Recent Advances in Lower Carboniferous Geology (Geological Society Special Publication No. 107)

Recent Advances in Lower Carboniferous Geology

Geological Society Special Publications Series Editor A. J. FLEET

G E O L O G I C A L SOCIETY SPECIAL P U B L I C A T I O N NO. 107

Recent Advances in Lower Carboniferous Geology

EDITED BY PETER STROGEN, IAN D. SOMERVILLE & GARETH LL. JONES Department of Geology, University College, Dublin, Ireland

1996 Published by The Geological Society London

THE G E O L O G I C A L SOCIETY The Society was founded in 1807 as The Geological Society of London and is the oldest geological society in the world. It received its Royal Charter in 1825 for the purpose of 'investigating the mineral structure of the Earth'. The Society is Britain's national society for geology with a membership of around 7500. It has countrywide coverage and approximately 1000 members reside overseas. The Society is responsible for all aspects of the geological sciences including professional matters. The Society has its own publishing house, which produces the Society's international journals, books and maps, and which acts as the European distributor for publications of the American Association of Petroleum Geologists, SEPM and the Geological Society of America. Fellowship is open to those holding a recognized honours degree in geology or cognate subject and who have at least two years' relevant postgraduate experience, or who have not less than six years' relevant experience in geology or a cognate subject. A Fellow who has not less than five years' relevant postgraduate experience in the practice of geology may apply for validation and, subject to approval, may be able to use the designatory letters C Geol (Chartered Geologist). Further information about the Society is available from the Membership Manager, The Geological Society, Burlington House, Piccadilly, London WlV 0JU, UK. The Society is a Registered Charity, No. 210161. Published by The Geological Society from: The Geological Society Publishing House Unit 7 Brassmill Enterprise Centre Brassmill Lane Bath BA1 3JN UK (Orders: Tel. 01225 445046 Fax 01225 442836)

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Contents

Preface

Mineralization, hydrocarbons and diagenesis JOHNSTON, J. D., COLLER, D., MILLAR, G. & CRITCHLEY, M. F. Basement structural controls on Carboniferous-hosted base metal mineral deposits in Ireland SHEARLEY, E., REDMOND, P., KING, M. & GOODMAN, R. Geological controls on mineralization and dolomitization of the Lisheen Zn-Pb-Ag Deposit, Co. Tipperary, Ireland HOLLIS, C. & WALKDEN, G. The use of burial diagenetic calcite cements to determine the controls upon hydrocarbon emplacement and mineralization on a carbonate platform, Derbyshire, England VEALE, C. & PARNELL, J. Metal-organic interactions in the Dinantian Solway Basin, UK: inferences for oil migration studies Carbonate buildups and Waulsortian mud-mounds PICKARD, N. A. H. Evidence for microbial influence on the development of Lower Carboniferous buildups AHR, W. M. & STANTON, R. J. JR. Constituent composition of Early Mississippian carbonate buildups and their level-bottom equivalents, Sacramento Mountains, New Mexico KIRKBY, K. C. & HUNT, D. Episodic growth of a Waulsortian buildup: the Lower Carboniferous Muleshoe Mound, Sacramento Mountains, New Mexico, USA JEFFERY, D. L. & STANTON, R. J. JR. Biotic gradients on a homoclinal ramp: the Alamogordo Member of the Lake Valley Formation, Lower Mississippian, New Mexico, USA SOMERVILLE, I. D., STROGEN, P., JONES, G. LL. & SOMERVILLE, H. E. A. Late Vis6an buildups at Kingscourt, Ireland: possible precursors for Upper Carboniferous bioherms RODRiGUEZ, S. Development of coral reef-facies during the Vis6an at Los Santos de Maimona, SW Spain Siliciclastic rocks MORENO, C., SIERRA, S. & SAEZ, R. Evidence for catastrophism at the FamennianDinantian boundary in the Iberian Pyrite Belt MAGUIRE, K., THOMPSON, J. & GOWLAND, S. Dinantian depositional environments along the northern margin of the Solway Basin GRAHAM, J. R. Dinantian river systems and coastal zone sedimentation in northwest Ireland Carbonate platforms and ramps RIZZI, G. & BRAITHWAITE, C. J. R. Cyclic emersion surfaces and channels within Dinantian limestones hosting the giant Navan Zn-Pb deposit, Ireland HORBURY, A. D. & ADAMS, A. E. Microfacies associations in Asbian carbonates: an example from the Urswick Limestone Formation of the southern Lake District, northern England

vii

1 23 35

51

65 83 97 111

127 145

153 163 183

207 221

vi

CONTENTS

GALLAGHER, S. J. The stratigraphy and cyclicity of the late Dinantian platform carbonates in parts of southern and western Ireland KELLY, J. G. Initiation, growth and decline of a tectonically controlled Asbian carbonate ramp: Cuilcagh Mountain area, NW Ireland STROGEN, P., SOMERVILLE, I. D., PICKARD, N. A. H., JONES, G. LL. & FLEMING, M. Controls on ramp, platform and basinal sedimentation in the Dinantian of the Dublin Basin and Shannon Trough, Ireland VANSTONE, S. The influence of climatic change on exposure surface development: a case study from the Late Dinantian of England and Wales

Basinal facies GURSKY, H.-J. Siliceous rocks of the Culm basin, Germany BELKA, Z., SKOMPSKI, S. & SOBON-PODGORSKA, J. Reconstruction of a lost carbonate platform on the shelf of Fennosarmatia: evidence from Vis~an polymictic debrites, Holy Cross Mountains, Poland NAYLOR, D., SEVASTOPULO, G. D. & SLEEMAN, A. G. Contemporaneous erosion and reworking within the Dinantian of the South Munster Basin REES, J. G., CORNWELL, J. D., DABEK, Z. K. & MERRIMAN, R. J. The Apedale Tufts, North Staffordshire: probable remnants of a late Asbian/Brigantian (Pla) Dinantian volcanic centre Faunas, floras and biostratigraphy MAKHLINA, M. KH. Cyclic stratigraphy, facies and fauna of the Lower Carboniferous (Dinantian) of the Moscow Syneclise and Voronezh Anteclise RUKINA, G. A. Sequence biostratigraphy of the Tournaisian-Lower Vis~an rocks of the Russian Platform JONES, G. LL. & SOMERVILLE, I. D. Irish Dinantian biostratigraphy: practical applications LEBEDEV, O. A. Fish assemblages in the Tournaisian-Vis~an environments of the East European Platform IVANOV, A. The Early Carboniferous chondrichthyans of the South Urals, Russia HARPER, D. A. T. & JEFFREY, A. L. Mid-Dinantian brachiopod biofacies from western Ireland SMITH, J. A palynofacies analysis of the Dinantian (Asbian) Glenade Sandstone Formation of the Leitrim Group, northwest Ireland Index

239 253 263

281

303 315

331 345

359 365 371 387 417 427 437

449

Preface To those of us who were students during the 1950s in the UK and Ireland the term Dinantian conjured up a picture of the Avon Gorge at Bristol, the great limestone uplands of Ingleborough in Yorkshire or the Burren of County Clare. To those who have since wandered into the petroleum industry the Lower Carboniferous was, and largely remains economic basement. Within this volume both viewpoints are, we hope, shown to be both parochial and commercially erroneous! The first European Dinantian Environments (EDE) Conference was held in Manchester in 1984. Speakers and attendees at the EDE II Conference, held at University College Dublin in September 1994, upon which this volume is based, came from points as far apart as western Ireland, Russia to the foot of the Ural Mountains, and the southern tip of Iberia. So much for parochialism. Almost all European countries with significant Dinantian strata were represented- Russia, Poland, Germany, Belgium, Spain, the United Kingdom, and Ireland. There was a most welcome addition to EDE - from the USA, as a result of which a new title for this decennial meeting has to be found! Many of the great steps forward in geology during the nineteenth century were made by geologists who had travelled widely and had seen rocks of very different character. These were demonstrably, because of the increasing success of palaeontologists in creating reliable biozonation schemes, of the same age. Today, despite the ease of modern travel, we are all too busy teaching, administrating and trying to keep up with our brighter research students to "travel extensively. This volume has attempted to remedy this situation arising from a relatively informal meeting of researchers from far-flung areas who are working on rocks of this one age. Those of us who have also worked on Lower Palaeozoic rocks are always struck by the relative ease with which correlations can be made, in some cases along the entire length of the Caledonian belt. Yet the younger Variscan (Hercynian) belt seems poorly-understood by comparison: to date there has been no overall, agreed synthesis of the Variscan of Europe. In part this is due to Mesozoic and Tertiary cover, but it is mainly the result of the intricate yet subtle nature of Late Palaeozoic, especially Carboniferous, events. The reasons for this complexity are but vaguely understood: compare, for example, the well-argued but contradictory structural interpretations of the Variscan orogen by Matthews (1984) and Matte (1986).

The very nature of the Variscan foldbelt is still a matter of debate. A major key to greater understanding of the evolution of this huge part of western Europe is a fuller knowledge of the stratigraphy and palaeontology of its Carboniferous rocks. The Dinantian is crying out for an even more refined biozonation to help us to sort out this complexity. In this matter we can only envy workers in the Jurassic for example! Apart from their intrinsic geological interest rocks of Dinantian age are of very considerable economic value, not least as hosts to major ZnPb-Cu-Ba deposits in Ireland, and Au-FeS2 deposits in the Iberian Pyrite Belt. The term 'Irish-type' is now applied world-wide to a major class of Zn-Pb deposits first described in the 1960s and 1970s from Ireland where, as in Iberia, exploration still continues apace. Apart from Navan, Europe's largest and one of the world's great Zn/Pb mines, further large mines are coming on stream at Galmoy and Lisheen in Tipperary. These and earlier discoveries have stimulated extensive exploration drilling. We would like here to acknowledge the many mineral exploration companies who have made drill cores available for academic research to workers from Ireland, the UK, the USA and elsewhere who are concerned with the origin of these enigmatic but valuable orebodies, and with matters Carboniferous in general. Research into the origin of these ores involves many sub-disciplines within geology: mineralogy, geochemistry, sedimentology, biostratigraphy and structural geology. It also incorporates a great deal of academic research from earlier years, raising once again the perennial question: what is 'purely academic' research, and what is its value? How much modern academic research will one day come to economic fruition? All of the modern Irish deposits have been located within a few hundred metres of the surface, and the search for deeper orebodies, not detectable by today's geophysical techniques, will require the erection of much more refined models for the origin and siting of these orebodies. Further, the Upper Palaeozoic rocks of Europe will increasingly become the target of the oil and gas explorationist. The wealth of academic data on Dinantian r o c k s - their sedimentology, biostratigraphy, tectonics and basin evolution- while no guarantee of success, will be an invaluable tool to exploration. The first section of the volume is concerned with the economic geology of the Dinantian, especially the mineralization and diagenesis of

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. vii-ix.

viii

PREFACE

Dinantian rocks in Ireland and Britain. It emphasizes the importance these limestones play as host rocks for major Zn-Pb deposits. In particular, two papers focus on the role of syn-depositional faulting and tectonic controls on mineralization pathways. There are also studies of the timing of oil and gas generation and migration through these rocks. The second section is more academic, and focuses upon the highly active field of research into Dinantian carbonate buildups. It highlights recent advances in our knowledge of the sedimentology, facies and fauna of Dinantian mud-mounds both in Europe and in North America. It is now becoming generally accepted that microbial activity had a major influence on buildup development and that many mudmounds are dominated by matrix peloidal fabrics. New work on the Muleshoe Mound, New Mexico has demonstrated that 'Waulsortian' buildups could form in higher energy environments than previously suspected and were subject to episodic erosional events which may have been controlled by relative changes in sea-level. Many late Vis6an buildups are now known to have greater skeletal diversity than earlier Waulsortian mounds and could form rigid frameworks aided by microbialite cementstone fabrics. The significance of calcareous algae in buildups is now becoming better understood, both as skeletal contributors and as water depth and photic indicators. The nature and scale of siliciclastic environments in the Dinantian are described and interpreted in the third section. Potential clastic reservoirs seem to be of limited extent and clearly subject to strong tectonic control. They appear in places to be stacked against basement margin faults and these sandstones may be important as vertical migration pathways; there is abundant evidence presented in the first section that this has occurred in some basins. In the fourth and fifth sections the controls on the formation and evolution of carbonate platforms, ramps and basins are discussed. In late Dinantian platform successions subtle changes detected by microfacies analysis and changes in faunas have been widely noted by many authors. The possible role of eustatic control on these changes has been widely discussed. At the same time acknowledgement is made of the equal and locally greater importance of tectonic influence on the evolution of platforms and ramps. Sequence stratigraphy for the Dinantian is still in its infancy as a result of the interplay of these two very different controls on sedimentation. Basinal

calciturbidites, pelagic carbonates and cherts, and the palaeooceanographic significance of the latter are also discussed. Reworked platform materials in Dinantian basinal sequences are described and their significance assessed; foundered platforms, no longer visible at surface, have been detected by these means. Finally, two papers deal with Dinantian vulcanicity, a topic that surely deserves more attention in the future. The final section on fauna and biostratigraphy emphasizes the importance of accurate dating in the correlation of Lower Carboniferous sections, which underpins all syntheses of basinal evolution and the timing of major tectonic pulses, transgressions and regressions. The problem of facies control on faunas and of distinct biozonation schemes for platform and basinal facies is apparent from many of the earlier papers. Several papers here highlight the advances made in micropalaeontological studies, particularly on biostratigraphic refinement using different fossil groups in an integrated approach. Also an increasing contribution is being made by research on microvertebrates (fish and sharks) and microproblematica, especially in successions where diagnostic taxa are sparse or absent. The editors record their thanks to the Geological Society and all the staff of the Geological Society Publishing House. They would also like to express their gratitude to the many sponsors of the conference and the ongoing support for Dinantian research from the mining industry in Ireland. The editors would also like to express their thanks to the following individuals who reviewed one or more of the papers for this volume, and helped to improve its quality: Tony Adams, John Ashton, Ron Austin, Fernando Barriga, Adrian Black, Paul Brenckle, Paul Bridges, Geoff Clayton, Jerry Davies, Jaraslov Dvorak, Howel Francis, Stephen Gallagher, John Graham, Pete Gutteridge, Stephen Habesch, James Hein, Hans-Georg Herbig, Ken Higgs, Murray Hitzman, Andy Horbury, Dave Hunt, Dave Johnston, John Kelly, Ben Kennedy, Martin Laloux, Oleg Lebedev, Alan Lees, Marie Legrand Blain, J. M. Leistel, Michael Lipiec, Kelly Maguire, Jim Marshall, John Miller, Ian Mitchell, Dave Mundy, Dave Naylor, John Nudds, Bernard Owens, Eva Paproth, Mike Philcox, Neil Pickard, Eddy Poty, Tony Ramsay, Pat Redmond, Robert Riding, Nick Riley, Pat Shannon, Andy Sleeman, Bob Stanton, Roger Suthren, Sue Turner, Brian Turner, Greg Webb, Brian Williams, Sally Young and Jiri Zidec.

PREFACE Finally, it is with great sorrow that the editors record that the senior author of the first paper of this volume, and also one of our reviewers, Dave Johnston, was killed in early October 1995 at the age of 36 in an accident while on fieldwork on coastal cliffs in the West of Ireland. Dave was an undergraduate of Trinity College Dublin until 1980 and completed his PhD at Monash, Melbourne in 1984, when he returned as a lecturer in the Department of Geology at Trinity. He has written extensively on the structural geology of Ireland and more recently had particular interests in the structural control on economic mineralization in the Irish Carboniferous. He will be greatly missed for his good company, his enthusiasm, and his enormous ability for generating stimulating ideas and models.

ix

References MATTE, P. 1986. Tectonics and Plate Tectonics model for the Variscan Belt of Europe. Tectonophysics, 126, 329-374. MATTHEWS, S. C. 1984. Northern margins of the Variscides in the North Atlantic region; comments on the tectonic context of the problem. In: HUTTON, D. H. & SANDERSON, D. J. (eds) Variscan Tectonics of the North Atlantic Region. Geological Society, London, Special Publication, 14, 71-85. Peter Strogen Ian D. Somerville Gareth L1. Jones

Basement structural controls on Carboniferous-hosted base metal mineral deposits in Ireland J. D. J O H N S T O N l*, D. C O L L E R 2, G. M I L L A R 2 & M. F. C R I T C H L E Y 2

l Geology Department, Trinity College, Dublin 2, Ireland 2 ERA-Maptec, 5 South Leinster Street, Dublin 2, Ireland Abstract: Explorationists have long recognized the importance of structural control of Irish

base metal deposits, but the kinematics and timing of the structures relative to the mineralization are subject to wildly differing interpretations. This paper presents a combination of structural, magnetic and gravity data that show that the fundamental structural controls are easterly to northeasterly-trending basement (Caledonian) structures which have been reactivated in dextral transtension by northeasterly to north-northeasterly extension during the Dinantian. In general, in the north of the country the ore-controlling faults dip south, while in the south of the country they dip north. The structures have evolved through a temporal kinematic history of early normal faulting, later oblique slip, and some show evidence of later reverse movements. Part of this evolution may reflect burial history, although it also reflects the transition from Dinantian transtension to Variscan (Hercynian) compression. The bulk of the mineralization appears to post-date the normal faulting, and pre-date Variscan compression. Mineralization is thought to be postcompactional and could have occurred during the dextral transtension, although some of the sulphides could post-date most of the transtensional movement. The Lower Carboniferous of Ireland is host to a large number of base metal deposits (Fig. 1). These range in size from small occurrences up to the Navan deposit of more than 70 million tonnes (Mt) of 12% combined Z n - P b (Table 1). Vein-hosted copper ores in red beds are abundant in the south of the country, while further north Zn predominates in carbonatehosted deposits. Metal ratios vary regionally, with Cu and Pb decreasing and Zn increasing northwards (Johnston 1996). This suggests that all the mineralization may be generated by the same regional phenomenon. All of the major deposits occur in the hanging walls of normal faults. Copper deposits at Aherlow, Mallow and Ballinalack are in compressional structures, and no faulting has been documented at Courtbrown or Carrickittle. These latter cases may reflect limited exploration, rather than absence of faults. Generally, individual faults trend between ESE (Silvermines) and NE (Keel). Local faults with lengths of hundreds of metres, controlling ore lenses (ESE at Silvermines and E at Lisheen), are made up of en echelon segments defining ENE to NEtrending fault zones many kilometres long, which are inferred to be expressions of *Dave Johnston lost his life whilst on fieldwork in Mayo during October 1995. His co-authors would like to pay tribute to Dave's great contribution to structural geology around the world over the last two decades.

controlling basement structures. The local faults, although generally clockwise of the trend of the zones, show a larger variation in orientation than the zones themselves. Although the empirical relationship is well known, only a few studies of the structural settings of the deposits have been published (Moore 1975; Coller 1984; Reilly 1986; Shearley et al. 1992, this volume). Russell (1968, 1972) and Russell & Haszeldine (1992) invoke N-S geofractures to explain the mineralization in the Irish Midlands. However, as pointed out by Leeder (1988), there is almost no field or geophysical evidence for the existence of such structures. Rather, both field and geophysical evidence suggest that reactivated Caledonian (E to NE-trending) basement structures were the key in focusing mineralizing fluids. There is extensive hydrothermal alteration around these faults where they cut Devonian and Carboniferous rocks. Alteration is less extensive in the basement, but may be observed at Navan. In this study, both structural and petrographical observations and magnetic and gravity data are used to locate basement structures. On the basis of fault and mineralized vein geometries, most of the deposits are seen to lie in dextral transtensional settings (oblique extension with a component of clockwise rotation), which are the result of northeasterly extension of east to northeast-trending basement structures. Each mineral deposit has been further localized by

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 1-21.

2

J. D. JOHNSTON E T AL.

200

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Fig. 1. Location map showing main mineral deposits and stratigraphical provinces. National Grid is in kilometres.

specific secondary structures: termination zones, bends in faults, or intersections of second or third order shears with the primary basement structures. There is abundant evidence (feeder veins, alteration zones, changing metal ratios, fluid inclusions) that hot brines came up these fault zones and deposited metals in the hanging-wall carbonates. Most of the deposits in the northern half of the Midlands lie in south-dipping faults,

while those in the southern half lie in the hanging walls of north-dipping faults. Generally, the mineralized structures are en echelon and leftstepping, trending clockwise of the underlying basement structure. The displacement prior to the mineralization is predominantly dip-slip, indicating that the ore-controlling structures are extensional, and it is postulated here that these are rooted in dextral transtensional shears in the basement.

BASE METAL

Table 1.

MINERAL

DEPOSITS, IRELAND

Tonnages and grades of the main Irish sediment-hosted base metal deposits

Deposit

Host

Tonnage (Mt)

Zn+Pb (Wt%)

Zn (Wt%)

Pb (Wt%)

Navan Aherlow Mallow Gortdrum Tatestown Oldcastle Keel Abbeytown Clougherboy Ballyvergin Moyvoughly N'town-Cashel Lisheen Silvermines Tynagh Galmoy Magcobar Ballinalack Rickardstown Courtbrown Garrycam Carrickittle Harberton Boston Hill Allenwood

N N N N N N N N N N N N W W W W W W W W W W S S S

>70 6.0 4.2 3.8 3.6 3.0 1.9 1.1 0.34 0.15 0.13 ? >20.0 17.7 9.9 6.7 5.0 3.5 3.5 1.0 1.4 <0.1 3.9 0.8 9

12.7 . . 8.7 -6.9 4.9 8.7 5.3 7.0 . . 8.0 6.3 13.5 8.9 8.1 11.9 . . 7.0 3.3 5.5 2.9 7.5 9.3 3.8 2.0

10.1 2.6 . . 7.7 1.0 --5.3 1.5 4.3 0.6 7.7 1.1 3.8 1.5 5.8 1.2 . . 6.5 1.0 3.1 3.1 12.0 1.5 6.4 2.5 5.8 0.4 10.9 1.0 . . 5.9 1.1 2.2 1.1 3.5 2.0 2.7 0.2 6.0 1.5 8.1 1.2 2.7 1.1 1.6 0.4

Zn/Pb

3.9 7.7 3.5 3.5 7.2 7.0 2.5 5.0

.

6.5 1.0 8.0 2.6 0.8* 10.9 . 5.4 2.0 1.8 13.5 4.0 6.7 2.6 4.0

Ag (gm/t)

Cu (Wt%)

BaSO4 (Wt%)

Discovery

3.5 33.8 27.5 23.0 37.0 -39.6 40.0 -17.1 --32.0 23.0 78.0 --

-0.9 0.7 1.2 -----1.2 -------

--14.0 ------

---------

--------------8.3 -85.0 ---36.1 -----

1970 1965 1973 1964 1975 1976 1964? 1785 1977 1960 1968 1972 1990 1963 1961 1986 1959 1969 ? 1980 1975 1965 1975 1975 ?

Based on data from Andrew et al. (1986), Bowden et al. (1992). W, Waulsortian; N, Navan Group & equivalents; S, late Chadian or younger; Mt, millions of tonnes. * Primary ore.

It is p o s s i b l e t h a t V a r i s c a n i n v e r s i o n o f s o u t h d i p p i n g n o r m a l faults h a s resulted in uplift a n d subsequent erosion of other mineral deposits.

Regional stratigraphical setting Irish Z n - P b d e p o s i t s are h o s t e d p r i m a r i l y in L o w e r C a r b o n i f e r o u s rocks; m o s t are in r o c k s o f C o u r c e y a n to C h a d i a n age, b u t in several deposits hydrothermal alteration and sulphides e x t e n d u p into A r u n d i a n rocks. S o m e o f the c o p p e r d e p o s i t s in the s o u t h o f the c o u n t r y are h o s t e d in D e v o n i a n elastic sediments. Stratig r a p h i c a n d p a l a e o g e o g r a p h i c a l settings o f the L o w e r C a r b o n i f e r o u s s e q u e n c e s are d e s c r i b e d in detail b y Phillips & S e v a s t o p u l o (1986) a n d P h i l c o x (1984). T h e Irish C o u r c e y a n has b e e n s u b d i v i d e d i n t o f o u r p r o v i n c e s b a s e d o n differences in the local succession: (1) the S o u t h M u n s t e r Basin; (2) the L i m e r i c k P r o v i n c e ; (3) the M i d l a n d s P r o v i n c e ; a n d (4) the m a i n l y p o s t C o u r c e y a n N o r t h W e s t P r o v i n c e (Fig. 1; Phillips & S e v a s t o p u l o 1986).

T h e S o u t h M u n s t e r B a s i n ( N a y l o r et al. 1989) is filled with a c a r b o n a t e - p o o r , c l a s t i c - d o m i n a t e d succession. F l u v i a t i l e red b e d s are o v e r l a i n b y m a r i n e s a n d s t o n e s a n d shales. Clastic (shale) d e p o s i t i o n c o n t i n u e d t h r o u g h o u t the C o u r c e y a n , with s t a r v a t i o n o f the b a s i n later in the E a r l y Carboniferous. L i t h o l o g i c a l units are w i d e s p r e a d a n d laterally p e r s i s t e n t in the c a r b o n a t e - d o m i n a t e d L i m e r i c k P r o v i n c e (Philcox 1984; Somerville et al. 1992; see also Fig. 1). T r a n s g r e s s i o n o c c u r r e d d u r i n g the early C o u r c e y a n . T h e b a s e o f the s e q u e n c e c o m p r i s e s n o n - m a r i n e red b e d s a n d m a r g i n a l m a r i n e s a n d s t o n e s a n d shales. O v e r l y i n g these are c a l c a r e o u s shales ( R i n g m o y l a n Shale F o r m a tion), the p r i n c i p a l h o s t s to small c o p p e r d o m i n a t e d deposits. T h e shales p a s s u p into m u d d y bioclastic l i m e s t o n e s ( B a l l y m a r t i n a n d B a l l y s t e e n L i m e s t o n e F o r m a t i o n s ) , w h i c h are in turn overlain by Waulsortian carbonate mud b a n k s (Lees 1964). T h e s e units h o s t the Silvermines, Lisheen, G a l m o y , C o u r t b r o w n and T y n a g h deposits. O v e r l y i n g the W a u l s o r t i a n are either shelf, or in s o m e a r e a s basinal, c a r b o n a t e s .

4

J . D . JOHNSTON E T A L .

In the Midlands Province, transgression took place later, during the mid-Courceyan (Sevastopulo 1981; Phillips & Sevastopulo 1986). A formal stratigraphy for the region was erected by Strogen et al. (1990). The region is characterized by lateral facies variation, and has been divided into sub-provinces by Philcox (1984) and Strogen & Somerville (1984). A widespread basal red bed lithology is overlain by a series of shallow-water micritic and oolitic limestones, collectively referred to as the Navan Group (temporally equivalent to the Ballymartin Limestone of the Limerick Province). At Navan Mine these are informally called the Pale Beds (host to the Navan and Tatestown deposits). The Pale Beds in turn are overlain by the argillaceous bioclastic limestone of the Moathill and Slane Castle Formations (lithologically equivalent to the Ballysteen Limestone). The majority of sulphides at Ballinalack and Garrycam are hosted by the Waulsortian micrites of the Feltrim Formation, which overlies the Slane Castle Formation. Overlying the Waulsortian is a sequence of basinal graded calcarenites and calcisiltites of the Lucan Formation, widely referred to as the Upper Dark Limestones. The late Chadian to Arundian was a period of tectonic activity, with syndepositional tilting apparent at some places (e.g. Navan; Ashton et al. 1986, 1992) and extensional deformation. The youngest host rocks for mineral deposits in the Midlands Province are Chadian and Arundian platform carbonates, which contain the Harberton Bridge and Allenwood deposits. In the North West Province, much of the lower part of the section is made up of very shallow water sediments, lithologically equivalent to the Pale Beds, with a slightly later Courceyan to Chadian transgression.

Regional structural setting The structural history of the Carboniferous in Ireland may be divided into two main events: Dinantian extension, and late Carboniferous to early Permian (Variscan-Hercynian) compression. Beginning in the Courceyan and continuing through the Dinantian, 6000 m of shallow-water sediments accumulated in the South Munster Basin, (Matthews et al. 1983; Price & Todd 1988; Naylor et al. 1989). South of a line drawn from Dingle to Dungarvan (Fig. 1), major south-dipping faults controlled sedimentation. In the Midlands, a slightly different structural style existed. A series of interconnected basins

developed, with north-dipping faults dominant in the south and south-dipping faults in the north. Typically, sediments adjacent to these faults display rapid facies and thickness changes from the late Chadian onwards (e.g. the Tynagh Fault; Philcox 1984). These faulting events were sometimes accompanied by chaotic breccias in the Midlands Province (Nolan 1986, 1989; Crowe 1986; Philcox 1989; Pickard et al. 1992), and in the NW Province (Philcox et al. 1989). During the Variscan Orogeny these basins were inverted. Details of the structure of the Variscan deformation are given by Dolan (1983), Sanderson (1984), Coller (1984), Cooper et al. (1984, 1986), Ford (1987), Rothery (1988a, b) and deBrit (1989), and only the salient points are summarized here. In the south of the country the rocks have undergone extensive shortening (50-60%) with upright folding. The deform-ation has been interpreted as being either thin-skinned and thrust-related (Cooper et al. 1984, 1986), or thick-skinned and related to dextral transpression (Sanderson 1984). While thrusting definitely occurs (e.g. Ford 1987), there is abundant field evidence for dextral transpression in the western part of the South Munster Basin. Folds are transected anticlockwise (i.e. the cleavage related to the development of the folds appears axial planar in cross-section, but in plan trends anticlockwise of the fold hinge). En echelon fold trains (arrays of right-stepping folds) occur, and pressure fringes (fibrous mineral fibres, interpreted to indicate the stretching direction) growing on pyrite flattened in the cleavage indicates sub-horizontal extension. There is an abrupt drop in strain north of a line from Dingle to Dungarvan (Gill 1962; Cooper et al. 1984). Rocks in this region have undergone gentle folding with shortening of less than 20%. However, evidence for dextral transpression continues northwards (Johnston 1993; Fitzgerald et al. 1994). Although there are stable blocks of almost undeformed limestones, approximately 50 kilometres across, these are bounded by corridors of deformation (Coller 1984). These are large scale ENE-trending dextral transpression zones, up to several kilometres wide, which are characterized by dipping beds, en echelon vein arrays, and weak stylolitic cleavages that transect gentle folds. Both the folds and the cleavages generally trend anticlockwise of the high strain zones (Coller 1984). Conjugate NNE-trending sinistral zones are consistent with a NNW-trending compression direction which is to be expected from regional tectonic reconstructions of the Hercynian (Ziegler 1988).

BASE METAL MINERAL DEPOSITS, IRELAND

Geophysical framework The deep structure of the Irish Midlands can be discerned through use of regional aeromagnetic and gravity data. The Geological Survey of Ireland holds an aeromagnetic data set which may be processed under licence. The magnetic map and its interpretation were summarized by Max et al. (1983) and Morris & Max (1995). Murphy (1952, 1960, 1974, 1981) collected and compiled gravity data over most of the country with an approximate 1 km spacing. With fairly simple processing, these data sets reveal a great deal about the deep (sub-Carboniferous) structure. Using these data, Williams & Brown (1986) and Brown & Williams (1985) identified pronounced NE-trending gravity anomalies, which they interpreted in terms of horsts of Ordovician volcanic rocks (highs) beneath the Carboniferous, separated by intervening grabens filled with Carboniferous sediments (lows). Exploration drilling has confirmed the ridge and trough

5

geometry of the Carboniferous in the Irish Midlands (Andrew, 1992). However, the Lower Palaeozoic inliers that expose the basement rocks do not always match the interpretations of Williams & Brown (1986) in detail. Nevertheless, their work identified the existence of NE-trending, steeply dipping lithological contacts in and beneath the basin. The magnetic data also show a similar Caledonian grain (Murphy 1981; Max et al. 1983). Filtering of the gravity data for the south of the country by Ford et al. (1991) has placed some constraints on the deep geology of the Munster Basin. In this study, the same regional aeromagnetic data sets and the Bouguer gravity data have been studied. The magnetic data are illustrated in Figs 2 and 3. The gravity data are illustrated in Fig. 2 of Readman et al. (1996), and a summary of the gravity data for the Irish Midlands is shown in Fig. 4. Comparison of the magnetic first derivative maps and the gravity map of Readman et al. (1996) and Murphy

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300000

Fig. 2. Residual magnetic anomaly map of the Irish Midlands in pseudo-relief, illuminated from NW. The NNW-trending grain is an artefact of the flight lines, and the bright ENE lines (e.g. between Tonduff and Rickardstown) are artefacts of splicing of different data sets.

6

J . D . JOHNSTON E T AL. 100000

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Fig. 3. Magnetic first derivative anomaly map of the [fish Midlands. Artefacts as in Fig. 2.

(1974) highlights regions of steep contacts, juxtaposing rocks of differing densities and magnetic susceptibilities. In many instances such contacts are tectonic. There are many examples where this has been confirmed in boreholes and Lower Palaeozoic inliers such as in the Loughshinny-Skerries region (Fig. 1).

Geology, geophysics and structure of the deposits A number of reviews of the deposits have been published in recent years (Hitzman & Large 1986; McArdle 1990; Andrew 1993; Johnston 1996) summarizing the main geological features of the deposits. Detailed descriptions of most of the deposits can be found in Andrew et al. (1986) and Bowden et al. (1992). Here, just the structural settings and those petrographic relationships relevant to constraining the timing of mineralization are described.

Gortdrum

In the south of the Irish Midlands, a group of copper deposits occur within the carbonate bearing sequence. Gortdrum (Fig. 5) is the only one of these deposits to have been mined. The sulphides occur in Courceyan shales and bioclastic limestones in the hanging wall of the ENE-trending, steeply NNW-dipping Gortdrum Fault (Steed 1986). The footwall is composed of Old Red Sandstone. The hanging wall is intruded by E-trending, hydrothermally altered mafic dykes and plugs. The deposit post-dates the host rocks, including the volcanic rocks, but pre-dates Variscan deformation. Mineralization is associated with both the main faults and an en echelon array of more E-W striking extensional faults (Fig. 5a). Both normal and thrust faulting is present. There is a compressional overlap of the main faults to the south of the deposit consistent with WNW compression and complementary to the dilation area (Fig. 5a) of E-trending en echelon normal

BASE METAL MINERAL DEPOSITS, IRELAND

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Fig. 4. A summary residual gravity map for part of the Irish Midlands, based on Brown & Williams (1985). Contours in gravity units. Hachured regions are lows. faults, which define a NW-trending zone consistent with NNE extension. This pattern of faulting is typical of a dextral shear zone fault complex. In section (Fig. 5b) it has the geometry of a negative flower (oblique extensional) structure. This geometry is thought to be the product of Dinantian dextral transtension. The fault zone is also characterized by a series of en echelon folds that show a strong clockwise rotation in strike into the fault zone at the deposit, indicating a large dextral shear displacement in the mineralized area. The folding, which is compressional and deformed consolidated rocks, is interpreted to be Variscan. This rotation into the footwall is also shown by the magnetic and gravity data. The principal geophysical elements are shown in Fig. 5c. The deposit lies on the southern margin of a NEtrending magnetic high in a region where the strike of the magnetic (basement) signature changes from NE- to E-trending (Figs 2, 3). A steep S-dipping magnetic gradient is coincident

with the main, N-dipping fault. First derivative processing of the magnetic data yields an anomaly that is slightly oblique to the fault trace (Figs 3, 5c). Major gravity linear features (enhanced by first derivatives) are NE-trending and reflect the distribution of lithologies in the folds adjacent to Gortdrum. The gravity features are oblique to both the magnetic fabric and the main fault (Fig. 5c), and they intersect the fault at the deposit. This suggests that a sedimentary basin developed oblique to the basement structure. The folds are likely to have formed in the cover during the Hercynian deformation, along and oblique to the basement structure. Navan

The geology of this deposit (over 70Mt) is summarized by Ashton et al. (1986, 1992). The ore body is subdivided into five stratigraphically stacked lenses hosted in the Pale Beds.

8

J. D. JOHNSTON E T A L .

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Fig. 5. (a) Structural map of the Gortdrum region. (b) Cross-section of Gortdrum, after Steed (1986). (e) Interpretation of main geophysical features of Gortdrum. Structural control on the mineralization is apparent as the ore lenses overlie steeply dipping and ENE-trending veins of sulphides, inferred to be feeder veins. The ore lenses are cut by the so-called 'Navan unconformity' (a major listric slide surface of Chadian age), which is overlain by the Boulder Conglomerate and Upper Dark Limestone. The ore body is also subdivided laterally into three zones by the

B, T and A faults, (Fig. 6a). The B and T faults are spoon-shaped, have a normal component of displacement and die out up-section. They trend oblique to, and occur between, the A and C faults. Minor folding occurs above the T fault in the Upper Dark Limestone. The B and T faults trend on average about ENE and dip to the SSE. Adjacent to the T fault and beneath the ore lenses there is a concentration of sulphide veins that trend ENE. Zone 2 lies in the footwall of the T fault, and the mineralization is less well developed in its hanging wall block. Significant rotation of bedding in the Pale Beds occurs in the hanging wall of the B and T faults, suggesting that they are rotational extensional normal faults. The Boulder Conglomerate represents the chaotic deposit produced by synsedimentary rotational slumping. Both the T and the B faults are truncated by the A - C fault complex (Fig. 6a). Folding in the Upper Dark Limestone that trends anticlockwise of the A and C faults intensifies in the vicinity of the A fault, which has a large consistent reverse component of displacement. As all of the faults offset the stratigraphic markers that define the ore lenses, they appear to offset the lenses. However, a case can be made for the ore lenses being selectively replacive, and post-dating movement on the B and T faults. While some mineralization terminates at the faults, at least in some cases, within the footwall block, grade increases towards the faults. Close examination of mineralization in the footwall of the T fault reveals that, at several levels, mineralization terminates several tens of centimetres short of the fault, and is not strictly truncated by it. The thickest ore in Zone 2 is in the hanging wall block of the B fault. Although displacement on the B fault is well constrained by stratigraphic displacements, ore lenses cannot be identified in the footwall, suggesting that the distribution of mineralization is influenced by the B fault. Where sulphides abut the B fault, the ore terminates against stylolitic surfaces defining the fault zone margin. In general terms, where an argillaceous unit lies in the hanging wall of a fault, sulphides terminate at the fault. Where the succession is shale poor, the ore extends across the faults. It is therefore possible that the mineralization occurred after much of the displacement had occurred on the B and T faults.

Fig. 6. (a) Geological map of Navan, modified after Ashton et al. (1986,1992). (b) Cross-section of Navan redrawn from Ashton et al. (1986). (e) Modified cross-section of Navan in which ore lenses are correlated by absolute position, rather than stratigraphical position (redrawn from Johnston (1996)).

BASE METAL MINERAL DEPOSITS, IRELAND

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10

J. D. JOHNSTON ET AL.

All of the ore lenses terminate beneath the unconformity, and clasts of mineralized Pale Bed lithologies occur within the Boulder Conglomerate. Where Zn-Pb ore lenses abut the unconformity, iron-rich Zn mineralization occurs above the unconformity (the Conglomerate Group Ore). The timing of the mineralization with respect to the unconformity is not clear-cut. Stratiform pyritic, sphalerite ore occurs in the matrix above the unconformity. Clasts of massive sulphides occur in the Boulder Conglomerate at the base of the Conglomerate Group Ore. Some carbonate clasts have been replaced in situ by sulphides (for example, many clasts show a concentric zonation of pyrite, sphalerite, dolomite from margin to centre). The replacement is highly selective: some bioclasts are preferentially replaced by sphalerite. In others, however, complex paragenetic sequences are preserved and some layers of sulphides appear to have been truncated by erosion (Ashton et al. 1986). This suggests that some of the mineralization occurred both prior to and during the slumping that generated the Boulder Conglomerate. The Navan ore body is located near the hinge region of a SW-plunging anticline, and overlies a linear SW-plunging basement horst that produces a regional magnetic high (Figs 3, 4). In this region the main magnetic trends range from N E - S W to almost E-W. Magnetic gradients from the first derivative processing, like the main axes of anomalies, follow the two main fault trends, NE and ENE, diverging at the deposit. South of the mine, a gravity high trends NNE, and swings in strike to ENE in the area of the deposit (Fig. 2 & Fig. 2 in Readman et al. 1996), similar but oblique to the main trend of the magnetic anomalies. To the SE, a large gravity low has been interpreted as a buried granite (Murphy 1952). The structural location of Navan and the generation of the ENE-trending feeder veins oblique to the main B and T faults is considered, therefore, to be related mainly to dilation created during regional extension. They may be en echelon rotational slump structures developed on a precursor of the NE-trending A - C fault complex which controlled the localization of the deposit. Reactivation and offset on the original structure was the result of dextral shear. The Tatestown-Scallenstown and the Clogherboy deposits are satellite deposits to Navan and are associated with similar strike swings in the main faults ('basement' structures) as Navan. They have similar lensoid geometries to the Navan ore bodies.

Keel and N e w t o w n Cashel

The Keel deposit (Slowey 1986) is hosted in an ENE-trending, S-dipping normal fault zone (Fig. 7). Mineralization is discordant and occurs in fractures related to the faulting. The host rocks range from Lower Palaeozoic slates to upper Navan Group rocks. Fault segments extend into the Waulsortian, and the wallrocks are extensively dolomitized adjacent to the faults. A series of N-dipping, NE-trending magnetic gradient boundaries occur on the northern margin of a major magnetic high which lies 10km to the south. At the Keel Fault, the anomalies parallel the strike of the adjacent Lower Palaeozoic inlier, trending ENE; away from the main fault line, they swing NE. Mineralization at Newtown Cashel was controlled by an en echelon array of E-W faults, forming a dilation zone at the termination or overlap area of the main ENE-trending Keel Fault (Fig. 7b). The faulting and dilation zone is a mirror image of the controlling faults at Silvermines (see below). The geometry of the fault system at Keel is consistent with dextral transtension. The Keel deposit occurs where the main fault on a regional scale has a minor strike swing; with dextral strike-slip motion this position is favourable for dilation and enhanced fracturing. The adjacent Newtown Cashel deposit (Crowe 1986) lies on the same ENE-trending structure as at Keel. This structure was active throughout the Dinantian, as there are lateral facies variations along and across the fault and a Chadian unconformity is associated with it (Crowe 1986). The fault pattern is reflected by a gravity lineament swing to E-W, and step off en echelon to the south. E-trending faults at Newtown Cashel deposit are also reflected by magnetic lineaments.

Silvermines and M a g c o b a r

The geology of the Silvermines and Magcobar deposits (Fig. 8) was summarized by Andrew (1986). The local geology is dominated by the ENE-trending, steeply north-dipping, obliqueslip, dextral-normal Silvermines fault (Fig. 8). Dolomitic breccias that are spatially associated with the tabular mineralization form the hanging wall; these have been interpreted as debris flows (Andrew et al. 1986) and also as hydrothermal alteration products (Hitzman & Large 1986). Observations by one of the authors (J.D.J.) suggest that elements of both hypotheses may be correct: some parts of the breccia

BASE METAL MINERAL DEPOSITS, IRELAND

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Fig. 7. (a) Keel cross-section, redrawn from SIowey (1986), showing the intimate relationship between faulting and mineralization. (b) Schematic map of the fault system Keel-Newtown Cashel fault system, showing a dextral transtensional setting with a termination zone at Newtown Cashel. units show clast alignment and sorting, parallel to bedding. Some of the clast margins are sutured and stylolitized, but others show rounding, and soft-sediment slump folds have been observed underground in the limestones adjacent to these breccias (Andrew et al. 1986). Superimposed on both the sedimentary breccias and the Waulsortian is a dolomite-matrix breccia. This comprises angular fragments of dolomitized wallrock with a filigree matrix of extremely fine-grained dolomite. In parts of this breccia the clasts can be matched and fitted together, while in other parts clasts have clearly moved relative to each other. This breccia appears to be hydrothermal in origin. Base

metal mineralization occurs both in immediately sub-Waulsortian stratiform lenses and in discordant feeder zones parallel to the fault. The ore lenses thicken towards the Silvermines fault, as do the stratigraphic units and the Zn/Pb ratios increase. There is an upper, tabular, predominantly stratiform ore body of barite, siderite and marcasite (Fig. 8). In the deeper parts of the ore bodies the mineralization occurred in veins which lie en echelon along the Silvermines Fault. The structural history has been documented by Coller (1984). A number of en echelon WNWtrending faults with predominantly normal displacement localize the mineralization. The

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Silvermines Fault is parallel to, and lies on the southern margin of, a regional magnetic high, and north of a smaller magnetic low. Silvermines ore bodies lie where the axis of the high swings from ENE to E. West of the deposits, the boundary to the magnetic high is approximately coincident with the Silvermines Fault. However, this boundary side-steps to the SE at Silvermines, suggesting an en echelon side-step or offset of the main fault (see inset on Fig. 8). Magnetic linears apparent on the first derivative map (Fig. 3) show this same offset. The axis of a NNE-trending gravity low (Murphy 1974) passes through the deposit. North of the fault this low is coincident with a regional syncline in the Carboniferous. However, in the vicinity of the deposit, this fold has been tightened and rotated into the Silvermines Fault, while the anomaly has not been deflected. The overall geometry suggests an ENEtrending dextral motion, but the component of motion on individual WNW-trending segments is extensional (i.e. normal). Mineralization was most intense at points of maximum normal throw. All structures show evidence of post-ore dextral reactivation (sub-horizontal slickensides are developed on some ore lenses). Thus, the faults appear to have been active during sedimentation and mineralization, and have been subsequently reactivated.

