Upper Devonian and l o ~ e r Carboniferous of south Cork, Ireland A re-assessment. Geol. Mijnbouw, 59:65 77. MacCarthy, I.A.J.. Gardiner. P.R.R. and Home. R.R.. 1978. The Lithostratigraphy of the Devonian Earlx Carboniferous succession in parts of Counties Cork and Waterford, Ireland Bull. Geol. Stir\. h . 2: 265 3O5. 'datthex~s, S.(. and Naylor, D., 1973. Lox~er Carboniferous couodolll faunas from south-west Ireland. Paleontology, 16:335 350. Naylor. D., 1966. The Upper Devonian and Carboniferous geology of the Old Head of Kinsale. (<,. { o r k Sci Proc. R. Dublin Sot., 2A: 229 249. Navlor. D., 1'069. Facies change in the Upper l)cvonian and Lower Carboniferous rocks t>l ~nuthcrn Ireland. Geol. J., 6: 307-328. Naylor, I).. 1975. Upper Dewmian Lower Carboniferous stratigraphy ahmg the south coasl ~1 Dunnlanus Bav. ('o.(.ork. Proc. R. lr. Acad..75B: 317 337. Naylor, 1). and Higgs, K., 1980. The geology of the coastline east of Kinsale Harbour, Count~ Cork. Bull. (;-eol. Surx. lr., 2:371 388. "qavlor, I)., tliggs, K. and Boland, M.A., 1977. Stratigraphy of the north flank of the Dunmanus s\ ncline; West Cork Bull. Geol. Surv. Ir., 2:143 157. ~)omkens, E., 1970. Depositional sequences and sand distribution in the postglacial Rhhnc delta ctmlplex. In: J.P. Morgan and R.H. Shaver (Editors), Deltaic Sedimentation Modern and Ancient. Soc Icon. Paleontol. Mineral., Spec. Publ., 15: 198-212. Rampino, M.R. and Sanders, J.E., 1981. Evolution of the barrier islands of southern Long Island, Ne~ York. Sedimentology, 28:37 47. Reilly, ],A. and Graham, J.R., 1976. The stratigraphy of the Roaringwater Bay area of south-v, esl Count\ Cork. Bull. Geol. Surv. It., 2: 1-13. Reineck, H.-E., 1972. Tidal fiats. In: J.K. Rigby and W.H. Hamblin (Editors), Recognition of Ancienl Sedimentar)~ Environments. Soc. Econ. Paleontol. Mineral.. Spec. Publ., 16: 147-. 159. Reineck, H.-E. and Singh, I.B., 1972, Genesis of laminated sand and graded rhythmites in storm-,and layers of shelf mud. Sedimentology, lg: 123 128. Reineck, H.-F. and Singh, 1.B., 1973. Depositional Sedimentary Enviromnents with Refers'nee to Ferrigenous Clastics. Springer, Berlin, 439 pp. Reineck, H.E. and Wunderlich, F., 1968. Classification and origin of flaser and leniicular bedding. Sedimentology, 11: 99-104. Reinson, G.E., 1979. Barrier Island Systems, In: R.G. Walker (Editor), Facies Models. (}eoscicnce Canada, Reprint Set., 1: 57-74. Rver, T.A., 1977. Patterns of Cretaceous shallow-marine sedimentation, Coalvill and Rockport arca~, Utah. Geol. Soc. Am. Bull., 88:177 188. Sanders, J,E. and Kumar, N., 1975. Evidence of shoreface retreat and in-place "'drowning" ~l'f Fire Island, Ne,* York. Geol. Soc. Am. Bull., 86:65 76. Schwartz, R.K.. 1975. Nature and genesis of some washover deposits, U.S. Army Coastal Eng. Rex ( ent., Tech. Mere., 61 : 1-69. Shepard. F.P., Emery, K.O. and La Fond, R., 1941. Rip currents, a process of geological importanLc..I. Geol., 49: 337-369. Sleeman, A.G., Reilly, T.A. and Higgs, K., 1978. Preliminary stratigraphy and palynology of five sections through the Old Head Sandstone and Kinsale Formations (Upper Devonian to Lower Carboniferous on the wesl side of Cork Harbour. Bull. Geol. Surv. Ir., 2:167 186. Stephens, D.G, Van Nieuwenhuise. D.S., Mullin, P., Lee, C. and Kanes. W.H., 1976. Destructive phaxe ~)f

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deltaic development: North Santee River Delta. J. Sediment. Petrol., 46: 132-144. Stubblefield, W.L., Lavelle, J.W., Swift, D.J.P. and McKinney. T.F., t975. Sediment response to the present hydraulic regime on the central New Jersey shelf, J. Sediment. Petrol., 45:337 358. Swift, D.P.J.. 1968. Coastal erosion and transgressive stratigraphy. J. Geol., 76:444 456. Swift, D.J.P., Sanford, R.B., Dill, C.E. and Avignone, N.F., 1971. Textural differentiation on the shoreface during erosional retreat of an unconsolidated coast, Cape Henry to Cape Hatteras, western North Atlantic shelf. Sedimentology, 16: 221-250. Swift, D.P.J., Sears, P.C., Bohlke, B. and Hunt, R., 1978. Evolution of a shoal retreat massif, North Carolina shelf: inferences from areal geology. Mar. Geol., 27: 19-42. Turner, J.S., 1952. The Lower Carboniferous rocks of Ireland. Liverpool Manchester Geol. J,, 1: 113 147. Twichell, D.C., McLennen, C,E. and Butman, B., 1981. Morphology and processes associated with the accumulation of the fine grained sediment deposit on the southern New England shelf. J. Sediment. Petrol., 51 : 269-280. Van der Zwan, C.J., 1980. Aspects of Late Devonian and Early Carboniferous palynology of southern Ireland. 111. Palynology of Devonian-Carboniferous transition sequences with special reference to the Bantry Bay area, Co. Cork. Rev. Palaeobot. Palynol., 30: 165-286. Van der Zwan~ C.T. and Van Veen, P., 1978. The Devonian-Carboniferous transition sequence in Southern Ireland: integration of Paleogeography and Palynology. Palinologia, 1:469 479. Van Gelder, A., 1974~ Sedimentation in the Marine Margin of the Old Red Continent, South of Cork, Ireland. Ph.D. Thesis, Univ. Utrecht, Utrecht, 76 pp. Van Gelder, A. and Clayton, G., 1978. New data on Early Dinantian (Early Carboniferous) stratigraphy and sedimentation in south Cork, Ireland. Geol. Mijnbouw, 57:25 32. Van Veen, P.M., 1981. Aspects of Late Devonian and Early Carboniferous palynology of Southern Ireland. IV. Morphological variation within Diducites, a new formgenus to accomodate camerate spores with cwo-layered outer wall. Rev. Palaeobot. Palynol., 31:261 287. Vos, R.G., 1976. Observations on the formation and location of transient rip currents. Sediment. Geol., 16: 15-19. Vos. R.G., I977. Sedimentology of an Upper Palaeozoic river, wave and tide influenced delta system in southern Morocco. J. Sediment. Petrol., 47: 1242-1260. Yeo, R.K. and Risk, M.J., 1981. The sedimentology, stratigraphy and preservation of intertidal deposits in the Minas Bay System, Bay of Fundy. J. Sediment. Petrol., 51: 245-260.