Tynagh

The Tynagh deposit occurs in the hanging wall of the Tynagh fault, with E-trending mineralized sections linked by short barren NE faults (Clifford et al. 1986). The role of the Tynagh Fault in the development of the deposit (Fig. 9) has been discussed by Moore (1975). Early faulting with a major normal component (greater than 600m) is reflected by thickness changes and slumping in the Courceyan sedimentary record. The throw is variable along the fault, with the ore concentrated in the vicinity of the maximum throw. Later reverse movement was the result of NW horizontal compression. This is the same orientation as the regional Variscan shortening direction, producing dextral transpression. Mineralization overgrows stylolites, which lie at a high angle to bedding adjacent to the fault, suggesting that significant burial and deformation pre-dated at least some of the mineralization. Detailed petrography reveals the occurrence of carbonate pressure fringes with horizontal and vertical fibres growing on sulphides. This observation suggests that at least some of the reverse movement, with associated high pore-fluid pressures, post-dated at least some of the mineralization. Late stage veins and slickensides indicate that later sinistral strike-slip motion of unknown magnitude occurred.

BASE METAL MINERAL DEPOSITS, IRELAND

13

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Fig. 9. (a) Map of Tynagh, showing E-trending mineralized faults linked by ENE barren faults in a dextral transtensional zone, redrawn from Clifford et al. (1986). (b) Cross-section of Tynagh redrawn from Clifford et (1986), showing fault control of mineralization.

Tynagh is on the northern margin of a magnetic high, and the main Tynagh Fault is coincident with a major magnetic boundary that extends ENE to the Moyvoughly prospect. First derivative structures diverge from ENE to NE in the vicinity of the deposit. Minor linears trend

al.

ENE parallel to the main mineralized zone, and oblique to the E - W controlling faults. Gravity features trend NE, with the deposit located on the north side of a linear low and to the south of a large gravity gradient associated with a regional high.

14

J. D. JOHNSTON ET AL.

The Tynagh deposit, like Silvermines, is a classic example of mineralization controlled by en echelon extensional faults that conform to an oblique extension with dextral shear on the main ENE fault (Fig. 9). Further evidence is provided by en echelon veins and slickenfibres. The reason for the dilation at this segment of the fault is either related to the main strike change in the basement (magnetic) fabric from near E-W to NE, or a dilational termination zone of the fault strand.

Lisheen and Galmoy Lisheen (Hitzman et al. 1992; Shearley et al. 1992, this volume) and Galmoy (Doyle et al. 1992 ) are both fault-controlled tabular deposits at the base of the Waulsortian. Both deposits lie on the NE-trending Killoran Fault Zone (Shearley et al. this volume). Other minor prospects have been discovered further north along this trend at Derrykearn and Glasha. The Killoran Fault is a dextral transtensional fault with individual E to ENE-trending strands (Shearley et al. this volume). At Lisheen, as at Silvermines and Tynagh, sulphides are spatially associated with an ENE-trending feeder fault zone. Aggregate normal displacement across the Killoran Fault at Lisheen is 210m. Its dip varies between 40 ° and 70 ° to the north. Three main ore lenses have been identified: the Main, Derryville and North zones. Most of the mineralization occurred within 30m of the base of the Waulsortian. Several sulphide bodies, containing iron sulphides, sphalerite and chalcopyrite, occur within oolitic limestones in the footwall at approximately the same elevation as the orebodies at the base of the Waulsortian. A series of parallel magnetic lineaments (060 °trending) define changes in gradient in a zone parallel to the main Killoran Fault between Lisheen and Galmoy. A magnetic high to the NE of Galmoy cannot be traced in the LisheenGalmoy region. All of the gradients dip south and a second derivative linear is coincident with the main mapped fault. In addition, NNWtrending cross-lineaments occur at both Lisheen and Galmoy. Two main 060°-trending linear features may be inferred from the Bouguer gravity data. They side-step with an en echelon overlap near Lisheen. Two oblique-striking (050 °) first derivative linear features may represent link faults in the overlap zone between the main linear features. The Killoran Fault lies parallel

to and just south of the interpreted boundary to the large gravity low associated with the Leinster Granite. To the north of the fault at the Derrykearn prospect there is a similarlytrending gravity high. The main 'basin margin' linear has minor strike swings at both Lisheen and Galmoy and a more significant swing at the Tonduff-Derrykearn prospects. Lisheen and Galmoy are both spatially associated with a second set of linear gravity features trending 030 °, oblique to the main fault trend. These are part of a regional-scale zone of features arranged in en echelon groups side-stepping westwards, that can be traced for 70 km to the southwest. As in the Tynagh deposit, steep stylolites adjacent to the fault are mineralized. This suggests that much of the normal displacement had occurred across the Killoran Fault prior to mineralization. Carbonate pressure fringes are also developed on the sulphides adjacent to the fault (as at Tynagh), indicating some post-ore fault movement. Both the Galmoy and Lisheen deposits occur in the footwall of a regional dolomitization front that affects much of the Lower Carboniferous limestones of SE Ireland. At Lisheen the sulphides and their associated dolomites crosscut, brecciate and clearly postdate this regional dolomite, which affects rocks from Courceyan to late Vis6an in age.

Discordant deposits hosted in the Waulsortian and supra- Waulsortian succession A cluster of small pipe-like, breccia-hosted deposits have been discovered in County Kildare. The largest of these are Harberton Bridge (Fig. 1; Emo 1986) and Allenwood (Andrew 1993). These prospects are hosted in pipes that crosscut the Waulsortian and extend into the overlying Chadian and Arundian carbonate sediments. There are no magnetic signatures associated with the pipe-like deposits in Kildare, but the gravity trends are oblique to the main faults. Several of the smaller occurrences are veinhosted. At Allenwood (Andrew 1993), E-trending faults are mineralized, and brecciation occurs in patches which coincide with bends on the faults. Mineralized faults in the prospect area step eastwards and northwards in an en echelon manner, consistent with dextral transtension. NW-directed thrusting is assumed to be post-mineralization and Hercynian in age.

BASE METAL MINERAL DEPOSITS, IRELAND

of each of the deposits shows that most of the deposits (Keel, Navan) that occur in the north of the basin are located on southerly dipping structures, while those in the south all dip to the north (Silvermines, Lisheen, Tynagh; Fig. 11). In the central part, faults dip in both directions. They usually have a normal throw with a mineralized hanging wall and barren footwall. All of the faults have some element of post-mineralization strike-slip reactivation. Many faults appear to have had predominantly dip-slip motions during an early, synsedimentary history. They almost all trend ENE to NE. In every major fault associated with the Irish deposits where sufficient exposure is available,

Discussion There is clear structural control of all of the major Irish Carboniferous-hosted sulphide deposits. Mineralization in the Pale Beds sequence either abuts a fault (Tatestown, Moyvoughly, Keel; Fig. 7) or an 'unconformity' (Navan; Fig. 6). All of the Waulsortian-related deposits are in the hanging walls of normal faults (Fig. 9). The Harberton Bridge, Allenwood and related deposits are in breccia pipes related to normal faulting. A fault map of the Midlands of Ireland (Fig. 10) reveals that the mineralized faults generally dip north in the southeast and to the south in the northwest. A cartoon of the structural geometries

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Fig. 10. Fold and fault map of the Midlands of Ireland.

16

J. D. JOHNSTON

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Fig. 11. Cartoon map of the structural settings of the Irish deposits. The deposits are located in their approximate relative positions, but not drawn to scale. McCoss constructions indicate movement vectors for deformation. All of the deposits are in dextral transtensional settings. This is the result of NE extension of an E to ENE-trending basement template. there is evidence of early dip-slip. The evidence to support this view is a combination of stratigraphical (e.g. formations thicken in the hanging wall at Tynagh, Silvermines and Lisheen) and textural (e.g. the orientation of fibres in the fault zones at Gortdrum, Tynagh, and Silvermines, of slickenlines and shear bands in the case of Navan, and early compacted veins at Ballinalack). In all cases the dip-slip motions were followed by strike-slip motions as indicated by the orientation of fibres, veins and faulting. In every case the strike-slip post-dates the mineralization. This is because

mineralization occurred in pure extensional orientations in en echelon systems (see below) of overall transtension. In localized occurrences, these structures were subsequently reactivated by reverse movements. Millar (1990) demonstrated the same general sequence of structural events on a regional basis in the North West Province, and Nolan (1986, 1989), de Brit (1988) and Johnston (1993) have identified the same sequence in the Irish Midlands. To a certain extent the division of structures into strike-slip and dip-slip is an

BASE M E T A L M I N E R A L DEPOSITS, I R E L A N D artificial one. Individual, en echelon, pure extension veins (dip-slip fractures) can link and rotate with depth to form oblique extension veins (strike-slip fractures) as is apparent at Silvermines (Fig. 8), where ESE-trending dip-slip

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17

faults rotate and link at depth to form an E N E dextral fault. In an attempt to semi-quantify the motions, McCoss constructions (McCoss 1986; Fig. 12a) for the Irish deposits were evaluated. It was assumed that the magnetic first derivative trend gave the deformation zone boundary. This is justified on the grounds that it is likely to give the orientation of the major basement structure, which in turn has controlled later deformation in the Carboniferous. Secondly, the best estimate of the orientation of extension fractures at the time of mineralization was used to define the strike of extension veins. Every one of the deposits sits in dextral transtensional zones (Fig. 11). These were the product of a N E extension of Carboniferous rocks deposited on a pre-existing Caledonian template. As pointed out by Millar (1990), the transition from dip-slip through strike-slip to thrusting throughout the Irish Carboniferous may simply reflect increasing burial depths (Fig. 12b). All of the textural evidence presented here suggests that the mineralization is related to NE-extension. At deeper structural levels, this m a y have been contemporaneous with NW-compression. Therefore, some of the copper deposits in the South Munster Basin (Fig. 1; Reilly 1986) may have formed at the same time as the dextral transtensional deposits in the Irish Midlands

Miaeralization,SouthMunsterBasin ~ ~ Thrusting, _ _

Fig. 12. (a) Figure illustrating McCoss construction (McCoss 1986). A line is constructed parallel to the margin of a shear zone. A circle is constructed with this line as tangent. A line is drawn parallel to the strike of pure extensional structures (veins, dip-slip faults) from the tangent point. A perpendicular is constructed. A diameter drawn from the intersection of the circle with the strike line to the intersection with the perpendicular gives the resultant vector of the shear zone. (b) Diagram to illustrate the switch from normal faulting to strike-slip faulting to thrusting with burial. The tensile part of the figure is only appropriate in some cases. During the Dinantian the greater horizontal stress axis was oriented NE and the smaller trended SE. Progressive burial resulted in increasing compression and the vertical stress increased linearly and slowly, while the horizontal stresses increased more rapidly. As a result, initial dip-slip switches to strike-slip and eventually thrusting. Miilar (1990) has suggested that all of these regimes may have been active simultaneously at different stratigraphic levels in the same place. However, as the Variscan deformation propagated northwards there must have been a general switch from transtension to transpression.

ement shear zone Basemen~t~tri~j//

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, Harberton,~ridge / Allenwood ~ t ~ _ i __ /Moyvoughly / tynagn/~ ](e-el --'~---" / ~ Newcastle .t~..,..\~------ Lisheen ~ ,'5"~ Galmoy ~ ~ ~ l]allinalack Silvermines T I ] /

"

Local mineral controllingfault Gravity anomaly Magnetic anomaly, revealing basement structure with inferreddip of structure Fig. 13. Idealized structural model for Irish-style deposits.

18

J . D . JOHNSTON E T AL.

during the mid to late Dinantian. However, it is generally accepted that the transpressional deformation associated with the Hercynian deformation post-dates the late Dinantian. Marine sedimentation persisted into the Namurian. Nevertheless, it is possible that the deformation was diachronous and the Irish Midlands may have been in transtension while the Munster Basin was in transpression. A generalized model for the structural controls of the deposits is shown in Fig. 13. Each of the deposits is in a unique setting related to the precise basement geometry of its location. However, they are all located adjacent to basement structures, in regions of strike swings along these structures. This is seen in the geophysical data as a divergence of the gravity and magnetic linears. This may reflect obliquity of lithological and structural grain, or obliquity of major and minor structures. Either way, complexity of the trends in the basement greatly enhanced the potential for dilatancy in the cover during reactivation, and favoured the generation of ore deposits.

Conclusions Most of the Irish carbonate-hosted base metal deposits appear to sit in dextral transtensional structural traps. The precise geometry of the traps varies from one deposit to another and the variations are the product of different basement structural geometries. Most of the deposits appear to have formed at intermediate burial depths, rather than on the sea floor. Precise timing is difficult to constrain, but the best estimates seem to be early Chadian for Navan, and late or post-Chadian for Lisheen. The fundamental structural controls are E- to NE-trending basement (Caledonian) structures, which have been reactivated in dextral transtension by N N E - to ENE-extension during the Dinantian. In general, in the north of the country, the ore-controlling faults dip south, while in the south of the country they dip north. The structures have evolved through a temporal kinematic history of early normal faulting, later oblique slip, and some show evidence of later reverse movements. Part of this evolution may reflect burial history, but it also reflects the temporal transition from Dinantian transtension to Variscan compression, which propagated northwards with time. The bulk of mineralization appears to post-date the normal faulting, but pre-date the Variscan compression.

We thank D. J. Sanderson, W. E. A. Phillips, D. Romer, G. D. Sevastopulo, J. Ashton, M. W. Hitzman and P. Redmond for stimulating discussions. This work would not have been possible without the gravity data set gathered by T. Murphy and access to the magnetic data set under license by the Geological Survey of Ireland. Reviews by M. W. Hitzman, M. J. Kennedy and P. Strogen greatly improved the manuscript.

References ANDREW, C. J. 1986. The tectono-stratigraphic controls to mineralization in the Silvermines area, County Tipperary, Ireland. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 377 417. - - 1 9 9 2 . Basin development chronology of the lowermost Carboniferous strata in the north-Central Midlands. In: BOWDEN, A., EARLS, G. V., O'CONNOR, P. & PYNE, J. (eds) The Irish Minerals Industry 1980 1990. Irish Association for Economic Geology, Dublin, 143-169. 1993. Mineralization in the Irish Midlands. In: PATTRICK, R. A. D. & POLYA, D. A. (eds) Mineralization in the British Isles. Chapman & Hall, London, 208-269. - - , GROWL, R. W. A., FINLAY, S., PENNELL,W. M. & PYNE, J. F. 1986. Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin. ASHTON, J. H., BLACK, A., GERAGHTY, J., HOLDSTOCK, M. & HYLAND, E. 1992. The geological setting and metal distribution patterns of Zn/Pb/Fe mineralization in the Navan boulder conglomerate. In: BOWDEN,A., EARLS, G. V., O'CONNOR, P. & PYNE, J. (eds) The Irish Minerals Industry 19801990. Irish Association for Economic Geology, Dublin, 171-210. - - , DOWNING, D. H. & FINLAY, S. F. 1986. The geology of the Navan Zn-Pb orebody. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 243 -280. BOWDEN, A. A., EARLS, G., O'CONNOR, P. G. & PYNE, J. F. 1992. The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin. BROWN, C. & WILLIAMS, B. 1985. A gravity and magnetic interpretation of the structure of the Irish Midlands and its relation to ore genesis. Journal of the Geological Society, London, 142, 1059-1075. CLIFFORD, J. A., RYAN, P. & KUCHA, H. 1986. A review of the geological setting of the Tynagh orebody, Co. Galway. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 419-439.

BASE M E T A L M I N E R A L DEPOSITS, I R E L A N D COLLER, D. W. 1984. Variscan structures in the Upper Palaeozoic rocks of west central Ireland. In: HUTTON, D. H. W. & SANDERSON, D. J. (eds) Variscan Tectonics of the North Atlantic Region. Geological Society, London, Special Publication, 14, 185-194. COOPER, M. A., COLLINS, D. A., FORD, M., MURPHY, F. X. & TRAYNER, P. M. 1984. Structural style, shortening estimates and the thrust front of the Irish Variscides. In: HUTTON, D. H. W. & SANDERSON, D. J. (eds) Variscan Tectonics of the North Atlantic Region. Geological Society, London, Special Publication, 14, 167-175. . . . . & O'SULLIVAN, M. 1986. Structural evolution of the Irish Variscides. Journal of the Geological Society, London, 143, 53-61. CROWE, R. W. A. 1986. The stratigraphic and structural setting of Zn-Ba-Pb mineralization at Newtown Cashel, Co. Longford. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 331-339. DEBRIT, T. J. 1989. Timing structural events and basement emplacement using extension veins and cements in the Carboniferous of north central Ireland. Irish Journal of Earth Sciences, 10, 13-31. DOLAN, J. M. 1983. A structural cross-section through the Carboniferous of northwest Kerry. Irish Journal of Earth Sciences, 6, 95-108. DOYLE, E., BOWDEN, A. A., JONES, G. V. & STANLEY, G. A. 1992. The geology of the Galmoy deposit. In: BOWDEN, A., EARLS, G. V., O'CONNOR, P. & PYNE, J. (eds) The Irish Minerals Industry 19801990. Irish Association for Economic Geology, Dublin, 211-226. EMO, G. T. 1986. Some considerations regarding the styles of mineralization at Harberton Bridge, Co. Kildare. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 461-469. FITZGERALD, E., FELLY, M., JOHNSTON, J. D., CLAYTON, G., FITZGERALD, L. J. & SEVASTOPULO, G. D. 1994. The Variscan thermal history of west Clare, Ireland. Geological Magazine, 131, 545-58. FORD, M., 1987. Practical application of the sequential balancing technique: an example from the Irish Variscides. Journal of the Geological Society, London, 144, 885-891. --, BROWN, C. & READMAN, P. 1991. An analysis and tectonic interpretation of gravity data over the Variscides of south-west Ireland. Journal of the Geological Society, London, 148, 137-148. GILL, W. D. 1962. The Variscan fold belt in Ireland. In: COL, K. (ed.) Some Aspects of the Variscan Fold Belt. Manchester University Press, 49-64.

19

HITZMAN, M. W. & LARGE, D. 1986. A review and classification of the Irish carbonate-hosted base metal deposits. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 217 238. --, O'CONNOR, P., SHEARLEY, E., SCHAFFALITZKY, C., BEATY, D. W., ALLAN, J. R. & THOMPSON, T. 1992. Discovery and geology of the Lisheen Zn-Pb-Ag prospect, Rathdowney trend, Ireland. In: BOWDEN, A., EARLS, G. V., O'CONNOR, P. & PYNE, J. (eds) The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin, 227-246. JOHNSTON, J. n. 1993. Three-dimensional geometries of veins and their relationship to folds: examples from the Carboniferous of eastern Ireland. Irish Journal of Earth Sciences, 12, 47-63. 1996. Regional fluid flow and the genesis of Irish Carboniferous base metal deposits. In: RICKARD, D. (ed.) Mineral Deposits of Europe. Wylie, London, in press. LEEDER, M. R. 1988. Recent developments in Carboniferous geology: a critical review with implications for the British Isles and N. W. Europe. Proceedings of the Geologists'Association, 99, 73-100. LEES, A. 1964. The structure and origin of the Waulsortian (Lower Carboniferous) 'reefs' of west-central Eire. Philosophical Transactions of the Royal Society of London, B247, 483-531. MCARDLE, P. 1990. A review of carbonate hosted base metal-barite deposits in the Lower Carboniferous of Ireland. Chronique de la Minidre, Recherche, 500, 3-29. McCoss, A. M. 1986. Simple constructions for deformation in transpression/transtension zones. Journal of Structural Geology, 8, 715-718. MATTHEWS, S. C., NAYLOR, D. & SEVASTOPULO, G. D. 1983. Palaeozoic sedimentary sequence as a reflection of deep structure in southwest Ireland. Sedimentary Geology, 34, 83-95. MAX, M. D., RYAN, P. D. & INAMDAR, D. D. 1983. A magnetic deep structural interpretation of Ireland. Tectonics, 2, 431-451. MILLAR, G. 1990. Fracturing in the Northwest Carboniferous Basin, Ireland. PhD Thesis, Queen's University of Belfast. MOORE, J. MCM. 1975. Fault tectonics at Tynagh mine, Ireland. Transactions of the Institution for Mining and Metallurgy, 84B, 141-145. MORRIS, P. & MAX, M. D. 1995. Magnetic crustal character in Central Ireland. Geological Journal, 30, 49-68. MURPHY, T. 1952. Measurement of gravity in Ireland - gravity study of Central Ireland. Dublin Institute of Advanced Studies, Memoirs, 3, 31 pp. 1960. Gravity anomaly map of Ireland. Dublin Institute of Advanced Studies, Geophysical Bulletin, 18. 1974. Gravity anomaly map of Ireland. Communications of the Dublin Institute for Advanced Studies, Bulletin, 32.

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J O H N S T O N E T AL.

- - 1 9 8 1 . Geophysical evidence. In: HOLLAND, C. H. (ed.) A Geology of Ireland. Scottish Academic Press, Edinburgh, 225-229. NAYLOR, D., SEVASTOPULO, G. D. & SLEEMAN, A. G. 1989. Subsidence history of the South Munster Basin, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous sedimentation in the British Isles. Occasional Publication, Yorkshire Geological Society, 6, 99-109. NOLAN, S. C. 1986. The Carboniferous Geology of the Dublin Area. PhD Thesis, University of Dublin. - - 1 9 8 9 . The style and timing of Dinantian synsedimentary tectonics in the eastern part of the Dublin Basin, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Occasional Publication, Yorkshire Geological Society, 6, 83-98. PHILCOX, M. E. 1984. Lower Carboniferous lithostratigraphy of the Irish Midlands. Special Publication, Irish Association for Economic Geology, Dublin. - - 1 9 8 9 . The mid-Dinantian unconformity at Navan, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Occasional Publication, Yorkshire Geological Society, 6, 67-82. , SEVASTOPULO, G. D. & MACDERMOT, C. V. 1989. Intra-Dinantian tectonic activity on the Curlew Fault, north-west Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Occasional Publication, Yorkshire Geological Society, 6, 55-66. PHILLIPS, W. E. A. & SEVASTOPULO,G. D. 1986. The stratigraphic and structural setting of Irish mineral deposits. In: ANDREW, C. J., GROWL, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 1-30.

PICKARD, N. A. H., JONES, G. LL., REES, J. G., SOMERVILLE, I. D. & STROGEN, P. 1992. Lower Carboniferous (Dinantian) stratigraphy and structure of the Walterstown-Kentstown area, Co. Meath, Ireland. Geological Journal, 27, 35-58. PRICE, C. A. & TODD, S. P. 1988. A model for the development of the Irish Variscides. Journal of the Geological Society, London, 145, 935-939. READMAN, P. W., O'REILLY, B. M., EDWARDS, J. W. F., SANKEY, M. J. 1996. A gravity map of Ireland and its surrounding waters. In: CROKER, P. F. & SHANNON, P. M. (eds) The Petroleum Geology of Ireland's Offshore Basins. Geological Society of London, Special Publication, 93, 9-16. REILLY, T. A. 1986. A review of vein mineralization in SW County Cork, Ireland. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M.

& PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 513-544. ROTHERY, E. 1988a. En-echelon vein array development in extension and shear. Journal of Structural Geology, 10, 63-71. 1988b. Transpression in the Variscan foreland: a study in east-central Ireland. Irish Journal of Earth Sciences, 10, 1-12. RUSSELL, M. J. 1968. Structural controls of base metal mineralization in relation to continental drift. Transactions of the Institution for Mining and Metallurgy, 77B, 11-28. - - 1 9 7 2 . The Geological Environment of Post Caledonian Base Metal Mineralization in Ireland. PhD Thesis, University of Durham. & HASZELDINE, R. 1992. Accounting for geofractures. In: BOWDEN, A., EARLS, G. V., O'CONNOR, P. & PYNE, J. (eds) The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin, 135-142. SANDERSON, O. J. 1984. Structural variation across the northern margin of the Variscides in NW Europe. In: HUTTON, D. H. W. & SANDERSON, D. J. (eds) Variscan Tectonics of the North Atlantic Region. Geological Society, London, Special Publication, 14, 149-166. SEVASTOPULO, G. D. S. 1981. Lower Carboniferous. In: HOLLAND, C. H. (ed.) A Geology of Ireland. Scottish Academic Press, Edinburgh, 147-172. SHEARLEY, E., HITZMAN, M., WALTON, G., REDMOND, P., DAVIS, R., KING, M., DUFFY, L. & GOODMAN, R. 1992. Structural control' on mineralization and dolomitisation, Lishe~.n Zn-Pb-Ag deposit, County Tipperary, Ire'.,nd. Geological Society of America, Abstro .s with Programs, 34, A354. , REDMOND, P. KING, M. & C jODMAN, R. 1996. Structural controls on rr',~eralization and dolomitisation, Lisheen Zn-" o--Ag deposit, Co. Tipperary, Ireland. This ," .ume. SLOWLY, E. P. 1986. The ~,l-Pb and barite deposits at Keel, Co. Lot ,.ord. In: ANDREW, C. J., CROWE, R. W ,., FINLAY, S., PENNELL, W. M. & PYNF J. F. (eds) Geology and Genesis of Minerr" ,3eposits in Ireland. Irish Association for Ec,- Jmic Geology, Dublin, 319-330. SOMERVi ,.E, I. D., STROGEN, P. & JONES, G. EL. 1992. The biostratigraphy of Dinantian limestones and associated volcanic rocks in the Limerick Syncline, Ireland. Geological Journal, 27, 201-220. STEED, G. M. 1986. The geology of the Gortdrum Cu-Ag-Hg orebody. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 481-499. STROGEN, P. & SOMERVILLE, I. D. 1984. The stratigraphy of Upper Palaeozoic rocks in the Lyons Hill area, Co. Kildare. Irish Journal of Earth Sciences, 6, 155-173.

BASE M E T A L M I N E R A L DEPOSITS, I R E L A N D , & JONES, G. LL. 1990. Stratigraphy and sedimentology of Lower Carboniferous (Dinantian) boreholes from west Co. Meath, Ireland. Geological Journal, 25, 103-137. WILLIAMS, B. & BROWN, C. 1986. A model for the genesis of Zn-Pb deposits in Ireland. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W.

M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 579-590. ZIEGLER, P. A. 1988. Evolution of the Arctic-North Atlantic and the western Tethys. American Association of Petroleum Geologists, Memoir, 43, 1-198.

Geological controls on mineralization and dolomitization of the Lisheen Zn-Pb-Ag deposit, Co. Tipperary, Ireland E. S H E A R L E Y ,

P. R E D M O N D ,

M. K I N G

& R. G O O D M A N

CSA Ltd, Parkview House, Beech Hill, Clonskeagh, Dublin 4, Ireland

Abstract:

The Lisheen deposit is a stratabound Zn-Pb-Ag deposit hosted in dolomitized Waulsortian limestone of Courceyan age. The deposit was discovered in April 1990 and the current resource is estimated to be 22 million tonnes at 11.5% Zn, 2% Pb and 16% Fe, with 26 g/t Ag. Lisheen is located along the Rathdowney Trend, a NE-SW trending, structurally controlled belt of carbonate rocks. Sulphides are present in the hanging wall of two major E to NE-trending normal faults, the Killoran and Derryville faults. These structures are large, northerly-dipping normal faults with displacements of approximately 200 m. The Killoran and Derryville structures form short fault segments of a major relay fault system that extends along the Rathdowney Trend. Faults are observed to have influenced sedimentation, and exercized a major control on the distribution of mineralization, hydrothermal alteration and dolomitization. Several phases of dolomitization, which affect Chadian rocks and temporally overlapped mineralization, help to constrain the timing of this mineralization to no earlier than the late Chadian. The Lisheen deposit is located close to the southeastern edge of the Rathdowney Trend, a 4 0 k m long, N E - S W trending, structurally controlled belt of carbonate rocks extending from Abbeyleix in County Laois to Thurles in County Tipperary (Figs 1 and 2). Lisheen is situated approximately 130km southwest of

Dublin and 10km northeast of the town of Thurles (Fig. 1). The Lisheen deposit forms the largest of several base metal deposits and subeconomic sulphide occurrences which occur in the Lower Carboniferous limestones of the Rathdowney Trend. It lies 8 k m southwest of the Galmoy deposit, which was discovered in

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From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 23-33.

24

E. SHEARLEY E T AL.

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1986 (Doyle et al. 1992). The Lisheen deposit is thought to be similar to the Silvermines zinclead deposit (10.7 million tonnes (Mr) at 7.4% Zn and 2.7% Pb), also a Lower Carboniferoushosted base metal deposit located approximately 45kin to the east, and a producer until 1982 (Andrew 1986). Base metal mineralization at Lisheen occurs principally in the hanging wall rocks of the Killoran and Derryville faults (Figs 3 and 4). These E-W trending and north-dipping normal faults, with maximum displacements of approximately 200 m, form part of a major 40 km long relay fault system that helps define the NE-SW Rathdowney Trend (Fig. 1). The deposit consists of three principal zones of mineralization: the Main, Derryville and North zones. Mineralization extends 2 km along the strike length of the fault system and approximately 1.5kin to the north. Mineralization formed as a series of massive to semi-massive stratabound sulphide lenses, and more irregular zones of breccia/vein style mineralization and finely disseminated sulphides. Individual sulphide lenses vary in thickness from <0.5m to >30m, and thicken and increase in metal grade towards the series of

normal faults at a depth of 170-210 m below the surface. Mineralization at Lisheen is predominantly hosted in the dolomitized Waulsortian Limestone complex, with subordinate mineralization hosted in the Lisduff Oolite Member of the Ballysteen Limestone Formation adjacent to the Killoran and Derryville faults (Fig. 4). Argillaceous bioclastic limestone (upper part of the Ballysteen Formation) forms the footwall to the Waulsortian-hosted mineralization over most of the deposit area. Although the sulphide bodies are both stratabound and flat-lying, they formed by replacement of the basal part of the Waulsortian sequence. Sulphide mineralization is hosted within a hydrothermal dolomite breccia, termed black matrix breccia (BMB). Where well developed, BMB forms stratabound lenses of matrix-supported breccia with angular to sub-rounded clasts of pre-breccia dolostone. This brecciation event pre-dates the mineralization and is thought to have increased the permeability of the dolomitized Waulsortian mudbank limestone and controlled the distribution of sulphide mineralization. This breccia

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26

E. SHEARLEY E T A L .

commonly forms an envelope to the massive sulphide mineralization. Several phases of dolomitization, which temporally overlapped mineralization, help to constrain the timing of mineralization to no earlier than the late Chadian.

Stratigraphy The southern and central areas of the Rathdowney Trend are underlain by the Lower Carboniferous (Courceyan) Waulsortian Limestone complex, which hosts the Lisheen deposit, and the stratigraphically lower Ballysteen Formation (often referred to as the Argillaceous Bioclastic Limestone). The Devils Bit Mountains, a Silurian inlier which is unconformably overlain by sandstones and conglomerates of Upper Devonian age, is located 7km to the northwest of the Lisheen area. The Upper Carboniferous sandstone and coal-bearing sequences of the Slieve Ardagh area lie approximately 8 km to the southeast (Fig. 1). The Lower Carboniferous succession at Lisheen (Fig. 5) is similar to that recognized elsewhere in south central Ireland (ShephardThorn 1963; Philcox 1984; Somerville & Jones 1985; Somerville et al. 1992). It records a marine transgression over Old Red Sandstone (ORS) alluvial plain deposits and the establishment of continuous marine, dominantly carbonate sedimentation. The initial transgression at Lisheen resulted in deposition of a shallow-water sequence of mixed carbonates and clastics of the Mellon House Formation (approximately 35m thick). This was followed later by deeper water shales and limestones of the Ringmoylan Formation (30-35m thick), the Ballymartin Formation (c. 50m thick), and the Ballysteen Formation (c. 370 m thick). A distinctive shale horizon, the Ballyvergin Mudstone Formation (2.5-6 m thick), occurs between the Ringmoylan and Ballymartin Formations. The transgression culminated in the deposition of the Waulsortian Limestone complex of the Limerick Limestone Formation (approximately 200m thick) during late Courceyan to early Chadian time. Waulsortian development was succeeded by an openmarine limestone shelf to ramp sequence of the Crosspatrick Formation (165-175 m thick), and an overlying shallow water, more restricted shelf sequence, the Aghmacart Formation (>200m thick), of Arundian age. The Ballymartin Formation conformably overlies the Ballyvergin Mudstone Formation and

consists of dark, thin-bedded, shelly calcarenites with interbedded argillites. Conformably overlying the Ballymartin Formation is the Ballysteen Formation, which, in the Lisheen area, can be subdivided into three members: the Lower Calcarenite Member, the Lisduff Oolite Member, and the Upper Calcarenite Member. The Lower Calcarenite Member (130-150m thick) consists of a sequence of fossiliferous packstones and grainstones. The overlying Lisduff Oolite Member (Sleeman 1990) is 70-80m thick and consists of well-sorted oolitic grainstones and crinoidal packstones and wackestones. The Upper Calcarenite Member (130-140m) consists of argillaceous, fossiliferous packstones with interbedded argillaceous wackestones. The upper 20-30m of this member consists of moderately fossiliferous, often nodular micrites interbedded with dark argillaceous wackestones and packstones, and is informally referred to as the Nodular Micrite Unit (equivalent to the Ballynash Member of Somerville & Jones 1985). Chert is common in the uppermost 5-6 m of the Nodular Micrite Unit, typically restricted to micrite nodules and bands of hydrothermal silica (up to 5cm thick) that occur near the contact with the overlying Waulsortian limestone sequence. Locally within the Lisheen area, close to the contact with the overlying Waulsortian, normally black argillites within the Nodular Micrite Unit are altered to green-grey and partially dolomitized. This is thought to be due to hydrothermal alteration associated with the Lisheen mineralizing event. Two very distinctive green tuff horizons, which rarely exceed 2 cm in thickness, occur within the Nodular Micrite Unit, approximately 13 m and 16 m below the top of the unit, and are important marker horizons both in the immediate vicinity of the Lisheen deposit and throughout the Rathdowney Trend. Lithologies in the Ballymartin and Ballysteen Formations are indicative of moderate to shallow water depths suggesting deposition on a broad shelf or ramp. The members within these formations indicate a number of sea-level changes through time. The Waulsortian Limerick Limestone Formation is the principal host rock for mineralization and in the Lisheen area ranges from 170-210 m in thickness. It consists of many coalesced mudbanks, with only minor development ofinterbank flank facies. The base of the Waulsortian complex appears to be approximately planar in the Lisheen area. The Waulsortian complex in the vicinity of the orebody is almost entirely dolomitized; however, the original Waulsortian textures and facies types can still be distinguished. The basal facies of the Waulsortian complex is generally 5 to

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10 m thick and consists of irregularly distributed crinoidal micrites, encrinitic calcarenites, and grey-green argillites. The majority of the Waulsortian complex comprises core facies which can be divided into two subfacies: 'veines bleues' and biomicrite. The 'veines bleues' subfacies (Lees et al. 1977) is distinguished by an abundance of sparry calcite and commonly contains abundant

stromatactis cavities. The biomicrite subfacies is composed of micrite mud with variable amounts ofbioclastic debris, primary fenestellid bryozoans and crinoids. The flank facies within the Waulsortian define the boundaries between individual mudbanks, and this facies is irregularly distributed throughout the Waulsortian complex at Lisheen. The flank facies consist of crinoidal

28

E. SHEARLEY E T AL.

micrites with minor to abundant argillite laminae, which give the rock a wavy laminated texture. The Waulsortian core facies is overlain by the cover facies, which at Lisheen is lithologically similar to the flank facies and consists of crinoidal micrite, commonly with a wavy laminated texture. The cover facies is termed the Upper Wavy Laminated Unit (Figs 3, 4 and 5) and is generally 10-20m thick. Adjacent to the Killoran and Derryville faults, however, the cover facies thickens dramatically to over 70 m, with the thickest development corresponding to the zones of maximum displacement on the main faults (Figs 3 and 4). Where the Upper Wavy Laminated Unit increases in thickness there is a corresponding decrease in the thickness of the Waulsortian complex core facies, and as a result the overall thickness of the entire Waulsortian complex remains relatively constant. Dips in the Upper Wavy Laminated Unit adjacent to the Killoran and Derryville faults range up to 45 ° . These high dips are typical of the flank facies of a mudbank complex. The Upper Wavy Laminated Unit is interpreted as being a slightly deeper-water facies of the Waulsortian complex, and its distribution suggests that it was deposited in fault-parallel sags that developed midway through the build-up of the Waulsortian Complex. The Crosspatrick Formation conformably overlies the Waulsortian and is approximately 165-175m thick in the Lisheen area. It consists of dark grey, cherty, argillaceous packstones and wackestones. The contact between the uppermost facies of the Waulsortian and the overlying Crosspatrick Formation is gradational. In the local fault-parallel depressions or sags, the lowermost part of the Crosspatrick Formation appears to be laterally equivalent to uppermost Waulsortian micrites and biomicrites. In the Lisheen area the lowermost part (70-80m) of the Crosspatrick Formation is often dolomitized.

Structure Two major E to ENE-trending normal faults, the Killoran and Derryville faults, bound the southern edge of the Lisheen deposit and are the primary controls on the distribution of the mineralized zones (Fig. 3). The hanging walls of these faults contain a number of subsidiary structures that appear also to have controlled alteration and mineralization. The major faults at Lisheen form part of a major, 40 km long, segmented fault system that extends along the Rathdowney Trend. Regionally, carbonate

rocks along the Rathdowney Trend dip at 5-8 ° to the SE. Within the ore body area, in the hanging walls of the main faults, dips are typically < 10°, but dip direction is variable in this structurally deformed area. The Killoran and Derryville faults trend ENE to E-W, have approximately 200 m maximum displacement, and have identified surface traces of 2.2 km and 2 km respectively. Both faults have subplanar surfaces, dip at 45-55 ° to the north in the zone of maximum displacement, and tend to steepen to >60 ° towards their termination zones. Displacement is interpreted to be transferred from the Killoran fault to the Derryville fault by means of a relay ramp (Walsh & Watterson 1989). Within the relay area, displacement along the Killoran fault decreases rapidly to the east, and displacement along the Derryville fault decreases rapidly to the west. The Killoran fault exhibits a decrease in displacement from approximately 200 m to 90 m over a distance of 500m, while the Derryville fault exhibits a decrease in displacement from approximately 210 m to 60 m over a distance of 450 m. Therefore a combined fault displacement of approximately 200m is maintained through the relay area. Bed dips in the relay area are typically 10-15 ° to the west. A number of roughly N-S trending, steeply dipping normal faults, with downthrows of 10-30 m to the west, also occur in the relay area between the faults. This evidence suggests that the displacement transfer between the Killoran and Derryville faults is accommodated by both normal faulting and increased bed dip. The Barnalisheen fault, an E - W trending, northward-dipping normal fault with a displacement in excess of 100m, is situated l km northwest of the Lisheen deposit (Fig. 3). A series of easterly dipping normal faults, with displacements of 15-25m, occur in the relay zone between the Killoran and Barnalisheen faults. In this area beds dip at 10-15 ° to the east. This relay zone has a major influence on the structure of the westernmost part of both the Main and North zones. The hanging wall of the Killoran and Derryville faults is cut by a number of subsidiary E - W trending normal faults that strike subparallel to the main faults (Fig. 3). These faults are vertical to steeply north-dipping, with apparent normal displacements in the order of 5-25m. Most probably merge with the major structures at depth (see faults labelled F-57, F-41, Figs 3 and 4). A number of approximately N-S trending subvertical normal faults, with displacements of the order of 10-40 m also occur in the hanging

THE LISHEEN ZN-PB-AG DEPOSIT walls of the Killoran and Derryville faults (Fig. 3). These N-S faults are thought to have formed as accommodation structures due to rapid lateral changes in normal displacement along the Killoran and Derryville faults. They control the extent of the supra-Waulsortian lithologies in the central and northern areas of the deposit (Fig. 3). Deposition of the Waulsortian Upper Wavy Laminated Unit is interpreted to have been influenced by movements on the Killoran and Derryville faults. This unit is interpreted as being a slightly deeper water facies of the Waulsortian complex, and its distribution suggests that it was deposited in fault-parallel sags. This suggests that movement on these faults was initiated midway through the buildup of the Waulsortian complex, during the late Courceyan. The major faults in the Rathdowney area are fairly steeply dipping (45-60°), do not appear to shallow at depth, and are thought to have formed due to the reactivation of Caledonian basement structures. As a result they may have acted as conduits between the Waulsortian and the ORS and underlying basement for the hydrothermal dolomitizing and later mineralizing fluids.

Regional ReplaciveDolomite

!

Regional

29

dolomitization

Waulsortian limestones and non-argillaceous beds within the overlying Crosspatrick Formation have been dolomitized over large areas of the Rathdowney Trend. The contact between regionally dolomitized and undolomitized Waulsortian occurs approximately 2km to the northwest of Lisheen. The contact is gradational, with dolomitization best developed in the upper portions of the Waulsortian. This contact forms a NE-SW trending regional 'dolomite front' that subparallels the structurally defined Rathdowney Trend (Hitzman et al. 1992) (Fig. 1). At the Lisheen deposit the entire Waulsortian sequence, and relatively non-argillaceous beds within the overlying lower Crosspatrick Formation, are strongly dolomitized. Locally, however, small portions of the basal Waulsortian facies are undolomitized or only weakly dolomitized. This appears to be due to the presence of thin argillite bands in the basal Waulsortian facies that may have acted as impermeable barriers to the dolomitizing fluids. In the northern part of the Rathdowney Trend, regional dolomitization has been observed to extend locally through the overlying Crosspatrick Formation and into the lower Aghmacart Formation.

I

Regional CoarseWhiteDolomite DarkGreyDolomlte

!

Black Matrix Breccia(BMB)

|

I !

Sulphide Mineralization

| > Fine Grained Color Banded Sphalerite >Fine Grained Iron Sulphide >Colloform Iron Sulphide >Colloform Sphalerite >Brecciation, Sulphide Replacement by Sphalerite - Galena - Tennantite >Chalcopyrite >Honeyblende Shalerite >Barite - Quartz

Hydrothennal FerroanDolomite

I

Late Dolomite/ Calcite

I .9 |

Time

Fig. 6. Sequence of dolomitization and mineralization at Lisheen.

,,

30

E. SHEARLEY ET AL.

Regional dolomite is composed of fine to medium crystalline, replacive dolomite, and texturally later coarsely crystalline, white dolomite (Fig. 6). The replacive dolomite is grey to buff in colour, and staining with potassium ferricyanide (Dickson 1966) indicates that it is non-ferroan. Individual crystals are euhedral to anhedral, range in size from 20-100#m, and display a xenotopic texture (Gregg & Sibley 1984). This dolomite appears to replace preferentially the microcrystalline components of the Waulsortian, with dolomite crystal size being partly controlled by the grain size of the precursor Waulsortian calcite. Selective replacement of certain Waulsortian components, such as bryozoa, is conspicuous in areas close to the regional dolomite front. The coarse crystalline component of the regional dolomite is white in colour, and staining indicates that it is non-ferroan. This dolomite principally occurs in vugs within the replacive dolomite, but also occurs within cross-cutting veins. Individual crystals are subeuhedral to euhedral and often display a saddle-shaped habit. Crystal size ranges from coarse to very coarse (0.5-5mm). This dolomite appears to replace preferentially the coarse-grained components of the Waulsortian, such as calcite filled cavities and bioclasts. The regional dolomite occurs within the Waulsortian on both the hanging wall and footwall sides of the Killoran-Derryville fault zone. At Lisheen there is no obvious relationship between the intensity of the replacive dolomitization and proximity to the faults that later controlled the distribution of hydrothermal alteration and mineralization. On a larger scale, however, the distribution of the regional dolomite subparallels the NE-SW Rathdowney Trend, and its distribution may be structurally controlled on a regional scale. It is thought that the regional dolomite may have formed as a result of pulses of heated, marine-derived formation waters, possibly from a foreland basin to the south of the carbonate shelf during the onset of the Hercynian Orogeny (Allen et al. 1992; Jones & Fitzsimons 1992). It is possible that the ORS sequence may have acted as an aquifer for the moderately hightemperature hydrothermal solutions, with early Hercynian tectonics providing the mechanism for the expulsion and migration of these brines from the ORS basin to the south. The occurrence of regional dolomite in the Crosspatrick Formation constrains the timing of the regional dolomitization event to no earlier than the Chadian.

Hydrothermal alteration Following regional dolomitization, significant local hydrothermal alteration occurred at Lisheen, consisting primarily of dolomitization (Beaty et al. 1991) and minor silicification. The regional dolomite has been replaced and brecciated by a series of hydrothermal dolomites, which are spatially associated with the massive sulphide bodies. A number of distinct hydrothermal dolomite stages have been recognized based on cross-cutting relationships observed in drillcore and in thin section. The principal stages of hydrothermal dolomitization at Lisheen are: early hydrothermal dark grey dolomite; black matrix breccia (BMB) dolomite; and hydrotherreal ferroan dolomite (Fig. 6).

Early hydrothermal dark grey dolomite The dark grey hydrothermal dolomite is finely crystalline, and staining indicates that it is nonferroan to slightly ferroan. Individual crystals are subhedral to euhedral, ranging in size from 40-100 #m and display a xenotopic texture. The dark grey dolomite appears to be restricted to the lower portion of the Waulsortian, and commonly occurs as clasts within the later dolomite breccias. In volumetric terms the dark grey hydrothermal dolomite is a very small component of the entire dolomite body at Lisheen. The age relationship between the dark grey dolomite and regional dolomite is unclear; the contact between the two appears gradational and no conclusive cross-cutting relationships have been observed. This suggests that the dark grey dolomite may be a textural variation of the regional replacive dolomite, with variations in crystal size accounting for the colour difference. The dark grey dolomite is cut by all other stages of hydrothermal dolomite.

Black matrix breccia ( B M B ) The BMB is a very distinctive hydrothermal alteration product, the recognition of which strongly influenced the discovery of Lisheen (Hitzman et al. 1992). Where well developed the BMB forms stratabound lenses of matrixsupported breccia, with angular to subrounded clasts of regional dolostone. The matrix of the breccia varies in colour from grey to black and is composed of non-ferroan to ferroan, fine crystalline dolomite, with individual crystals ranging in size from <20#m to 50#m. Scanning electron microscopy of the matrix of the breccia indicates

THE LISHEEN ZN-PB-AG DEPOSIT that there is 25-40% void space between the dolomite crystals. The argillite content of the matrix is only significant where the breccia cuts argillaceous horizons, and in close proximity to the massive sulphide lenses the matrix of the breccia commonly contains disseminated sulphides. The dark coloration of the fine crystalline dolomite that forms the matrix of the BMB appears to be due to a combination of factors, including: crystal size, void space between the crystals, argillite content, and the presence of disseminated sulphides. Well developed BMB contains in the order of 40-60% clasts by volume, and megascopically appears matrix-supported. Regional dolostone clasts are the most common, but clasts of early dark grey hydrothermal dolomite and of undolomitized or partially dolomitized Waulsortian Limestone also occur. Clasts are angular to subrounded and vary in size from several hundred microns to >1 m in diameter. Subrounded clasts often exhibit embayed edges and are partly replaced by the fine crystalline black matrix dolomite. Thin section examination using cathodoluminescence petrography also shows that many clasts contain microveinlets and disseminations of the same fine crystalline dolomite that forms the matrix of the breccia. Well developed BMB is restricted to previously dolomitized portions of the Waulsortian, and occurs as irregular stratabound lenses that commonly form a 'halo' surrounding the massive sulphide. It may extend up to 10m above the massive sulphide lenses, and may also be developed between and below sulphide lenses. The contact between the massive sulphide and the overlying BMB is generally sharp, with only minor disseminated and veinlet sulphides occurring above the massive sulphide. Moderately to well-developed BMB extends up to 500m beyond the edge of the economically significant mineralization on the northern edge of the deposit, and up to 200m on the western and eastern edges of the deposit. On the southern margin of the deposit, BMB (and massive sulphide) abuts the Killoran and Derryville faults and locally extends up the fault planes (Fig. 4). Only very minor brecciation occurs on the southern side of these faults. Locally, BMBtype brecciation is also developed in the footwall Lisduff Oolite in close proximity to the Killoran and Derryville faults, associated with sulphides. The lack of significant BMB in the Waulsortian on the footwall side of the main faults, and the geometry of the breccia body, suggest that the brecciation probably took place after significant movement had occurred on the main faults. Most

31

of the Waulsortian-hosted mineralization at Lisheen occurs in BMB. The mineralization post-dates the brecciation and replaced both matrix and clasts, and relict breccia textures are commonly observed in the massive sulphides.

Hydrothermal ferroan dolomite Hydrothermal ferroan dolomite occurs in the Ballysteen Limestone adjacent to the Killoran and Derryville fault zones, within the dolomitized Waulsortian, and in the hanging wall Lisduff Oolit. Hydrothermal ferroan dolomite occurs as both fine crystalline replacive dolomite and as medium to coarse, white, cross-cutting veins and veinlets. This dolomite stage appears to be late in the paragenetic sequence, and cuts regional dolomite, early dark grey hydrothermal dolomite, and BMB.

Other hydrothermal alteration Locally, minor silicification in the form of silica nodules and small zones of pervasive silicification occurs throughout the Lisheen mineralizing system, and is most common in the upper Ballysteen Limestone and the lower Crosspatrick Formation. These zones of silicification are occasionally associated with zones of reddening within the adjacent dolostones and limestones. Silicified zones are cut by the hydrothermal dolomite, indicating that this phase of dolomitization post-dates the silicification. Late carbonate veins cross-cut both the sulphides and the hydrothermal breccias at Lisheen. These consist of both calcite and dolomite veins up to several centimetres wide, which are usually white or pink in colour. These veins may post-date the main phase of hydrothermal activity and may be related to later tectonic activity.

Sulphide mineralization The Lisheen deposit consists of three principal zones of sulphides, the Main, Derryville and North zones, which occur at the base of the dolomitized Waulsortian complex, and two relatively minor sulphide bodies in the Lisduff Oolite on the footwaU side of the Killoran and Derryville faults (Fig. 3). The Main zone is the largest in terms of tonnage, followed by the Derryville and North zones. Mineralization occurs as massive to semi-massive sulphide lenses, more irregular zones of stratabound

32

E. SHEARLEY E T A L .

vein-style sulphides, and finely disseminated sulphides. Within the three zones, individual stratabound sulphide lenses vary in thickness from <0.5 m to >30 m, and thicken and increase in metal grade towards a series of normal faults. Alteration and metal zonation indicate that the main Killoran and Derryville faults, and a number of subsidiary faults in the hanging walls of the main faults, acted as conduits for the hydrothermal mineralizing fluids. In the Main zone the base of the Waulsortian ore body can be subdivided into at least three sulphide lenses, adjacent to the Killoran, F-57 and F-41 faults (Fig. 4). In proximity to the Killoran fault, mineralization occurs as a thick (up to 30m), wedge-shaped massive sulphide body at the base of the Waulsortian Limestone which extends along the hanging wall surface of the Killoran fault. The thickness of the massive sulphide body decreases rapidly in a northerly direction, from 30 m immediately adjacent to the Killoran fault, to 5 m over a horizontal distance of 180 m. In this area the single massive sulphide lens, as well as thinning in a northerly direction, also splits into two lenses separated by BMB and/or dark grey hydrothermal dolomite. Further to the north of the Killoran fault, stratiform massive sulphide lenses occur on the northern side of the roughly E-W trending F-57 and F-41 faults. The thicknesses of these lenses also decrease from south to north (Fig. 4). The Derryville and North zones are likewise composed of a number of individual sulphide lenses at the base of the Waulsortian Limestone, which are separated by E-W trending faults and also decrease in thickness from south to north. Sulphide thickness appears to be principally controlled by E-W trending faults. However, in some areas of the deposit minor N-S trending faults also appear to control sulphide thicknesses. This fault pattern of E-W and N-S faults subdivides the orebody into over 30 faultbounded blocks, each with variations in mineral thickness and, to a lesser extent, grade. Metal grades have better lateral continuity than sulphide thickness and do not vary to the same degree within individual sulphide lenses. In general, thicknesses are greater on the downthrown side of both E-W and N-S structures. Other variations in mineral thickness may be due to irregularities in the replacive nature of the mineralizing process. The thickest sulphide bodies at Lisheen occur immediately adjacent to the main KiUoran and Derryville faults in the areas of maximum displacement. The geometry of the Lisheen deposit suggests that faulting exercized a primary control on sulphide dis-

tribution, and that individual faults served as conduits for mineralizing fluids. The bulk of the mineralization at Lisheen is hosted in BMB. The sulphides appear to have replaced both the matrix and clasts within the breccias, and relict breccia textures are commonly observed in massive sulphide. Locally, sulphides also replace undolomitized Waulsortian Limestone and more argillaceous horizons at the base of the Waulsortian Limestone. The mineral assemblage at Lisheen consists of pyrite-marcasite-sphalerite-galena, with minor chalcopyrite, tennantite, silver, arsenopyrite, and various lead sulphosalt minerals. Gangue minerals are principally dolomite, with pyrite, marcasite, some barite, siderite, calcite, and very small amounts of quartz and illite. Chalcopyrite, tennantite, arsenopyrite and barite are most abundant near the Killoran and Derryville faults, especially within the Lisduff Oolitehosted mineralization. In general the ore forms three textural types: fine-grained disseminated ore, coarser grained colloform banded ore, and complex fine and coarser grained replacement ore (Hitzman et al. 1992). The earliest sulphides consist of fine-grained colour-banded sphalerite spheres and finegrained iron sulphides. Later massive iron sulphides (interbanded pyrite-marcasite-sphalerite) form concentrically zoned colloform masses, and commonly enclose early sphalerite spheres. A distinctly later generation of sphalerite has a coarse-grained banded texture, and is colourzoned from dark to light amber. Fracturing and brecciation also occurred during the mineralization sequence. This varies from a fracturing of sulphides to the development of a fine-grained sulphide-matrix breccia containing fragments of sphalerite and colloform iron sulphide. Late veins that cut the banded sphalerite and intermineral breccias are filled with dolomite, calcite, sphalerite and ferroan carbonate.

Conclusions The Lisheen deposit is a stratabound replacement ore body, hosted in dolomitized Waulsortian limestones, which lies along the Rathdowney Trend, a structurally deformed belt of carbonate rocks extending between Abbeyleix and Thurles. Major faults, with normal displacements of up to 200 m, form short fault segments of a major 40 km long relay fault system that subparallels and defines this trend. Fault geometry influences sedimentation within the Waulsortion Upper Wavey Laminated Unit of Courceyan age. The distribution of this

THE LISHEEN ZN-PB-AG DEPOSIT unit appears to be controlled by the major E - W trending structures, and is interpreted as having been deposited in fault-parallel sags or depressions that were initiated during the Courceyan. The structural geometry also exercized a control on the distribution of mineralization, alteration and dolomitization. Structures formed before mineralization, and provided conduits for local hydrothermal dolomitizing solutions and mineralizing fluids. Major conduits for mineralizing solutions are located along the major E - W trending structures, particularly at points of maximum throw. There is no evidence of significant post-mineralization movements on these structures. Waulsortian limestones were converted to regional dolostone over a large part of southeastern Ireland, prior to mineralization. It is thought that the regional dolomite may have formed as a result of pulses of heated, marinederived formation waters, possibly from a foreland basin to the south of the carbonate shelf during the onset of the Hercynian Orogeny. It is possible that the ORS sequence may have acted as an aquifer for the moderately high temperature hydrothermal solutions, with early Hercynian tectonics providing the mechanism for the expulsion and migration of these brines from the ORS basin to the south. The major fault systems, which probably formed due to the reactivation of Caledonian basement structures, appear to have acted as conduits between the ORS/underlying basement and the Waulsortian for the hydrothermal dolomitizing and later mineralizing fluids. Regional dolomite has been recorded from late Chadian rocks, around 200 m above the top of the Waulsortian sequence, in the Rathdowney Trend area. As later hydrothermal dolomitization and mineralization both cross-cut the regional dolomite, this constrains the time of mineralization to no earlier than the lower Chadian.

References ALLEN, J. R., BEATY, D. W., STURTEVANT, R. G., HITZMAN, M. W. & SHEARLEY, E. 1992. The origin of regional dolomite in the Waulsortian of Southeast Ireland: Implications for time of ore deposition. Abstracts of the Geological Society of America annual meeting, Cincinnati, Ohio, 24, A354. ANDREW, C. J. 1986. The tectono-stratigraphic controls to mineralization in the Silvermines area, County Tipperary, Ireland. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in

33

Ireland. Irish Association for Economic Geology, Dublin. 377-417. BEATY, D. W., ALLAN, J. R. & HITZMAN, M. W. 1991. The relationship of dolomitization and mineralization of the Lisheen deposit, south central Ireland. Abstracts of the Geological Society of America, 23, A172. DICKSON, J. K. 1966. Carbonate identification and genesis as revealed by staining. Journal of Sedimentary Petrology, 36. 491-505. DOYLE, E., BOWDEN, A. A., JONES, G. V. & STANLEY, G. A. 1992. The geology of the Galmoy zinc-lead deposits, Co. Kilkenny. In: BOWDEN. A. A., EARLS, G., O'CONNOR, P. G. & PYNE, J. F. (eds) The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin, 211-226. GREGG, J. M. & SIBLEY, D. F. 1984. Epigenetic dolomitization and the origin of xenotopic dolomite texture. Journal of Sedimentary Petrology, 54, 908-931. HITZMAN, M. W., O'CONNOR, P., SHEARLEY, E., SCHAFFALITZKY, C., BEATY, D. W., ALLEN, J. R. & THOMPSON, T. 1992. Discovery and geology of the Lisheen Zn-Pb-Ag prospect, Rathdowney Trend, Ireland. In: BOWDEN, A. A., EARLS, G., O'CONNOR, P. G. & PYNE, J. F. (eds) The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin, 227-246. JONES, G. LL. & FITZSIMONS, J. 1992. A pure dolomite deposit northeast of Cahir, Co. Tipperary. In: BOWDEN. A. A., EARLS, G., O'CONNOR, P. G. & PYNE, J. F. (eds) The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin, 327-334. LEES, A. NOEL, B., & BOUW, P. 1977. The Waulsortian "reefs" of Belgium: a progress report. Memoirs of the Institute of Geology, University of Louvain, 29. 289-315. PHILCOX, M. E. 1984. Lower Carboniferous Lithostratigraphy of the Irish Midlands. Irish Association for Economic Geology, Dublin. SHEPHARD-THORN, E. R. 1963. The Carboniferous limestone succession in north-west County Limerick, Ireland. Proceedings of the Royal Irish Academy, 62B, 267-294. SLEEMAN, A. 1990. Lexicon entries for formations depicted on the Geological Survey of Ireland 1 : 1000000 compilation of Rathdowney District. Unpublished lexicon, Geological Survey of Ireland, lexicon. SOMERVILLE, I. D. & JONES, G. LL. 1985. The Courceyan stratigraphy of the Pallaskenry Borehole, County Limerick, Ireland. Geological Journal, 20, 377-400. --, STROGEN, P. & JONES, G. LL. 1992. The Biostratigraphy of Dinantian Limestones and associated volcanic rocks in the Limerick Syncline, Ireland. Geological Journal, 27, 201-220. WALSH, J. J. & WATTERSON, J. 1989, Displacement gradients on fault surfaces. Journal of Structural Geology, 11,307-316.

The use of burial diagenetic calcite cements to determine the controls upon hydrocarbon emplacement and mineralization on a carbonate platform, Derbyshire, England CATHY

HOLLIS 1 & GORDON

WALKDEN

Department of Geology and Petroleum Geology, University of Aberdeen, Meston Building, Kings College, Aberdeen AB9 2UE, UK 1 Present address." Badley Ashton and Associates, Winceby House, Winceby, Horncastle, Lincolnshire LN9 6PB, UK

Abstract: Late diagenetic calcite cements in the Upper Dinantian limestones of the Derbyshire Platform are contemporaneous with both hydrocarbon emplacement and Mississippi Valley-type (MVT) mineralization. Calcite cementation began during the progressive burial of the Derbyshire Platform and the surrounding basins, principally within fractures generated during the waning effects of Upper Carboniferous extension. Six burial calcite cements can be recognized in dilationai vein systems. Successive veins contain progressively more mature hydrocarbon inclusions, and calcite cements are intergrown with fluorite, baryte, galena and sphalerite in increasing quantities. Compacting DinantianNamurian shales in basins adjacent to the platform offer the most likely sources of fluids, trace elements and hydrocarbons. Fluids entered the platform along major fault systems, and circulated using smaller fracture systems, precipitating calcite. The final phases of calcite cementation and the main phase of MVT mineralization coincided with the onset of the Variscan Orogeny. A model is now established relating fluid flow to Variscan tectonic events in northern Britain. Published mineralogical and diagenetic studies of the Derbyshire-East Midland Platform have focused either upon mineralization and mineralizing fluids (Ford 1968; Ineson & Ford 1982) or upon the early and late diagenetic history of the area as revealed by intergranular cements (Walkden & Berry 1984; Walkden & Williams 1991). This study is the first to integrate these fields by developing a complete model for the geochemical evolution of the Derbyshire Platform and the surrounding basins, in line with the regional tectonic history of the area and the control this imposed upon fluid flow.

Geological and tectonic history The Derbyshire Platform (Fig. 1) developed as a stable block within the Pennine Basin during Lower Carboniferous back-arc extension, surrounded by the more rapidly subsiding Edale, Staffordshire and Widmerpool Basins (Leeder 1988; Fraser et al. 1990; Fig. 2). Present day exposure on the platform is mainly of Asbian and Brigantian shelf limestone (Aitkenhead et al. 1985), which was deposited in cyclic sequences in response to minor sea level fluctuations (Walkden 1974, 1987). Carbonate production halted at the end of the Dinantian as extension waned. Regional thermal subsidence became the

overriding control on sedimentation and, although north-south extension continued into the late Westphalian, its effects were greatly reduced (Guion & Fielding 1988). Namurian to Westphalian prograding fluviodeltaic systems filled the surrounding basins and buried the Derbyshire Platform throughout the Upper Carboniferous (Kelling & Collinson 1992). The onset of the Variscan Orogeny in the midStephanian led to basin inversion across Northern Britain as a compressional regime was established, and N W - S E and N E - S W trending normal faults (Fig. 2) were reactivated, often giving them a strike-slip component (Smith & Smith 1989). The area was probably completely exposed by the mid-Triassic (Walkden & Williams 1991). There is no evidence to suggest that the Derbyshire Platform was reburied in the Mesozoic as has been suggested for the East Midlands Shelf (Green 1989).

Diagenesis, mineralization and hydrocarbon emplacement Diagenesis The diagenetic history of the Derbyshire Platform carbonates is characterized by four distinct calcite cement zones, identified in syntaxial overgrowth

From STROGEN, P., SOMERVILLE,|. D. & JONES, G. LE. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 35-49.

36

C. HOLLIS & G. WALKDEN

sequences using cathodoluminescence (CL) (Fig. 3; Walkden & Williams 1991). Early nonluminescent Zone 1 and thin bright luminescent Zone 2 cements were both precipitated in the meteoric phreatic environment. Dull luminescent Zone 3 calcite occluded most remnant porosity during shallow burial of the platform (Walkden & Williams 1991; Bingham 1992). They are succeeded by deep burial Zone 4 cements, which post-date dolomitization (Hollis 1995) and are contemporaneous with both P b - Z n - B a - F mineralization and hydrocarbon emplacement. Only shallow burial Zone 3 and deep burial Zone 4 cements are considered in this study; these have been further subdivided in this work (Table 1).

200°C (Coleman et al. 1989). The source of sulphur for mineralization is uncertain, but the range of/534S values and the absence of equilibria within sulphide and sulphate minerals suggests that sulphur was probably sourced from sulphate-rich formational fluids within the platform limestone (Coleman et al. 1989). Mineralization is primarily concentrated along Variscan normal faults and strike-slip faults (Fig. 2) and joint systems (predominantly N E - S W and N W - S E trending). Where mineralizing fluids encountered impermeable horizons, such as basalt horizons or

(I) "0

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Upper Dinantian limestone (including dolomitized beds) on the Derbyshire Platform hosts formerly economic galena and fluorite mineralization, with baryte gangue and minor sphalerite. The mineralization has previously been considered to be of Mississippi Valley-type (MVT) since it is carbonate-hosted and thought to have formed from fluids at temperatures up to

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.

Caledonian basement faults

Fig. 2. Dinantian palaeogeography of the Derbyshire-East Midlands Shelf (after Walkden & Williams 1991). WP, Welsh Platform; WBM, Wales-Brabant Massif; DP, Derbyshire Platform; EMCS, East Midlands carbonate shelf; GB, Gainsborough Basin; EB, Edale Basin; SB, Staffordshire Basin; WB, Widmerpool Basin. shale beds, metasomatic replacement of limestone formed mineral deposits concordant or semi-concordant with bedding (Ford 1968).

Hydrocarbons

the Derbyshire Platform (Quirk 1987; Coleman et al. 1989). Mechanisms for the release and

transport of trace elements within the fluid are poorly understood, however, and the timing of these events poorly constrained in terms of the burial history of the area.

Hydrocarbon accumulations are common across the Derbyshire Platform, and are well documented (Pering 1973; Ewbank et al. 1993). No economic hydrocarbon deposits have been found, but traces of oil and solid hydrocarbon are common within remnant intergranular and shelter porosity and are also distributed along vein and fracture systems.

F l u i d source

Fluids responsible for mineralization on the Derbyshire Platform are thought to have been basinal brines, expelled from Upper Carboniferous shale basins adjacent to the platform (Quirk 1987; Jones & Plant 1989). The basins were buried to a depth suitable for hydrocarbon generation and contained sufficient concentrations of trace elements to mineralize completely

Fig. 3. Non-luminescent Zone 1 (a) and bright luminescent Zone 2 (b) syntaxially overgrown by poreoccluding, dull brown luminescent Zone 3 (c) calcite. The whole cement sequence has been cross-cut by Zone 4 calcite veinlet (d). Cathodoluminescence. Millers Dale Station. Complete scale bar is 0.5 mm.

Crystal shape

Columnar Columnar

Columnarblocky

Blocky

Blocky

Blocky

Cement zone

3A 3B

4A

4B

4C

4D

Bright orange

Bright orange

Dull brownorange, Concentric subzonation

Dull orange

Dull brown Dull brown

CL

Unit

Unit

Unit undulose

Undulose

Undulose Undulose

Extinction

Straight

Straight

Straight

Bent

Bent-straight Bent

Twin and cleavage planes

Dull lime green

Dull blue-green

Dull blue-green

Dull blue-green

Dull yellow-brown Dull yellow-brown

Cement fluorescence

Micro-dolomite

Rare greenbrown

Rare fluid

Micro-dolomite

Green brown

Quartz Fluorite Quartz Pyrite Fluorite Quartz Pyrite Sphalerite Baryte Fluorite Sphalerite Baryte Fluorite -

Fluid Micro-dolomite Micro-dolomite

Other minerals

Other inclusions

Green

White

Hydrocarbon inclusion fluorescence

Table 1. Summary of petrographic characteristics of individual cement phases and interpretation of their environment of precipitation.

Deep burial and uplift

Deep burial and uplift

Deep burial

Shallow burial Intermediate burial Deep burial

Burial environment

Z

>" t"

t-" © t-"

¢3

BURIAL DIAGENESIS, DERBYSHIRE

Techniques Burial calcite cement Zones 3 and 4, within intergranular pore systems and small (<5cm wide) fracture systems, were described from polished sections of Asbian and Brigantian limestones sampled from across the Derbyshire Platform (Fig. ~). Calcite samples were also collected across metre-scale mineralized strikeslip and normal fault systems. Basic petrographic description of calcite cements and mineral assemblages employed cathodoluminescence (CL) and epifluorescence (blue light) techniques. The cements were further categorized using stable isotope and trace element geochemistry, and fluid inclusion microthermometry. Full analytical details are given in Hollis (1995). The carbon and oxygen isotopic compositions of 105 individual cements and whole rock limestone samples were analysed. Veins with only one cement zone were sampled (around 10mg) using a tungsten-tipped dentist's drill, but several cements, in particular Zone 4D, frequently do not occur in large enough concentrations to be isolated for isotopic analysis. Eighteen whole rock limestone samples were selected away from calcite veins to remove any influence of vein calcite cements on whole rock limestone composition. Results were corrected using standard procedures (Craig 1957) and are expressed in the standard %0 notation, relative to the PDB standard. Individual cement zones were analysed for Mn and Fe using a Cameca Camebax microprobe. The instrument permits the specimens to be viewed in CL, enabling precise correlation of the results with specific zones. A total of 296 analyses were made from 15 samples representing all six cement zones. Trace elements with concentrations below the detection limit of the microprobe, but likely to be associated with burial cements during mineralization and hydrocarbon emplacement, were analysed using a Cameca IMS 4f ion microprobe. A total of 85 analyses of Ba, Cu, F, Pb, Zn, La and Ce were made from ten samples. Fluid inclusion analysis was carried out on eight vein calcite cements, and a total of 54 measurements were made from two-phase fluid inclusions. To try to reduce any error through stretching and leakage of inclusions, representative samples of each cement were selected, where possible, from vein systems showing no evidence of refracturing or later tectonic activity. Measurements were not attempted in fault-fill calcite. Homogenization temperatures (T~) values were reproducible within individual

39

inclusions to within 3°C, and temperatures of final melting (TM) to within 0.5°C.

Burial cement sequence

Cement petrography Calcite vein cements are ubiquitous in Asbian and Brigantian limestones, and offer the best opportunity for burial cement identification and correlation across the platform. Cements were studied within 0.5-100cm wide veins which were dominantly orientated in a NE-SW and N W - S E direction. Vein systems are largely extensional, although late-stage veins may show some evidence of oblique movement, for example where inclusions of wall rock within the calcite cement are sequentially offset. Crosscutting relationships between stylolites and individual cement zones within veins indicate that pressure dissolution took place throughout burial cementation. Using CL, epifluorescence and the crosscutting relationships between vein systems, six principal burial calcite cement zones can be identified (Table 1). Although Zone 4 cements are clearly distinguished from Zone 3 calcite in syntaxial overgrowth sequences, vein cements suggest that the change from shallow to deep burial cementation was more gradual. Cement nomenclature is therefore based upon subdivisions of the Zone 3 and Zone 4 cements of Walkden & Williams (1991), and characteristics of individual cements are summarized in Table 1. A paragenetic sequence can be established from the relationship of the cements with hydrocarbon and MVT mineralization (Fig. 4). Zone 3A calcite is limpid and contains no hydrocarbon inclusions, clearly distinguishing it from turbid Zone 3B calcite. Zone 4A is similarly turbid, but is separated from Zone 3B by the absence of cement fluorescence and the dull green fluorescence of its hydrocarbon inclusions. Zone 4B is characterized by its concentric subzonation in CL, and is also frequently intergrown with MVT minerals. Zone 4C is invariably intergrown with fluorite, baryte or sphalerite, and is limpid, with only rare hydrocarbon and microdolomite inclusions. Zone 4D cements are not intergrown with either MVT minerals or hydrocarbon. They frequently post-date strong dissolution of both the host limestone and earlier cements, indicating that Zone 4D cements precipitated after the migration of an undersaturated or acidic fluid through the limestone.

40

C. HOLLIS & G. WALKDEN Dolomite Pyrite Silica Zone 3A Baryte Zone 3B

:

Zone 4A Sphalerite Zone 4B

i m m

Fluorite

i

Zone 4C

I

m m I m

j

m

[3

U

m

~

V---qDV7 ~ [---l [Z[S]

Zone 4D m Hydrocarbon i Stylotisation m and fracturing

m

m

m

Increased time and depth of burial Fig. 4. Paragenetic sequence for burial calcite cements on the Derbyshire Platform. Overlap of bars represents diachroneity of cements. Gaps illustrate several phases of refracturing and cementation within an individual cement zone.

Zone 3B to 4A calcite cements (Fig. 5) are found in extensional vein systems, within which calcite precipitation followed dissolution of the vein walls or earlier, refractured, calcite cements. Zones 3A and 3B precipitated as columnar crystals on vein margins, whilst later cements have formed a massive, sparry vein fill. Zone 4B calcite (Fig. 6) is found mainly within extensional vein systems, but some Zone 4B vein cements contain inclusions of wall rock that show oblique offset with respect to each other ......... ~

~

and the vein margins. Zone 4C calcite (Fig. 7) is found almost exclusively in veins that contain evidence of oblique offset along their length. This is also true where Zone 4D calcite has precipitated in larger veins, although Zone 4D precipitated most commonly along sub-micrometre veinlets. Burial calcite cements in veins can be correlated with burial calcite cements in intergranular pores, and cements precipitated along stylolites that were jacked open at elevated hydraulic pressure. Zone 3A calcite precipitation preceded the onset of pressure dissolution, but all other cements may be found along

O

js

!.

Fig. 5. Dull brown luminescent Zone 3B calcite vein (a), with arrowed vein margins, refractured by dull orange luminescent, Zone 4A cemented vein (b). The whole sequence has subsequently been refractured by bright orange luminescent, Zone 4D calcite that postdates dissolution of the earlier cements along vein margins (e). Cathodoluminescence. Dow Low Quarry. Complete scale bar is 1 mm.

Fig. 6. Dull orange-brown luminescent, concentrically subzoned Zone 4B calcite (a), intergrown with hydrocarbon (b). Cathodoluminescence. Jug Holes. Complete scale bar is 1 mm.

BURIAL DIAGENESIS, DERBYSHIRE

41

associated with other minerals within host-rock porosity. Occasional intergrowth of fluorite with Zone 3B calcite marks the onset of MVT mineralization. Fluorite, baryte and sphalerite are increasingly intergrown with Zones 4A, 4B and 4C (Fig. 8), and the distribution of mineralization in cross-cutting veins suggests that mineralization reached a maximum during Zone 4B and Zone 4C cementation (Fig. 4).

Hydrocarbons Fig. 7. Bright orange luminescent Zone 4C calcite (a), intergrown with fluorite (b) along vein system. Vein cross-cuts and offsets non-luminescent Zone 1 and dull brown luminescent Zone 3 pore-fill cements (c). Hope Quarry. Cathodoluminescence. Complete scale bar is 0.5 mm. stylolites. Zones 3A-4A commonly precipitated in shelter porosity beneath brachiopod shells, but this is not seen in later cements. Zone 3A to 4D calcites were recognized within syntaxial overgrowth sequences and as pore-fill cements, and Zone 4D calcite was also found along intercrystalline boundaries of earlier cements. On petrographic evidence, Zones 3A and 4D are equivalent to the Zone 3 and Zone 4 cements, respectively, of Walkden & Williams (1991).

Mineralization Successive calcite cements are increasingly intergrown with galena, fluorite, baryte and sphalerite along vein systems (Fig. 4), although burial calcite cements are less commonly

Fig. 8. Dull orange luminescent Zone 4A calcite (a) with intergrown dolomite (b), cross-cut by bright orange luminescent Zone 4C calcite (c) with intergrown baryte (d). Cathodoluminescence. Middle Peak Quarry. Complete scale bar is 1 mm.

The onset of hydrocarbon migration onto the Derbyshire Platform is marked by white fluorescent, primary hydrocarbon inclusions in Zone 3B calcite in both vein and pore systems (Fig. 9). These are succeeded by dull green fluorescent inclusions in Zone 4A-4C cements. Epifluorescence provides an indication of hydrocarbon maturity, with fluorescence decreasing as maturity increases (Hagemann & Hollerbach 1985), indicating increasingly mature hydrocarbons in successive phases of calcite cementation. Larger accumulations of hydrocarbon are nonfluorescent and common along vein systems and within intergranular and mouldic pores (Fig. 10). They follow the precipitation of Zone 4B calcite, and accumulated in the final stages of hydrocarbon migration. Zone 4D frequently cross-cuts solid hydrocarbon in pore systems, indicating that it post-dates hydrocarbon emplacement.

Geochemistry Geochemical data for individual vein cements are summarized in Table 2.

Fig. 9. Dull yellow brown fluorescent Zone 3A calcite (a) with white fluorescent hydrocarbon inclusions (b). UV fluorescence (blue light). Dow Low Quarry. Complete scale bar is 0.5mm.

-10.3 -12.3 -7.6 -8.9 -14.0 -4.8 -6.1 -5.8 -2.7 -8.5 -9.9 -5.8 -7.5 -9.5 -6.2 -7.5 -7.8 -6.5

2.3 0.7 3.7 1.1 -3.0 4.1 0.4 -2.0 3.7 2.0 0.3 3.3 2.3 0.7 2.9 1.1 -0.7 2.8

613C

57.3 35.6 80.6 106 82.2 125 136 124 153 176 148 200 168 117 207 -

TH 5.8 0.75 13.0 4.4 0.17 18.1 16.4 14.3 17.4 7.7 0.16 22.2 9.4 0.66 23.1 -

TM 163 0.0 610 412 0.0 1320 1962 20 8120 2496 100 7410 3248 20 8220 2768 360 6780

Mn 136 0.0 440 211 0.0 750 1604 0.0 7380 1706 0.0 6890 1694 0.0 7820 1526 0.0 4980

Fe 165 9.91 741 415 16.7 1180 102 18.7 327 108 23.5 302 567 77.6 3170 156 41.3 363

F 70.0 23.7 138 133 8.47 362 157 5.01 45.2 67 5.01 260 338 25.3 1980 104 30.8 260

Zn 12.2 4.99 19.9 18.3 2.01 49.1 56.9 1.39 3.59 16.2 1.40 78.4 85.3 3.11 767 20.9 5.82 45.0

Cu

Ba 0.77 0.24 2.2 1.18 0.14 3.66 0.44 0.01 1.89 0.88 0.09 880 86.1 0.41 940 3.63 0.30 34.7

11.2 0.0 25.1 12.5 1.83 49.6 14.9 5.05 72.8 11.5 1.15 30.4 34.9 0.89 315 10.1 1.64 35.7

Pb

3.31 0.25 8.34 1.83 0.09 12.40 5.29 0.0 31.1 2.86 0.09 14.7 4.93 0.06 16.8 4.54 0.26 17.7

La

2.76 0.25 6.68 1.64 0.14 8.81 4.85 0.0 25.9 3.59 0.09 24.9 5.22 0.08 16.4 5.53 0.28 23.6

Ce

613C and 180 are in the %o notation, relative to the PDB standard. Homogenisation temperatures TH are given in °C and are not pressure corrected. Fluid salinities are calculated from final melting temperatures TM and are presented as wt% equivalent NaC1. All trace element concentrations are given in parts per million.

4D

4C

4B

4A

3B

6t80

3A

Mean Min Max Mean Min Max Mean Min Max Mean Min Max Mean Min Max Mean Min Max

Zone

Table 2. Summary of the geochemistry of burial calcite cements on the Derbyshire Platform

BURIAL DIAGENESIS, D E R B Y S H I R E

43

cement zones on the Derbyshire Platform, most noticeably between Zones 3B and 4A calcite (Fig. 1 la). Similarly, Cu, Zn, La and Ce concentrations all increase between Zones 3B and 4A, whilst F concentrations increase between Zones 3A and 3B calcite (Figs l lb, c,d). All mineralizing elements reach a m a x i m u m concentration within Z o n e 4C calcite cements.

Stable isotopes Fig. 10. Hydrocarbon (a) in coral chambers, following precipitation of dull brown subzoned Zone 3 calcite (b). Remnant porosity infilled by bright orange luminescent Zone 4 calcite (c), partially replacing dolomite (d). Cathodoluminescence. Middle Peak Quarry. Complete scale bar is 1 mm.

Trace elements A n overall increase in average concentrations of measured M n and Fe is seen through successive

The 613C compositions of all vein calcite cements from the Derbyshire Platform range between - 3 . 0 and +4.1%0o, and 6180 ranges from - 1 4 . 0 to -2.7%o. It is difficult to differentiate clearly between separate calcite cement zones isotopically (Fig. 12a). Vein calcite cements from Upper D i n a n t i a n - L o w e r N a m u r i a n limestone horizons within the surrounding shale basins show little variation from platform vein cements (613C = - 5 . 4 to +6.0%o; 6 1 8 0 = - 1 0 to -4.6%o; Fig. 12c). Calcite cements from fault systems on the Derbyshire Platform have lighter 6180 values

40007

800

~3ooof~

.o= 600

o , ~

"~,., 20001 ~ooo g o

! ~

a

,_ . . . . . . . . . . . . . . . . . .

4oo

~ 200 o Manganese . . . . .

Iron

b

•~

~

f

4A

4B

~

Fluorine . . . . .

4C

4D

3B

4A

4B

4C

Lanthanum . . . . . Cerium

]

d[ ~

..,

,

0 . 3A

4D

Zinc

60 40 20

o.oo 3A

3B

,.

6.00 5.00 .£= 4.00 3.00 2.00 8 1.O0 rj

o 3A

.

. 3B

Copper . . . . .

.

. 4A

4B

4C

4D

Barium ......... Lead

Fig. 11. Average trace element concentrations within burial calcite cements on the Derbyshire Platform. (a) Average concentrations of manganese and iron in individual burial calcite cements. Greater than 95% of the Mn data and 90% of the Fe data is within two standard deviations (2s.d.) of the mean. Detection limits: Mn, 250 ppm, Fe, 180 ppm. (b) Average concentrations of fluorine and zinc in individual burial calcite cements. Greater than 90% of the data are within 2s.d. of the mean. Detection limits: F, 0.2 ppm; Zn, 0.2 ppm. (e) Average concentrations of lanthanum and cerium in individual burial calcite cements. Greater than 90% of the data are within 2s.d. of the mean. Detection limits: La, 0.003 ppm; Ce, 0.03 ppm. (d) Average concentrations of copper, barium and lead in individual burial calcite cements. Greater than 90% of the data are within 2s.d. of the mean. Detection limits: Cu, 0.06ppm; Ba, 0.004ppm; Pb, 0.3 ppm.

44

C. HOLLIS & G. WALKDEN

+0.7 to +4.1%o; 6 1 8 0 = - 1 4 to -8.50%0) than platform vein calcite cements (Fig. 12b), or corresponding whole rock values ( 6 1 3 C = - 1 . 2 to +3.6%0; 6~80= -10.2 to -0.5%0). (613C =

-4

[] Zone 3A

-2

O Zone 3B A Zone 4A

r

-16

"12 1

~ O ~'1~.

--27o --4

81SO PDB

X Zone 4D

-6

-16

8 o

.

-2 ,~

[]

b)

-4 -6

r -16

-12

NF

•

SS

H VC

6

-2 ~ •

c)

•

81sO PDB •

8180 PDB

Average homogenization temperatures (TH) generally increase within successive burial calcite cements, but decrease between Zone 4B and Zone 4C calcite, partly as a result of a wide range of homogenization temperatures (Table 2). Fluid salinities, calculated from temperatures of final melting, also increase with time and burial depth. In 3A calcite a maximum salinity of 13.02 wt% NaC1 equiv, is reached, increasing to 23wt% NaC1 equiv, in Zone 4C (Table 2).

O Zone 4B + Zone 4C

O a)

Fluid inclusions

+ DP • EB

[-4

•

SB

[ -6

•

WB

Fig. 12. Stable isotope compositions of burial calcite cements from the Derbyshire Platform and the surrounding basins. (a) Carbon-oxygen isotope values for burial calcite cements. (b) Carbon-oxygen isotope values for fault-fill calcite and vein calcite. VC, vein calcite (Zones 3A-4D); NF, normal fault fillcalcite; SS, strike-slip fault fill calcite. (e) Carbon-oxygen isotope values for calcite cements from the Derbyshire Platform and the surrounding basins. DP, Derbyshire Platform (Zone 3A-4D); EB, Edale Basin; SB, Staffordshire Basin; WB, Widmerpool Basin. The large open rectangle on all plots represents whole-rock Dinantian platform limestone (data from Hollis 1995). Zone 3A4D data in Fig. 12a is repeated in Figs. 12b and 12c.

Discussion and interpretation

Fluid composition Fluid inclusion results indicate that calcite cements were precipitated from fluids that reached salinities of up to 23 wt% NaC1 equiv., with an increase in maximum fluid salinity in successive calcite cement zones. Stable isotope results, which should also provide some indication of fluid composition, do not vary systematically (Fig. 12a). The carbon and oxygen isotopic compositions of vein calcite cements on the platform and within the basins resemble whole-rock Dinantian limestone values, although ~513C values are depleted and 8180 values enriched in 3B and 4A calcite, relative to other vein cements. Fluid inclusion data indicate that calcite cementation took place from formational fluids that became increasingly saline with time and, therefore, with successive burial. Deep burial cements would be expected to show highly negative 6180 values, consistent with temperature-dependent fractionation of oxygen isotopes. Walkden & Williams (1991) noted the reverse, however, with 180-enrichment during Zone 4 cementation following the release of 180-enriched pore waters from the surrounding clastic basins during clay diagenesis. Correspondingly, they recorded a relative 13C-depletion in Zone 4 calcite relative to Zone 3, which they related to the release of ~3C-depleted CO~ from the basins, following kerogen decarboxylation. Light 13C will also be released during shale diagenesis and the transformation of smectite to illite. The similar isotopic composition of both platform and basin vein calcite cements (Fig. 12c) suggests a basinal origin for the fluids. There is a trend towards heavier/5180 values from Zone 3A to 4A cementation, but this trend is reversed during precipitation of Zone 4B to 4D

BURIAL DIAGENESIS, DERBYSHIRE cements. Zone 3B and 4A were precipitated at temperatures where hydrocarbon maturation and clay diagenesis would be widespread within the basins, releasing 180-enriched fluids onto the Derbyshire Platform. It is also Zones 3B and 4A that exhibit the most negative ~513C values. The decrease is minor, however, with the most negative value reaching a 613C value of -2.0%0. Dissolution of wall rock and precursor cements (Fig. 5) under low to moderate fluid-rock ratios during fluid migration altered the original fluid composition, so that vein cements reflect the composition of whole-rock Dinantian limestone. This had a greater effect on ~513Cthan ~5180values, given the lower volume of carbon within the system than oxygen, and hence ~513C compositions were buffered to near-marine values. In addition, stylolitization was an ongoing process throughout the period of burial diagenesis, locally buffering migrating fluids as earlier cements and the host limestone became cannibalized along the pressure contact. Zones 4B to 4D were precipitated under the continuing influence of fluid-rock interaction and, with TH above 150°C, post-dated the main stages of lSO-enriched and 13C-depleted fluid expulsion from the basins. The extent of fluid-rock interaction means cements are now difficult to differentiate on the basis of their isotopic compositions.

Release of trace elements Concentrations of fluorine first increase within Zone 3B calcite. Average concentrations of both Mn and Fe increased during precipitation of Zone 4A calcite, coupled with an increase in the concentration of Cu, Zn, La and Ce (Fig. 11). Basins surrounding the Derbyshire Platform have been invoked as a source of trace elements for mineralization, and therefore it should be possible to relate changes in trace element composition on the platform to events within the basins. Trace elements may be released to basinal pore fluids by the breakdown and transformation of clay minerals and maturation of organic matter. The dehydration and breakdown of hydrous clays, especially the conversion of smectite to illite, is commonly proposed as a mechanism for trace element release in clastic sediments (Boles & Franks 1979). Within black shales, elements may also be incorporated as complexes within the matrix of organic compounds, or absorbed onto organic matter, and can be released to pore fluids with the onset of hydrocarbon maturation and the breakdown of

45

organic compounds (Spears & Amin 1981; Manning 1986). Spears & Amin (1981) suggested that Pb and Cu concentrate within the organic fraction of shales in the Derbyshire basins, whilst Zn is found within illite. Zn may also be incorporated within organic compounds, and will partition, along with Cu, into petroleum during metal transport in carbonate systems (Manning 1986). Ba may substitute for K in alkali feldspars, whilst Fe is released during the transformation of smectite to illite (Boles & Franks 1979; Jones & Plant 1989). Mn can substitute for both Fe and Mg in clay minerals, and may also form oxide coatings on minerals which can later be reduced, releasing Mn 2+ (Francois 1988). Ce is commonly associated with Mn in sediments, whilst La may be concentrated within the smectite lattice (Fleet 1984). Elevated F concentrations (>300ppm) are first detected in Zone 3B calcite (Fig. l lb) coinciding with the first phases of hydrocarbon emplacement and fluorite precipitation on the Derbyshire Platform. This relationship suggests that F was released to fluids from organic compounds during the initial stages of hydrocarbon maturation. This is supported by fluid inclusion data that indicate that Zone 3B cements were precipitated at temperatures between 82 and 125°C, placing the clastic basins well within the oil window at this time. Cu, Zn and Pb are also held within organic compounds, and concentrations of these elements were not elevated (>50, 150 and 13 ppm respectively) within calcite cements until Zone 4A calcite precipitation (Figs l lb, d). This suggests that either F was released during earlier stages of maturation than Cu, Zn and Pb, or that these elements were maintained within petroleum rather than aqueous phases during fluid transport (Manning 1986), preventing the uptake of Cu, Zn and Pb during precipitation of earlier calcite cements. Increased concentrations of both Fe and Mn (Fig. 11 a) between Zone 3B and Zone 4A calcites are accompanied by increases in the concentrations of La and Ce (Fig. 11c). Fluid inclusion and stable isotope data (Table 2) have already shown that clay diagenesis within the basins, especially the transformation of smectite to illite, was extensive within the basins prior to and during Zone 4A cementation. Transformation of smectite may have released both Fe and La to fluids, whilst the breakdown of Mn-bearing minerals and the reduction of Mn-oxides could account for the mobilization of both Mn and Ce. Therefore, the composition of Zone 4A calcite

46

C. HOLLIS & G. WALKDEN

not only demonstrates the continued release of metals from organic matter within the basin, but also incorporates the first elements released during clay diagenesis within the basins. Barium concentrations remained low until Zone 4C cementation (Fig. l ld), indicating that it was mobilized during the latter stages of diagenesis. Feldspar dissolution should have been complete in the shales prior to Zone 4C precipitation, inferring either that Ba was not sourced from the basinal shales, or that transport of Ba was inhibited earlier in the history of the basins. The latter is difficult to prove, but an alternative, stratigraphically higher, source of Ba was the Upper Namurian fluviodeltaic sands, which would have been at an earlier stage of diagenesis during Zone 4C cementation. Fluid migration downwards onto the Derbyshire Platform would be restricted during burial due to the Namurian shale caprock separating the Derbyshire Platform from the sandstones. The onset of Variscan compression, folding and faulting during Zone 4C cementation, however, could have provided suitable fluid migration pathways onto the platform as the overlying sediments were deformed and fractured.

t.ar~nt~erous

D

INI w

I O,,rm~.,

ISl . . . . .

Temperature and timing of burial diagenesis Having established a timing for hydrocarbon maturation, clay diagenesis and trace element release from the basins, it is possible to determine a timing for hydrocarbon migration, mineralization and burial calcite cementation from the basins in relation to tectonic events on the Derbyshire Platform. This is done by use of a burial history curve (Fig. 13), constructed from vitrinite reflectance data (Coleman et al. 1989, modified by Walkden & Williams 1991). Coleman et al. (1989) invoked an increased geothermal gradient during the Carboniferous, consistent with continued extension, Dinantian and Westphalian volcanism, and the 'thermal blanket' effect of the overlying radiogenic Namurian shales. The Derbyshire Platform reached a maximum burial depth of approximately 2.5km in the mid-Stephanian, prior to the onset of the Variscan Orogeny and basin inversion. Hydrocarbons were emplaced on the Derbyshire Platform during precipitation of Zone 3B to Zone 4B calcite, between temperatures of 82 to 200°C, indicating that the basins lay within the oil window during this time (Fig. 13).

Mesogenesis Zone 1 (Eogenesis)

Triassic

Zone 2" ~ ~

1 ln~,as~g.. ~t\ V~/Z/

mineralis.ion

Telogenesis

-II Zone

4^ igl /zo. 4B

Onset Varisean compression 4kin J

'

150o4=

zone 1"-'7--1 4Z~aDe 41)

Fig 13. Summary of the burial history and cementation history of the Derbyshire Platform (modified after Walkden & Williams 1991). Z, Zone; D, Dinantian; N, Namurian; W, Westphalian; S, Stephanian. See text for explanation.

BURIAL DIAGENESIS, DERBYSHIRE Organic maturation and clay diagenesis in the basins was releasing trace elements to pore fluids prior to Zone 4A cementation, and extrapolation of the burial history curve suggests that hydrocarbon generation began in the midNamurian. Hydrocarbon migration ceased prior to precipitation of Zone 4C and Zone 4D cements. Petrographic data indicate that mineralization was most extensive during precipitation of Zone 4B and 4C calcite. F, Zn, Cu, Ba and Pb all reached maximum concentrations during the precipitation of Zone 4C calcite, indicating that precipitation of this cement coincided with the main phase of mineralization on the Derbyshire Platform. Within Zone 4C cements, homogenization temperatures become more scattered, ranging from above 200°C to as low as 117°C. If Zone 4C and 4D calcite precipitation began during maximum burial of the platform and continued during Variscan uplift, a range of fluid temperatures might be expected as the geothermal gradient decreased (Fig. 13) and cooler, less saline fluids were squeezed onto the platform from overlying sediments. By the incorporation of both the trace element and the fluid inclusion data with the burial history curve, it can be concluded that hydrocarbon emplacement and Zone 3A to 4B precipitation and the early stages of mineralization took place during progressive burial of the Derbyshire Platform, whilst Zone 4C-4D and main phase MVT mineralization began during maximum burial and continued into the early stages of uplift.

Fluid flow pathways It has been established that individual phases of cementation cannot be differentiated on the basis of their isotopic composition, and that they do not generally show the light 6180 values that would be expected during burial calcite diagenesis (Fig. 12). Calcite cements sampled from fault systems on the Derbyshire Platform do show low 6~80 values, however, and do not resemble the isotopic composition of the host limestone. If hot, mineralizing fluids from the basins entered the platform along fault systems, both these phenomena can be explained. Fluids could be expelled rapidly into faults with little opportunity to cool and similarly to react with the host limestone. The distribution of burial calcite cements in small vein systems across the Derbyshire Platform indicates that fluids circulated away from fault systems along propagating

47

fracture systems, cooling, reacting with the wall rock, and mixing with fluids expelled during pressure dissolution. The concentration of MVT mineralization along major normal and strikeslip fault systems offers strong support for faultcontrolled fluid flow onto the Derbyshire Platform. Petrography has shown that there is strong tectonic control of the distribution of vein calcite cements on the Derbyshire Platform away from fault systems. Zones 3A-4A are found along extensional veins, and whilst Zone 4B cements are found primarily in extensional veins, they may show oblique movement along their length. Precipitation of Zones 3A-4A cements took place during progressive burial of the Derbyshire Platform, in veins formed by the waning effects of Carboniferous extension. Zones 4C and 4D were emplaced during maximum burial and uplift of the Derbyshire Platform, and oblique offset along these veins is common due to the increasing influence of compressional tectonics.

Mechanisms for mineralization and hydrocarbon emplacement Coleman et al. (1989) argued that fluid expulsion from the basins and mineralization of the Derbyshire Platform was largely compressional, following the onset of the Variscan Orogeny. The main phase of mineralization on the Derbyshire Platform coincided with the precipitation of Zone 4C calcite, and it has been shown that Zone 4C calcite cementation was controlled by Variscan tectonism. This suggests that compression provided the driving force for the expulsion of large volumes of fluids and trace elements from the basins, and a route for fluid migration along reactivated fault systems. The intergrowth of Zone 3B to 4B calcites with fluorite, baryte, galena and sphalerite, as well as hydrocarbons, indicates that the onset of fluid release from the basins occurred much earlier, however. A process for expelling mineralizing brines from the basins to the platform prior to the onset of the Variscan Orogeny must therefore be established. Compaction-driven fluid flow from the basins to the platform margins (Noble 1963) is unlikely, since it would not be possible to maintain fluid release into the deep burial environment. Furthermore, it is unlikely that sufficient energy would be available to squeeze mineralizing fluids up fault systems and across the Derbyshire Platform.

48

C. HOLLIS & G. WALKDEN

Fluid may be expelled along fault systems by the periodic rupture of overpressured sediments (Cathles & Smith 1983). Deposition and early burial of basal Namurian shales may have been sufficiently rapid to allow overpressuring to develop, but the absence of any detailed diagenetic and fracture studies within the basins leaves the idea conjectural. An alternative mechanism for driving fluids through fault systems is by seismic pumping (Sibson et al. 1975). Fluids could be squeezed from the basins onto the platform by continued movement along extensional fault systems, and this would permit mineralization and calcite cementation on the platform with repeated movement along fault systems. The mechanism was probably not sufficient to drive more than a small quantity of the available fluids onto the platform, and this could account for the scattered distribution of mineralization in Zone 3B to 4B vein systems. As the effects of N-S extension waned in the late Westphalian, the increasing influence of compressional tectonics probably became the overriding control on fluid expulsion. This provided the driving force for large-scale expulsion of mineralizing fluids from the basins onto the platform along reactivated fault systems.

Conclusions The data presented above demonstrate the complex relationship between calcite cementation, hydrocarbon emplacement and mineralization on the Derbyshire Platform, in combination with Variscan tectonic events in northern Britain. These events have been related petrographically and geochemically as follows. Hydrocarbon generation within Upper Dinantian and Lower Namurian basinal shales began in the mid-Namurian. The basins and the platform lay within the oil window during precipitation of Zone 3B to early Zone 4C cements, and hydrocarbons became progressively mature with increasing depth of burial. Hydrocarbon migration ceased immediately prior to the onset of the Variscan orogeny. Galena-fluorite-baryte-sphalerite mineralization on the platform began during precipitation of Zones 3B and 4A calcite. F, Cu, Zn and Pb were released to the fluids during maturation of organic matter within the basins. Mn, Fe, La and Ce were expelled during clay diagenetic reactions, including the transformation of smectite to illite. Ba was probably released from feldspars in Upper Carboniferous sandstones with the onset of Variscan compression.

Fluid flow onto the platform was faultcontrolled. Fluids circulated away from the faults on the platform along dilational fracture systems and were progressively cooled and buffered by wallrock interaction at low fluidrock ratios. Calcite cementation and mineralization took place along fault and vein systems, on interaction with sulphate-rich formational fluids in the limestone. Limestone was replaced by ore minerals where fluids concentrated beneath permeability barriers, such as volcanic horizons and stylolites. Zones 3B-4B were precipitated in extensional fracture systems during progressive burial of the platform. Fluids were pulsed onto the platform along major fault and fracture systems by seismic pumping during the waning phases of Carboniferous extension. The onset of Variscan compression led to fault reactivation, basin inversion and dewatering of the basins onto the platform. This led to the major phase of mineralization on the Derbyshire Platform, coinciding with the precipitation of Zone 4C calcite. Zones 4C and 4D calcite were precipitated at maximum burial and during uplift of the Derbyshire Platform. This research was conducted under a NERC research studentship, GT/91/GS/01 (CH). We would like to thank T. Fallick and the staff at the Scottish Universities Research and Reactor Centre for providing facilities for stable isotope analysis. Ion microprobe and electron probe microanalysis took place at Edinburgh University, with valuable advice from J. Craven, P. Hill and S. Kearns. Borehole material was sampled from British Geological Survey boreholes at Keyworth. J. Hendry read and provided helpful comments on earlier versions of this manuscript. This paper has benefited from helpful reviews by J. Marshall and P. Redmond.

References AITKENHEAD, N., CHISHOLM,J. I. & STEVENSON,I. P. 1985. Geology of the Country around Buxton, Leek and Bakewell. Memoir of the British Geological Survey, Sheet 111. BINGHAM, G. C. 1992. The Origin and Interaction of Diagenetic Fluids in the Derbyshire-East Midland Shelf. PhD Thesis, University of Aberdeen. BOLES, J. R. & FRANKS, S. G. 1979. Clay diagenesis in Wilcox sandstones of Southeast Texas: implications of smectite diagenesis on sandstone cementation. Journal of Sedimentary Petrology, 49, 55-70. CATHLES, L. M. & SMITH, A. T. 1983. Thermal constraints on the formation of Mississippi Valley-Type lead zinc deposits and their implications for episodic basin dewatering and deposit genesis. Economic Geology, 78, 983-1002.

BURIAL DIAGENESIS, DERBYSHIRE COLEMAN, T. B., JONES, D. G., PLANT, J. A. & SMITH, K. 1989. Metallogenic Models. In: PLANT, J. A. & JONES, D. G. (eds) Metallogenic Models and Exploration Criteria for Buried Carbonate-hosted Ore Deposits-A Multidiseiplinary Study in Eastern England. British Geological Survey, Keyworth and Institute of Mining and Metallurgy, London, 123-133. CRAIG, H. 1957. Isotopic standards for carbon and oxygen correction factors for mass spectrometric analysis of carbon dioxide. Geochimica et Cosmochimica Acta, 12, 133-149. EWBANK, G., MANNING, D. A. C. & ABBOTT, G. D. 1993. An organic geochemical study of bitumens and their potential source rocks from the South Pennine Orefield, Central England. Organic Geochemistry, 20, 579-598. FLEET, A. J. 1984. Aqueous and sedimentary geochemistry of the rare earth elements. In: HENDERSON, P (ed.) Rare Earth Element Geochemistry. Elsevier, 343-373. FORD, T. D. 1968. Epigenetic mineralisation of the Carboniferous limestone. In: SYLVESTERBRADLEY, P. C. & FORD, T. D. (eds) Geology of the East Midlands. Leicester University Press, 112-137. FRANCOIS, R. 1988. A study on the regulation of the concentrations of some trace elements (Rb, Sr, Zn, Pb, Cu, V, Cr, Ni, Mn and Mo) in Saanich Inlet sediments, British Columbia, Canada. Marine Geology, 83, 285-308. FRASER, A. J., NASH, D. F., STEELE, R. P. & EBDON, C. C. 1990. A regional assessment of the intraCarboniferous play of Northern England. In: BROOKS, J. (ed.) Classic Petroleum Provinces. Geological Society, London, Special Publication, 50, 4 1 7 - 4 4 0 .

GREEN, P. F. 1989. Thermal and tectonic history of the East Midlands shelf (onshore UK) and surrounding regions assessed by apatite fission track analysis. Journal of the Geological Society, London, 146, 755-773. GUION, P. D. & FIELDING, C. R. 1988. Westphalian A and B sedimentation. In: BESLEY, B. M. & KEELING, G. (eds) Sedimentation in a Synorogenic Basin Complex: the Upper Carboniferous of NW Europe. Blackie, Glasgow, 153-177. HAGEMANN, H. W. & HOLEERBACH, A. 1985. The fluorescence of crude oils with respect to their thermal maturation and degradation. Organic Geochemistry, 10, 473-480. HOLEIS, C. E. 1995. Burial Diagenetic Events, Hydrocarbon Emplacement and Mineralisation in Dinantian Limestones of Northern Britain. PhD Thesis, University of Aberdeen. INESON, P. R. & FORD, T. D. 1982. The South Pennine Orefield; its genetic theories and eastwards extension. Mercian Geologist, 8, 285-304. JONES, D. G. & PLANT, J. A. 1989. Geochemistry of shales. In: PLANT, J. A. & JONES, D. G. (eds) Metallogenic Models and Exploration Criteria

49

for Buried Carbonate-hosted Ore DepositsA Multidisciplinary Study in Eastern England. British Geological Survey, Keyworth, and Institute of Mining and Metallurgy, London, 65-94. KEELING, G. & COLLINSON, J. D. 1992. Silesian. In: DUVF, P. MCL. D. & SMITH, A. J. (eds) The Geology of England and Wales. Geological Society, London, 239-263. LEEDER, M. R. 1988. Recent developments in Carboniferous geology: a critical review with implications for the British Isles and NW Europe. Proceedings of the Geological Association, 99, 73-100. MANNING, D. A. C. 1986. Assessment of the role of organic matter in ore transport processes in low-temperature base metal systems. Transactions of the Institution of Mining and Metallurgy, 9 5 B , 195-200. NOBLE, E. A. 1963. Formation of ore deposits by water of compaction. Economic Geology, 55, 1145-1156.

PERING, K. L. 1973. Bitumens associated with lead, zinc and fluorite ore minerals in North Derbyshire, England. Geochimica et Cosmochimica Acta, 37, 401-417. QUIRK, D. 1987. Structure and Genesis of the South Pennine Orefield. PhD Thesis, University of Leicester. SIBSON, R. H., MOORE, J. MCM. & RANKIN, A. H. 1975. Seismic p u m p i n g - a hydrothermal fluid transport mechanism. Journal of the Geological Society, London, 131, 653-659. SMITH, K. & SMITH, N. J. P. 1989. Deep Geology. In: PLANT, J. A. & JONES, D. G. (eds) Metallogenic Models and Exploration Criteria for Buried Carbonate-hosted Ore Deposits-A Multidisciplinary Study in Eastern England. British Geological Survey, Keyworth, and Institute of Mining and Metallurgy, London, 53-64. SPEARS, D. A. & AMIN, M. A. 1981. Geochemistry and mineralogy of marine and non-marine Namurian black shales from the Tansley borehole, Derbyshire. Sedimentology, 28, 407-417. WAEKDEN, G. M. 1974. Palaeokarstic surfaces in Upper Vis6an (Carboniferous) limestones of the Derbyshire Block, England. Journal of Sedimentary Petrology, 44, 1232-1247. 1987. Sedimentary and diagenetic styles in Late Dinantian carbonates of Britain. In: MILLER, J., ADAMS, A. E. & WRIGHT, V. P. (eds) European Dinantian Environments. John Wiley, Chichester, 131-155. - & BERRY, J. R. 1984. Syntaxial overgrowths in muddy crinoidal limestones: cathodoluminescence sheds new light on an old problem. Sedimentology, 31, 251-267. -& WILLIAMS, D. O. 1991. The diagenesis of the late Dinantian Derbyshire-East Midlands carbonate shelf, central England. Sedimentology, 38, 643-670.

Metal-organic interactions in the Dinantian Solway Basin, UK: inferences for oil migration studies C. V E A L E

& J. P A R N E L L

School o f Geosciences, The Queen's University, Belfast B T 7 1 N N , U K Abstract: Carboniferous basins within the British Isles show a variety of metal-organic fluid interactions, including the precipitation of thoriferous bitumen nodules. These represent the interaction between a thorium-rich fluid, such as groundwater, and migrating hydrocarbons. They are found in the Dinantian of the Northwest Irish Basin and the Solway Basin. In the latter case, nodules are found in coarse-grained clastics laid down in a variety of depositional environments. The source of the thorium was the nearby Caledonian CriffelDalbeattie Granite. Studies of sedimentological parameters of both nodule-bearing and barren beds, and the structure of the basin, show that factors controlling the distribution of nodules include the maturity and porosity of the host rock, with an enhanced maturity and high porosity preferable, as well as the depositional environment. In the case of the Dinantian of the Solway Basin, distal shallow-marine sediments and channelized deposits host the nodules. Chemical age dating of the thoriferous inclusions within the nodules points to the possible use of these nodules in assessing the timing as well as spatial distribution of hydrocarbon migration within sedimentary basins. Results indicate a correlation with previous models of Mesozoic hydrocarbon migration.

Many Carboniferous basins in the British Isles exhibit evidence for interactions between migrating hydrocarbons and metalliferous groundwaters. The products of these interactions include, for example, metal-rich organic cores in reduction spheroids (Harrison 1970), bitumen residues in sulphide deposits (Pering 1973; Monson & Parnell 1992), hydrocarbon fluid inclusions in vein gangue minerals (Hollis & Walkden this volume; Moser et al. 1992), and light gas anomalies in ore deposits (Ferguson 1988). In several regions of northern Britain, the interaction of hydrocarbons with groundwaters draining off Caledonian plutons caused the formation of uranium-rich and thorium-rich oil residues (Miller & Taylor 1966; Parnell 1988, 1995; Eakin 1989; Parnell & Monson 1990). Many of the Caledonian plutons are anomalously rich in radioelements (Gallagher et al. 1971; Simpson et al. 1979; Stephens & Halliday 1984) and, as they had been widely unroofed in upland regions during Devonian/Lower Carboniferous times, they were a ready source of uranium, thorium and a range of other metals. Typical radioelement-bearing minerals to be found in granitic rocks include monazite, zircon, xenotime, thorianite and thorite. Thoriferous bitumen nodules occur in sandstones of several Carboniferous basins and are believed to represent a product of interaction between hydrocarbons and metalliferous groundwaters (Parnell & Monson 1990; Parnell et al. 1990). Dinantian basins that exhibit large

numbers of thoriferous bitumen nodules include the Northwest Irish Basin (Parnell & Monson 1990) and the Solway Basin (Fig. 1). This paper describes the occurrence and distribution of

Nol

Fig. 1. Map showing the main Carboniferous basins of the British Isles. Dinantian nodule-bearing basins are labelled; the Northwest Irish Basin and the Solway Basin.

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 51-63.

52

C. VEALE & J. PARNELL

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Geological setting The Solway Basin forms the western extension of the Carboniferous Northumberland Trough and occupies the Solway Firth, with onshore exposures on both the northern and southern margins. The basin is bounded by reactivated Caledonian structures, trending WSW-ENE, with the oldest sedimentary rocks being of Upper Devonian age. Dinantian sequences are exposed in four outliers along the northern margin of the Solway Firth (Fig. 2). From the west these are the Rerrick Outlier, the Castle Point Outlier, the Portling Outlier and the Kirkbean Outlier.

The deposits within these outliers are of variable lithology, mainly conglomerates, conglomeratic sandstones, sandstones and limestones deposited in alluvial fan, fluvial and shallow-marine settings. The westernmost',, three outliers are all in contact with Silurian metasediments of the Southern Uplands Block. This is also true of the Kirkbean Outlier for the m6st part, although in places the Dinantian rocks of this outlier also abut both Upper Old Red Sandstone lavas and the late Caledonian Criffel-Dalbeattie Granite, dated at 397 4-2 Ma (Halliday et al. 1980). The stratigraphy of the Dinantian outliers is summarized in Fig. 3.

Rerrick Outlier

The Wall Hill Sandstone Group comprises three formations: White Port, Sheep Bught and Abbey Head Formations. These are coarse-grained, pale, fluviatile conglomerates and sandstones lying unconformably on the Silurian, and separated from the overlying Orroland Group by a zone of faulting and barytes veining.

METAL-ORGANIC INTERACTIONS: DINANTIAN SOLWAY BASIN

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53

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Fig. 3. Schematic representation of the Dinantian sediments in the Solway Basin. Starred formations are nodule-bearing. Brackets indicate unsampled formations. LG, Lochenling Formation; RB, Rascarrel Burnfoot Formation; CM, Castle Muir Formation; BN, Black Neuk Formation; SH, Scar Heugh Formation; OL, Orroland Lodge Formation; BH, Barlocco Heugh Formation; DC, Dropping Craig Formation; SD, Spouty Dennans Formation; HM, Hanged Man Formation; AH, Abbey Head Formation; SB, Sheep Bught Formation; WP, White Port Formation; AB, Arbigland Beds; TS, Thirlstane Sandstone; PW, Powillimount Beds; GL, Gilfoot Beds; SN, Southerness Beds; BC, Basal Cementstone. Deegan (1973) tentatively correlated the Castle Point and Portling Outliers with the Lochenling Formation. Craig & Nairn (1956) correlated a portion of the Orroland Group with the Southerness Beds. The remainder of the Rerrick Outlier, representing two lithostratigraphic groups, comprises a series of conglomeratic red beds formed in an alluvial fan with occasional shallow-marine influences (Deegan 1973). Deegan (1973) also recorded evidence of three cycles reflecting depositional environment. Each cycle comprises basal alluvial fan conglomerates followed, due to a decrease in debris supply, by lower fan deposits and occasional shallowmarine limestones. A resurgence in fluvial activity heralds the onset of the next cycle, starting again with basal conglomerates. The cycles start at the base of the Hanged Man Formation of the Orroland Group, the Black Neuk Formation also of the Orroland Group, and the Lochenling Formation of the Rascarrel Group. While the second cycle is believed to have been the result of gentle subsidence, characterized by the gradational boundary of the Black

Neuk Formation, the onset of the third cycle at the base of the Lochenling Formation was a sudden event resulting in a coarse angular breccia (Deegan 1973). While the Wall Hill Sandstone Group lies unconformably upon the Silurian rocks of the Southern Uplands Block, the Orroland and Rascarrel Groups are faulted against these metasediments. The bounding fault here is the North Solway Fault Zone. This has been dated as active during the Jurassic by U-Pb isotopic methods (Miller & Taylor 1966) and by the presence of uraniferous hydrocarbons (Parnell 1995), dated chemically by an iterative method developed by Bowles (1990) which relies on the decay of uranium and thorium to lead in uranium-bearing minerals (see below). However, this Jurassic movement probably indicates a phase of reactivation of the basin-bounding fault, which in turn was probably controlled by

54

C. VEALE & J. PARNELL

earlier Caledonian-trending structures that gave rise to the NE-SW oriented shape of the basin (Chadwick & Holliday 1991).

Castle Point and Portling Outliers These two isolated outliers have been tentatively correlated by Deegan (1973) as being equivalent to the Lochenling Formation of the Rascarrel Group (Fig. 3) and field observations of the lithologies (conglomerates, conglomeratic sandstones, coarse-grained sandstones and occasional limestones) bear this out. The strata in these two sequences thus equate with the top of the Rerrick succession (Fig. 3) and they are also faulted against the Silurian rocks of the Southern Uplands Block.

Kirkbean Outlier The easternmost area of Dinantian exposure at Kirkbean (Fig. 3) is very different in nature. Craig (1956) described the succession as dominated by shallow marine deposits with intermittent terrestrial input, deposited onto a tectonically stable basin margin. This lack of evidence for tectonic control on sedimentation is in direct contrast with the other outliers discussed above. At the base of the succession is the Basal Cementstone, comprised of carbonaceous siltstones, sandstones and fine-grained limestones. These outcrop inland and were not sampled in this study of coastal exposures. Overlying this are the Southerness Beds comprised of mostly thin limestones separated by four thin, flaggy sandstone bands. These pass up into the Gilfoot Beds, dominated by reddish sandstones and conglomerates with minor shales and marine bands above a basal calcareous sandstone. The following Powillimount Beds are mainly grey sandstones and shales with fine-grained cementstones (massive, clayey limestones) and algal oolitic limestones. Above this the Thirlstane Sandstone is a strongly cross-bedded, quartzrich medium-grained sandstone with prominent overturning and contortion of beds. The Arbigland Beds are separated from the Thirlstane Sandstone by a reverse fault and comprise a number of sequences of shales, limestones and sandstones.

Methodology A sandstone sample of 500 to 750 g was taken from 19 of the formations cropping out along

the northern shore of the Solway Firth (Fig. 3). The exceptions, due to problems of access, were the White Port and Abbey Head Formations of the Wall Hill Sandstone Group, the Castle Muir Formation of the Rascarrel Group, and the Basal Cementstone of the Kirkbean succession. Chips of the samples were screened microscopically for the presence of thoriferous bitumen nodules. Rock-chips containing nodules were examined using a JEOL 733 electron microprobe. For this, samples were mounted on aluminium stubs, carbon-coated and studied in backscattered mode using an accelerating voltage of 25 kV and a beam current of 30-70 nA. A number of nodules were analysed in the electron microprobe, using energy-dispersive methods, to determine the uranium, thorium and lead content of some of the inclusions within them. Bowles (1990) devised a method, based on uraninite, enabling individual grains to be dated from the relative proportions of these three elements. The validity of the method is based upon two observations concerning uraninite chemistry. Firstly, in modern uraninite virtually all analysed lead is radiogenic, avoiding overestimation of the age. Secondly, uraninite retains radiogenic lead within its crystal lattice, thus minimizing leaching and the resultant underestimation of the age. From the known decay rates of U and Th, formulae can be constructed to determine the age of the grains. By applying this method to thorite, it is possible to build up a map of the chemical ages of individual inclusions within a single nodule.

Fig. 4. Thoriferous bitumen nodule from the Solway Basin showing a crenulate margin (arrowed) indicating its replacive nature. Powillimount Beds, Kirkbean Outlier. Scale bar is 10#m.

METAL-ORGANIC INTERACTIONS: DINANTIAN SOLWAY BASIN

Petrography Thoriferous bitumen nodules from the Dinantian of the Solway Basin are sub-millimetre in size, with an average diameter of approximately 200#m. A crenulate margin observed in some nodules (Fig. 4) indicates their replacive nature.

¢:

55

The inclusions within the bituminous matrix are predominantly of thorium silicate (thorite) with occasional substitution by uranium, phosphorus and calcium, in line with other mineralogical investigations of such inclusions (Parnell & Monson 1990). The inclusions exhibit a number of styles. The most common is a random distribution of

! IF

Fig. 5. Backscattered electron micrographs of thorite inclusions. (a) Typical pattern of discrete, micron-scale inclusions (white). Portling Bay, Portling Outlier. Scale bar is 100#m. (b) Nodule containing domains of inclusion-rich material separated by areas of inclusionfree bitumen. Portling Bay, Portling Outlier. Field width is 540 #m. (c) Thoriferous bitumen nodule formed around a detrital thorite grain (T). Portling Bay, Portling Outlier. Scale bar is 10 #m. (d) Radially arranged groups of inclusions (arrowed). Powillimount Beds, Kirkbean Outlier. Scale bar is 100 #m. (e) Inclusions of thorite that fill almost the entire volume of the nodule. Portling Bay, logged section Bed Z, Portling Outlier. Scale bar is 10#m.

56

C. VEALE & J. PARNELL

discrete, micron-scale inclusions (Fig. 5a). Other patterns include the segregation of inclusions into domains separated by areas of inclusionfree bitumen (Fig. 5b), large fragmented grains (Fig. 5c), and groups of inclusions arranged in circular or oval-shaped, radially arranged patterns (Fig. 5d). Apart from the final style, so far only recorded in the Solway Basin, these inclusion patterns are consistent with other studies on thoriferous bitumen nodules (Parnell & Monson 1990). The inclusion density is highly variable, ranging from sparse (Fig. 5a) to cases where the inclusions almost fill the entire volume of the nodule, leaving only a thin rim of bitumen around them (Fig. 5e).

Nodule distribution In previous work on thoriferous bitumen nodules (Parnell & Monson 1990), it has been noted that the distribution of nodules is unpredictable and only rarely appears to follow trends such as bedding. The Dinantian of the Solway Basin provides a good opportunity to examine nodule distribution within an established sedimentological and structural framework to explore this feature. Possible factors that explain the nature and selective distribution of the nodules are gross lithology of the nodule-bearing beds, mineralogy and maturity, as well as the structural setting of the Solway Basin. The distribution of nodule-bearing formations is indicated in Fig. 3. In the Rascarrel Group nodules were restricted to the Lochenling Formation. Both the small outliers at Castle Point and Portling were nodule-bearing, and further discoveries were made in the Powillimount Beds and Southerness Beds of the Kirkbean Outlier.

Gross lithology The foreshore at Portling Bay, within the Portling Outlier, exposed a total of approximately 15m of sediments. This succession yielded the largest number of nodules (see the detailed lithological log in Fig. 6). The section was sampled regularly, with a total of 28 siltstones, sandstones, conglomeratic sandstones and conglomerates being examined for nodules. Nine nodule-bearing beds were discovered: six sandstones, two conglomerates and a single conglomeratic sandstone (Fig. 6). This distribution shows a reasonably significant preference

for sandstone over other lithologies, but it should be noted that there were eight further barren sandstone bodies in the succession. A further point of interest is the apparent bunching of nodule-bearing beds that appears to take place with two main groups (samples J, K, M & N and W, X, Z & AA), as well as the somewhat isolated occurrence in bed C. At other localities only a representative sandstone-grade sample was taken, so little discriminatory information was available.

Mineralogy Point-counting data were compiled from a thin section from each of the formations sampled. These were plotted on a QFL ternary diagram (Fig. 7a). All of the nodule-bearing beds fall in the top third of the diagram (minimum 65% quartz), and half of them contain over 80% quartz, indicating a good correlation with mineralogically more mature lithologies. It should be noted that all the Kirkbean lithologies were more mature than the arkosic sediments from the Rerrick Outlier, as well as being much finer-grained and representing much more distal deposits.

Maturity A basic parameter of maturity, the feldspar/ quartz ratio, was then determined for each sample and plotted in Fig. 7b. Two distinct groups of nodule-bearing beds can be distinguished. The more quartzose (more mature) group comprises the two nodule-bearing formations (the Powillimount and Southerness Beds) from Kirkbean as well as a sample from Portling which has a very high carbonate content, thus distorting the feldspar/quartz ratio. The other less mature group includes the nodule-bearing horizons from the Lochenling Formation, and the Castle Point and Portling Outliers.

Discussion Nodule formation There are a number of records of solid bitumen envelopes surrounding thorium-bearing mineral phases (e.g. Rasmussen & Glover 1990; Parnell & Monson 1990). Thorium emits alpha-radiation and the effect of this upon hydrocarbons is to cause the cross-linking of straight-chain molecules to form more complex solid polymers (Charlesby 1954; Colombo et al. 1964).

METAL-ORGANIC INTERACTIONS: DINANTIAN SOLWAY BASIN

57

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58

C. VEALE & J. PARNELL

Q TS~ • Nodule-bearingBeds SN'Gs~DOpp~'c'~XOBarren Beds Th-enriched fluid

'":"-"" !:

t- -

F

L I

6

IM,nrh,la.l~,.'~rin~, l~.~,rlS

5

I

'-

:

(. :"

:. i:: ::.

Fig. 8. Schematic diagram to illustrate the method of formation of the majority of thoriferous bitumen nodules in the Solway Basin, by groundwater interaction with hydrocarbons (HC) (after Parnell & Monson 1990).

i3 u_ 2 1

0 0

0.1

0.2

0.3

0.4

0.5

0.6

Feldspar/Quartz Fig. 7. Mineralogy and maturity studies conducted in the formations sampled from the Solway Basin. (a) Quartz-feldspar-lithic fragment compositional plot. Abbreviations as in Fig. 3. (b) Feldspar:Quartz ratio. As a parameter of maturity, this discriminates nodule occurrence into two main groups: a more mature group comprising the samples from the Kirkbean Outlier with a very low F : Q ratio, and a less mature group from the Lochenling Formation and the Castle Point and Portling Outliers.

In the context of thoriferous bitumen nodules, this process occurs when migrating hydrocarbons meet a thorium source. This can be either discrete grains of detrital minerals or some thorium-bearing fluid (Fig. 8). The original source for this thorium in the Solway region is likely to be the Caledonian Criffel-Dalbeattie Granite, immediately to the north of the study area. Similar bodies have often been found to be anomalously rich in radioelements (Stephens & Halliday 1984; O'Connor 1986). In cases where the thorium-bearing phase is authigenic, the exact form that the thoriferous fluid takes is unclear since thorium is stable as an insoluble oxide, ThO2, over a large range of Eh and pH

conditions (Parnell & Monson 1990). It only becomes soluble as a sulphate at conditions effectively unknown in sedimentary systems. It is conceivable that the thorium is transported in the presence of organic matter as an organic anion complex (Dementyev & Syromyatnikov 1965). Parnell & Monson (1990) recorded thoriumbearing inclusions within fracture-bound bitumen which proves the ability of thorium to precipitate as an authigenic mineral within bitumen, although again, the exact chemical mechanism of the precipitation process is as yet unknown. Since the area over which alpha particles can travel is limited, only a small volume of hydrocarbon is affected and the remainder is free to continue migrating, often leaving the nodules as the sole evidence for hydrocarbon migration through the succession. During the polymerization process resulting from the irradiation, there is a considerable volume increase as the hydrocarbon swells (Bauman & Glantz 1957). The presence of unaffected hydrocarbon can aid the swelling process, as the hydrocarbon reaches and passes the gel point, by effectively 'feeding' the reaction (Born 1963). This expansion can place great stress on the surrounding grains, which may then fail, leading to a pattern of radiating microfractures in them (Fig. 9a). Conversely, during the final stages of solidification there is a decrease

METAL-ORGANIC INTERACTIONS: DINANTIAN SOLWAY BASIN

59

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

i:!!i:,;iii:iii

•~, ~

~

~. ,,;. "~

,

,

Fig. 9. Backscattered electron micrographs of thoriferous bitumen nodules. (a) Radial microfractures (arrowed) developed around a nodule. These are caused by stress generated during polymerization-induced expansion. Castle Point Outlier. Scale bar is 100 #m. (b) Internal cracks (arrowed) within a nodule caused by contraction on solidification of the hydrocarbon. Portling Bay logged section Bed M, Portling Outlier. Scale bar is 100#m. (e) Two-phase bitumen nodule. The margin of the original nodule is arrowed. Powillimount Beds, Kirkbean Outlier. Scale bar is 100 #m. (d) Two-phase bitumen nodule showing an image-intensity difference between bitumen phases, possibly as a result of differences in maturity. The original nodule (1) is paler than the second hydrocarbon phase (2). Portling Bay, Portling Outlier. Scale bar is 10 ~m. in volume which may lead to the internal cracks seen within some nodules (Fig. 9b). After solidification the entire nodule remains weakly radioactive and, if a second phase of migrating hydrocarbon occurs, a two-phase bitumen nodule can develop (Fig. 9c). In twophase nodules, the core is formed from an original nodule, full of thoriferous inclusions, which is surrounded by a rim of inclusion-free bitumen. The rim represents the second phase of migrating hydrocarbon that enveloped the nodule, was irradiated by it and solidified around it. In some two-phase nodules a difference in image intensity between the bitumen phases can be seen in photomicrographs taken in backscattered mode (Fig. 9d). This may be due to the radiation from the thorite

apparently increasing the maturity of the hydrocarbon in the central core area (Sassen 1984)• The majority of nodules discovered in the Dinantian of the Solway Basin have discrete, micron-scale inclusions (e.g. Figs 4, 5a, 5b). This points to the interaction of hydrocarbons with a thorium-bearing fluid as the dominant mechanism of formation of thoriferous bitumen nodules in this region• It is only a minority of nodules that contain larger, fragmented remains of detrital mineral grains. The fragmentation process is also believed to be the result of the expansion of the hydrocarbons during polymerization• Rasmussen & Glover (1990) proposed that liquid hydrocarbons could enter structural weaknesses such as cleavage planes and subsequently force these apart.

60

C. VEALE & J. PARNELL

4

2 ii 1

0 100

II 200

300

400

500

S00

700

800

900

1000

Chemical Age (Ma)

Fig. 10. Histogram of chemical ages calculated from electron microprobe analyses of nodules from the Portling Outlier, Solway Basin. Dating Aggregate chemical age data for a number of nodules from Portling Bay in the Portling Outlier, calculated by the method of Bowles (1990) outlined earlier, are summarized in Fig. 10. There is a bimodal distribution of ages, which reflects the two methods of formation of the inclusions. The older peak at around 400-500 Ma represents the detrital thorite grains, which possess the signature age of the granite or possibly its precursor. The younger peak, between 200 and 300 Ma, represents the authigenically-formed inclusions, when thorium was fixed from its host fluid as thorite. Since this precipitation is caused by the interaction with hydrocarbons, this age effectively dates the timing of the hydrocarbon migration. Thus, thoriferous bitumen nodules can provide temporal as well as spatial information on the migration of hydrocarbons through sedimentary basins. The use of Bowles' technique on thoriferous minerals is at an early stage, and more complex arithmetic is involved compared with the calculation of ages of uraninite. This is likely to be responsible for the spurious values such as those in the 600-700Ma and 900-1000Ma ranges. Further refinement of the technique may be expected. These chemical ages of 200-300 Ma obtained from the Dinantian of the Solway Basin correlate well the model of Barrett (1988), which envisaged hydrocarbon generation in Dinantian strata of the Solway Basin commencing in the early Permian (Fig. 11). Nodule distribution Within the study area, nodules have developed in both the proximal alluvial fan sediments of the Rerrick, Castle Point and Portling Outliers

and the more distal predominantly shallowmarine sandstones of the Kirkbean succession. The nodules within the alluvial fan sediments are restricted to the easternmost deposits of these outliers, the Lochenling Formation of the Rascarrel Group, and the two smaller outliers at Castle Point and Portling, with Portling having the highest concentration of nodules of any locality. Conversely, within the Kirkbean succession, the two nodule-bearing formations are separated by a barren formation and the succession is topped by nodule-free formations. The discrimination into two groups of nodulebearing formations on the basis of feldspar/ quartz ratio (Fig. 7b) suggests that the difference in relation between the sediments and the basin margin between the Rerrick-Castle Point-Portling Outlier and the Kirkbean Outliers may be of some significance with regard to nodule formation and distribution. This in turn reflects the difference in source areas and the distance from those source areas. The structure of the Solway Basin has been well described (Barrett 1988; Chadwick & Holliday 1991; Chadwick et al. 1993a, b). Barrett (1988) proposed a model based upon a series of half-grabens of alternating polarity, switching about a set of transfer zones trending NNW-SSE across the basin, with one such transfer zone separating the Rerrick and Kirkbean successions. This would allow the development of a faulted margin and associated alluvial fan deposits on the northern margin in the Rerrick area, and the predominantly shallow marine Kirkbean succession immediately to the east could have been laid down on the stable, northern margin of the adjacent half graben. However, Chadwick et al. (1993b) suggested that the Maryport-Stublick-Ninety Fathom Fault system, on the southern margin of the Solway Basin, was the main basin-defining structure, and that the northern margin was the product of a series of en echelon synsedimentary structures, including the North Solway Fault. Deegan (1973) stated that the deposits of the Rerrick Outlier are proximal and are sourced from the north, and the Southern Uplands Block and the Criffel-Dalbeattie Granite in particular. However, the dominantly shallow-marine sediments of the Kirkbean succession, which are much more distal in nature, are also considerably finer-grained for the most part than the sandstones of the Rerrick, Castle Point and Portling Outliers. This has implications for the migration of oil through these finer-grained deposits in terms of pore size and pore throat

METAL-ORGANIC INTERACTIONS: DINANTIAN SOLWAY BASIN Dev I E. Cart) I L. Carlo l, Permian I 360 310 260

Trias

I 210

Jurassic 160

I

:.._._:'_:.

---.__.. _

~

"--__._.____

_

61

Cretaceous 110

Recent 60

_ ......::-;-

......

0 rnybp

50:°

Permo-Trias basin margin ~

Base Gourceyan-~clian

""" Perr~ ~, ,,~.~de[~e.enlre/

......

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

~-''"

150"C

Non-generating I 7 km

Onset Piakl , ~

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

Oil generation window in Permo-Trias depocentre ~ Oil generation window in P e ~ T r i a s basin margin Periods of tectonic inversion of unknown intensi~y

Fig. ll. Timing of hydrocarbon migration in the Solway Basin (after Barrett 1988). The Dinantian sequence studied was generating hydrocarbons in the early Permian, agreeing with the chemical ages in Fig. 10.

diameter, but migration has clearly taken place through at least two of the formations at Kirkbean. This has perhaps been aided by the enhanced maturity and porosity displayed by these more distal deposits. In contrast, the proximal arkosic deposits of Rerrick, Castle Point and Portling would be poor hydrocarbon migration pathways due to their negligible porosity. Indeed, apart from the eastern end of the Rerrick succession and the two smaller outliers, there is no evidence that any such migration has taken place through these sequences. When Deegan (1973) inspected the Lochenling Formation, he discovered that one feature produced by the sudden onset of the third cycle of alluvial fan conglomerates was the development of deep channelization. Channelized

deposits are regarded as potentially good sites for nodule development (Parnell & Monson 1990) due to their open fabric, high porosity and potential for fluid flow. Further effects of subsequent fluid flow, such as grain dissolution, could enhance the porosity and permeability of these particular beds. This channelization could also help to explain the grouped distribution of nodule-bearing beds in the logged section at Portling Bay, tentatively correlated with the Lochenling Formation by Deegan (1973). Conclusions Thoriferous bitumen nodules were discovered in the Dinantian sediments of the Solway Basin.

62

C. VEALE & J. PARNELL

As in previous studies of thoriferous nodulebearing basins, the distribution of nodules was found to be very selective throughout the succession, with only a small number of beds containing nodules, and a variable number of nodules occurring in each nodule-bearing bed. Several factors play a part in determining whether or not a specific horizon can host nodules, even when all the obvious necessary components of formation such as a thorium source, high porosity clastic host-rock and a history of hydrocarbon migration are present. In the case of the Solway Basin, there are two separate occurrences of nodules. Firstly, in the easternmost outlier at Kirkbean the development of nodules correlates with the increased maturity and porosity of the more distal shallow-marine sediments. Secondly, further west in the easternmost part of the Rerrick succession, the primary factor appears to be the deep channelization of a limited number of beds, allowing enhanced fluid flow. In terms of the sedimentological parameters considered above as possible controls on nodule development, no single indicator can be used to discriminate between nodule-bearing and barren beds, although a marked preference for quartz arenites with enhanced porosity and maturity has been noted. This is to be expected since the process is dependent upon the migration of hydrocarbons. The development of thoriferous bitumen nodules is evidence that hydrocarbon migration has occurred in the Solway Basin, and their areal distribution indicates the spatial extent of that migration. However, the absence of nodules may also reflect a lack of interaction between the constituent components and not necessarily a lack of hydrocarbon migration. The radiometric dating of the thorite involved in the irradiation of the hydrocarbons provides constraints on the timing of hydrocarbon migration. This will prove useful in assessing whether any hydrocarbons remain trapped in suitable structures in a basin. The authigenic inclusions within nodules from Portling Bay in the Portling Outlier gave a most frequent age of around 200-300Ma, which is consistent with Barrett (1988), who predicted oil generation from basal, oil-prone algal source rocks in the Permian, and Parnell (1995), who derived Mesozoic chemical ages for thoriferous inclusions within bitumen nodules. A postgraduate award from the European Social Fund to C.V. is gratefully acknowledged, as is the technical assistance of the staff of the Queen's University Electron Microscope Unit and G. Alexander.

References BARRETT, P. A. 1988. Early Carboniferous of the Solway Basin: A tectonostratigraphic model and its bearing on hydrocarbon potential. Marine and Petroleum Geology, 5, 271-281. BAUMAN, R. & GLANTZ, J. 1957. The effect of copolymer composition on radiation cross-linking. Journal of Polymer Science, 26, 397-399. BORN, J. W. 1963. Elastomeric materials. In: CARROL, J. G. & BOLT, R. O. (eds) Radiation Effects in Organic Matter. Academic Press, New York, 245-288. BOWLES, J. F. W. 1990. Age dating of individual grains of uraninite in rocks from electron microprobe analyses. Chemical Geology, 83, 47-53. CHADWICK, R. A. & HOLLIDAY, D. W. 1991. Deep crustal structure and Carboniferous basin development within the Iapetus convergence zone, northern England. Journal of the Geological Society, London, 148, 41-53. - - , EVANS, D. J. & HOLLIDAY, D. W. 1993a. The Maryport fault: the post-Caledonian tectonic history of southern Britain in microcosm. Journal of the Geological Society, London, 150, 247-250. --, HOLLIDAY,D. W,, HOLLIDAY,S. & HULBERT, A. G. 1993b. The evolution and hydrocarbon potential of the Northumberland-Solway Basin. In: PARKER, J. R. (ed.) Petroleum Geology of NW Europe. Proceedings of the 4th Conference. Geological Society, London, 717-726. CHARLESBY, A. 1954. The cross-linking and degradation of paraffin chains by high-energy radiation. Proceedings of the Royal Society of London, A222, 60 74. COLOMBO, U., DENTI, E. & SIRONI, G. 1964. A geochemical investigation upon the effects of ionizing radiation on hydrocarbons. Journal of the Institute of Petroleum, 50, 228-237. CRAIG, G. Y. 1956. The Lower Carboniferous Outlier of Kirkbean, Kirkcudbrightshire. Transactions of the Geological Society of Glasgow, 22, 113 132. & NAIRN, A. E. M. 1956. The Lower Carboniferous Outliers of the Colvend and Rerrick Shores, Kirkcudbrightshire. Geological Magazine, 93, 249-256. DEEGAN, C. E. 1973. Tectonic control of sedimentation at the margin of a Carboniferous depositional basin in Kirkcudbrightshire. Scottish Journal of Geology, 9, 1-28. DEMENTYEV, V. S. & SYROMYATNIKOV,N. G. 1965. Mode of occurrence of thorium isotopes in ground waters. Geochemistry International, 2, 141 147. EAKIN, P. A. 1989. Isotopic and petrographic studies of uraniferous hydrocarbons from around the Irish Sea Basin. Journal of the Geological Society, London, 146, 663-673. FERGUSON, J. 1988. The nature and origin of light hydrocarbon gases associated with mineralization in the Northern Pennines. Marine and Petroleum Geology, 5, 378 384.

M E T A L - O R G A N I C I N T E R A C T I O N S : D I N A N T I A N S O L W A Y BASIN GALLAGHER, M. J., MICHIE, U. McL., SMITH, R. T. 8~; HAYNES, L. 1971. New evidence of uranium mineralization in Scotland. Transactions of the

Institute of Mining and Mineralogy. Section B. Applied Earth Science, 80, 150-173. HALLIDAY, A. N., STEPHENS, W. E. & HARMAN, R. S. 1990. Rb-Sr and O isotopic relationships in three zoned Caledonian granitic plutons, Southern Uplands, Scotland: evidence for varied sources and hybridization of magmas. Journal of the Geological Society, London, 137, 329-348. HARRISON R. K. 1970. Hydrocarbon-bearing nodules from Heysham, Lancashire. Geological Journal, 7, 101-110. HOLLIS, C. & WALKDEN, G. 1996. The use of burial diagenetic calcite cements to determine the controls upon hydrocarbon emplacement and mineralization on a carbonate platform, Derbyshire, England. This" volume. MILLER, J. M. & TAYLOR, K. 1966. Uranium mineralization near Dalbeattie, Kirkcudbrightshire. Bulletin of the Geological Survey of Great Britain, 25, 1-18. MONSON, B. & PARNELL, J. 1992. Metal-organic relationships from the Irish Carboniferous. Chemical Geology, 99, 125-137. MOSER, M. R., RANKIN, A. H. & MILLEDGE, l-t. J. 1992. Hydrocarbon-bearing fluid inclusions in fluorite associated with the Windy Knoll bitumen deposit, UK. Geochimica et Cosmochimica Acta, 56, 155-168. O'CONNOR, P. J. 1986. Uranium mineralization of the Irish Caledonides. In: ANDREWS, C. J., CROWE, R. W. A., FINLAY, S., PENNELL. W. & PYNE, J. F. (eds) The Geology and Genesis" of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 157-175.

63

PARNELL, J. 1988. Mineralogy of uraniferous hydrocarbons in Carboniferous-hosted mineral deposits, Great Britain. Uranium, 4, 197-218. 1995. Hydrocarbon migration in the Solway Basin. Geological Journal, 30, 25 38. -& MONSON, B. 1990. Sandstone-hosted thoriumbitumen mineralization in the Northwest Irish Basin. Sedimentology, 37, 1011 1022. 8,: TOSSWILL, R. J. 1990. Petrography of thoriferous hydrocarbon nodules in sandstones, and their significance for petroleum exploration. Journal of the Geological Society, London, 147, 837-842. PERING, K. L. 1973. Bitumens associated with lead, zinc and fluorite ore minerals in North Derbyshire, England. Geochimica et Cosmochimica Acta, 37, 401-417. RASMUSSEN, B. & GLOVER, J. E. 1990. The diagenetic and economic significance of composite grains of monazite and hydrocarbon in Western Australian arenites. Journal of the Geological Society', London, 147, 843-850. SASSEN, R. 1984. Effects of radiation exposure on indicators of thermal maturity. Organic Geochemistry, 5, 183-186. SIMPSON, R. R., BROWN, G. C., PLANT, J. & OSTLE, D. 1979. Uranium mineralisation and granite magmatism in the British Isles. Philosophical Transactions o[" the Royal Society of London, 291A, 385 412. STEPHENS, W. E. & HALLIDAY, A. N. 1984. Geochemical contrasts between late Caledonian granitoid plutons of northern, central and southern Scotland. Transactions of the Royal Society of Edinburgh. Earth Sciences, 75, 259-273.

Evidence for microbial influence on the development of Lower Carboniferous buildups NElL

A. H. P I C K A R D

Institute for Biology and Geology, University o f Tromso, Dramsveien 210, N-9037 Tromso, Norway Abstract: Following a series of mass extinction events towards the end of the Devonian (Frasnian), true skeletal reef-building organisms were effectively lost from latest Devonian (Famennian) and Lower Carboniferous (Dinantian) marine ecosystems. Nevertheless, nonskeletal, mud-dominated buildups (mud-mounds) are common throughout this geological interval, and microbial activity (fungi, algae and/or bacteria) is considered to have played a significant role in the development of these buildups. Evidence of microbial activity in Dinantian buildups may be divided into two broad categories: (a) direct evidence, such as the development of stromatolitic and thrombolitic macrostructures, presence of cryptic incrustations and preservation of calcimicrobes such as Renalcis; (b) cryptic evidence, in the form of peloidal 'mud' which forms the matrix of many Dinantian buildups. The matrix peloidal micrites closely resemble modern microbially mediated (algal and bacterial) carbonate precipitates. Precipitation of the peloidal fabrics is considered to have occurred within a microbial mat that bound the accreting buildup surface, thereby allowing steep depositional slopes to develop in deep water buildups and protecting shallow water buildups from erosion. Extensive cavity systems, often developed in Dinantian buildups, may also contain stromatolitic mats and micritized cavity walls. These fabrics indicate the additional presence of internal heterotrophic microbial communities and suggest that a diverse range of microbes were present in Dinantian buildups. Controls on the development of Dinantian buildups are likely to have been multiple. Nutrient-rich (mesotrophic) water masses, in combination with relatively low sedimentation rates, may have promoted buildup initiation, while regional palaeogeography and tectonic events played a role in their location. Tournaisian buildups commonly developed in a wide range of water depths on carbonate ramps, whereas Vis6an buildups, particularly in northwestern Europe, were often restricted to relatively shallow water over tectonically controlled topographic highs. Evolution of new organisms during mid-Carboniferous time led to a diversification of buildup types. However, microbial fabrics identical to those developed in Dinantian buildups are still present in many late Carboniferous and Permian phylloid algal, Palaeoaplysina and bryozoan buildups. The role of microbial activity in the development of these late Carboniferous and Permian buildups may have been greatly underestimated.

The evolution or succession of carbonate reefs throughout Phanerozoic time has received much attention and a voluminous literature exists (see Heckel 1974; James 1983; Copper 1988, 1989). It is well known that at certain periods during the Phanerozoic, true reefs (herein defined as carbonate buildups possessing a rigid skeletal framework) were absent from marine ecosystems. During such periods of arrested reef development (cf. Copper 1988), carbonate buildups were generally non-skeletal, muddominated structures which were still capable of developing significant topography on the seafloor. In the literature these structures are commonly referred to as 'mud-mounds'. Indeed, even during periods of 'climax' reef development (cf. Copper 1988), mud-dominated buildups were often present in deeper water settings, prime examples being the deep-water

Devonian buildups of North Africa (Brachert et al. 1992; Wendt 1993; Wendt et al. 1993), which were contemporaneous with shallowwater stromatoporoid-coral reefs, or the deepwater thrombolitic mud-mounds of the Upper Jurassic (Leinfelder et al. 1993), a period of otherwise shallow-water framework reef development (Wood 1993). With the collapse of marine communities during the Late Devonian (Frasnian), organisms capable of constructing skeletal framework reefs were effectively lost from marine ecosystems (West 1988). As a consequence true skeletal reefs were virtually absent from the Famennian (Dreesen et al. 1985; Buggisch 1991) and Dinantian, although there are a few notable exceptions (Adams 1984; Fang & Hou 1987; Webb 1989; Mundy 1994; Rodriguez this volume). Nonetheless, non-skeletal buildups

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 65 82.

66

N. A. H. P I C K A R D

Fig. 1. Peloidal and clotted microfabrics in Waulsortian buildups. (a) Peloidal (P) and homogeneous (H) mud fabrics in Waulsortian facies. Note how peloid allochems (arrowed) within the peloidal micrite can be readily identified by their larger size, more dense character and sharp margins. Late Courceyan, Navan, Co. Meath, Ireland. Scale bar 300 #m. (b) Detail of peloidal micrite from same thin section as (a). The buildup matrix is composed of irregular microcrystalline peloids. These matrix peloids are cemented by a thin rim of inclusion-rich (greyish) calcite microspar (small arrows). Larger microcavities are occluded by clear calcite spar cements (c). Peloid allochems (p) and bioclast fragments (b) are present within the peloidal mud fabrics. Scale bar 200 #m. (c) Peloidal micrite developed above a fenestrate bryozoan frond (f). The peloid micrite forms the matrix sediment to several intraclasts (i) and bioclast fragments (b). The intraclasts also possess fine peloidal and clotted mud fabrics.

MICROBIAL INFLUENCE ON BUILDUPS existed throughout the Dinantian and persisted into the late Carboniferous and Permian, particularly on the northern margin of Pangaea (Davies et al. 1989a). Lower Carboniferous buildups are diverse structures, and several different types of buildups have been recognized (see West 1988; Webb 1994; Bridges et al. 1995). A wide range of bioclasts are generally present in the buildups: bryozoans, crinoids, corals, sponges, foraminiferans and algae are present in most structures, but their relative contribution to any single buildup is highly variable. For instance, trepostome bryozoans are particularly important in several Vis6an buildups such as the Nant-yGamar buildup in North Wales (Bancroft et al. 1988) and the Codroy buildups described by Dix & James (1987) from Newfoundland, but they formed only a relatively minor component in other buildups, and may simply represent opportunistic colonizers. Similarly, palaeoberesellids (probable green algae; Skompski 1987) are important constituents of several late Asbian buildups of northern England (Horbury 1992). Regardless of skeletal contribution, one feature common to many Dinantian buildups is their fine-grained or 'muddy' appearance at outcrop. The origin of the mud in these structures has long been debated, and two mechanisms have been proposed: (a) trapping and baffling by organisms living on the buildups, in particular fenestrate bryozoans and crinoids (Pray 1958; Wilson 1975; King 1986); and (b) generation within some form of microbial mat (Lees & Miller 1985, 1995; Miller 1986; Bridges & Chapman 1988; Somerville et al. 1992a; Pickard 1992, 1993; Bridges et al. 1995; Gutteridge 1995). Although bryozoans and crinoids may have locally trapped and baffled sediment (McKinney et al. 1987; Ausich & Meyer 1990), they are by no means ubiquitous in Dinantian buildups and are often just as abundant in flank and off-buildup sediments (Lees & Miller 1985, 1995). Several features indicate that the mud matrix, characteristic of these buildups, was generated in situ and that the buildups possessed a stable surficial layer. Firstly, contemporaneous

67

off-buildup sediments, especially in the case of deep-water buildups, are typically thinner and contain significantly higher quantities of fine siliciclastic sediment, suggesting that the carbonate mud was generated locally. Secondly, in the case of shallow water buildups, the fine-grained nature of the matrix is apparently at variance with inter-buildup and off-buildup sediments, indicative of mobile substrates and relatively high water energies (Somerville et al. 1992a; Pickard 1992; Gutteridge 1995). This suggests a stabilizing biotic influence. Thirdly, Dinantian buildups commonly possessed steep depositional slopes (up to 50°), again indicative of some form of stabilizing mat. Speculation abounds that microbes played an important part in generating the mud and stabilizing the surface of Dinantian buildups (see above references), but little evidence has been provided from the most important element, the mud itself. This paper highlights what evidence there is for a microbial influence on the generation of Dinantian buildups.

Evidence for a microbial origin Evidence of microbial activity in Dinantian buildups may be divided into two broad categories: (a)

(b)

direct evidence, as witnessed by the devel-

opment of stromatolitic and thrombolitic macrostructures (cf. Kennard & James 1986), the presence of cryptic incrustations either on bioclasts or directly on the mud matrix, and through the preservation of calcimicrobes such as Renalcis; indirect or cryptic evidence within the mud which comprises the bulk of the buildup sediment.

Indirect evidence. the buildup m u d m a t r i x

At outcrop the fine-grained, micritic nature of many Dinantian buildups is immediately obvious, with mudstone and wackestone fabrics

Note how the matrix peloids have not collapsed into the shelter cavity beneath the fenestrate bryozoan frond (arrows), implying that they must have either been 'supported' in some form of mat and/or have been cemented very early (see Somerville et al. 1992a, fig. 4h for a similar feature developed in a Vis6an buildup). The inclusionrich (greyish) calcite microspar cementing the matrix peloids may be an early marine cement. Late Courceyan Waulsortian buildup, Navan, Co. Meath, Ireland. Scale bar 200 #m. (d) Early stromatactoid cavity (cav) developed in Waulsortian facies displaying both peloidal (P) and homogeneous (H) mud fabrics. Internal micritic sediment possessing bioclast fragments (Int) partly fills the cavity. Note how parts of the cavity wall are formed by peloidal micrite (arrowed). Same locality as in (c). Scale bar 500#m.

68

N. A. H. PICKARD

Fig. 2. Peloidal matrix sediment in Visban buildup facies. (a) Peloidal micrite cemented by yellow inclusion-rich calcite microspar (appears grey on photograph) forming the matrix of an Arundian buildup. Note how the inclusion-rich microspar forms thin rims around microcavities (small arrows) in the peloidal matrix, while radiaxial calcite cements line larger cavities (cav). Detrital peloid allochems (possibly faecal pellets) are readily distinguished from the matrix peloids by their dense character and sharp margins (large arrow). Arundian buildup, Salmon Hill, north Co. Dublin, Ireland. Scale 1ram. (b) Peloidal matrix sediment forming macroscopic clots reminiscent of those developed in thrombolites. The macroclots are composed of peloidal micrite (pm) and are cemented by radiaxial calcite cement. Note the thin micritic lining on the cavity walls (arrowed). Asbian buildup facies, Isle of Man, Great Britain. Scale bar I mm.

dominating the facies. However, when viewed in thin section the matrix mud is often seen to be heterogeneous, with patches of uniform micrite interspersed with areas composed of small (30-200 t~m), irregular peloids of microcrystalline calcite possessing diffuse margins (Figs 1-3). Where packing of these peloids is more intense, clotted microfabrics can result. The peloids are generally cemented by a finegrained equant calcite microspar, which is normally inclusion-rich and has a distinctive yellow-brown colour (Figs 1 and 2). The terms matrix peloids or peloidal micrite will be used here to define these 'carbonate muds', which often form the matrix of Dinantian buildups. The matrix peloids are readily distinguished from detrital peloid allochems (cf. McKee & Gutschick 1969; Wright & Tucker 1990), which are generally larger, denser and possess sharply defined margins (see Figs l a-c, 2a). The peloidal micrites are considered to be primary structures because early cavity systems, commonly referred to as stromatactis or stromatactoid cavities, may be developed within them (Figs lc~t, 2). Marine cements lining the early

cavities often nucleated directly on the peloidal micrite, again indicating the early nature of these fabrics (Figs l c, 2a, 2b). A discussion on the origin of these early cavity systems lies outside the scope of this paper, and the reader is referred to Wallace (1987) and Pickard (1992) for reviews.

Problems of terminology: peloidal textures and mud-mounds The peloid and clotted fabrics present in many Dinantian buildups pose a problem in textural classification. Strict adherence to the two standard carbonate textural classification schemes of Folk (1962) and Dunham (1962) would classify the peloidal textures respectively as pelsparites or peloidal grainstones. The Dunham classification, in particular, has connotations to water energy which are completely at variance with the peloid fabrics described here, because the matrix peloids are not considered to have been individual allochems (see discussion below). There is no evidence to suggest that the matrix peloids were unbound and reworked on the surface of the accreting buildups (Pickard 1992) and they cannot be considered true allochems.

Fig. 3. Microbial incrustations in Vis~an buildups. (a) Microbially encrusted bioclast fragments (c, coral; f, fenestrate bryozoan; X, a recrystallised bioclast of unknown affinity). Note the transition between the microbial incrustation and micritized cements (arrowed) and multiple generations of internal sediment. Arundian buildup facies, Clonalvy Quarry, north Co. Dublin, Ireland. Scale bar 2 mm. (b) Irregular laminated 'stromatolitic' mat developed within a cavity system. Early generations of radiaxial marine cements line the cavity and nucleated directly on to peloidal matrix fabrics (large arrow). A hiatus in cementation enabled microbial infestation and micritization of the cements (m) while a laminated cavity-fill developed: irregular crinkly micrite laminae (small arrows) alternate with peloid-rich laminae. Bioclast debris is very rare in the cavity-fill, which can be classified as a 'skeletal stromatolite' (sensu Riding 1991b). Note how the micritic laminae of the cavity-fill pass laterally into the micritized cements (x). Arundian buildup, Clonalvy Quarry, north Co. Dublin, Ireland. Scale bar 3 mm. (e) Microbial incrustation (M) directly coating microthrombolites composed of peloidal micrite (P). The incrustation consists of thin, discontinuous, crinkly micritic laminae alternating with microspar laminae. Inclusion-rich radiaxial calcite cement coats the microbial incrustations. Arundian buildup facies, Salmon Hill, north Co. Dublin, Ireland. Scale bar 2 mm.

70

N. A. H. PICKARD

Problems also arise when the term 'mudmound' is applied to such buildups as the matrix is not strictly a true carbonate mud (as defined in the microfacies classification schemes mentioned above). Pickard (1992, p. 1082) highlighted these problems when describing a series of Brigantian buildups in the Midland Valley of Scotland and rejected the term 'mud-mound' for the less contradictory term 'buildup'. An opposite tack was taken by Horbury (1992), who termed the late Asbian structures he studied 'peloidal cementstone{-palaeoberesellid}' buildups, although the matrix peloid fabrics he described are identical to those figured here from other Vis6an buildups (see Somerville et al. this volume). Nevertheless, these buildups appear 'muddy' at outcrop, and the term 'mudmound' ought to be used as a general field description, pending detailed microfacies studies. Origin o f the peloidal micrites Peloidal and clotted micrites have long been associated with microbial carbonates. Riding (1991a) termed such fabrics 'cryptic microbial carbonates' and they have been described from numerous fossil buildups (e.g. Schwarzacher 1961; Scoffin 1971; Alberstadt et al. 1974; F1/igel & Steiger 1981; Mountjoy & Riding 1981; Pratt 1982; Pratt & James 1982; Bourque & Gignac 1983; Reid 1987; Dickson et al. 1987; Bridges & Chapman 1988; Sun & Wright 1989: Somerville et al. 1992a, b; Kelly & Somerville 1992; Guo & Riding 1992a; Leinfelder et al. 1993; Gutteridge 1995; Lees & Miller 1995). Peloidal fabrics have also been described from many Holocene reefs where they generally form cement crusts in intra- or interskeletal pores (Friedman et al. 1974; Macintyre 1985; Friedman 1993). Both inorganic (Macintyre 1984, 1985; Ai'ssaoui 1988) and microbial (Marshall 1983; Chafetz 1986) controls on the formation of these peloidal cements have been proposed. The extensive development of matrix peloids in many Dinantian buildups suggests that they must have formed at, or close to, the accreting buildup surface. Even in buildups associated with high-energy off-buildup sediments, no sedimentary structures indicative of current reworking are present in the peloidal micrites. The peloids were never 'loose' sedimentary grains, which suggests that they formed in some sort of stable surficial mat. This implies, albeit circumstantially, a microbial origin for these fabrics. Perhaps the most compelling evidence for microbial controls on the formation of the peloidal micrites comes from studies of modern microbial mats (Dalrymple 1965;

Friedman et al. 1973; Monty 1976; Buczynski & Chafetz 1991; Chafetz & Buczynski 1992) and travertines (Chafetz & Folk 1984; Folk et al. 1985; Pedley 1992; Guo & Riding 1992b, 1994). Both algae and bacteria are known to be able to induce precipitation of calcium carbonate, given the correct geochemical microenvironments. Aragonite and high-Mg calcite may form within and/or upon the mucilaginous envelopes of coccoidal cells and sheaths enclosing cyanophyte trichomes (Riding 1975; Monty 1976; Pratt 1982; Pentecost & Riding 1986; Riding 1991b). Monty (1976) recorded peloidal fabrics within pustular cyanobacterial colonies from Shark Bay, Western Australia, while Buczynski & Chafetz (1991) and Chafetz & Buczynski (1992) demonstrated the importance of bacteria in the precipitation of carbonate peloids in modern algal mats. The bacterially mediated precipitates consisted of submicron-sized aggregates or bundles of crystals, whose gross morphology ranged from rods, dumbbells and spheres to more angular forms. Viewed in thin section, such precipitates are likely to resemble the peloidal fabrics described here. The diffuse margin of the Carboniferous matrix peloids probably results from the thickness of the rock slice, as a single peloid margin would necessarily be comprised of many submicron-sized crystals. Interestingly, the mineralogy of the precipitates produced in Buczynski & Chafetz's (1991) experiments varied with the viscosity of the media on which the bacteria were grown, viscous gel (agar) media producing calcite, and liquid media aragonite. Bacterially induced precipitation of calcium carbonate has been noted by many authors (see Guo & Riding 1994 for extensive review). Precipitation may be promoted by biochemically controlled processes such as oxidation of organic matter, reduction of sulphates and nitrate, ammonification and photosynthesis (Lalou 1957; Greenfield 1963, 1965; Dalrymple 1965; Krumbein 1979; Krumbein & Giele 1979; Morita 1980; Guo & Riding 1994). Mucilage secreted by bacteria (extracellular polymeric substances) together with the bacteria cell surface can act as sites for the accumulation of calcium ions, thereby promoting the precipitation of calcium carbonate (Greenfield 1963, 1965; Morita 1980; Guo & Riding 1994). Moreover, decomposition of the organic matter associated with microbes can release ammonia, which at high pH can promote calcite precipitation (Greenfield 1963, 1965; Dalrymple 1965; Golubic 1976), thereby enabling calcite .precipitation to continue some distance below the sediment surface.

MICROBIAL INFLUENCE ON BUILDUPS The peloidal micrites described here also bear a striking resemblance to the microbially mediated fabrics present in modern and Pleistocene travertines (Chafetz & Folk 1984; Folk et al. 1985; Guo & Riding 1992b, 1994; Pedley 1992), a point highlighted by Guo & Riding (1992a), in their comparison of Recent travertines and Upper Permian reefs. The peloidal micrites that form the matrix 'mud' in many Dinantian buildups are considered to have formed within a surficial microbial mat or biofilm (cf. Pedley 1992), which produced a cohesive surface to the accreting buildup. The biofilm was capable of supporting the finegrained matrix sediment on high depositional slopes and apparently resisted erosion in turbulent environments (Pickard 1992) in a manner similar to microbial mats in modern environments (Scoffin 1970; Monty 1976; Burne & Moore 1987; Riding 1991a).

D i r e c t evidence

Stromatolites and cryptic laminated incrustations have been documented from several Dinantian buildups (Wolfenden 1958; Mundy 1980, 1994; Dix & James 1987; Bridges & Chapman 1988; Somerville et al. 1992a; Pickard 1992; Lauwers 1992). Riding (1991a) subdivided stromatolites into those formed by the calcification of microbes (skeletal stromatolites) and those formed by the trapping of sediment grains by a microbial mat (agglutinated stromatolites). Most of the stromatolitic and encrusting fabrics described from Dinantian buildups conform to the definition of skeletal stromatolites, although agglutinated stromatolites have been described from shallow water, Brigantian-aged buildups in the Midland Valley of Scotland (Pickard 1992). Most of the skeletal stromatolites are composed of alternating laminae of dense micrite and calcite microspar (Fig. 3). The micritic laminae are often impersistent and may be crinkled. Thin layers of peloidal micrite can also be developed within the skeletal stromatolites (Fig. 3b; see also Somerville et al. 1992a). Encrusting organisms are often associated with these skeletal stromatolites, and include attached encrusting foraminiferans (e.g. Tetrataxis), serpulid worms and vermetid gastropods, indicating that the stromatolites and encrustations formed cohesive, if not lithified mats and coatings. Skeletal stromatolites may directly coat peloidal matrix sediment (Fig. 3c; see also Dix & James 1987) supporting an early, primary origin for

71

these fabrics. It also suggests that an irregular microtopography sometimes existed on the surface of accreting buildups, not dissimilar to that envisaged during the development of thrombolites (cf. Kennard & James 1986; Riding 199 la). Indeed, thrombolitic fabrics have been recorded from several Lower Carboniferous buildups (Fig. 3c; Adams 1984; Dix & James 1987; Webb 1987; Horbury 1992) and the thrombolite heads may themselves be coated by microbial incrustations and encrusting organisms (Fig. 3c; Webb 1987; Somerville et al. 1992a). Peloidal micrites are commonly associated with these Dinantian thrombolites (Fig. 3c; Dix & James 1987; Webb 1987; Horbury 1992) and with thrombolites in general. Moreover, thrombolites are themselves considered to have formed from the calcification of coccoid cyanobacterial cells (Monty 1976; Pentecost & Riding 1986; Kennard & James, 1986; Riding 1991a). Stromatolitic mats have also been recorded within internal (stromatactoid) cavity systems (Fig. 3b) indicating microbial communities capable of heterotrophic growth. In some cases it is difficult to separate these internal laminated fabrics from those of 'micritized cements' (Figs 3a, b; Somerville et al. 1992a) or micritized cavity walls (Figs l c, 2b), features that are common in many Dinantian buildups (Lees et al. 1985; Lees & Miller 1985, 1995; Miller 1986; Bridges & Chapman 1988; Somerville et al. 1992a). Furthermore, some authors have suggested that the early marine cements present in many internal cavity systems also have a microbial origin (see below). Further evidence of microbial activity in Dinantian buildups comes from the occasional preservation of the calcimicrobe Renalcis (Adams 1983, 1984; Somerville et al. 1992a; Somerville & Strogen 1992). Renalcis is generally considered to have formed from the precipitation of microcrystalline calcite within coccoid cyanobacteria colonies (Pratt 1984; Riding 1991b).

E a r l y m a r i n e ?microbial m e d i a t e d c e m e n t s

Inferred marine cements can form a significant proportion of the rock volume in Lower Carboniferous buildups. They line intra- and interskeletal voids, stromatactoid cavities and early submarine fissure systems (Figs 2 & 3). Single or multiple generations of isopachous radiaxial, fascicular optic and coarse columnar calcite crystals typify these cements, although botryoidal growth forms have also been

72

N. A. H. PICKARD

described (Gillies 1987; Dix & James 1987; Lees & Miller 1995). Controversy persists over whether radiaxial and fascicular optic calcite are primary morphologies or secondary neomorphic fabrics produced from the recrystallization of former acicular high-Mg calcite (or even aragonite) cements (Sandberg 1985; Kendall 1985). Several authors have even suggested that such cements (or their original precursors) may themselves have been influenced by microbial activity (Monty 1982, 1984; Mamet & Boulvain 1988). Videtich (1985) showed that the inclusion-rich nature of Tertiary to Recent prismatic and radiaxial calcite cements from Enewetak Atoll were the result of extensive microbial infestation, particularly by fungi. Moreover, she went on to conclude that micritization by endolithic organisms may well be as important in cements as it is in sediment grains and could account for the inclusion-rich character of many ancient marine cement generations. Similar microbial infestation and micritization has also been recorded from subaerial crusts by Kahle (1977), who termed the process 'sparmicritization'. In Dinantian buildups, early marine cement generations may alternate with thin, sometimes impersistent, micritic laminae, which represent either microcrystalline cements or micritization caused by infestation of the cements during a hiatus in crystal growth. As stated above, it is often difficult to differentiate between 'cavity stromatolites', laminated incrustations and these micritized cements (Figs 3a, b). The yellow, inclusion-rich calcite microspar which commonly cements the matrix peloids could also have been precipitated from marine water. Beach (1993) recorded similar yellow, inclusion-rich calcite spar in Pliocene carbonates cored from the Great Bahama Bank which he considered to be of marine origin. Early stabilization of peloidal micrites in Dinantian buildups, through the precipitation of the inclusion-rich calcite microspar, may have played an important role in the formation of internal cavity systems (see Fig. l c). When compared with the Tournaisian Waulsortian buildups, cavity systems are generally not as extensive in Vis6an buildups, where peloidal micrite forms a significant percentage of the mud matrix (Somerville et al. 1992a). Early cementation of the matrix peloids in Vis6an buildups would have inhibited both internal erosion and dissolution of the matrix, two processes which are considered to be important in the development of stromatactoid cavity systems (Miller 1986; Wallace 1987; Pickard 1992).

Discussion The peloidal and clotted micrites present in many Dinantian buildups are considered to have formed within a surficial microbial mat. Rare preservation of the calcified coccoid cyanobacteria Renalcis offers the only direct evidence of microbial activity within the matrix fabrics, and the exact nature of the benthic microbial communities responsible for generating these 'muds' remains unknown. However, both algae and bacteria are likely to have been important constituents. That the microbial mats influenced sediment stability is witnessed by the steep depositional slopes attained on the margins of many Dinantian buildups, and by the fact that while some buildups developed in relatively turbulent environments the peloids were not reworked (Pickard 1992). Meadows et al. (1994) demonstrated how extracellular polymeric substances secreted by bacteria can influence the slope stability of loose sand. The viscous mucus produced by the single celled bacterium Pseudomonas atlantica effectively glued the sediment grains together. Meanwhile, experiments by Buczynski & Chafetz (1991) using bacteria cultivated from marine algal mats, illustrated how the viscosity of the growth media influenced the mineralogy of carbonate precipitates induced by the bacteria. High-viscosity media favoured the precipitation of calcite. The original mineralogy of the Dinantian matrix peloids is not known. However, former aragonitic bioclasts, within the peloidal matrix, are normally preserved as calcite-cemented moulds, indicating that skeletal aragonite was susceptible to early dissolution (see Fig. 6b). That the peloids still retain a microcrystalline character is perhaps indicative of a primary calcitic mineralogy. Could the peloidal micrites have been formed within viscous extracellular polymeric substances secreted around bacteria and/or algae? Precipitation could have occurred within the viscous extracellular material, on the surface of microbes or even within the microbes (algae or bacteria) themselves. It is also conceivable that it continued during the subsequent decomposition of the organic material during burial (Fig. 4). Skeletal stromatolites and cryptic incrustations played an important role in the development of some Dinantian buildups. They are best developed in shallow-water Vis6an buildups where, in conjunction with skeletal elements, they locally formed biogenic frameworks (Fig. 3a; Wolfenden 1958; Mundy 1980, 1994;

MICROBIAL INFLUENCE ON BUILDUPS

Cavities Surficial microbial mat Zone of degradation and decay of the microbial mat - incipient cavity development Zone of extensive cavity development and sediment stabilisatlon by submarine cementation

Adams 1984; Webb 1989; Somerville et al. 1992a, this volume). Internal stromatolitic fabrics together with micritized cements and cavity walls attest to the presence of microbial communities capable of heterotrophic growth, similar to various extant algae and bacteria found in modern reef cavities (Kanwisher & Wainwright 1967; Schroeder 1972; Riding 1975; Chafetz 1986). In conclusion, a complex and diverse assemblage of autotrophic and heterotrophic microbes was probably present in all Dinantian buildups, while the skeletal contribution was very variable and probably related to the depositional setting (i.e. depth, degree of water turbulence, nutrient supply, etc.) of individual buildups. A simplistic model for the buildup growth and mud generation is presented in Fig. 4.

73

Fig. 4. Model for mud generation. The surface of the accreting buildup has been subdivided into three layers. (a) A living surface layer of microbes (algae, fungi and/or bacteria) and secreted extracellular polymeric substances which bound the sediment producing a cohesive surfacial mat. Calcite precipitation could occur within the viscous extracellular substances, on the surface of microbes or even within the microbes themselves. Contribution of skeletal debris is highly variable. Skeletal stromatolites and encrusting microbial mats locally coat bioclasts and the buildup surface (see Fig. 3). The surfacial mat may locally develop thrombolitic fabrics. (b) Zone of breakdown and decay of the mat. Given the correct geochemical environments calcite precipitation could continue into this zone. Decay of the extracellular polymeric substances could have released organic acids; very early carbonate dissolution in Dinantian buildups has been recorded by several authors (Miller 1986; Gillies 1987; Pickard 1992; Lees & Miller 1995) and may have played an important role in the initiation of stromatactoid cavities. As the binding influence of the mat decreased through decay, the sediment would have been more susceptible to internal erosion and collapse, which also plays an important role in cavity development (Wallace 1987). This may also have allowed influx of marine phreatic water and thus precipitation of theinclusion-rich microspar cement, thereby stabilizing the sediment. Each of the above processes may have been occurring in close proximity to each other and would have continued into the third zone. (c) Zone of extensive cavity formation and sediment stabilization through submarine cementation. Micritized cements and skeletal stromatolites within the cavities attest to abundant heterotrophic microbial activity within the cavity systems.

Controls on buildup initiation Throughout Phanerozoic time, periods of arrested reef development have tended to follow major episodes of extinction, during which organisms capable of constructing framework reefs were lost from marine ecosystems (Heckel 1974; James 1983; Copper 1988, 1989). The extinction events at the end of the Frasnian Age removed most of the main reef frame builders (e.g. stromatoporoids and tabulate corals) from the marine ecosystems of the Famennian and Dinantian Ages. Buggisch (1991) suggested that the extinction events were the result of oceanic turnover which bathed shallow shelf regions with cold, oxygendeficient, nutrient-rich waters. Nutrient-rich waters are known to have detrimental effects

74

N. A. H. PICKARD

on modern reef systems (Hallock & Schlager t986), and may have had similar effects on ancient oligotrophic 'reef' communities such as the late Devonian stromatoporoid-coral reefs (Buggisch 1991; Wood 1993). Although several controls have been proposed for the distribution and development of early Dinantian Waulsortian buildups, including development around sea-floor seeps (Hovland 1990), or location in areas protected from tropical storms (King 1990), there is a growing consensus that the distribution of Waulsortian buildups is more likely to reflect sites of upwelling, nutrientrich water (Gutschick & Sandberg 1983: Wright 1991, 1994; Wood 1993; Bridges et al. 1995). Indeed, Wright (1994) speculated that the upwelling of cold, nutrient-rich waters onto early Dinantian carbonate ramps (see below) may have favoured inorganic precipitation of ooids, deterred the growth of coral reefs, and promoted the growth of heterotrophs such as crinoids, bryozoans and brachiopods. It has been argued that Lower Carboniferous nutrient-enriched (mesotrophic) tropical waters might have suppressed the evolution of oligotrophic reef-building organisms, thereby promoting the development of mud-mounds (Wood 1993; Wright 1994). Whereas this could account for the preponderance of non-skeletal buildups in the Lower Carboniferous, it does not explain what triggered the initial colonization of the sea floor by the microbial communities responsible for buildup development. Pickard (1992), while discussing the origin of shallow water Brigantian-aged buildups from the Midland Valley of Scotland, suggested that decreased sedimentation rates, increased illumination and water turbulence, or a critical combination of these, may have provided the micro-environmental conditions required for initial microbial colonization. Indeed, several authors have suggested that low rates of sediment accumulation would favour the development of microbial communities (Sun & Wright 1989; Leinfelder et al. 1993). Decreased sedimentation rates would be favoured during transgression and particularly during periods of maximum flooding. The Tournaisian Waulsortian buildups of northwestern Europe developed during a regional transgression, albeit one that was punctuated by short regressive episodes. In the Dublin Basin, Waulsortian facies (Feltrim Limestone Formation) overlie a thick sequence of rocks that shows a progressive deepening of depositional environment, reflecting the continuing northward transgression of the early Carboniferous seas (Somerville et al. 1992b). Lower sedimentation

rates related to this deepening may have triggered the development of the Waulsortian facies. Likewise, Hennebert & Lees (1991), in their study of the Lower Carboniferous ramp of southwest Britain, considered that the Waulsortian buildup encountered in the Knap Farm Borehole at Cannington was initiated during a period of maximum flooding. Upper Carboniferous 'Waulsortian-like' buildups of Arctic Canada also appear to have developed during transgressive episodes, and those cropping out in northwestern Ellesmere Island are developed in strata that overlie the evaporite sequences of the Otto Fiord Formation (Davies et al. 1989b). Flooding of the evaporite basin may have produced nutrient-rich waters which, in combination with slow sedimentation rates during the transgressive phase, may have provided conditions favourable for the development of microbial buildups.

Evolution of Carboniferous buildups

In northwestern Europe the evolution of Dinantian buildups was closely linked with the regional tectonic framework (Fig. 5). Following the early Carboniferous transgression, carbonate ramps dominated early Dinantian (Courceyan) depositional environments (Ahr 1989; Wright & Faulkner 1990; Wright 1994). Large Waulsortian buildup complexes developed on the distal parts of the ramps (Fig. 5; Bridges & Chapman 1988; Somerville et al. 1992b; Bridges et al. 1995) and formed in a wide depth range. However, a pulse of extensional tectonism in latest Tournaisian time (early Chadian, see Jones & Somerville this volume) resulted in the break-up of the ramp systems and the development of more 'block and basin' sedimentation (Gawthorpe et al. 1989; Adams et al. 1990; Pickard et al. 1994; Bridges et al. 1995). Vis6an carbonate buildups tended to be restricted to the shallower water that persisted across the tectonic highs (Fig. 5; Bridges et al. 1995). Peloidal and clotted micrites are particularly well developed in Vis6an buildups, and generally comprise the bulk of the matrix sediment (Somerville et al. 1992a; Pickard 1992; Gutteridge 1995; Somerville et al. this volume). Cryptic incrustations and skeletal stromatolites are also more prevalent in Vis6an buildups. However, both peloidal micrites and (albeit less frequently) skeletal stromatolites have also been described from the shallower-water Waulsortian buildups (Bridges & Chapman 1988; Somerville et al. 1992b; Bridges et al. 1995; Lees & Miller 1995). Interestingly, Lees & Miller (1995) noted

MICROBIAL INFLUENCE ON BUILDUPS

75

Late Dinantian

~~._/+ + + + ~ "++÷

++++ + + + +

+++++++~

I/lick ~ a l a

- , + + + + + + + + + + + + + +

~ + + + + + + .

+ + + + + + + + + + + + + " + + + + + + + + + + + + + + +

+ + + + + + + + + + + + + .

Carbonate platforms develop over structural

Early Dinantian

highs +/-

~flux

of coarse dlclclodlcs •

,

,

Fig. 5. Sketch showing the generalized evolution of Dinantian buildups in northwestern Europe. Tournaisian Waulsortian buildups of Belgium, SW England and Ireland developed on carbonate ramps, while Vis6an buildups were generally restricted to structural highs following a major late Tournaisian extensional event. Modified from Bridges et al. (1995).

that peloidal mud fabrics are absent in Waulsortian rocks possessing the deepest Phase A grain assemblages. They are, however, common (if not ubiquitous) in the shallowest Waulsortian buildups (Fig. 1; Phases C and D). It is well documented that skeletal grain assemblages in Waulsortian facies change with decreasing water depth (Phases A-D; Lees & Miller 1985, 1995). Is it not, therefore, reasonable to expect a change in the benthic microbial communities responsible for the in situ generation of mud with increasingly shallower water conditions? Certainly, the benthic microbial communities responsible for generating the peloidal and clotted micrites appear to be confined to the shallower-water spectrum of Dinantian buildups. The benthic microbial communities responsible for the development of the shallow-water Waulsortian facies appear to have readily

exploited the shallow-water settings developed on structural highs following the latest Tournaisian tectonic event. Influx of fine siliciclastic sediment into basinal areas increased during the Vis6an, and large deep-water buildups appear to have been restricted to areas starved of siliciclastic sediment. Examples include Arundian and Asbian-aged buildups in northwestern Ireland (Schwarzacher 1961; Kelly & Somerville 1992), which developed on the hanging walls of major tilt blocks, or Brigantian buildups of central England (Gutteridge 1995), which developed on the margin of an intrashelf basin. Interestingly, Kelly & Somerville (1992) considered their deeper-water Arundian buildups to be a continuum of Waulsortian-type facies into the Vis6an. Throughout the Vis~an, carbonate buildups were notably absent from the areas of high subsidence and sediment influx, such as parts of the Dublin Basin (Pickard et al. 1994),

76

N.A.H.

PICKARD

Fig. 6. Photomicrographs of peloidal matrix fabrics from Upper Carboniferous buildups. (a) Homogeneous (H) and peloidal micrite (P) fabrics developed in a fenestrate bryozoan-Tubiphytes 'Waulsortian-like' buildup. Late Moscovian, Central Spitsbergen. Scale bar 300 #m. (b)-(d) Kasimovian-aged phylloid algal buildups, central Spitsbergen. (b) Peloidai micrite in a phylloid algal buildup. The irregular diffuse-margined peloids are cemented by inclusion-rich calcite microspar, which forms a fringe around the peloids (small arrows). Note fenestrate bryozoan frond (fb) and mould of the former aragonitic phylloid algal plates (large arrows). Scale bar 300 #m. (e) Irregular cavities developed in peloidal micrite matrix sediment (area depicted in (d) is outlined). Scale bar 750 #m. (d) Close-up of (e) illustrating the peioidal and clotted nature of the matrix sediment. A thin fringe of inclusion-rich cement coats all the peloids (arrows). Scale bar 300 #m.

MICROBIAL INFLUENCE ON BUILDUPS Edale Basin (Gutteridge 1991), and the Craven Basin of northeastern England (Gawthorpe 1986; Gawthorpe et al. 1989; Adams et al. 1990). The mid-Carboniferous heralded the development of several different types of carbonate buildups as new organisms evolved (West 1988). Of particular note are the phylloid algae and several problematic organisms such as Tubiphytes and Palaeoaplysina. The term phylloid was first coined by Pray & Wray (1963) and literally means 'leaf-like'. They applied it to Late Palaeozoic calcareous algae possessing a leaf-like geometry but with insufficient internal features preserved for a generic identification. From well preserved specimens, phylloid algae are known to have varying affinities; both codiaceans and ancestral coralline red algae have been identified (Wray 1977; Riding & Guo 1991). Although very rare, phylloid algae do occur in late Dinantian buildups (Somerville et al. this volume). However, phylloids became important constituents of late Carboniferous and early Permian buildups (Toomey & Winland 1973; Toomey et al. 1977; Dawson & Carozzi 1986). Peloidal micrites continued to be an important component of the matrix sediment in many Upper Carboniferous and Permian buildups, and are common in phylloid algal, Palaeoaplysina and bryozoan buildups which crop out on Svalbard (Fig. 6). Indeed, microbial activity may have played a far greater role in the development of Upper Carboniferous and Permian buildups than has previously been recognized. Buildups possessing features almost identical to the classic Waulsortian (cf. Lees 1988) persisted into the late Carboniferous and Permian and are common along the northern margin of Pangaea. These 'Waulsortian-like' buildups are normally rich in fenestrate bryozoans and the problematic sponge Tubiphytes (cf. Riding & Guo 1990) and have been described from Arctic Canada (Davies et al. 1989a, b; Beauchamp 1989) and Svalbard (Fig. 6a). On the northern margin of Pangaea, deep-water bryozoan mud-mounds were unaffected by the late Lower Permian (Artinskian) climatic cooling event, which led to the disappearance of shallower-water, algaldominated (phylloid and Palaeoaplysina) buildups (Beauchamp 1993, 1994). Microbial activity continued to be influential in the development of these temperate, deep-water buildups.

Conclusions It is considered here that there is compelling evidence for significant microbial influence on

77

the development of many Carboniferous buildups. Evidence for microbial (algal, fungal and/ or bacterial) activity is preserved in the form of (a) peloidal micrites in the buildup matrix sediment, (b) encrusting mats, stromatolites and thrombolites, and (c) micritized cements and internal stromatolites in early submarine cavity systems. Much of the carbonate mud within the Lower Carboniferous buildups is considered to have formed in situ in a surficial microbial mat which stabilized the accreting buildup surface, allowing steep depositional slopes to develop at the margins of deeperwater buildups, and cohesive, erosion-resistant buildups to develop in shallow water. Periods of rapid transgression and/or low sedi-mentation rates, in combination with nutrient-rich (mesotrophic) water, could have promoted sea floor colonization of the benthic microbial communities that were responsible for the initiation and growth of Lower Carboniferous non-skeletal, mud-dominated buildups. I. Somerville is thanked for encouraging me to put these ideas on paper and for critically commenting on an early version of the manuscript. Thoughtful and constructive reviews by R. Riding and B. Stanton are also gratefully acknowledged. I am indebted to H. Lunde and A. Jensen (Chr. Hansen Express Bilder) and G. Granaas (IBG, UiTo) for their help in putting together the photographic plates. Work on the Late Palaeozoic buildups of Svalbard was initially supported by a Royal Society Overseas Fellowship and now by the Norwegian Research Council (Norges Forskningsrfid). Additional funding has been provided by Amoco, British Gas, Elf Norge A/S, the Norwegian Petroleum Directorate (Oljedirektoratet), Norsk Hydro, Norske Shell, Saga Petroleum and Statoil. I would like to thank these companies for their continuing support of the Late Palaeozoic buildup project. This work has benefited from countless discussions with fellow Carboniferous buildup enthusiasts.

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WEST, R. R. 1988. Temporal changes in Carboniferous reef mound communities. Palaios, 3, 152 169. WILSON, J. L. 1975. Carbonate Facies in Geologic History. Springer-Verlag, New York. WOLFENDEN, E. B. 1958. Paleoecology of the Carboniferous reef complex and shelf limestones in northwest Derbyshire, England. Geological Society of America Bulletin, 69, 871-898. WOOD, R. 1993. Nutrients, predation and the history of reef-building. Palaios, 8, 526-543. WRAY, J. L. 1977 Calcareous Algae. Developments in Palaeontology and Stratigraphy, 4, Elsevier, Amsterdam. WRIGHT, V. P. 1991. Comment on 'Probable influence of early Carboniferous (Tournaisianearly VisGan) geography on the development of Waulsortian and Waulsortian-like mounds'. Geology, 19, 413. 1994. Early Carboniferous carbonate systems: an alternative to the Cainozoic paradigm. Sedimentary Geology, 93, 1-5. - & FAULKNER, T. J. 1990. Sediment dynamics of Early Carboniferous ramps: a proposal. Geological Journal, 25, 139-144. -& TUCKER, M. E. 1990. Carbonate Sedimentology. Blackwell Scientific Publications, Oxford.

Constituent composition of Early Mississippian carbonate buildups and their level-bottom equivalents, Sacramento Mountains, New Mexico WAYNE

M. A H R & ROBERT J. S T A N T O N JR.

Dept. of Geology & Geophysics, Texas A & M University, College Station, Texas 77843-3115, USA Abstract: The Alamogordo Member of the Lake Valley Formation (Lower Mississippian Tournaisian to early Vis6an) was deposited on a homoclinal ramp at depths below the fair-weather wave-base and mainly below the photic zone. Depth increased toward the present southwest, and the level-bottom biota of the Alamogordo Member is present in assemblages that parallel the depth gradient (Jeffery & Stanton this volume). Waulsortianlike mounds are abundant in the Alamogordo, Nunn and Tierra Blanca Members of the Lake Valley Formation on the distal portion of the homoclinal ramp. These mounds contain essentially the same biota as the Waulsortian buildups of Europe. Alamogordo mounds are low, tabular features composed mainly of spiculiferous micrite, while Nunn and Tierra Blanca mounds are more domical and are composed largely of fenestrate bryozoan cementstone. Lithology and biota of coeval, level-bottom beds and mounds differ significantly because of habitat differences (cavities and hard substrates in and on mounds), and because of taphonomy, an example of which is the preservation of fenestrate bryozoan sheets in radiaxial fibrous calcite cements on mounds as compared with comminuted and dispersed fenestrate hash in level-bottom muds. Lake Valley mound constituents, although essentially identical to those in the European Waulsortian, do not occur in regular growth 'phases' like those described by Lees & Miller (1985). The difference is interpreted to mean that the New Mexican mounds lacked enough vertical relief above the sea floor in deep water to have been affected by any environmental gradients that may have existed across the entire platform. Carbonate buildups of Mississippian age (approximately equivalent to the Dinantian Series of Europe) have been classified collectively as 'Waulsortian mounds' by many geologists. This broad generalization has proven to be misleading and incorrect. Today it is known that there are many different types of Mississippian buildups around the globe, some with and some without Waulsortian characteristics (Webb 1994). In North America and northwestern Europe, Waulsortian buildups tend to occur mainly on distal parts of ramps of Early Mississippian (Tournaisian and early Vis6an) age. Shelf-margin buildups of this age, and buildups of younger Mississippian (late Vis6an) age (see Somerville et al. this volume) are generally not like those of the 'type' Waulsortian as described by Lees (1988). Waulsortian buildups are characterized mainly by mudstone-cementstone fabrics, comparatively low macrobiotic diversity, and a lack of framebuilding organisms, although some fenestrate bryozoan colonies in cementstone fabrics may approach framestone character. In contrast to the moderately low diversity of larger fossils, Waulsortian buildups may have rich assemblages of microfossils, a wide variety of mud and cement textures and fabrics, and

commonly, stromatactis cavities. Studies on the constituent composition of Tournaisian-early Vis6an buildups in northwestern Europe, particularly in Belgium, Britain and Ireland, indicate that skeletal and non-skeletal mound constituents, along with diagenetic characteristics such as micritization, may form assemblages, which have been called 'phases' by Lees & Miller (1985). These phases are similar to microfacies of other workers. The vertical arrangement of the phases and of 'relays' of constituent types within the phases and sequences of phases, have been related primarily to relative depth. This is based mainly on inferred depth of light penetration (photic zone) in the sea, but also other environmental parameters such as turbidity, temperature, salinity, oxygen, nutrients, and wave and current activity. The importance of these environmental parameters during mound growth should be corroborated in the depositional microfacies and sedimentary structures of coeval, level-bottom beds, and inferences about the environment of deposition could thus be confirmed or denied. An analysis of mounds and associated level-bottom beds has been difficult in northwestern Europe because outcrops there are limited in areal extent (quarries and gullies) and discontinuous. Such an analysis, however, is

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 83-95.

84

W. M. AHR & R. J. STANTON JR.

possible in the strata and mounds deposited on the distal portion of a homoclinal ramp now exposed in the Sacramento Mountains of New Mexico along more than 30kin of continuous outcrops. This study has five objectives: (1) to describe the general lithological and palaeontological characteristics of nine Lower Mississippian mounds in the Lake Valley Formation in the Sacramento Mountains; (2) to describe distribution patterns and systematic groupings of constituents within these mounds; (3) to describe the general lithological and palaeontological characteristics of the coeval, level-bottom beds; (4) to compare the lithology and biota of the level-bottom beds with those of the mounds during their growth; (5) to compare briefly the Sacramento Mountain mounds and level-bottom beds on the basis of these characteristics, with some time-equivalent examples from northwestern Europe.

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Setting The Mississippian strata of the Sacramento Mountains have been studied extensively during the past century. The current structural and lithostratigraphic relationships were established by Pray (1961), and the regional context for these strata has been investigated by Armstrong (1962), Kottlowski (1963,1965) and Lane (1974). Mounds and level-bottom beds in the study area crop out along the fault-scarped west front of the Sacramento Mountains, near the town of Alamogordo, New Mexico, from Indian Wells Canyon in the north to Deadman Canyon in the south (Fig. 1). The mounds were first mentioned by Laudon & Bowsher (1941, 1949), and the parallel between them and the European Waulsortian mounds was pointed out by Pray (1958). They are restricted to the Alamogordo, Nunn and Tierra Blanca Members of the Lake Valley Formation (Figs 2 and 3). Mound growth was limited to the Alamogordo Member north of Marble Canyon, but extended into the Nunn and Tierra Blanca Members from there southward. Most of the mounds 'rooted' in the Alamogordo Member, but a few began to grow during deposition of the upper part of the Andrecito Member. The Alamogordo Member and correlative phases of mound growth are in Faunal Unit 3 of Lane & Ormiston (1982), which is equated with the Lower typicus Zone of Lane et al. (1980). Mound growth during the Nunn and Tierra Blanca Members extended upward into Faunal Units 4L, 4U and 5 of Lane & Ormiston (1982),

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corresponding to the Upper typicus and anchoralis-latus Zones (Fig. 2). The cement stratigraphy and diagenetic history of these strata were described by Meyers (1974). Ahr (1989a) established that the Mississippian platform in the Sacramento Mountains was a ramp, and showed that the Waulsortian mounds were localized to palaeobathymetric highs of both tectonic and depositional origin. The level-bottom and mound biota have been compared by Ahr & Stanton (1994). Lithological and biotic gradients in the levelbottom strata have been studied by Jeffery & Stanton (this volume).

COMPARISON OF RAMP AND M O U N D CHARACTERISTICS Laudon & B o w s h e r 1949

rock unit

~

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(25 × 45 mm) thin sections: (a) 100 thin sections are from a 21m core in the upper part of Muleshoe Mound (upper Tierra Blanca Member equivalent) taken by Shinn et al. (1983), at intervals ranging from 5 cm to 46 cm; (b) 54 thin sections are surface samples from three vertical transects across the lower and middle part of the cliff face of Muleshoe Mound (mainly lower Tierra Blanca and Nunn Member-equivalents), and collected by W. D. Jackson at approximately 1.5 m intervals (these two sets of samples represent nearly a complete vertical transect through the mound at Tierra Blanca-Nunn levels); (c) 103 thin sections from Alamogordo mounds (88 from 3 × 10 cm core plugs drilled at 1.5m intervals along vertical transects on nine mounds in Indian Wells and Marble Canyons and in Deadman Branch of Alamo Canyon, and 15 large sections ( 7 . 5 × 1 5 c m ) made from surface samples of only the Alamogordo phase of Muleshoe Mound). Thin sections were studied at magnifications ranging from 25 to 100 diameters, depending on the size and complexity of the constituent being observed. The presence or absence of mound constituents was noted and a visual estimate of relative abundance in each thin section was made, always at 25 diameters magnification. The abundance of constituents in a thin section was ranked on a scale ranging from 1 for most abundant to 4 if simply 'present'; ranks of 1 to 3 are considered most reliable and significant.

Lane, Sandberg & Ziegler 1980

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Data

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Mounds The analysis of the Alamogordo Member is based on data from measured sections at Alamo Peak, Indian Wells Canyon, Marble Canyon,

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86

W. M. AHR & R. J. STANTON JR.

Deadman Branch of Alamo Canyon, and San Andres, Dog and Deadman Canyons (Fig. 1). Lithology, sedimentary structures and macrofossils were described and samples were collected from each major bed at each section. Sample spacing is irregular because bed thickness ranges from less than 10cm to over 1 m, and multiple samples were taken from beds thicker than 1 m in order to determine intrabed lithologic variations. Large (7.5 × 15cm) thin sections were prepared from 83 representative samples. The thin sections were studied in the same way as those from the mounds, except that information about depositional texture, presence or absence of cavities, geometry of cavities, geopetal fabrics, degree of skeletal-grain fragmentation, and grain orientation was also available because of the larger size of the thin sections. The level-bottom strata of the Nunn and Tierra Blanca Members were not studied because they are not continuous over the entire study area and because lateral facies and thickness changes obscure the fundamental aspects of the ramp at mound initiation and during mound growth.

Mound lithology In the earliest careful examination of the mounds, Pray (1958) described them as bioherms with a core facies ofaphanitic and sparry calcite within a meshwork of intact and partially comminuted fenestrate bryozoans, grading laterally and abruptly into a flank facies of coarse crinoidal debris. The lime mud ('aphanitic calcite') in the core facies was interpreted to be primary and autochthonous, probably derived from disintegrated algal sheaths or mats. The accumulation of sediment through the baffling and trapping effects of fenestrate bryozoans and crinoids was also considered. This is a hypothesis that was important in the past (Wilson 1975) but is not now in favour. Pray (1958) also noted from the geometric relationship between core and flank facies at Muleshoe Mound that water depth may have been 100 m or more during mound growth, anticipating the conclusion by Lees & Miller (1985) that Waulsortian mounds commonly grew in water depths below the photic zone. The sparry calcite cement of the mound core facies occurs as several types and generations, and patchy submarine cementation may have stabilized the surface during mound growth (Schaefer 1976; Cowan 1980). Radiaxial-fibrous calcite is commonly associated with stromatactis and other cavities, as well as with large fenestrate sheets.

Schaefer (1976) and Cowan (1980) proposed that many of the stromatactis cavities were formed through erosion of sediment below hardground patches in the mound substrate. Cement makes up a considerable fraction of the upper part of the Nunn and Tierra Blanca mounds. For example, it has been estimated to comprise nearly 90% of the core drilled through the upper 21 m of Muleshoe Mound (Shinn et al. 1983). The detailed nature of mound growth has been studied primarily at Muleshoe Mound. There, Jackson & DeKeyser (1984) proposed that during cyclical and shallowing-upward mound growth, localized growth centres ('point-sources') shifted with time on the mound surface, and recent workers (Hunt et al. 1994; Kirkby & Hunt this volume) have recognized five separate and distinct episodes of mound growth. A broader survey of the mounds along the outcrop belt indicates that mound growth was not uniform from mound to mound, and was spatially and temporally complex within the mounds. In general, Alamogordo mounds are relatively broad, low domical to tabular, less than 10 m thick, and include composites of small masses only 2 - 3 m in diameter and thickness. Nunn mounds, in contrast, may be limited in lateral extent, steep-sided and conical because of both greater upward growth potential and subsequent failure and erosion of the mound flanks. Tierra Blanca growth phases are commonly much broader because of both aggradational and progradational growth from the nucleus of the preceding Nunn mound. Combined thickness of the Nunn/Tierra Blanca mound phases increases southward, and is nearly 100 m in Muleshoe Mound. Alamogordo mounds consist of: peloidal, clotted and poorly bedded lime mud; poorly sorted, poorly aligned skeletal allochems, including lenticular concentrations of skeletal debris; and radiaxial-fibrous and other forms of calcite cement as early cavity-lining and intergranular cement, and as later diagenetic pore-filling. Cavities present in the mound rock are of diverse origin, but are largely synsedimentary, on the basis of geopetal fabrics consisting of peloids, multiple generations of internal micrite, and skeletal debris. Most of the peloids are uniform in size and appearance and are particularly common as geopetal fillings in shelter voids and large constructed cavities. The peloids are similar to the microbial peloids described by Chafetz (1986). Pickard (1993, this volume) illustrated nearly identical peloids from Visban mounds in Scotland and Ireland and argues for a microbial

C O M P A R I S O N OF R A M P A N D M O U N D C H A R A C T E R I S T I C S origin. Most of the skeletal allochems are f r a g m e n t e d with the notable exception of corals, ostracodes, and spicules. M o u n d s of N u n n and Tierra Blanca age also contain peloidal m u d s t o n e a n d wackestone, but they differ from A l a m o g o r d o m o u n d s in that: (a) primary cavities are larger and more numerous; (b) submarine cements and fenestrate sheets are m u c h m o r e a b u n d a n t ; (c) micrite and spicules are less abundant; and (d) skeletal allochems grade from poorly oriented to increasingly r a n d o m l y oriented, and lenses of crinoidal debris are m u c h less a b u n d a n t .

Mound biota Biotic constituents in thin sections from the m o u n d s are listed in Table 1. The percentage

87

values in each c o l u m n represent the percent of thin sections from that m o u n d phase in which the constituent is present. Crinoids, fenestrate bryozoans (recorded as both fenestrate sheets and fenestrate hash), and ostracodes were the most ubiquitous and a b u n d a n t organisms on the m o u n d s . Regular echinoids, recognizable by their distinctive spines, were the other macrofossil that was c o m m o n on the m o u n d s , but m o r e so during the A l a m o g o r d o than during the N u n n / T i e r r a Blanca phase of m o u n d growth. Other macroorganisms were u n c o m m o n to rare, a n d m a n y differed significantly in a b u n d a n c e on the two groups of m o u n d s . A m o n g the micro-organisms, the foraminifer Earlandia and the probable red alga Mametella were particularly a b u n d a n t and widespread. Other micro-organisms were m u c h less c o m m o n , and m a n y were

Table 1. Frequency of occurrence (%) and rank of 26 key biotic constituents in mound thin sections Constituent

Crinoid plates and spines All ostracodes Fenestrate hash Sponge spicules Echinoid spines Mollusc and brachiopod shells Trilobite fragments Hyalosteliid spicules Mametella Earlandia Fenestrate sheets Ramose bryozoans Calcispheres Sphaerinvia Moravamminids Gastropods Corals Encrusting bryozoans Geopetal peloids Plurilocular forams Micritized grains Peloids Filaments, undifferentiated Cryptalgal coatings Globochaetes Girvanella Number of thin sections

Level-bottom beds of Alamogordo

Mounds of Alamogordo age

Muleshoe Mound, Nunn & Tierra Blanca phases

%

Rank

%

Rank

%

Rank

99 98 89 87 82 64 53 46 31 25 17 16 13 12 11 7 6 5 4 2 2 0 0 0 0 0

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 20 22 22 22 22 22

100 96 100 72 85 46 27 47 75 60 62 43 42 5 35 10 12 18 9 16 1 47 5 4 15 17

1 3 1 6 4 11 15 9 5 8 7 12 13 23 14 21 20 16 22 18 26 9 23 25 19 17

100 97 100 3 58 57 23 14 49 78 86 65 24 5 5 9 7 32 37 9 1 78 29 0 2 40

1 3 1 23 8 9 16 17 10 5 4 7 15 21 21 18 20 13 12 18 25 5 14 26 24 11

83

103

154

The constituents were chosen because they are among the ones listed as key components in the European Waulsortian mounds by Lees & Miller (1985). Each column gives the percentage of thin sections in which the constituents occur, and the rank order of the constituents.

88

W. M. AHR & R. J. STANTON JR.

markedly different in abundance in the two groups of mounds. Differences in the biota of the Alamogordo and Nunn/Tierra Blanca mounds are indicated in Table 2. The most marked differences are: (a) the greater relative abundance in the Nunn and Tierra Blanca mounds of large fenestrate bryozoan sheets, ramose bryozoans, encrusting bryozoans, Girvanella and filaments in association with peloids and, particularly, geopetal peloids; and (b) the greater relative abundance in the Alamogordo mounds of both hyalosteliid and undifferentiated sponges and calcispheres, Mametella, moravamminids, plurilocular foraminifers and Globochaetes. Fenestrate hash is ubiquitous in both sets of thin sections, but its density is greater in those from the Alamogordo mounds.

Lithological and biotic changes during mound growth The temporal gradations in lithological characteristics during mound growth represent fundamental changes in the mechanism of mound growth and in the environment. The low composite structure of the Alamogordo phase mounds, with pods of lime mud and scree of skeletal fragments between the pods, suggests that the mounds had little upward growth potential, and that even though situated on the outer ramp, were subject to bottom currents that winnowed and shifted loose skeletal debris into lows between the component mounds. The subsequent change in constructional geometry, from the low domical form of composite pods to a more vertically oriented

Table 2. Relative abundance of constituents in Alamogordo age mounds v. those in level-bottom beds and the mound phases of Nunn and Tieraa Blanca age at Muleshoe Mound

Crinoid plates and spines All ostracodes Fenestrate hash Sponge spicules Echinoid spines Mollusc and brachiopod shells Trilobite fragments Hyalosteliid spicules Mametella Earlandia Fenestrate sheets Ramose bryozoans Calcispheres Sphaerinvia Moravamminids Gastropods Corals Encrusting bryozoans Geopetal peloids Plurilocular forams Micritized grain Peioids Filament, undifferentiated Cryptalgal coating Globochaetes Girvanella

Level-bottom

Alamogordo mounds

V.

V.

Alamogordo mounds

Nunn/Tierra Blanca phases

= = = > = > > -<< << << << << >> << < < << << << * << << << << <<

= = = >> > < = >> > < < < > = >> = > < << > * < << >> >> <<

>>, <<, values differ by more than 100%; >, <, values differ by more than 20%, but less than 100%; =, values differ by less than 20%; *, values too small to be significant.

COMPARISON OF RAMP AND MOUND CHARACTERISTICS and steep-flanked shape in the Nunn and Tierra Blanca phase, reflects a change in the ability to construct a rigid and high-relief buildup. This transition is coincident with marked increases in peloids, cavities and fissures, and cement. The peloids suggest microbial sediment production and probably binding as well. Cavities and fissures of diverse size and shape indicate early lithification of the mudmound into a rigid structure. The early lithification was probably the result both of the microbial activity and particularly of early cementation. The cavities were in part constructional, in part erosional around patches of lithified mound surface; the fissures were tensional, pull-apart features within the upward-growing and oversteepened mound. This oversteepening is also documented by the slumping that is common on the flanks of large mounds (like Muleshoe) during the later stages of mound growth. The lithological differences were apparently not determined primarily by changes in the biota, because the most common organisms were not sufficiently abundant to have played either a significant sediment-producing or a constructional, bindstone or framestone-building role in mound growth. Biotic changes during mound growth (Table 2) may be interpreted in terms of (a) temporal differences in the organisms available to inhabit the mounds, (b) the environment as it controlled the presence of specific organisms, and (c) how it controlled the taphonomic processes affecting their preservation in the fossil record. Most of the organisms are identified to generic and higher taxonomic levels, and none had a geological range shorter than the time span in which the Alamogordo Member and the mounds were deposited. Consequently, biotic differences are best explained by environmental and taphonomic differences. Substrate differences are of prime importance. The relatively greater abundance of large fenestrate sheets, ramose and encrusting bryozoans, Gir~anella, and filaments in the Nunn/ Tierra Blanca phase of mound growth is best explained by the greater extent of hardground, which provided a better substrate for these attached benthic organisms than the muddier, probably softer, substrate of the Alamogordo mounds. In addition, the ramose and sheet bryozoans, as they grew up from the surface and were encased in early cement, would have enhanced the local relief and contributed to the construction of cavities within the substrate. Regular echinoids were relatively more common in the Alamogordo mounds, which is in accordance with their general occurrence on

89

soft substrates in the Late Palaeozoic. Sponges were a significant contributor to the micrite matrix of the Alamogordo mounds, but were much less important during later mound growth when microbial peloid production was relatively so much more important. Apparently, the mucilaginous microbial surface of the Nunn/ Tierra Blanca mounds was not congenial to sponge settlement and growth, and peloids were not an important by-product of the decomposition of sponges. Bathymetry must be considered as a potentially significant environmental variable in determining the biotic differences during mound growth. As noted above, the interbeds and lenses of coarse ( > 2 m m diameter) packstones and grainstones composed of crinoid and bryozoan hash within the Alamogordo mounds indicate that these mounds were affected to a great extent by bottom currents and possibly by storm-generated waves. Nunn and Tierra Blanca mounds, with their great volume of cementstone which anchored skeletal constituents early in constructional history, could have developed in shallower water than the Alamogordo mounds, as hydrological pumping regularly forced volumes of water through the mainly bryozoan skeletal networks enabling the great amount of cement to precipitate. The presence of mappable discontinuities in the Nunn and Tierra Blanca phases of Muleshoe mound, interpreted to be erosional or non-depositional hiatuses by Hunt et al. (1994) and Kirkby & Hunt (this volume), lends credence to this possibility, although the absence of micritization, ooids, and calcareous green algae is not consistent with a euphotic environment above fair-weather wave-base.

Level-bottom lithology The Alamogordo Member consists of resistant beds of skeletal lime mudstone and wackestone up to 1 m thick, separated by thin argillaceous beds less than 5cm thick. The Member is separated into 14 beds or bed sets that are continuous from north to south throughout the 32km long outcrop belt. The individual beds thin down-ramp, and thus the total thickness of the Member decreases persistently to the south from a maximum of about 11 m in the northern part of the study area to about 3 m in the vicinity of the mounds occurring farthest down dip, and to less than 2 m in the southernmost outcrops. The contacts with the underlying shaly, silty Andrecito Member and the overlying argillaceous lime packstones and grainy wackestones of

90

W. M. AHR & R. J. STANTON JR.

the Nunn Member are mainly gradational. In that part of the ramp in which mounds occur, the micrite matrix is made up largely of siliceous sponge spicules preserved as mouldic calcite replacements. Chert nodules are common in these beds, presumably formed from silica derived from solution of the spicules. Peloids and pore-filling cement are rare; cavities are absent. Depositional fabric is characterized by widely dispersed bioclasts in the lime mud matrix, and by the absence of both geopetal structures and of sedimentary structures indicative of hydraulic action, such as shell concentrations, ripples, flute casts or cross-beds. Consequently, the sediments are considered to be autochthonous. Bioturbation is common, although it is not obvious in most hand specimens.

Level-bottom biota The fossils present in the level-bottom beds are listed in Table 1 and are ranked by their rela'!ve abundance as measured by percentage of ;hin sections in which they occur. The most common organisms that were living on the level-bottom were crinoids, ostracodes, fenestrate bryozoans (preserved as fine fragmental hash), sponges (preserved as dissociated spicules) and regular echinoids. The spicules are a major component of the micrite matrix, indicating that sponges may well have been the most abundant organism present on the sea floor. Of the more common and larger taxa, on!y ostracode valves and small corals generally occur as whole specimens. Multicomponent skeletons such as crinoids are largely disarticulated, and virtually all skeletal elements except small ostracodes, spicules, and durable skeletons such as small corals and crinoid plates are finely comminuted. Evidence of grain solution is absent and bioerosion is rare, consisting only of 4 0 # m diameter borings in some skeletal allochems. Predation and scavenging are interpreted as the dominant mechanisms of skeletal breakage and disarticulation. In addition to the sedimentological evidence presented above, the lack of breakage, abrasion and orientation of individual shells argue for the autochthonous origin of both sediment and fossils. Productivity on the sea floor was low, as indicated by the mudstone to wackestone texture of the level-bottom strata and the sparse distribution of fossils - 13 of the 23 taxa logged in this study occur in less than 25% of the samples.

Comparison of level-bottom and mound lithology and biota The lithology and biota of the level-bottom strata of the Alamogordo Member differ in significant ways from those of the mounds. In a number of respects, the change is a progressive one from the level bottom to the low mounds of the Alamogordo phase of mound growth, and then to the Nunn/Tierra Blanca phase. Perhaps the most important lithological gradation was from a soft, spiculitic micrite on the level sea floor to a sediment on the mounds that was decreasingly spiculitic and increasingly peloidal, clotted, and lithified by microbial activity and cement. Cavities were essentially absent in the level-bottom sediment, but were increasingly abundant within the mounds. Also associated with these primary cavities, stromatactis, commonly containing bundles of hyalosteliid spicules, were increasingly common during mound growth. The level-bottom strata are well-bedded in contrast to the mound rock, but sedimentary structures within the beds, such as the shell concentrations present within the Alamogordo phase mounds, are absent. The overall abundance of organisms was low in both level-bottom and mound settings, as indicated by the mudstone to wackestone lithology. However, it was approximately 50% greater on the mounds than on the level-bottom. This difference is expressed by greater diversity on the mounds (23 constituents) than on the level-bottom (19 constituents), and by a 33% lower density of fossils in level-bottom thin sections than in mound thin sections. The relative abundance of organisms on the sea floor or on the mounds during the different phases of mound growth can be estimated from their frequency of occurrence within the thin sections. In all settings, a relatively small number of constituents were widely distributed, and many constituents were rare. A slightly greater proportion of level-bottom than mound constituents were cosmopolitan (22% v. 17% occurring in more than 80% of thin sections in the respective sets). To explain these differences, sedimentation rate, environmental heterogeneity and skeletal production must all be considered. Fossils are volumetrically more abundant in the mounds than in the level-bottom beds, even though the sediment accumulation rate was an order of magnitude greater on the mounds. Thus, even though the fossils are a minor component of the mounds in volumetric terms, production was

COMPARISON OF RAMP AND MOUND CHARACTERISTICS much greater there than on the level-bottom. The slightly greater proportion of cosmopolitan organisms on the level-bottom is explained by the environmental uniformity of the levelbottom. During the slow accumulation of sediment, any of the organisms present might have been found throughout the area, even though density at any time would have been low. Thus the relatively slow sedimentation rate and the effect of bioturbation, which mixed the slowly accumulating sediments, produced a strongly homogenized time-averaged record. The abundance of some organisms changed progressively from the level-bottom through the phases of mound growth. For example, crinoids, ostracodes, and fenestrate bryozoans were the dominant and ubiquitous organisms on both the level-bottom and the mounds, but were more abundant on the mounds. The sponge population also decreased from the level-bottom through the successive phases of mound growth. The trend during mound growth of increasing abundance of bryozoans of all types, cryptalgal coatings, Girvanella and filaments began on the level bottom, with the lower abundance or even absence of these constituents there, reflecting the progression from a soft, lime-mud substrate to an increasingly firm and cemented substrate. Concentrations of hyalosteliid spicules and the ostracode Kirkbya are largely restricted to cavities in the later mound phase, suggesting that cavities were their preferred life site (Coen et al. 1988). The organisms of potential bathymetric significance-Mametella, calcispheres, Globochaetes, cryptalgal coatings and Girvanellaincrease in abundance from the level-bottom to the Alamogordo mounds, and then decrease in abundance during later mound growth. The effect of preservational or taphonomic processes must be considered, as well as habitat preferences, in analysing the differences in biota between the level-bottom and contemporaneous Alamogordo mounds and between the mound subsets. The effect of taphonomy is particularly evident in the relative abundances of fenestrate bryozoan sheets and hash in mound and level-bottom samples. Fenestrate hash is ubiquitous and indicative that fenestrate bryozoans lived in both mound and level-bottom settings. The much greater abundance of fenestrate sheets in the mounds is attributed to better preservation there because of more rapid burial (sedimentation rate was an order of magnitude greater than on the levelbottom), substrate cementation which held the sheets in fixed positions, less predation and scavenging by crustaceans (as suggested by the

91

lower frequency of occurrence of trilobites on the mounds), and less bioturbation. In general, however, taphonomic processes appear to have been similar in the mound and level-bottom settings. Grain solution, erosion of skeletal components and mechanical abrasion are absent in both settings, and individuals of specific microfossils are equally well preserved in the different settings. Consequently, differences in assemblages are interpreted largely in terms of habitat differences.

Comparison of mounds and level-bottom beds in the Sacremento Mountains with analogues in Europe Lower Mississippian mounds in the Sacramento Mountains differ in tectonosedimentary setting, depositional environment and biota from upper Mississippian (Asbian and Brigantian Stages of Britain and Ireland) buildups in Texas, Utah, and Derbyshire, UK. Upper Mississippian mounds commonly occur on rimmed or drop-off shelf margins and have shallow-shelf grainstones and packstones with taxonomically diverse skeletal biota as level-bottom equivalents (Ahr 1989a). Lower Mississippian buildups, on the other hand, are commonly found on distal parts of ramps in deeper-water settings (Ahr 1989b), and they are not associated with slope-breaks, shallow-water biota, or grainstone/packstone level-bottom beds with taxonomically diverse skeletal allochems. The Sacramento Mountain mounds and levelbottom beds are comparable in age and depositional setting to the European Waulsortian as described by Lees & Miller (1995) in Britain, Ireland, and Belgium, but it is difficult to make a detailed comparison for the following reasons: (a) data on the European buildups are limited to the mounds, rather than mounds plus contiguous level-bottom beds, because of limited outcrop exposures in Europe; (b) information on mound constituents is presented in the form of relays or relay indices rather than as primary data; (c) the methods of data processing have varied in different studies, especially in correspondence analyses where the operator must select components without following a prescribed selection routine.

Age and environmental setting European Waulsortian and Sacramento Mountain mounds and level-bottom beds occur on

92

W. M. AHR & R. J. STANTON JR.

distal parts of ramps and are Tournaisian to early Vis~an in age. Many European buildups are interpreted to have begun developing at water depths below the photic zone, which is defined according to the presence/absence of fossil phototrophs and micritization (Lees & Miller 1985; Miller 1986). In this context, the shallowest Waulsortian phase, Phase D, indicates an origin within the photic zone, and in some cases, above fair-weather wave-base (autochthonous ooids). Fossil green algae and micritization are rare to absent in the Sacramento Mountain mounds and sediments, and Phase D consitituents are essentially absent from all mounds. The developmental history of the Nunn and Tierra Blanca mounds in the Sacramento Mountains is punctuated by hiatuses, but mounds in the Alamogordo Member represent comparatively continuous deposition. Nunn-Tierra Blanca mound lithology and biota may change across hiatuses, especially in the large mounds like Muleshoe and Sugarloaf. This pattern of interrupted growth is commonly attended by phase inversions or omissions, a characteristic that either is not common and important, or has not been recognized, in the European Waulsortian sections. Component assemblages (phases) in Alamogordo mounds do not regularly occur in predictable patterns, such as Phase A at mound base and Phase C at mound top. A depthdependent sequence of biotic assemblages does, however, exist in Alamogordo Member levelbottom beds (Jeffery & Stanton this volume), and may represent the closest parallel in New Mexico to the phase array described from European Waulsortian mounds.

Lithology and biota European and New Mexican buildup lithologies are similar, especially the Alamogordo Memberequivalent mounds and the European Waulsortian buildups. Nunn and Tierra Blanca-equivalent mounds appear to consist of much greater volumes of cement than their European counterparts, and they also appear to be marked by more numerous and much larger cavities, fissures and neptunian dykes than the European buildups. However, similar features are recorded from the Waulsortian buildups in the Dublin Basin, Ireland (Somerville et al. 1992). The full significance of these internal cavities for the biotic make-up of the New Mexican mounds is not clear. Peloidal muds and 'polymuds' (Lees & Miller 1995) are volumetrically more abundant in the European buildups. Alamogordo Member

mounds contain much more poorly-bedded micrite as neomorphic microspar, and NunnTierra Blanca mounds tend to be cementstone buildups, with peloidal and polymuds confined to cavity fillings. European and New Mexican buildups, along with the Sacramento Mountains level-bottom beds, contain mainly the same taxa; however, the New Mexico mound constituents do not occur in regular and continuous phases from mound base to top. As previously mentioned, the phase arrays commonly exhibit omissions, inversions, or repetition of phases on vertical transects through mounds of all ages. To test for parallelism between New Mexican and European constituent arrays, several data processing methods were employed.

Data processing Petrographical data from the New Mexican mounds were first processed with an early version of the Jaccard similarity coefficient program developed by Hennebert & Lees (1985), to seek 'relays' within Alamogordo mounds, younger mounds, and all mounds. Relays consist of assemblages in which some constituents occur in overlapping vertical ranges from one assemblage to the next, perhaps indicating a gradual shift in constituent composition which, in turn, represents a gradient in the palaeoenvironment. The results were inconclusive at all stratigraphic levels because the calculated matrix of similarity coefficients did not reveal relays. Constituent assemblages from samples ranging from base to top of all mounds in the Alamogordo, and from the Nunn-Tierra Blanca-equivalent portions of Muleshoe Mound, were grouped into phases based on key constituent types identified in the Waulsortian buildups of Europe by Lees & Miller (1985). This also proved to be unsatisfactory because the Phase A to D progression of Lees & Miller (1985) could not be found in any of the Alamogordo or other mounds in the Sacramento Mountains. Instead, samples typically exhibited phase inversions, such as Phase C directly to Phase A, or Phase B to Phase A, or vertical omissions such as Phase A directly to Phase C, even in Alamogordo samples taken from mounds without apparent discontinuities. The irregularities in phases exist throughout the mounds rather than on a sampleby-sample basis, and inversions may be interlayered with omissions, or vice-versa, suggesting that the absence of ordering in the vertical progression of phases is not due to a missing

COMPARISON OF RAMP AND MOUND CHARACTERISTICS section. There is, however, a general tendency for more Phase C constituents to occur in the upper parts of all mounds. Correspondence analysis (Hennebert & Lees 1991) is used by Lees & Miller (1995) to produce a graphical array of constituents or samples such that the array is construed to represent an environmental gradient in a manner that parallels the relays of Phases A through D in European Waulsortian buildups. Relay indices are constructed to indicate the parallelism (Lees & Miller 1995). The Hennebert & Lees (1991) correspondence analysis program was run on data from Alamogordo mounds, Nunn-Tierra Blanca mounds separately, and on data from all the Sacramento Mountain mounds combined. Correspondence analysis runs were made with New Mexican data as subsets compared with an earlier 1988 dataset on European mounds. Regardless of which constituents were chosen for program runs of New Mexico-only data, Alamogordo as well as Nunn-Tierra Blanca mound constituent groups did not produce a correspondence 'arch' like those illustrated by Lees & Miller (1995). Instead, when New Mexico data were run in correspondence analysis for comparison against the European dataset of 1988, the New Mexico results produced a point cluster that fits in the central portion of the European array ('arch'). The significance of this result is not clear because of these equivocal results. Finally, unranked presence-absence data were run in the database program 'Paradox for Windows' to search for patterns of occurrence in constituent composition, but the results did not reveal consistent patterns. Therefore, the analysis presented in this paper is based only on primary, non-ranked, presence-absence data. Constituents from the level-bottom beds studied by Jeffery & Stanton (this volume) occur in assemblages similar to the phases in European Waulsortian mounds described by Lees & Miller (1995). It appears that environmental gradients existed along the Sacramento Mountain ramp and that they influenced the biotic composition of the Alamogordo Member level-bottom sediments. These gradients probably represent continuous change along the sloping sea floor, leading to the conclusion that the parallel patterns seen in the European Waulsortian phases represent gradients- probably of depth, light, and agitation, from inception to termination of buildup growth. The absence of regular progressions in component phases in Alamogordo mounds suggests that the small mounds did not develop enough bathymetric relief in deep water (below storm

93

wave-base and possibly below the photic zone) to have been influenced by environmental gradients. The absence of Phase D components, the phase inversions and omissions, the hiatuses in mound growth, along with the massive production of fenestrate bryozoan cementstones in the Tierra Blanca equivalent mounds, indicate that the larger mounds developed in a vigorous and variable hydrological regime, but one that may not have been above fair-weather wave-base. Such hydrological activity was probably related to thermohaline, geostrophic, or storm currents.

Conclusions The Alamogordo Member of the Lake Valley Formation (Lower Mississippian-Tournaisian to early Vis~an) was deposited on a homoclinal ramp at depths below fair-weather wave-base and mainly below the photic zone. Depth increased toward the present southwest, and biotic assemblages that parallel the depth gradient have been identified (Jeffery & Stanton this volume). Waulsortian-like mounds are abundant in the Alamogordo, Nunn and Tierra Blanca Members of the Lake Valley Formation on the distal portion of the homoclinal ramp. The mounds in all three Members contain essentially the same biota as the Waulsortian buildups of Europe. Alamogordo mounds are low, tabular features composed largely of spiculiferous micrite; Nunn and Tierra Blanca mounds are more domical and are composed largely of fenestrate bryozoan cementstone. Comparison of the Alamogordo mound biota with Alamogordo level-bottom biota revealed that the mound biota functioned as dwellers rather than constructors, bafflers, or sediment producers. Mound biota was recruited in part from the level-bottom, but the two differ in both taxonomic composition and abundance because of differences in habitats (cavities and hard substrates in and on mounds), and because of differences in preservation due to variations in taphonomy, an example of which is the preservation of fenestrate bryozoan sheets in radiaxial fibrous calcite cements on mounds as compared with the comminuted and dispersed fenestrate hash in level-bottom muds. Lake Valley mound constituents, although essentially identical to those in the European Waulsortian, do not occur in regular growth 'phases' like those described by Lees & Miller (1985). The difference is interpreted to mean that the New Mexican mounds lacked enough vertical relief

94

W. M. A H R & R. J. S T A N T O N JR.

a b o v e the sea-floor in deep water to have been affected by a n y regional e n v i r o n m e n t a l gradients t h a t m a y have existed o n the p l a t f o r m . We thank G. Webb and M. Connolly for their constructive comments on the original manuscript, and A. Lees and P. Bridges for their helpful reviews. We are indebted to G. Shinn and W. D. Jackson for lending us thin sections from Muleshoe Mound. The petrographic study of mound thin sections was performed by W. M. Ahr as part of a Fulbright Research Scholarship to Belgium. The generous help of A. Lees and the late R. Conil as well as the Fulbright Program and CIES are gratefully acknowledged. Field study was supported in part by funds from the Ray C. Fish Professorship.

R e f e r e n c e s

AHR, W. M. 1989a. Sedimentary and tectonic controls on the development of an Early Mississippian carbonate ramp, Sacramento Mountains area, New Mexico. In: CREVELLO, P. D., WILSON, J. L., SARG, J. F. & READ, J. F. (eds) Controls on Carbonate Platform and Basin Development. Society of Economic Paleontologists and Mineralogists Special Publication, 44, 203-212. - - 1 9 8 9 b . Comparative sedimentology of Waulsortian reefs at Waulsort, Belgium and Alamogordo, New Mexico. Geological Society of America Abstracts with Programs, 21, A292. & STANTON, R. J. JR. 1994. Comparative sedimentology and paleontology of Waulsortian mounds and coeval level-bottom sediments, lower Lake Valley Formation (lower Mississippian), Sacramento Mountains, New Mexico. Verhandlungen der geologischen Bundesanstalt, Wien, 50, 11-24.

ARMSTRONG, A. K. 1962. Stratigraphy and paleontology of the Mississippian System in southwestern New Mexico and adjacent southeastern Arizona. New Mexico Bureau of Mines and Mineral Resources Memoir, 8, 1-99. CHAFETZ, H. S. 1986. Marine peloids: a product of bacterially induced precipitation of calcite. Journal of Sedimentary Petrology, 56, 812-817. COEN, M., MICHIELS, D. & PARISSE, E. 1988. Ostracodes dinantiens de I'Ardenne. M~moires de I'institut GOologique de I'UniversitO de Louvain, 34, 1-20. COWAN, P. E. 1980. Diagenesis of line mud, Mississippian-age bioherms, Sacremento Mountains, New Mexico. Masters Thesis, Stoney Brook, State University of New York. HENNEBERT, M. & LEES, A. 1985. Optimized similarity matrices applied to the study of carbonate rocks. Geological Journal, 20, 123-131. - & 1991. Environmental gradients in carbonate sediments and rocks detected by correspondence analysis: examples from the Recent of Norway and the Dinantian of southwest England. Sedimentology, 38, 623-642.

HUNT, D., KIRKBY, K., SIMO, J. A., & VANDEN BERGH, T. C. V. 1994. Growth phases, internal surfaces and reservoir compartmentalization in Waulsortian buildups: Muleshoe Mound, Lower Carboniferous, New Mexico. Abstracts, American Association of Petroleum Geologists Annual Meeting, Denver, Colorado, 177. JACKSON, W. D. & DEKEYSER, T. 1984. Microfacies analysis of Muleshoe mound (Early Mississippian), Sacramento Mountains, New Mexico: a point-source depositional model. Parts I and II. West Texas Geological Society Bulletin, 23, 6-10. JEFFERY, D. L. & STANTON, R. J. JR 1996. Biotic gradients on a homoclinal ramp: The Alamogordo Member of the Lake Valley Formation, Lower Mississippian, New Mexico, USA. This volume. KIRKBY, K. C. & HUNT, D. 1996. Episodic growth of a Waulsortian buildup: The Lower Carboniferous Muleshoe Mound, New Mexico, USA. This volume. KOTTLOWSKI, F. E. 1963. Paleozoic and Mesozoic strata of southwestern and south-central New Mexico. New Mexico Bureau of Mines and Mineral Resources Bulletin 79, 1-100. - - 1 9 6 5 . Sedimentary basins of south central New Mexico. American Association of Petroleum Geologists Bulletin, 49, 2120-2139. LANE, H. R. 1974. Mississippian of southeastern New Mexico and West Texas: A wedge-on-wedge relation. American Association of Petroleum Geologists Bulletin, 58, 269-282. - & ORMISTON, A. R. 1982. Waulsortian facies, Sacramento Mountains, New Mexico: Guide for an international field seminar. In: BOLTON, K., LANE, H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting of the Waulsortian facies: El Paso, Texas. E1 Paso Geological Society, 115-182. , SANDBERG, C. A. & ZIEGLER, W. 1980. Taxonomy and phylogeny of some Lower Carboniferous conodonts and preliminary standard post-Siphonodella Zonation. Geologica et Paleontologia, 14, 117-164. LAUDON, L. R. & BOWSHER, A. L. 1941. Mississippian formations of the Sacramento Mountains, New Mexico. American Association of Petroleum Geologists Bulletin, 25, 2107-2160. - & 1949. Mississippian formations of southwestern New Mexico. Geological Society of America Bulletin, 60, 1-88. LEES, A., 1988. Wauisortian 'reefs': The history of a concept. M~moires de l'institut GOologique de l'Universit~ de Louvain, 34, 43-55. & MILLER, J. 1985. Facies variations in Waulsortian buildups, 2: Mid-Dinantian buildups from Europe and North America. Geological Journal, 20, 159-180. -& -1995. Waulsortian banks. In: MONT¥,C., BOSENCE, D. W. J., BRIDGES, P. H. & PRATT, B. (eds) Carbonate Mud-Mounds: their Origin and Evolution. International Association of Sedimentologists Special Publication, 23, 191-271.

C O M P A R I S O N OF R A M P A N D M O U N D C H A R A C T E R I S T I C S MEYERS, W. J. 1974. Carbonate cement stratigraphy of the Lake Valley Formation (Mississippian), Sacramento Mountains, New Mexico. Journal of Sedimentary Petrology, 44, 837-861. MILLER, J. 1986. Facies relationships and diagenesis in Waulsortian mudmounds from the Lower Carboniferous of Ireland and N. England. In: SCHROEDER, J. H. & PURSER, B. (eds) Reef Diagenesis. Springer-Verlag, Berlin, 311-334. PICKARD, N. A. H. 1993. Depositional controls on Lower Carboniferous microbial buildups, eastern Midland Valley of Scotland. Sedimentology, 39, 1081-1100. - - 1 9 9 6 . Evidence for microbial influence on the development of Lower Carboniferous buildups. This volume. PRAY, L. C. 1958. Fenestrate bryozoan core facies, Mississippian bioherms, southwestern United States. Journal of Sedimentary Petrology, 28, 261-273. - - 1 9 6 1 . Geology of the Sacramento Mountains escarpment, Otero County, New Mexico. New Mexico Bureau of Mines and Mineral Resources Bulletin, 35, 1-144.

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SCHAEFER, P. J. 1976. Microfacies and Cementation History of a Mississippian Mud Mound. MSc Thesis, Stony Brook, State University of New York. SHINN, E. A., ROBBIN, D., LIDZ, B. & HUDSON, J. H. 1983. Influence of deposition and early diagenesis on porosity and chemical compaction in two Paleozoic buildups: Mississippian and Permian age rocks in the Sacramento Mountains, New Mexico. In: HARRIS, P. M. (ed.) Carbonate Buildups, a Core Workshop. Society of Economic Paleontologists and Mineralogists Core Workshop No. 4, 182-223. SOMERVILLE, I. D., STROGEN, P. & JONES, G. LL. 1992. Mid-Dinantian Waulsortian buildups in the Dublin Basin, Ireland. Sedimentary Geology, 79, 91-116. , & SOMERVILLE, H. E. m. 1996. Late Vis6an buildups at Kingscourt, Ireland: possible precursors for Upper Carboniferous bioherms. This volume. WEBB, G. E. 1994. Non-Waulsortian Mississippian bioherms: a comparative analysis. Canadian Society of Petroleum Geologists Memoir, 17, 701 712. WILSON, J. L. 1975. Carbonate Facies in Geologic History. Springer-Verlag, New York.

Episodic growth of a Waulsortian buildup: the Lower Carboniferous Muleshoe Mound, Sacramento Mountains, New Mexico, USA KENT

C. K I R K B Y 1 & D A V E

HUNT 2

1Department of Geology and Geophysics, University of Wisconsin-Madison, 1215 West Dayton Street, Madison, Wisconsin 53706, USA Present address." Department of Geology and Geophysics, 108 Pillsbury Drive S.E., Minneapolis, Minnesota 55455-0219, USA 2 Department of Earth Sciences, University of Manchester, Oxford Road, Manchester M I 3 9PL, UK

Abstract: Muleshoe Mound is a composite Waulsortian buildup that crops out along the western escarpment of the Sacramento Mountains, New Mexico. Exceptional exposures of Muleshoe Mound have enabled a detailed stratigraphic model to be developed, subdividing the buildup into five separate stratal units. Each stratal unit differs in facies, geometry and symmetry, reflecting changes in the character and intensity of different environmental conditions (energy, carbonate production, oxygenation and accommodation space). Hiatal surfaces mark periods of mound crisis and/or sea-level fluctuations. These hiatal surfaces highlight the episodic nature of the mound's growth. Episodic growth acted as an important control on this buildup's development, facies and early diagenesis. Features and patterns of the Muleshoe Mound conflict with the traditional interpretation of the Waulsortian as quiet, stable, deep-water buildups. Muleshoe Mound appears to have grown in appreciable currents where accommodation space was a crucial control on growth. Previous studies of the Lower Carboniferous Waulsortian facies and related strata have concentrated on their fauna and detailed textural fabric (e.g. Lees et al. 1985; Lees & Miller 1985, 1995; Ahr & Stanton this volume; Jeffery & Stanton this volume). This study focuses on the field study of stratal patterns and geometry of the well-known Muleshoe Mound, Sacramento Mountains, New Mexico (Fig. 1). This led to a recognition that Waulsortian buildups can be divided into a succession of separate stratal units. Each stratal unit is characterized by a distinctive geometry, facies association and distribution. Field and diagenetic studies of the surfaces that separate stratal units indicate that the mounds do not form a depositional continuum. Rather, each unconformity-bounded stratal unit represents a separate recolonization of antecedent bathymetry after a period of hiatus and erosion, to form a composite Waulsortian buildup. We suggest that episodic growth is a characteristic of Waulsortian buildups and acted as an important control of their facies, diagenesis and ultimately reservoir compartmentalization (Kirkby 1994).

Geological setting Muleshoe Mound and its correlative flanking strata form part of a Lower Carboniferous

(Tournaisian-Vis6an) carbonate ramp succession which is represented by the Lake Valley Formation (Fig. 2). Buildups within these fossiliferous strata were recognized and described by Laudon & Bowsher (1941). They interpreted the buildups to have developed in a relatively low-energy environment, but under the influence of southwards directed currents to account for the asymmetry they observed. The 'grainy' and spar-rich nature of the buildups in the Sacramento Mountains is well known (Pray 1958; Shinn et al. 1983). These cementstones and the asymmetry so characteristic of Muleshoe's 'upper' mound (Fig. 2b) suggests growth in a current-influenced, relatively high-energy environment, conflicting with traditional views of the Waulsortian as a deep-water low-energy facies (e.g. 200-300m, Lees et al. 1985). Four members of the Lake Valley Formation correlate with Waulsortian buildups; the Alamogordo, Nunn, Tierra Blanca and Table Top Members (Lane et al. 1982; Fig. 2). Like many of the Lake Valley buildups, Muleshoe Mound was initiated during deposition of the Alamogordo Member (Fig. 2) (see Ahr & Stanton this volume; Jeffery & Stanton this volume). The location and development of buildups within the Alamogordo is a reflection of antecedent depositional and compactiongenerated topography, as well as more

From STROGEN, P., SOMERVILLE, I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 97 110.

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Fig. 1. Map illustrating the extent of Mississippian (Andrecito-Tierra Blanca) exposures in the Sacramento Mountains. Superimposed is the interpreted extent of the Tierra Blanca 'lobe' on the upper ramp. The eastern margin of this lobe was controlled by an elongate chain of Waulsortian buildups developed along a zone of subtle tectonic uplift (Hunt 1994). The rose diagrams summarize the orientation of elongate crinoid stems and clinoform orientations from the Mississippian. The rose diagrams are a representative selection of over 2500 readings (from Hunt & Allsop 1993).

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localized tectonically generated relief (Ahr 1989; Hunt 1994; Hunt e t al. 1995). In the north of the platform, a northerly trending and elongate chain of composite Waulsortian mounds (of Alamogordo-Tierra Blanca age) formed a rampart that acted to focus currents (Hunt & Allsop 1993; Hunt 1994; Hunt e t al. 1994, 1995; Fig. 1). These currents in turn controlled the development of a giant lobe of skeletal sands: the Tierra Blanca 'shelf' of Meyers (1975). Mounds developed within the confines of the Tierra Blanca lobe tend to have a sheet-like form and developed relatively

99

little depositional relief. In this high-energy current-swept area, mounds have relatively high width:height ratios (typically 5: 1-10: 1). These more elongate mounds are generally current-scoured and exhibit strong progradation on their southern flanks (Fig. 3). They are associated with voluminous 'flanking' strata and strong basin-floor aggradation which acted to facilitate lateral progradation of Waulsortian facies. Mound aggradation appears to have been limited by some sort of accommodation control throughout this upper ramp area. The suite of relatively large (80-100m relief) buildups that characterizes the mid-part of the Lake Valley ramp is the focus of interest here. These are developed 3-6 km basinwards (south) of the toe-of-slope of the southerly prograding Tierra Blanca lobe (Fig. 1). This is a setting that is directly comparable to many other Waulsortian mounds (e.g. Belgium-Lees e t al. 1985; SW England-Lees 1982; Ireland-Murphy 1988; Montana-Cotter 1965, Smith 1977; AlbertaKirkby 1994). In the mid-part of the Lake Valley ramp, the five large exposed composite buildups have an average spacing of 1.6km (Fig. 1), and have width to height ratios of 1.5:1 to 4:1. Aggradation of these mounds is similar to that of the time-equivalent Tierra Blanca lobe to the north (< 100 m), and suggests a regional control on accommodation and aggradation; we speculate that this may represent wave base.

Muleshoe Mound Muleshoe Mound is 100m high, with a width of 400-500m. Recent erosion has exhumed threequarters of the buildup, resulting in a close-tocentre exposure through the buildup (SW face; Fig. 4a). Two sections expose the transition between the Waulsortian facies of the buildup's 'core' facies and its flanking facies on the shelfward (NW) and basinward (SE) sides of the buildup (Fig. 4b, c). The mound has previously been informally divided into lower and upper mound intervals on the basis of associated flank strata (Pray 1975; Figs 2b, 5b). Here, we further subdivide the upper mound interval into four stratal units (Msu II-V; Figs 4 & 5). Muleshoe's five stratal units (Msu I-V) were initially recognized and distinguished by variations in core geometry and flank stratal patterns. Stratal truncations occur at the top of each stratal package, and beds of the succeeding package downlap onto these truncation surfaces. Flank strata of successive units also exhibit marked changes in bedding orientation.

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Fig. 3. Two sections through the composite Deadman Waulsortian buildup, Deadman Branch, Alamo Canyon, near Alamogordo (see Fig. 1). This buildup, like Muleshoe Mound, is divisible into five stratal units (here Dmsu l-V), and these five stratal units are thought to be correlative. (a) The southern flank of D e a d m a n Mound is characterized by spectacular progradation of Dmsu Ill IV. Progradation of flanking strata on this accretionary southern (basinwards) side of the buildup facilitated coeval lateral progradation of Waulsortian facies. (b) A strike-section through this buildup, in contrast, reveals that the contemporaneous southwestern side of the buildup was a largely static and aggradational bypass type margin at the same time that the southern boundary was prograding (Dmsu Ill-IV). The boundary between stratal units Dmsu III and IV is a prominent onlap surface in strike-section associated with major megabreccia units, but this is a more subtle boundary in dip-section (a). The base of Dmsu V is locally associated with strong erosional truncation (a) and passes laterally to a sharp-based unit associated with an abrupt increase of grainsize (b), a relationship seen across most of the Tierra Blanca lobe on the upper ramp (line drawings from Hunt & Allsop 1993).

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MULESHOE BUILDUP, NEW MEXICO

101

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Fig. 4. (a) Line drawing of the southwest face of Muleshoe Mound illustrating bedding patterns and division into stratal units. (5) Line drawing of the northwestern face of Muleshoe Mound showing its division into stratal units. Note large channel and megabreccia deposits in the foreground. (c) Line drawing of the spectacular southeast flank of Muleshoe Mound. Megabreccia deposits near the centre of the drawing were described by Pray (1958) and Meyers (1975).

102

K. C. KIRKBY & D. HUNT

i::.

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On the flanks of the mound, the boundaries between the stratal units are sharp, scoured surfaces which are locally overlain by breccias (Figs 4b, c & 6). The distribution of megabreccias is not random, as previously thought, but is associated with hiatal surfaces; they occur at the base and top of stratal units (Fig. 6). The outcrop expression of these surfaces varies from obvious features (such as the surfaces capping units Msu I and II; Fig. 4) to subtle surfaces that can be obscured by surficial weathering. Erosion, downlapping stratal patterns and diagenetic evidence combine to show that these surfaces represent significant hiatal breaks.

cavity-fill may comprise up to 40% of the core volume, and consists of early marine cements, skeletal grainstones, and minor amounts of argillaceous carbonates. Although micritedominated, the core also contains irregular layers and lenses of crinoid packstones and grainstones which are often erosively based and normally graded. In spite of this heterogeneity, there does not appear to be any lateral facies segregation other than the rapid gradation into inter-mound wackestones and mudstones deposited under normal marine conditions (Alamogordo Member).

Hiatal surface. The hiatal surface overlying Muleshoe Mound, stratal unit I (Msu I) Geometry. The oldest mound growth revealed in Muleshoe's exposures is a broad, relatively low relief (35-40m high, 200m wide) mound that comprised Pray's (1975) 'lower mound' (Figs 2, 4a, 6). Facies. Massive micrite-rich core facies grade laterally into off-mound skeletal wackestones and mudstones of the Alamogordo Member with no intervening grain-supported flank facies. Core facies are composed of skeletal wackestones and mudstones riddled by a network of irregular stromatactoid cavities. Stromatactis

Msu I is only preserved in the mound centre as a sharp, irregular and locally erosive surface which separates relatively massive core facies of Msu I from thin-bedded crinoidal grainstones and packstones of Msu II (Fig. 4b). The upper ten metres of relief developed by Msu I is bounded by this erosion surface. Below this, the surface was truncated by erosion that occurred in association with the depositional hiatus at the end of Msu II.

Muleshoe Mound, stratal unit H (Msu II) Geometry. This second stratal unit unconformably succeeds its precursor and consists of an

MULESHOE BUILDUP, NEW MEXICO areally-restricted building-up of the mound (to 40 m high and 120 m across). It is composed of intercalated lenses of massive core facies and bedded flank facies (Figs 4a, 5b). Gently dipping (5°-10 °) beds comprise the base of Msu II, and grade up into more steeply dipping (25 ° to >45 °) strata which are locally deformed (towards the toe-of-slope) in association with the hiatus following the growth of Msu II.

Facies. Core facies in the second stratal unit are similar to those of its precursor, with the addition of abundant neptunian dykes in the younger, steeply dipping, core strata. These dykes, originally described by Pray (1964), are filled by multiple generations of steeply dipping micrites and/or crinoid grainstones. Many neptunian dykes exhibit composite fills that reflect repeated opening and filling (Pray 1964). Msu II exhibits no lateral segregation of core and flank strata. Flank strata intercalate and encase lenses of micrite-rich core facies. These flank strata are dominantly crinoidal grainstones and packstones. They are similar to crinoidal flank strata in the overlying stratal units (Msu III-V), but contain isopachous, (early) marine cements, which are relatively rare in the later, off-mound flank strata. Hiatal surface. Msu II and Msu III are separated by a sharp surface which makes a pronounced recessive break in the outcrop face (Fig. 4a). This surface erodes the top of Msu II, the flanks of Msu I and II, and has also removed the youngest inter-mound Alamogordo strata up to 150 m from the buildup, over which a fairly extensive sheet-like breccia unit is found (e.g. Figs 4b, 6). Core strata of the third stratal unit (Msu III) downlap onto this eroded surface, and small clasts of Msu I-II core facies are incorporated into the basal flank facies of Msu III.

Muleshoe Mound, stratal unit III (Msu III) Geometry. This third stratal unit marks a significant change in the buildup growth, from aggradational to progradational stratal patterns (Fig. 6). Msu III consists of stacked, massive core facies that, laterally, pass abruptly into thin-bedded flank strata. The contact between massive and bedded facies is strikingly abrupt a subhorizontal boundary in Msu III (Fig. 6), which may represent some sort of current base and/or top of stratified anoxic bottom waters. Mound height increased slightly during Msu III, but most growth was progradational. At the end

103

of Msu III the mound complex formed a roughly symmetrical dome, over l l 0 m high and 300 m across.

Facies. Core strata in the third stratal unit exhibit distinct lateral facies transitions. Central core facies are dominated by cementstones and skeletal grainstones. The cementstones have a framework of large, aligned, in situ fenestrate bryozoans with thick rims of isopachous marine cement and irregular patches of micrite. Skeletal grainstones are composed of crinoid and fenestrate bryozoan debris, with irregular, in situ patches of micrite. Thick massive beds of crinoid-rich packstones with small stromatactis cavities form the core margin and massive proximal flank facies, and grade rapidly into thin-bedded flank facies. Flank facies are only exposed on the southern side of the buildup, where crinoid grainstones and matrix-poor packstones dominate the section.

Hiatal surface. On the well-exposed southern flank of the buildup, these flank strata are truncated by the overlying hiatal surface (Figs 4c, 6). Flanking strata of Msu IV downlap onto this surface and exhibit a 70 ° difference in dip direction from the underlying flanking beds of Msu III (Fig. 7). Erosion also removed correlative Msu III flank strata on the northern side of the buildup, and channels and thick lenses of megabreccia overlie the surface on both sides of the mound (Figs 4b, c, 5b).

Muleshoe Mound, stratal unit I V (Msu IV) Geometry. In contrast to the symmetrical aggradation and progradation of Msu I-III, progradation in Msu IV is strikingly asymmetric (Fig. 5). Preserved core strata are only developed on the southern (basinward) side of Muleshoe Mound, where they form three overstepping lenses which intercalate with, and abruptly pass into, bedded flank strata with a basinward rising contact (Figs 3-5). Correlative strata on the northern flank are entirely composed of bedded flank facies (Figs 4b & 5). Facies. Erosion related to Upper Carboniferous subaerial exposure and recent exhumation of Muleshoe have removed the mound crest strata of this growth phase. Preserved core-centre facies are marine-cemented, fenestrate bryozoan grainstones similar to those of the underlying Msu III core strata. Towards the core margin these are

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Fig. 5. (a) Five component stratal units (Msu I-V) of Muleshoe Mound. (b) The distribution of massive Waulsortian 'core' facies, bedded flank strata and megabreccias in Muleshoe Mound. (e) The asymmetry of flank facies in Msu IV, excluding megabreccias. In Msu IV core facies are only developed on the basinwards (south) side of the buildup. On the shelfward (northern) side, the only preserved correlative strata are bedded flank strata. Grain-supported flank strata dominate the southern side of the buildup, whilst matrix-supported beds are more abundant on the northern side, Pie diagrams record the relative abundance of different depositional textures in three measured sections.

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MULESHOE BUILDUP, NEW MEXICO succeeded by micrite-rich skeletal packstones with a silicified fenestrate bryozoan-rich fauna which includes large cylinder- and vase-shaped growth forms. These core margin strata occur at the upper slope of the mound margin, and overlie and grade into steeply-dipping (up to 35 °) flank strata. They comprise a previously unrecognized, depositionally unstable facies association, which is distinguished by multiple generations of neptunian dykes and an in situ 'brecciated' fabric, with large (up to 7 m) blocks of micriterich core facies that have been rotated (<50 °) towards the buildup flanks. The transition from core-centre to core-margin is best developed in the lower core lens. It is difficult to track the upper two core lenses into the complex mound centre because the exposure slope parallels bedding and is largely covered. Exposed core facies in these upper lenses are dominated by silicified micriterich skeletal packstones and core-margin facies similar to those of the lower lens. Asymmetry in Msu IV is not limited to the core facies, but also occurs in the more volumetrically-important bedded flank strata. Grain-supported strata dominate the southern (basinward) side of the buildup, while matrixsupported beds are more common on the northern flank (Fig. 5c). Scour, graded beds and small intraclasts are also more abundant in the southern flank deposits, indicating greater depositional energy. The combination of flank and core asymmetry probably reflects selective mound growth under the influence of currents and/or wave energy. Spectacular megabreccias occur at the base of Msu IV, mantling and deforming the lower

(a) n=14

NI

(b) \

l

N~

107

buildup slopes and adjacent 'level-bottom' strata (Figs 4a, b, c & 6). These megabreccias are typically composite units, deposited by several individual mass flows. They contain blocks of core and core-margin facies in a crinoid-rich matrix. Those on the south (basinward) side of the buildup were reported by Pray (1958), and are described in detail by Meyers (1975). These deposits consist of overstepping lenses of breccia, composed of large blocks (with lengths up to 5 m) of core facies in a crinoid-rich matrix. A discontinuous exposure (or related lens) of megabreccia truncates the underlying Alamogordo strata to the southeast, including the southern flank of an underlying small buildup (equivalent to the lowermost Muleshoe Mound, Msu I; Figs 4c, 6). On the north side of Muleshoe, a large channel complex with associated megabreccia deposits unconformably overlies and locally truncates the Alamogordo strata (Figs 4b, 6) (Kirkby 1994). The channel and megabreccia deposits are up to 15m thick, and contain blocks of core facies up to 10m long in a crinoid grainstone matrix. Although these deposits cannot be traced around to the mound's south flank (due to recent erosion and faulting), they appear to be roughly correlative to the megabreccia deposits on the mound's south side. To our knowledge, the allochthonous nature of these northern deposits has not been recognized previously.

Hiatal surface. The surface that caps the fourth stratal unit is more subtle than the other hiatal surfaces, as erosion of underlying

(c

n=40 avg:95

Fig. 7. (a & b) Rose diagrams showing dip directions on the southeast side of Muleshoe Mound, in the area of stratigraphic section 1-S (see Fig. 6). (a) Msu III, flank facies and (b) Msu IV, flank facies. Flank strata in stratal units III and IV exhibit a 70° difference in their dip directions. (c) Rose diagram of oriented bryozoans in life position measured across the crest of Muleshoe Mound (Msu III-IV). Whereas we expected a concentric orientation of bryozoan growth forms around the circumference of the buildup, a strong preferred NNW-SSE orientation was measured, closely matching the trends of oriented fauna seen within coeval 'level-bottom' strata. This tends to confirm the important control that the N-S current regime of the platform had on the growth and development of Muleshoe Mound.

108

K. C. KIRKBY & D. HUNT

strata was minimal. A clast-bearing interval was deposited on both sides of the mound, and the succeeding flank strata of Msu V downlap onto the top of this unit.

calcareous shales with a sparse fauna and rare Chondrites. The lap-out of these strata against the topography of both the platform and its buildups indicates a reduction of water energy and oxygenation, related to nutrient upwelling and/or a relative sea-level rise.

Muleshoe Mound, stratal unit V (Msu V) Geometry. The youngest episode of Muleshoe's growth (Msu V) is also an asymmetrically prograding package with a single lens of core facies developed on the southern side of the mound complex (Figs 4c, 6). This lens is downset from the crest of the mound, and is also characterized by a basinwards descending coreflank transition (Figs 4c, 6). Crest deposits were eroded and the northern flank is composed entirely of bedded flank strata (Fig. 4b). Facies. Core facies are similar to the micriterich core and core-margin facies of the upper two lenses of Msu IV. Flank facies are also similar, and Msu IV and Msu V are primarily distinguished on the basis of stratal patterns and differing cement histories.

Hiatal surface. Flank strata of Msu V are either truncated by the overlying surface, or rapidly grade into restricted, laminated strata of the off-mound Table Top Member. Surface exposures at Muleshoe are not conclusive, as the bedded flank strata are less resistant than those of the underlying stratal units. However, Table Top and Arcente Formation strata do intercalate with flank strata in mounds seven kilometres north of Muleshoe, in Alamo Canyon (Hunt & Allsop 1993; Meyers pers. comm. 1993; Hunt 1994). This relationship is important, as it implies the presence of pronounced basin oxygen-stratification. For whilst the buildups are grain-supported, spar-rich, and have abundant fauna, correlative basin-floor facies are dominantly mud-supported and internally laminated. They contain only a sparse fauna and are mostly unbioturbated except for occasional Zoophycus and Chondrites burrows, an association characteristic of oxygen depletion (Ekdale & Mason 1988). The growth of Muleshoe Mound is interpreted to have been terminated by being smothered by a thickening blanket of lowenergy, oxygen-poor bottom waters, as strongly stratified waters impinged on the platform. The 'drowning' of both the platform and its buildups is recorded by strata of the Arcente Formation (Fig. 2). These strata are typically dark (olive black), clean lime mudstones and interbedded

Discussion Muleshoe Mound is a composite Waulsortian buildup which grew episodically rather than continuously. It is divisible into five stratal units on the basis of geometry, facies and stratal terminations. These five stratal units record Muleshoe's intermittent growth in a currentinfluenced environment. Strong asymmetry in the upper two stratal units (Msu IV-V) reflects the strong N-S current regime that influenced the platform's development. This is recorded by the large-scale architecture and distribution of core and flank facies in Msu IV and Msu V, as well as by the alignment of fauna such as bryozoans (within core facies; Fig. 7) and crinoid stems in the flank facies (Fig. 1). Each of the five stratal units differs in terms of its geometry, distribution and symmetry (Fig. 6). The relative proportion of core and flank facies varies between stratal units, suggesting that each developed under differing conditions. Flank facies comprise a significant proportion of the mound growth in Msu IV and V, when production of crinoidal sediment was at an acme. The boundary separating relatively massive core facies (with abundant early marine cements) from well-bedded crinoidal flank facies (with interbedded argillaceous drapes) is abrupt. This boundary is interpreted to represent a pycnocline, separating welloxygenated circulating waters above from relatively still dysaerobic waters below. Bioturbation within the oxygenated water column contributed to the massive character of the Waulsortian facies. The boundary between core and flank facies remained subhorizontal during the deposition of Msu III, but then climbed and descended basinwards during deposition of Msu IV and Msu V, respectively (Fig. 6). The downshift of Msu V core facies from the crest of Muleshoe Mound, and the downwards translation of its core-flank transition, is interpreted to reflect the deposition of these strata during a period of relative sea-level fall. The strata of Msu V are correlated northwards with an erosively-based crinoidal sand package developed across the Tierra Blanca 'shelf' (e.g. Dmsu V, Fig. 3).

MULESHOE BUILDUP, NEW MEXICO Consideration of features found in association with the boundaries of stratal units (such as argillaceous drapes over the mound crest, extensional fracturing of core facies, toe-ofslope deformation and in situ brecciation, slope collapse, loss of diagenetic (cathodoluminescence) zones, silicification, megabreccia deposition, and strong scour and erosional truncation on the buildups flanks) clearly indicates that the surfaces separating the five stratal units represent significant depositional hiatuses. Thus, each of the upper four stratal units of Muleshoe Mound (Msu II-V) represents a discrete recolonization of the antecedent relief. The deposition of laminated argillaceous sediments, in association with the absence of encrusting fauna over the mound crest at hiatal surfaces, suggests that lowenergy, poorly oxygenated environmental conditions prevailed during times of mound crisis. We interpret these depositional hiatuses to be associated with periods when the crest of Muleshoe Mound was smothered by lowenergy, poorly oxygenated (dysaerobic-anaerobic) bottom waters in association with relative sea-level rises and/or nutrient upwelling. Recognition of the episodic growth of Muleshoe Mound is important for several reasons. The mound stratigraphy encodes a wealth of information on basin development that is not evident within the adjacent 'levelbottom' strata. Features and growth patterns of Muleshoe and other buildups in the Sacramento Mountains also conflict with their traditional interpretation as quiet, stable, deep-water buildups. Rather, the mounds appear to have grown in appreciable currents with intermittent high energy, at a depth where accommodation space played a critical role on mound growth. The identification of five stratal units within Waulsortian buildups on the upper ramp (e.g. Fig. 3), thought to be coeval with those of Muleshoe Mound, further suggests that there was a regional control on both accommodation and the development and architecture of Waulsortian buildups in the Sacramento Mountains. This is thought to be the wave base. Episodic growth has also been recognized in other Waulsortian mound suites. In the subsurface hydrocarbon-bearing Waulsortian buildups of west-central Alberta, the hiatal surfaces between growth episodes can act as a fundamental control of reservoir compartmentalization in core and flank strata (Kirkby 1994). The example provided by Muleshoe Mound provides a framework by which less well-exposed or subsurface examples of Waulsortian buildups can be interpreted.

109

Conclusions Muleshoe Mound is a composite buildup which grew episodically. Mound accretion took place during times when the mound crest was influenced by relatively well-oxygenated and high-energy waters. Episodes of mound growth are separated by depositional hiatuses associated with drapes of argillaceous sediments across the mound crest. Periods of mound crisis and depositional hiatus are thought to have occurred during times when a blanket of anoxic bottom waters smothered the mound crest. Such episodic growth is a characteristic of Waulsortian buildups. On the flanks of the buildup, depositional hiatuses are characterized by loss of diagenetic zones, silicification, toe-of-slope deformation (folding and in situ brecciation), and ultimately slope collapse, erosion and megabreccia deposition. The first and second stratal units of Muleshoe Mound are dominantly aggradational, whilst Msu III-V are largely progradational. This change in accretion style is thought to be controlled by upwards growth of the buildups to some sort of upper accommodation limit, interpreted as the wave base. Symmetrical aggradation and progradation is characteristic of Msu I-III, whereas accretion is markedly asymmetric in Msu IV-V. This is interpreted to reflect growth of the mound under the influence of the N-S current regime which characterized the platform. The downstepped position of Msu V and the descending nature of its core-flank transition reflects deposition during a period of relative sea-level fall. This study was sponsored by grants from the American Chemical Societies Petroleum Research Foundation, the American Chemical Society and Union Pacific Resources (to K.C.K.), and the Natural Environment Research Council of Britain (fellowship to D.H.); their support is greatly appreciated and acknowledged. This paper benefited from encouragement and the timely reviews of I. Somerville and P. Strogen, w i t h additional comments from A. E. Adams- thanks. During time in the field this study greatly benefited from the collaboration and comments of T. Simo, L. Pray, C. Van den Burgh, W. Ahr, B. Stanton, D. Jeffery and T. AUsop. We must pass on particular thanks to L. Pray for taking time out of his 'retirement' to be with us in the field, as well for constant questioning and lively debate of the work presented here. His tireless work and enthusiasm for the Mississippian of the Sacramento's has been an inspiration to us.

110

K. C. K I R K B Y & D. H U N T

References AHR, W. M. 1989. Sedimentary and tectonic controls on the development of an early Mississippian carbonate ramp, Sacramento Mountains, New Mexico. In: CREVELLO, P. D., WILSON, J. L., SARG, J. F. & READ, J. F. (eds) Controls on Carbonate Platform and Basin Development. Special Publication, Society of Economic Paleontologists and Mineralogists, 44, 203-212. --& STANTON, R. J. JR. 1996. Constituent composition of Early Mississippian carbonate buildups and their level-bottom equivalents, Sacramento Mountains, New Mexico. This volume. COTTER, E. 1965. Waulsortian-type carbonate banks in the Mississippian Lodgepole Formation of central Montana. Geological Journal 73, 881-888. EKDALE, A. A. & MASON, T. R. 1988. Characteristic trace-fossil associations in oxygen-poor environments. Geology, 16, 720-723. HUNT, D. 1994. Architecture and sequential development of a Mississippian carbonate platform as exemplified by the Sacramento Mountains, New Mexico. Annual Convention Program Abstracts 1994. American Association of Petroleum Geologists, Denver, 177. -& ALLSOP, T. 1993. Mississippian Strata Of The Sacramento Mountains, New Mexico." Sketches of Stratal Patterns, Palaeogeography, and Some Preliminary Interpretations. University of Manchester. , KIRKBY, K. C., ALLSOP, T., VAN DEN BURGH, T. C. V., SIMO, J. A., PRAY, L. C. & SWARBRICK, R. E. 1995. A One-Day Fieldguide To Mississippian Strata of Muleshoe canyon, Sacramento Mountains, New Mexico, USA. University of Manchester. - - , SIMO, J. A. & VAN DEN BERGH, T. C. V. '1994. Growth phases, internal surfaces and reservoir compartmentalization in Waulsortian buildups: Muleshoe Mound, Lower Carboniferous, New Mexico. Annual Convention Program Abstracts 1994. American Association of Petroleum Geologists, Denver, 177. JEFFERY, D. L. & STANTON, R. J. JR. 1996. Biotic gradients on a homoclinal ramp: The Alamogordo Member of the Lake Valley Formation, Lower Mississippian, New Mexico, USA. This volume. KIRKBY, K. C. 1994. Origin and Evolution of Waulsortian Mound Systems: Pekisko Formation, West Central Alberta, and Lake Valley Formation, New Mexico. PhD Thesis, University of MadisonWisconsin. LANE, H. R., ORMISTON, A. R., ET AL. 1982. Waulsortian facies, Sacramento Mountains, New Mexico: guide for an international field seminar, March 2-6, 1982. In: BOLTON, K., LANE, H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. El Paso Geological Society and the University of Texas at El Paso, 115-182.

LAUDON, L. R. & BOWSHER, A. L. 1941. Mississippian formations of the Sacramento Mountains, New Mexico. Bulletin of the American Association of Petroleum Geologists, 25, 2107 2160. LEES, A. 1982. The paleoenvironmental setting and distribution of the Waulsortian facies of Belgium and southern Britain. In: BOLTON, K., LANE, H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. E1 Paso Geological Society and the University of Texas at E1 Paso, 1-16. & MILLER, J. 1985. Facies variations in Waulsortian buildups. Part 2. Mid-Oinantian buildups from Europe and North America. Geological Journal, 20, 159-180. -& 1995. Waulsortian banks. In: MONTY, C., BOSENCE, D. W. J., BRIDGES, P. H. & PRATT, B. (eds) Carbonate Mud-Mounds: their Origin and Evolution. Special Publication, International Association of Sedimentologists, 23, 191-271. - - , HALLET, V. & HIBO, D. 1985. Facies variation in Waulsortian buildups. Part 1. A model from Belgium. Geological Journal 20, 133-158. MEYERS, W. J. 1975. Stratigraphy and diagenesis of the Lake Valley Formation, Sacramento Mountains, New Mexico. In: PRAY, L. C. (ed.) A Guidebook to the Mississippian Shelf Edge and Basin Facies Carbonates, Sacramento Mountains and Southern New Mexico Region. Dallas Geological Society, 45-65. MURPHY, F. X. 1988. Facies variations within the Waulsortian Limestone Formation of the Dungarvan Syncline, southern Ireland. Proceedings of the Geologists Association, 99, 205 219. PRAY, L. C. 1958. Fenestrate bryozoan core facies, Mississippian bioherms, southwestern United States. Journal of Sedimentary Petrology, 28, 261-273. 1964. Limestone clastic dikes in Mississippian bioherms, New Mexico. Geological Society of America Special Paper, 82, 154. 1975. Muleshoe and San Andres biohems, supplemental notes and road log. In: PRAY, L. C. (ed.) A Guidebook to the Mississippian Shelf Edge and Basin Facies Carbonates, Sacramento Mountains and Southern New Mexico Region. Dallas Geological Society, 129-134. SHINN, E. A., ROBBIN, D. M., LIDZ, B. H. & HUDSON, J. H. 1983. Influence of deposition and early diagenesis on porosity and chemical compaction in two Paleozoic buildups: Mississippian and Permian age rocks in the Sacramento Mountains, New Mexico. In: HARRIS, P. M. (ed.) Carbonate Buildups, a Core Workshop. Society of Economic Paleontologists and Mineralogists Core Workshop, 4, 182-222. SMITH, D. t . 1977. Transition from deep to shallow water carbonates, Paine member, Lodgepole Formation, central Montana. In: COOK, H. (ed.) Deep Water Carbonate Environments. Special Publication, Society of Economic Paleontologists and Mineralogists, 25, 187-201.

Biotic gradients on a homoclinal ramp: the Alamogordo Member of the Lake Valley Formation, Lower Mississippian, New Mexico, USA DAVID

L.

JEFFERY & R O B E R T J. S T A N T O N JR.

Dept. o f Geology & Geophysics, Texas A & M University, College Station, Texas T X 77843-3115, USA Abstract: The Lower Mississippian Alamogordo Member of the Lake Valley Formation in south-central New Mexico represents a well exposed carbonate-dominated homoclinal ramp with little evidence of sediment transport. Biotic components present and evident in thin section have distinct distributions in the level-bottom facies in relation to their shoreward and basinward ramp positions. Assemblages (I-IV) are established along the depth gradient from shallow to deep water, and are useful in palaeobathymetric interpretation; they are similar in composition to the depth-related phases recognized in Dinantian mounds by Lees & Miller (1985). The biotic assemblages shift laterally in successive beds of the Alamogordo Member, indicating variations in sea level during deposition. As a result, deepening and successive shallowing during a transgressive-regressive cycle can be tracked within the Alamogordo Member. The ramp slope is interpreted to have been of the order of 0.5°, based on the maximum depths for green and red algae having been l l0m and 250m respectively. The systematic positions of several problematical taxa (Asphaltinella, Sphaerinvia, stacheiins and salebrids) are discussed.

Bathymetry is one of the most basic parameters of any depositional setting, but it is difficult to determine in the analysis of Palaeozoic carbonate strata. Sedimentological depth criteria are often difficult to interpret because of diagenesis, limitations of outcrop, and differences between modern and fossil sediment producers. Texture is commonly correlated with depth, in that coarser-grained and better-sorted sediments are more likely to be produced and deposited in shallower water, but coarse biogenic sediment can be produced and accumulated in deep, quiet water, and fine sediment can accumulate in shallow but low-energy settings. Sedimentary structures such as graded event beds and bioturbation are poorly correlated with depth, and criteria that depend upon recognition of normal and storm wave-base are poorly constrained. Although sequence stratigraphy elucidates general trends of relative sea-level change between cycles, it has not contributed new and quantitative bathymetric criteria. Palaeontological bathymetric criteria have received relatively little attention in recent years compared with sedimentological criteria. In general, we are unable to estimate more than a few of the parameters of the complex and multidimensional palaeoenvironment. The work of Lees and associates (e.g. Lees 1982; Miller & Grayson 1982; Lees et al. 1985; Lees & Miller 1985; Hennebert & Lees 1991) on Lower Carboniferous carbonates, however, has provided a model for the palaeoenvironmental

interpretation of biotic and lithological components in terms, primarily, of bathymetry. Depthrelated, assemblages recognized by Lees et al. (1985), on the basis of the biotic distribution patterns in the Waulsortian Furfooz mound in southern Belgium, integrate a range of palaeontological and sedimentological criteria within a well-established biostratigraphical framework. These assemblages were used by Lees & Miller (1985) to define phases, with the same depth significance, which have been applied by other workers to determine the developmental history of other Waulsortian mounds (e.g. Bridges & Chapman 1988; Ahr 1989a). The statistical procedures employed by Lees & Miller (1985) for Mid-Dinantian mounds, and by Hennebert & Lees (1991) for a Dinantian ramp, to order the component grains into sequences that are interpreted as being from deepest to shallowest water depths, provide another interpretive tool in addition to the basic data of assemblage distributions and phases. The objective of this paper is to examine the distributions of biotic components in sediments deposited on a homoclinal ramp (Ahr 1973) in the Sacramento Mountains of south-central New Mexico. The superb continuous outcrops and the tight stratigraphic control on age relationships of samples provide the opportunity (a) to establish independent, level-bottom, depth-correlated biotic distribution patterns and assemblages, and (b) to compare these with the biotic distributions and depth-related phases

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 111-126.

112

D. L. JEFFERY & R. J. STANTON JR.

and relays described by Lees and co-workers from Dinantian ramps and Waulsortian mounds. The analysis incorporates all those fossils present and evident in thin section. This database corresponds closely to that which Lees and associates have found to be of interpretive value in their depth-related phases and relays. The biostratigraphical and spatial control for their analyses and conclusions was strongest for the Waulsortian mound at Furfooz, Belgium; that for the ramp example in southwest England (Hennebert & Lees 1991) is much less constrained. Thus, comparison of our results with theirs will be largely in relation to their mound results. We have applied their statistical procedure to our data in order to generate relays, but have not included these here because the potential added insight was not sufficient considering the complexity and subjectivity of the method. Two questions arise in a comparison of ramp and mound biotic distributions. The first is that, although the components present in the ramp strata are also present in the associated mounds, differences in relative abundances and even presence or absence indicate that environmental and community characteristics differed between mound and level-bottom (Ahr & Stanton 1994; Ahr & Stanton this volume). Thus, it may be that the depth criteria described by Lees & Miller (1985) are uniquely applicable to mounds, and provide only a comparative data-base for non-mound settings. The second question is the extent and detail to which the components and phases in the mounds are depth-diagnostic. The distributions of many of the components in the Belgian prototype example at Furfooz closely follow time lines, cross facies boundaries, and have broad bathymetric ranges from mound crest to off-mound. For example, the first appearance of moravamminids as 'locally important' in the Furfooz mound is essentially contemporaneous and is across a bathymetric range of about 130 m from mound crest to the base of the mound flank (Lees et al. 1985, fig. 7). This suggests that the distributions of individual components were more strongly controlled by the sequence of developmental events during mound growth than by bathymetry. The distribution of phases also appears to be determined more strongly by stratigraphy than by bathymetry. For example, terminal deposition of Phase C in Furfooz mound was essentially contemporaneous on mound crest and flank, across a depth range of about 80m (Lees & Miller 1985, Fig. 7); this was

in contrast to the total depth range of Phase C deposition of about 30m (Lees & Miller 1985, Fig. 1).

Setting The Lower Carboniferous Caballero and Lake Valley Formations in the Sacramento Mountains of south-central New Mexico lie disconformably on Devonian strata (Fig. 1). They represent several deep-water onlap-ofllap sequences on a southward-dipping, homoclinal outer ramp. The Alamogordo Member of the Lake Valley Formation was deposited during the maximum flooding stage of one of these sequences. It consists primarily of thick-bedded, poorly sorted, cherty mudstones to packstones. Wackestone is the dominant lithology. There is little or no evidence of grading or other sedimentary structures indicative of sediment transport; bioturbation within individual beds is pervasive. Rare grainstones are present on the flanks of the contemporaneous Waulsortian mounds which were

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ALAMOGORDO RAMP BIOTIC GRADIENTS scattered on the ramp• Waulsortian phases have been described in mounds on this ramp that are in part correlative with these strata, although the shallowest phase (D) is not present within them (Lees & Miller 1985; Ahr 1989a; Ahr & Stanton 1994)• The general geology of the Sacramento Mountains and the stratigraphy of these units are described in Pray (1961), Ahr (1989b) and Bolton et al. (1982)• Details of the stratigraphy and sedimentology of the Alamogordo Member will be presented in a later paper• The Alamogordo Member conformably overlies the progradational Andrecito Member, except at several localities where upper Andrecito beds are gently folded and truncated, probably as a result of local penecontemporaneous small-scale deformation. The Andrecito Member consists of interbedded thin strata of limestone and argillaceous marl. The Alamogordo Member is overlain conformably by the progradational Nunn Member, consisting of interbedded limestone and marl. The lithologies of all these units are locally modified in the

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proximity of the Waulsortian mounds, most of which began to grow during Alamogordo deposition, and a few during latest Andrecito deposition. The Alamogordo Member lies within Conodont Faunal Unit 3, the Lower typicus Zone (Fig. 1; Lane 1982). It was deposited off the southwesternmost shore of the Transcontinental Arch, at low latitude (Fig. 2; Ahr 1989a). The Alamogordo Member is exposed in nearly continuous outcrop on the western escarpment of the Sacramento Mountains for a distance of approximately 32 km (Fig. 3). This represents a north-south, down-dip transect on a homoclinal ramp (Ahr 1989b). The Alamogordo Member thins from 10.5m in the north to 1.5m in the south. It consists of 14 beds or bed sets which are continuous and correlatable along the entire transect; beds are separated by argillaceous partings, and thin gradually from north to south (Figs 4, 5; Jeffery & Stanton 1993). Two beds (1 and 2) represent the transition from the slightly more argillaceous, grainy, and recessive

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Andrecito lithology to the darker lime mud and wackestone Alamogordo lithology. They are separated by a nodular, recessive interbed. Then come two thick, massive, resistant, double beds (3-4, 5-6), which are the most prominent beds at any location. These are followed by four thinner, massive, resistant beds (7-10). These

beds are less resistant but still recognizable in the southernmost exposures, where the interval weathers to a rubbly slope. These are followed in turn by: a recessive, argillaceous interval (11); a massive, resistant bed set (12) which amalgamates to form only one layer to the south; a second recessive, argillaceous interval (13); and a massive, resistant three-bed set (14) in the northern part of the transect, which amalgamates southward to two and then to one bed.

Methods Twenty measured sections were sampled and correlated along the outcrop ramp transect (Fig. 3). Field observation along nearly the entire outcrop, as it was walked out to trace beds between sections, added an additional 60 control points. Nearly 200 large (7.5× 15cm) thin sections were examined for fossil and rock components. The fossils were ranked according to their volumetric abundance, independent of the amount of matrix present; the ranking thus expresses only the relationship between fossil

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components, as in Lees & Miller (1985). Categories of relative abundance are: present (less than 2%); low (2 to 10%); intermediate (10 to 20%); and high (20% and greater). Where several components are of high abundance, their ranking is noted as first (being highest), second, and so on. Twenty components were used in the analysis: crinoid/echinoid plates, echinoid spines, fenestellid bryozoan sheets, fenestellid hash, stick/ramose bryozoans, encrusting bryozoans, brachiopods, corals, gastropods, molluscs other than gastropods, ostracodes, trilobite fragments, simple foraminifers (Earlandia), plurilocular foraminifers, sponge spicules (primarily hexactinellid), Sphaerinvia and Asphaltinella (probable worm tubes), salebrids, stacheiins (Mametella?), and green algae. These comprise all the body fossils present in these strata and evident in thin section. They are similar to those used by Lees and co-workers in their studies of Dinantian mound and ramp strata (Lees et al. 1985; Lees & Miller 1985; Hennebert & Lees 1991). They differ, however, in several respects, as will be discussed later. Rock type was characterized by percentage of

mud/spiculitic matrix between grains, and classified using Dunham's (1992) textural scheme. Thin sections in which spicules represent more than 50% of the matrix are classified as spiculites (Fig. 6K).

Components

The dominant fossil components are hexactinellid sponge spicules, crinoid/echinoderm debris, fenestellid bryozoan hash, and ostracodes. The most abundant component in the samples is generally either sponge spicules or crinoid/ echinoderm debris. In most thin sections, fenestellid bryozoan hash is second or third in abundance; ostracodes (occasionally), brachiopods or stacheiins (rarely) are third in abundance. Fossil components that have widespread and scattered distributions throughout the ramp transect include crinoid/echinoderm debris, fenestellid bryozoan hash, ostracodes, brachiopods, Sphaerinvia, trilobite fragments, Asphalt# nella, Earlandia, molluscs (non-gastropod), gastropods, corals, fenestellid sheet bryozoans,

116

D. L. JEFFERY & R. J. STANTON JR.

Fig. 6. Photomicrographs of fossil components within the Alamogordo Member. Scale bars for a, e, and j are 125 #m. All others are 250/zm. (a-c) Sphaerinvia. (a) well preserved with geopetal sediment; note the circular structures around the parietal ridges. (b) specimen with two open chambers. (c) elongate specimen. (d-e) salebrids. (f-h) stacheiins. (i-j) chlorophyte thalli. (k) hexactinellid-rich spiculitic wackestone. (I-m) Asphaltinella.

stick/ramose and encrusting bryozoans, and echinoid spines. The distribution and relative abundance of major components on the ramp is shown in Fig. 7. In samples from the northern half of the study area, sponge spicules (most notably hexactinellid) are generally most abundant, followed closely by crinoid/echinoderm debris. In most beds, sponge spicules are the dominant component from Lead Canyon northward (Fig. 3). The down-ramp limit

of abundant sponge spicules has been noted in each bed (letter H in Fig. 8). In several beds, the southernmost samples contain high abundances of sponge spicules, so the southern limit of abundance is not marked in Fig. 8 (beds 7 and 11). This reflects the wide ecological tolerance of sponges (Bergquist 1978), and emphasizes their unreliability as depth indicators. Other components distributed primarily in the shallower part of the ramp include salebrids,

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Fig. 7. Distribution of components on the Alamogordo ramp. Line thickness indicates relative abundance; dashed line indicates rare occurrences. The components indicated in bold are organisms upon which the Alamogordo assemblages are based. stacheiins, foraminifers and green algae. Stacheiins (possible red algae) occur in nearly every sample northward from San Andres Canyon. Salebrids are widely distributed from San Andres Canyon northward, but are rare, occurring as only one or a few specimens per slide. Plurilocular foraminifers (Endospiroplectammina venusta?) are present only in samples from the northernmost sections. Green algae are only present as fragments (Fig. 6i,j), in samples from the youngest bed (Fig. 8, bed 14). Green algae do not extend as far down-slope as do plurilocular foraminifers. In samples from the southern half of the study area, crinoid/echinoderm debris is generally most abundant, followed closely by fenestellid bryozoan hash; ostracodes are dominant in a few samples, co-occurring with a high proportion of Earlandia. Sponge spicules are much less abundant than in samples from the up-dip part of the ramp (Fig. 8). Other common components include corals, trilobites, molluscs, ramose bryozoans, encrusting bryozoans and lithistid(?) sponges.

Systematic positions and palaeoecological interpretations of problematical organisms Several of the organisms that were common on the Alamogordo Ramp (Sphaerinvia, Asphalt# nella, stacheiins and salebrids) have uncertain

affinities at the phylum or even kingdom level. Sphaerinv& (Fig. 6a, b, c) has been described as a hollow sphere, but specimens in the Alamogordo Member are commonly elongate and sometimes branching. It consists of a doublewalled hollow tube less than 1 mm in diameter. The outer layer, of tangential calcite crystals, is relatively thick. The inner layer is dark, thinner, and has a polygonal inner cross section. At each corner of the polygonal inner wall, parietal spikes radiate out through the outer wall. In specimens that are very well preserved, circular nodes surround each radiating spike• This fabric of elements of the inner wall radiating through the outer wall suggests that the organism may have lived on the outside of the tube, first laying down the inner layer, then depositing the layer between the radiating spikes of the inner layer. Sphaerinvia has been considered to be a charophyte and a marine calcareous alga (Mamet 1991)• However, its distribution independent of depth on the ramp and its presence in deep, probably aphotic sediments, together with a lack of transported sediments on the Alamogordo Ramp, demonstrate that it is not a landderived charophyte. This precludes an algal affinity• Additionally, its microstructure is considered too complex to be a sponge (A. Pisera pers. comm. 1994). Possible affinities with the Cnidaria (possibly heterocorals, as suggested by Vachard 1980), worm tubes, or perhaps some kind of reproductive cyst, among others, remain

118

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alga, following Mamet (1991). Its restricted distribution on the Alamogordo Ramp, correlated with depth and thus incident light, supports the relationship with red algae. Modern red algae may live down to 250m in clear waters (Flfigel 1982). The low-energy, sediment-starved, slow deposition of the Alamogordo Member suggests that the water was not turbid, and that light penetration was probably relatively deep. Thus, this value would be a reasonable estimate of the maximum depth of its occurrence during deposition of the Alamogordo Member. Salebrids (Fig. 6d, e) are small, < l m m clusters of dark, thin-walled tubes that generally contain a central tube surrounded by nonpartitioned tubes and vesicles (Bogush & Brenckle 1982). Salebrids are extremely rare in the Alamogordo Member, limited to sediments in the up-ramp portion of the study area. They are generally much more abundant in thin sections from Waulsortian mounds, and their distribution may be dependent upon proximity to mounds. Their systematic position is unclear (Bogush & Brenckle 1982), but it has been suggested, on the basis of wall structure, that they are allied to the Bryozoa (B.Mamet pers. comm.).

ALAMOGORDO RAMP BIOTIC GRADIENTS

Assemblages The major components have diverse distribution patterns on the ramp, but can nevertheless be grouped into discrete assemblages and characterized by a limited number of components (Fig. 7).

Assemblage I (deepest water). Crinoid/echinoderm debris+fenestellid hash+ostracodes. Most of the fossils in Assemblage I are ubiquitous throughout the ramp. Some samples are dominated by large, thick-shelled ostracodes plus numerous Earlandia; others contain only crinoid/echinoderm debris. Brachiopods, Sphaerinvia, trilobite fragments, Asphaltinella, molluscs, gastropods, corals, fenestellid sheet, stick/ramose, and encrusting bryozoans, and sponge spicules are present sporadically. Assemblage H. Assemblage I + stacheiins, rare salebrids and commonly abundant sponge spicules. Stacheiins are the main diagnostic component. Sponge spicules, although the most abundant component except in a few samples (where they are displaced in abundance by crinoid/echinoderm debris or fenestellid hash) are not diagnostic because they are so widely distributed in all the assemblages and are cosmopolitan on the ramp. Although most sponge spicules are preserved as calcite, may were originally siliceous and went through a mouldic stage during which they were destroyed or distorted by compaction (Meyers 1977). Chert beds and nodules within the Alamogordo Member coincide with the high proportion of spicules (Fig. 6k), and with the strata containing Assemblage II. The spicules were probably the source of the chert silica (Meyers 1975). Fenestellid sheets are uncommon in our samples, but are more abundant in Assemblage II than in the other assemblages.

Assemblage III. Assemblage II + plurilocular foraminifers.

Assemblage IV (shallowest). Assemblage III + green algae. This assemblage, indicative of a photic environment and approaching wave-base, is poorly represented in the Alamogordo Member. Assemblage IV, although based on the occurrence of only a few fragments of green algae, provides a useful reference point for bathymetric analysis.

119

Distribution of assemblages during ramp deposition The distributions of assemblages vary considerably from bed to bed within the Alamogordo Member (Fig. 8). The components of Assemblage I are ubiquitous on the ramp, but the assemblage itself, as defined by the absence of the components diagnostic and characteristic of the other assemblages, is confined to the southern, down-dip part of the ramp. The most upramp occurrences of Assemblage I are at Deadman Branch of Alamo Canyon in bed 8, and at Mule Canyon in beds 10 and 12. Assemblage II is not present south of San Andres Canyon as based on the distributions of stacheiins and salebrids, the diagnostic components. In addition, sponge spicules, which are particularly abundant in Assemblage II, are much less common in the deeper part of the ramp. Assemblage III is restricted to the up-dip part of the ramp. The down-dip limit is in the vicinity of Goat Springs and Marble Canyon in beds 2, 3, and 14; the assemblage is limited to Indian Wells Canyon in beds 4, 5, 6, 11 and 13, and is absent in the study area in beds 7, 8, 9, 10, and 12. Plurilocular foraminifers, the diagnostic fossil, are generally rare and consequently have patchy distributions within the areas delineated as representing Assemblage III occurrence. Assemblage IV only occurs in bed 14, in Indian Wells Canyon. The general trend in distribution of the assemblages during deposition of the Alamogordo Member is an up-ramp shift from the base to a high in beds 7-10, with a maximum in bed 8; a down-ramp shift in bed 11; another upramp shift in bed 12; and then a down-ramp trend to the top of the Member, with the sole occurrence of Assemblage IV in bed 14, in the most up-ramp part of the study area.

Bathymetry of assemblages The bathymetric significance of the assemblages in the Alamogordo Member is based on the deepest occurrences of the red algal stacheiins (Mamet 1991) in Assemblage II, of plurilocular foraminifers in Assemblage III, and of green algae in Assemblage IV. The algal occurrences are of prime importance; two other criteria (micritization, and sponge spicules) are valuable, but lead to less quantitative, more relative bathymetric inferences. Micritization has proven to be a useful indicator of relative

120

D. L. JEFFERY & R. J. STANTON JR.

depth in modern sediments (Swinchatt 1969), but is a poor absolute depth indicator because it occurs in aphotic waters at depths as great as 871m (Hook et al. 1984). Consequently, and because it is absent from our samples, we have not used it in this study. The significance of hexactinellid sponge spicules is problematic because of the wide ecological tolerances of living representatives of this group. Modern glass sponges, for example, live at depths of several metres to thousands of metres, and appear to be primarily indicative of very low energy (low nutrient) environments (Tabachnick 1991). On the other hand, in the geological record, spicules are preferentially preserved in deep-water sediments. This broad bathymetric range exists in the Alamogordo Member, judging from the presence of spicules across the full ramp transect. Bathymetric interpretation of the assemblages is best based on the lower depth limit of calcareous green algae (which depends upon red wavelengths of light) at about l l0m in modern oceans (James & Ginsburg 1979), and on the lower depth limit of calcareous red algae (dependent on blue wavelengths of light) at about 250 m (Flfigel 1982). It has been suggested that Mississippian plurilocular foraminifers range to a depth of about 140m (Gutschick & Sandberg 1983). This depth, being slightly greater than the estimate for green algae, agrees with the distribution of plurilocular foraminifers extending a short distance farther down-ramp than green algae. However, it must be kept in mind that the algal depth limits are maximum values. These estimates yield a maximum depth of about l l0m for the transition between Assemblages III and IV, and a maximum depth of about 250m for the transition between Assemblages I and II. The maximum depth of the transition between Assemblages II and III cannot be constrained better than the range of 110 m to 250 m on the basis of the algal evidence. The horizontal distance between the downramp limits of Assemblage IV and Assemblage II is 15km in bed 14. If the bathymetric difference was approximately 140 m, the dip of the ramp would have been on the order of 0.5 °. The distance between the down-ramp limit of Assemblages IV and II is unknown, but apparently greater in most of the other beds than in bed 14 (Fig. 8). Thus, our estimate of 0.5 ° is a maximum value. Using this estimate, the total depth range on the ramp as exposed in the study area would have been about 300m, with the shallowest portion at l l 0 m and the

deepest portion at 410m. Using this same dip, and a depth at the up-dip end of the transect during deposition of bed 14 of 110 m, and somewhat deeper during the rest of the deposition of the Alamogordo Member, the shallow end of the transect would have been on the order of 10km to 15km from shore. This distance is a minimum since there is generally a change in slope, or 'slope crest', formed due to the sigmoidal geometry of depositional units (Burchette & Wright 1992). It is probable that the Alamogordo outer ramp sediments were deposited on the seaward side of the ramp crest, so ramp dip probably decreased shoreward. The distance to shore would have been greatest at the time of maximum transgression, and least during the final, progradational part of the Alamogordo deposition. These estimates of distance to shore are fraught with assumptions, but provide a starting point for further refinement of an important palaeoenvironmental parameter.

Bathymetric changes during deposition of the Alamogordo Member Bathymetric changes during a net gain (transgressive), and subsequent loss of accommodation (highstand to lowstand?) during deposition of the Alamogordo Member can be estimated from the shifting of assemblages in the successive beds. The maximum landward, up-ramp, positions of the assemblages (during deposition of beds 8 and 12) represent maximum flooding surfaces. The lateral shifts in the assemblages do not form a single, smooth landward-thenseaward migration of facies, but several minor transgressive-regressive cycles. Beds 1-7 represent late transgressive systems tract deposition, with bed 8 representing maximum flooding. Highstand systems tract progradation of the offlapping, downlapping wedge probably began updip while the study area remained starved. The basinward progression of the progradational wedge is reflected in beds 13 and 14 and the overlying Nunn Member. It is assumed that falling sea level was associated with the shallowing represented by beds 13 and 14 because the minor thicknesses of these beds could not account for the amount of shoaling represented by the shifting assemblages. In our model of deposition on the Alamogordo Ramp (Fig. 9), the depth-related assemblages occur as microfacies bands perpendicular to depositional dip. Mud-mound topography is portrayed to result in local occurrences of shallower assemblages. Fluctuations in relative

ALAMOGORDO RAMP BIOTIC GRADIENTS sea-level (due to eustacy, uplift, subsidence) shift the assemblage bands shoreward and basinward, and result in mud-mound drowning or shallowing. These fluctuations could account for some of the complex successions of depth-related characters as noted by Ahr (1989a) and Ahr & Stanton (1994, this volume).

Discussion

Comparison of Alamogordo Member assemblages with those in Waulsortian phases The assemblages recognized in the Alamogordo Member differ from those in the Waulsortian phases established by Lees & Miller (1985; 1995) in several respects (it should be noted that their data are a composite relay of many data sets compiled through correspondence analysis, whereas ours represent raw data). The analysis of the Waulsortian mounds by Lees & Miller (1985) incorporates both biotic and lithological components; that of the Alamogordo Member, only biotic components. Of the lithological components in the Waulsortian mounds, ooids, micritized components and coated grains are not present within the Alamogordo Member. This is probably because an environment of adequate energy to form ooids or coated grains was not present in the study area, and such sediments were not transported down-ramp into the study area. Intraclasts are present in the flank beds of mud-mounds, but all of our samples were taken

I

121

as far from any mud-mound as possible, and intraclasts do not occur. Differences in the biotic components in the Alamogordo Member and in the Waulsortian mounds are in part attributed to differences in nomenclature. Asphaltinella in the Alamogordo Member is probably synonymous with serpulid worm tubes of Lees & Miller (1985). Stacheiins correspond to Lees & Miller's (1985) aoujgaliids. Several components (moravamminids and filaments) are present in European samples (Lees et al. 1985; Lees & Miller 1985; Fig. 10), but are absent from the Alamogordo Ramp samples (Fig. 7). Moravamminids are primarily European, Lower Carboniferous organisms (Vachard 1991) and may not be present in western North American faunas. The alga shown in Fig. 6i, however, may be allied to the Beresellidae, which is considered to be closely related to, or belonging to, the order Moravamminida (Vachard 1991; Lees & Miller 1995). Filaments are absent in the Alamogordo levelbottom biota, but do occur in the Alamogordo mud-mounds of equivalent age (Ahr & Stanton this volume). Filaments are also not reported from the European ramp sediments in which the grain components have been studied in detail (Hennebert & Lees 1991). This may indicate that the organism that formed them is restricted to mud-mound environments. Sphaerinvia and salebrids were not reported in the Waulsortian phases of Lees & Miller (1985), but subsequently they have been reported as present but rare (Lees & Miller 1995). Sphaerinvia is widely distributed in the level-bottom sediments of the Alamogordo Member. In

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122

D. L. JEFFERY & R. J. STANTON JR. in the occurrence of salebrids and stacheiins (aoujgaliids of Lees & Miller 1985), the components upon which Assemblage II is based. Another difference is the presence of filaments and cyanophytes in Phase B (Lees & Miller 1995). Their interpretation of Phase B being at photic depths concurs with our interpretation of Mametella as a stacheiin alga occurring at photic depths, rather than as an aoujgaliid sponge. Assemblage III and Phase C are similar in being based on the presence of plurilocular foraminifers. They differ mainly because aoujgaliids and gastropods first appear near the base of Phase C (in the upper portion of Phase B), whereas gastropods are present in all Alamogordo assemblages, and stacheiins are present in Assemblage II. Assemblage IV and Phase D are similar due to the presence of calcareous green algae. They differ in that coated grains, micritization and ooids are absent in the Alamogordo Member.

contrast, in the European ramp strata described by Hennebert & Lees (1991), it has a very low relay index, perhaps indicating restriction to deeper water. Salebrids are rare in the Alamogordo mud-mounds but are more common there than in the level-bottom strata. The assemblages characteristic of Waulsortian phases are similar to those in the Alamogordo Member but with some significant differences. Assemblage I and Phase A are similar in the abundance of crinoid/echinoderm debris and the common occurrence of organisms with wide environmental tolerances, including ostracodes, brachiopods, bryozoans, molluscs, simple foraminifers (Earlandia), and trilobites. They differ significantly, however, in the abundance of fenestellid sheets and sparry cement in Phase A, and the rareness of these components in Assemblage I. This difference is clearly mound v. ramp-related, and exists between the Alamogordo Member and mounds in the Sacramento Mountains as well as in this comparison with the Belgian mounds (Ahr & Stanton 1994, this volume). They also differ because echinoid spines, serpulid tuL~s and gastropods are not present in Phase A, but are present in Assemblage I. Assemblage II and Phase B are similar in the abundance of hexactinellid sponges (inclusive of hyalosteliids). They differ significantly, however,

Comparison of depth interpretations of Alamogordo assemblages and Waulsortian phases The grouping of components into phases and their interpreted depth and light distributions

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ALAMOGORDO RAMP BIOTIC GRADIENTS within Belgian Waulsortian mounds are depicted in Fig. 10 (Lees & Miller 1985, 1995). These phases have proven valuable in the analysis of Waulsortian mounds not only in Europe but in North America (Lees & Miller 1985; Ahr 1989a), and, to the extent that water depth can be determined during mound growth, are useful in estimating magnitudes and fluctuations in subsidence and sea level during sedimentation. The bathymetric interpretation of Waulsortian phases depends on the delineation of the 'photic zone', based on the lowest occurrence of micritization, which is assumed to be caused by algae and to occur as deep as 220 m in modern environments (Fig. 10). The depth values inferred for deeper phases are based on the stratigraphic distance of each phase below the lowest occurrence of micritization. Thus, the lowest occurrence of micritization is the lower limit of Phase D (220 m); the lowest occurrence of plurilocular foraminifers, stratigraphically 30 m below the limit of micritization, is the lower limit of Phase C (250m); and the lowest occurrence of hyalosteliid spicules, 60 m below the limit of micritization, is the lower limit of Phase B (280 m). Uncertainties in this procedure include the extent to which micritization is uniquely caused by algae, and thus is indicative of the photic zone (see Lees et al. 1985), and the possibility that depth values based on sediment thicknesses are in error as a result of compaction, diagenesis, subsidence, uplift, and sea-level fluctuation during deposition. Aoujgaliids in the Waulsortian phases have been considered to be sponges because of their poor association with the < 110 m limit for green algae. If aoujgaliids are actually red algae (stacheiins), however, then they could have extended as much as 140m deeper than the l l 0 m limit, to 250m, which would provide a more dependable depth value than 280 m for the first appearance of hyalosteliid sponges. Although the assemblages in the Alamogordo Ramp and in the Waulsortian mounds are similar, depth inferences for the two entities differ for a number of reasons. Firstly, the criteria and procedures differ in the two studies. Secondly, differences in age and location may significantly affect the biota. Most of the Waulsortian mounds that have been studied previously (Lees & Miller 1985; Lees et al. 1985) are European, and younger than the Alamogordo Member. Whereas the earlier phases of the European mounds are Tournaisian and may be contemporaneous with the Alamogordo Member based on conodont zones, the later,

123

shallower phases in the European mounds range across several conodont zones and are younger than the Alamogordo Member. Furthermore, the different floral and faunal realms to which North America and Europe belonged might help to explain the differences in components (e.g. Mamet 1992; Ross 1981). Closely tied to this point is the fact that first appearances of many components and boundaries between phases in the European mounds closely parallel conodont zones, whereas the Alamogordo Member assemblages within each bed are contemporaneous. Lastly, the samples from the Alamogordo Member are from level-bottom strata rather than from mounds, and thus may reflect significantly different environments (Ahr & Stanton 1994, this volume). For example, several taxa that are typically restricted to the shallower spectrum of Waulsortian mounds (gastropods and sponges) are ubiquitous on the Alamogordo Ramp. Interestingly, Ahr & Stanton (1994) found that, in the Waulsortian mounds of the Sacramento Mountains, spicules are more common in the deeper mound growth phase, and gastropod abundance is not significantly different between level-bottom and mound. This may signify differences in preservation or habitat between European and North American Waulsortian facies. It may also indicate that some factor discouraged or inhibited the habitation of the early stages of European mounds by sponges or gastropods. Hennebert & Lees (1991) described the biotic gradients on a Dinantian ramp in southwest England that contains largely the same components as those recognized in the Waulsortian mounds by Lees et al. (1985) and Lees & Miller (1985). The goal of Hennebert & Lees (1991) was to demonstrate the effectiveness of correspondence analysis in recognizing general trends in grain components within a depositional system, especially where there is a limited number of samples. Their data yield a relay of components correlated with a depth gradient across a ramp. This relay is similar to the distribution of components on the Alamogordo Ramp, but more detailed comparison is not possible because of marked differences in the quality of geographic and temporal control and in the analytical procedures used.

Conclusions The Alamogordo Member of the Lake Valley Formation is exposed in the Sacramento Mountains of south-central New Mexico in

124

D. L. JEFFERY & R. J. STANTON JR.

continuous outcrop for 32 km. Individual beds in the Member extend the full length of the transect, which represents a dip profile of a homoclinal ramp. Criteria for determining palaeobathymetry are established by determining gradients of biotic components and component assemblages on the ramp. Assemblages on the ramp, from deepest to shallowest, are characterized by: I

crinoid/echinoderm debris + fenestellid hash + ostracodes; II Assemblage I+stacheiins, rare salebrids and commonly abundant sponge spicules; III Assemblage II +plurilocular foraminifers; IV Assemblage III + green algae.

Depth limits for these assemblages are based on the occurrences of green algae in Assemblage IV and probable red algae (stacheiins) in Assemblage II, and on the probable maximum depth of occurrence for these algae of l l 0 m and 250m, respectively. As a result, Assemblage IV lived in water less than 110m deep, and Assemblage II in water less than or equal to 250 m deep. Using these depth values and distances along the ramp of assemblage limits, the slope of the Alamogordo Ramp was about 0.5°; the shallowest depth, at the north end of the study area, was about 110m, and the water depth estimate for the ramp-basin floor transition in the southern portion of the study area was about 410 m. Sea-level fluctuations during deposition of the Alamogordo Member are recorded by the lateral shifts in assemblages from bed to bed within the Alamogordo Member. The lower portion of the Alamogordo Member (beds 1-7) represents a transgressive systems tract. The middle portion of the Alamogordo Member (beds 8-10 and 12) represents maximum flooding and the onset of highstand systems tract deposition. The upper portion of the Alamogordo Member (beds 11, 13, 14) represents a highstand systems tract and shallowing. The assemblages on the Alamogordo Member ramp are similar to those in the phases in the Waulsortian mound at Furfooz, Belgium, and depth interpretations are also similar (Lees et al. 1985; Lees & Miller 1985, 1995). Differences are explained by differences in the mound v. ramp habitats, in nomenclature, in the systematic placement of problematic taxa, and in the choice of bathymetric criteria used. Comparison of the biotic gradient on the Alamogordo Member ramp with that on the Dinantian ramp of southwest England (Hennebert & Lees 1991)

is more generalized because of the broad spatial and temporal distribution of samples there, and because bathymetric inferences there are only relative. The excellent exposures and continuous bedding of the Alamogordo Member in the Sacramento Mountains provide a tightly constrained temporal and spatial framework within which quantitative as well as relative bathymetric interpretations can be made. The biotic gradients and assemblages in the Alamogordo Member provide depth criteria for recognizing palaeobathymetry of level-bottom strata, and provide a standard for the comparison of levelbottom and adjacent mound biotas and developmental histories. Support from the Paleontological Society, the Texaco Philanthropic Foundation, the Texas A&M University Department of Geology and Geophysics, and the Ray C. Fish Professorship are gratefully acknowledged. We thank W. Ahr, S. Bachtel, J. Miller, N. Pickard and G. Webb for their stimulating discussions and valuable reviews of the manuscript.

References

AHR, W. M. 1973. The carbonate ramp: an alternative to the shelf model. Transactions of the Gulf Coast Association of Geological Societies Annual Convention, 23, 221-225. 1989a. Comparative sedimentology of Waulsortian reefs at Waulsort, Belgium and Alamogordo, New Mexico. Geological Society of America Annual Meeting Abstracts with Programs, 21, A292. - - 1 9 8 9 b . Sedimentary and tectonic controls on the development of an Early Mississippian carbonate ramp, Sacramento Mountains area, New Mexico. Society of Economic Paleontologists and Mineralogists Special Publication, 44, 203-212. 81, STANTON, R. J. JR. 1994. Comparative sedimentology and palaeontology of Waulsortian mounds and coeval level-bottom sediments of the lower Lake Valley Formation (Lower Mississippian) in the Sacramento Mountains (New Mexico, USA). Abhandlungen der Geologischen Bundesanstalt, Wien, 50, 11-24. - & -1996. Constituent composition of Early Mississippian carbonate buildups and their levelbottom equivalents, Sacramento Mountains, New Mexico. This volume. BERGQUIST, P. R. 1978. Sponges. University of California Press, Berkeley & Los Angeles. BOGUSH, O. I. and BRENCKLE, P. L. 1982. Salebridae- a new family of uncertain affinity from the Lower Carboniferous of the USSR and USA: In: YUFEREV, O. V. (ed.) Stratigraphy and Palaeontology of the Devonian and Carboniferous. Akademiya Nauk SSSR. Sibirskoe Otdelenie. Institut Geologii i Geofiziki. Trudy 483, 103-118.

ALAMOGORDO

R A M P BIOTIC G R A D I E N T S

BOLTON, K. LANE H. R. & LEMONE, D. V. (eds) 1982. Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. E1 Paso Geological Society and the University of Texas at E1 Paso. BRENCKLE, P. 1977. Mametella, a new genus of calcareous red algae (?) of Mississippian age in North America. Journal of Paleontology, 51, 250-255. BRIDGES, P. H. & CHAPMAN, A. J. 1988. The anatomy of a deep-water mudmound complex to the southwest of the Oinantian platform in Derbyshire, UK. Sedimentology, 35, 139-162. BURCHETTE, T. P. & WRIGHT, V. P. 1992. Carbonate ramp depositional systems. Sedimentary Geology, 79, 3-57. BYRD, T. M. 1985. Facies Analysis of the Caballero Formation and the Andrecito Member of the Lake Valley Formation in the Northern Sacramento Mountains, Otero County, New Mexico. MSc Thesis,Texas A&M University. DUNHAM, R. J. 1962. Classification of carbonate rocks according to depositional texture. In: FRIEDMAN, G. T. (ed.) Classification of Carbonate Rocks. American Association of Petroleum Geologists Memoir, 1, 108-121. FLOGEL, E. 1982. Microfacies Analysis of Limestones. Springer Verlag, Berlin. GUTSCHICR, R. C., & SANDBERG, C. A. 1983. Mississippian continental margins of the conterminous United States. In: STANLEY, D. J. & MOORE, G. T. (eds) The Shelfbreak: Critical Interface on Continental Margins. Society of Economic Paleontologists and Mineralogists Special Publication, 33, 79 96. HENNEBERT, M. & LEES, A. 1991. Environmental gradients in carbonate sediments and rocks detected by correspondence analysis: examples from the Recent of Norway and the Dinantian of southwest England. Sedimentology, 38, 623 -642. HOOK, J. E., GOLUBIC, S. & MILLIMAN, J. D. 1984. Micritic cement in microborings is not necessarily a shallow-water indicator. Journal of Sedimentary Petrology, 54, 425-431. JAMES, N. P. & GINSBURG, R. N. 1979. The seaward margin of Belize barrier and atoll reefs. International Association of Sedimentologists Special Publication, 3, 1-161. JEFFERY, D. L. & STANTON, R. J. JR. 1993. Extensive and continuous parallel bedding as evidence for stable level-bottom deposition on the distal portion of a ramp. Geological Society of America Annual Meeting Abstracts with Programs, 25, A160. LANE, H. R. 1982. The distribution of Waulsortian facies in North America as exemplified in the Sacramento Mountains of New Mexico. In: BOLTON, K., LANE, H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. E1 Paso Geological Society and the University of Texas at El Paso, 96 114.

--,

125

SANDBERG, C. A. & ZIEGLER, W. 1980. Taxonomy and phylogeny of some Lower Carboniferous conodonts and preliminary standard post-Siphonodella zonation. Geologica et Paleontologica, 14, 117-164. LAUDON, L. R. & BOWSHER, A. L. 1949. Mississippian formations of southwestern New Mexico. Geological Society of America Bulletin, 60, 1 88. LEES, A. 1982. The paleoenvironmental setting and distribution of the Waulsortian facies of Belgium and southern Britain. In: BOLTON, K., LANE, H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies, El Paso Geological Society and the University of Texas at El Paso, 1-16. & MILLER, J. 1985. Facies variations in Waulsortian buildups, 2: Mid-Dinantian buildups from Europe and North America. Geological Journal, 20, 159 180. & -1995. Waulsortian Banks. In: MONTY, C., BOSENCE, D. W. J, BRIDGES, P. H. & PRATT, B. (eds) Carbonate Mud-Mounds: their Origins and Evolution. International Association of Sedimentologists Special Publication, 23, 191-271. --, HALLET, V. & HIBO, D. 1985. Facies variation in Waulsortian buildups, 1: a model from Belgium. Geological Journal, 20, 133-158. MAMET, B. 1991. Carboniferous calcareous algae. In: RIDING, R. (ed.) Calcareous Algae and Stromatolites, Springer-Verlag, Berlin, 370-451. 1992. PalOog6ographie des algues calcaires marines carbonif~res. Canadian Journal of Earth Sciences, 29, 174-194. MEYERS, W .J. 1975. Stratigraphy and diagenesis of the Lake Valley Formation Sacramento Mountains. In: PRAY, L. C. (ed.) Mississippian ShelfEdge and Basin Facies Carbonates, Sacramento Mountains in Southern New Mexico Region. Dallas Geological Society, 45-66. 1977. Chertification in the Mississippian Lake Valley Formation, Sacramento Mountains, New Mexico. Sedimentology, 24, 75-105. MILLER, J. & GRAYSON, R. F. 1982. The regional context of Waulsortian facies in northern England. In: BOLTON, K., LANE, H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. El Paso Geological Society and the University of Texas at E1 Paso, 17-33. PRAY, L. C. 1961. Geology of the Sacramento Mountains escarpment, Otero County, New Mexico. New Mexico Bureau of Mines and Mineral Resources Bulletin, 35, 1-144. RAILSBACK, L. B. 1993. Original mineralogy of Carboniferous worm tubes: evidence for changing marine chemistry and biomineralization. Geology, 21,703-706. ROSS, J. R. P. 1981. Biogeography of Carboniferous ectoproct bryozoa. Palaeontology, 24, 313-341. SWINCHATT, J. P. 1969. Algal boring: a possible depth indicator in carbonate rocks and sediments. Geological Society of America Bulletin, 80, 1391-1396.

126

D. L. J E F F E R Y & R. J. STANTON JR.

TABACHNICK, K. R. 1991. Adaptation of the hexactinellid sponges to deep-sea life. In: REITNER, J. & KEUPP, H. (eds) Fossil and Recent Sponges, Springer-Verlag, Berlin, 378-386. VACHARD, D. 1980. T&hys et Gondwana au Pal~ozo~que Sup6rieur, les donnees Afghanes; biostratigraphie, micropal6ontologie, palbog6ographie. Doc. et Trav. Igal, Paris, 2, 1-464.

1991. Parathuramminides et moravamminides (Microproblematica) de l'Emsien Sup+rieur de la Formation Moniello (Cordilleres Cantagriques, Espagne). Revue de Palkobiologie, 10, 255-299.

Late Vis~an buildups at Kingscourt, Ireland: possible precursors for Upper Carboniferous bioherms I A N D. S O M E R V I L L E ,

PETER

& H. E. A N N E

STROGEN,

GARETH

LL. J O N E S

SOMERVILLE

Dep a rtmen t o f Geology, University College Dublin, Belfield, Dublin 4, Ireland Abstract: Two late Vis6an (Asbian-early Brigantian) buildup complexes occur in the Kingscourt Outlier in Ireland, near the top of the Mullaghfin Formation, a shallow-water, grainstone unit. These massive buildups at Ardagh and Cregg accumulated on the margins of a carbonate platform bordering a deep-water basin. Both have a buildup facies of finegrained, peloid-rich, algal lime mudstones and wackestones, interbedded with coarser grained intraclastic, skeletal packstones and grainstones (interbuildup facies). Microbial structures are well developed in the buildup facies, principally domal stromatolites, thrombolites and oncoidal fabrics of cyanophytes (Ortonella and Girvanella) with encrusting foraminifers (Aphralysia and Tetrataxis). Algal structures include rhodoliths (Solenopora) and fragments of Ungdarella and stacheiids, with less abundant chlorophytes (Koninckopora); in the interbuildup facies, Koninckopora is more abundant. Near the top of the Ardagh buildup is an unusual development of phylloid algal boundstone composed of the possible ancestral coralline red alga Archaeolithophyllum. This boundstone also contains encrusting bryozoans and foraminifers, and directly overlies abundant Brigantian in situ fasciculate rugose corals. This upper unit appears to be associated with a rapid shallowing event, which stimulated the development of a waveresistant rigid framework. Laterally, bedded intraclastic skeletal packstones and grainstones, with thin interbedded clays and shales, pass into the lower part of the massive Ardagh buildup, but appear to onlap and drape the upper part of the buildup. Buildup clasts occur within the interbuildup facies and proximal flank facies, indicating synsedimentary cementation. Late Vis6an buildups of the Ardagh type highlight an important period of diversification of colonial rugose corals and calcareous algae associated with the development of widespread shallow-water carbonate platforms worldwide. This allowed the emergence of new algal groups, particularly the red algae (Archaeolithophyllum and Ungdarella) and the palaeoberesellid green algae (Kamaenella), which subsequently dominated Upper Carboniferous wave-resistant bioherms.

The early Carboniferous marks a period of recovery after the late Devonian reef collapse, with buildups dominated by mud-mounds (West 1988). Recent analysis of Lower Carboniferous bioherms (Webb 1994), however, has shown that their distribution and succession was not the result of evolution from a single primitive 'reef' community, but controlled by regional tectonostratigraphic settings and environments of deposition. Mud-mounds are particularly well developed in the late Tournaisian (Courceyan and early Chadian in the British Isles, equivalent to the early Mississippian of North America) where they are referred to as Waulsortian buildups (Lees 1964; Cotter 1965; Miller & Grayson 1972, 1982; Sevastopulo 1982; Lees et al. 1985; Miller 1986; Bridges & Chapman 1988; Somerville et al. 1992b). However, in the late Vis6an (Asbian and Brigantian Stages, equivalent to the late Mississippian of North America) many buildups had rigid skeletal, cementstone

or microbialite frameworks (Parkinson 1957; Schwarzacher 1961; Orme 1971; Mundy 1980, 1994; Bancroft et al. 1988; Webb 1994). Although fenestrate bryozoans, crinoids and hyalosteliid sponges are locally abundant in Waulsortian mud-mounds, none are deemed to be important in reef construction (West 1988). Many workers consider that in the absence of a recognizable framework they accumulated by microbially-induced precipitation of lime mud. In the early Vis+an (late Chadian and Arundian), Waulsortian-type mud-mounds developed sporadically (Kelly & Somerville 1992; Somerville et al. 1992a), but it was not until the late Vis6an (Asbian and Brigantian) that a second major development of buildups occurred (Nevill 1958; Ramsbottom 1969; Stevenson & Gaunt 1971; Bridges et al. 1995; Gutteridge 1995). Many of these later buildups occurred in shallower water depositional environments than the Waulsortian mud-mounds, which were initiated in deeper

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 127-144.

128

I. D. SOMERVILLE E T A L .

water distal ramp settings (Lees 1982; Lees & Miller 1985; Lees et al. 1985; Somerville et al. 1989, 1992b). This paper presents data from two late Vis6an buildup complexes in the Kingscourt area at Ardagh and Cregg, Co. Meath, Ireland (Fig. 1). The massive buildup complex at Ardagh is exposed in a large working quarry, which reveals the internal structure and geometry of the buildup and its relationship to enclosing bedded limestones. The buildup complex at Cregg, 8 km south of Ardagh, is a natural exposure, showing one main buildup with two smaller satellites.

Geological setting and history of research The buildups at Ardagh and Cregg were first reported by Jackson (1955) who noted that the mud-mounds had a specialized fauna of

brachiopods, bivalves, gastropods and goniatites, but a paucity of colonial rugose corals. He considered that the buildup complex at Cregg (his Lower Ardagh 'Reef' Limestone) was older than that at Ardagh (his Upper Ardagh 'Reef' Limestone), but the buildup complexes are now considered to be approximately coeval (Strogen et al. 1995). They occur near the top of the Mullaghfin Formation (Fig. 2), a 500m thick, shallow-water crinoidal-algal-intraclastic packstone/grainstone unit with pseudobreccias and palaeokarstic horizons. New coral and foraminiferal/conodont evidence (Strogen et al. 1995) has established that the two buildups are of late Asbian age, with the top of the Ardagh buildup being early Brigantian. Both buildup complexes represent accumulation at the margin of the shallow-water Ardagh Platform, which passes south into the deeper water basinal facies of the Lucan and Loughshinny Formations (Fig. 3).

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Methods of study Systematic sampling of both buildups was undertaken to examine the vertical and lateral variation in lithofacies and fauna. A vertical sampling interval of approximately 2 m was used in both buildup complexes. Sampling of the adjacent proximal bedded limestones at Ardagh was also carried out. Approximately 100 thin sections and polished slabs were collected from Ardagh (70 from the buildup and 30 from the bedded limestones) and 75 thin sections and polished slabs from Cregg. However, only petrographical detail from the much larger and more complete Ardagh buildup complex is presented here. A semi-quantitative log of the entire Ardagh complex was compiled (Fig. 4), based on visual estimates of allochem distribution in oriented thin sections of samples taken, in a composite vertical profile from base to top, encompassing measured sections in the three quarry bench faces and natural scarps. A separate study of the conodont faunas from the buildups and associated bedded limestones is being carried out by H.E.A.S. and has provided further faunal, palaeoecological and sedimentological data. All thin sections and fossils are housed in the UCD Geology Department collection.

131

Ardagh buildup complex Location This buildup complex is well exposed at Ardagh (N 835955), 3 km east of the town of Kingscourt. Quarry activity currently utilizes three benches cut into a prominent knoll-like hill (Fig. 5). In the quarry the mud-mound facies is at least 200m wide and over 300m long, but extends for a further 225 m to the southwest beyond the quarry. The complex is over 95 m thick, with neither the lower nor upper contacts exposed. Field mapping indicates that the buildup complex forms an elongated domiform ridge trending NE-SW and is bounded to the east by a normal fault of similar trend, where Namurian shales are downthrown against the buildup. The northwestern margin of the buildup is exposed in the west face of the middle bench. There, thickly bedded limestones with thin clay interbeds dipping 30-45 ° NW abut abruptly (and higher up drape over) massive buildup facies (Fig. 6).

Lithofacies The buildup complex is exposed in stratigraphic order, in the lower, middle and upper benches (Figs 4, 7). Note that thicknesses cited in the text

Fig. 5. Oblique aerial view of Ardagh Quarry looking southwest showing the low domiform topographic expression of the buildup complex. 1, lower bench and roadside scarp; 2, middle bench; 3, upper bench and eastern scarp; m, massive mud-mound; b, bedded limestones; F, Fault.

132

I. D. S O M E R V I L L E E T AL.

Fig. 6. View of the west face of the middle bench at Ardagh looking SW, showing the contact of bedded limestones (b) dipping to the NW and the massive pinnacle (m) of mud-mound facies (see Fig. 5). Note the thinning of the thick limestone beds towards the buildup as indicated by the continuity of the thin interbedded clay and shale horizons, and the onlap and drape of the upper bedded limestones over the buildup (between localities 10 and 12 in Fig. 7).

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LATE VISEAN BUILDUPS, IRELAND refer to measured heights above the exposed base of the buildup complex (see Fig. 4).

Lower bench. The lowest 2m of the exposed buildup in the roadside scarp east of the quarry (Loc. 1, Fig. 7) are coarse, fossiliferous packstones (coquinas) with abundant brachiopods, bivalves and gastropods (Fig. 8a). The succeeding 35 m is a uniform peloid-rich wackestone/packstone interval (Fig. 8b) with rare concentrations of skeletal material and scattered solitary corals. Stromatactoid cavities with radiaxial fibrous (RFC) and bladed calcite cements (Fig. 8c) are developed rarely (e.g. at 16m), but small irregular spar-filled cavities (Fig. 8b) occur throughout.

Middle bench. Approximately 35 m of massive limestone are exposed in the west face of the middle bench (Loc. 12, Fig. 7) overlying strata of the lower bench with a 2 m exposure gap. Between 39-48m (Fig. 4) are similar peloid-rich wackestone/packstone buildup facies to those of the lower bench, but with more abundant skeletal material, especially calcareous algae (Fig. 9h). However, between 48-67m the complex is dominated by mostly massive interbuildup facies consisting of (i) coarse-grained, intraclastic, skeletal packstones/grainstones rich in algae (Figs 8d, e), brachiopods and gastropods, (ii) finegrained well-sorted peloidal, skeletal packstones and (iii) intraclastic limestone breccias (e.g. at 49m) (Fig. 8f). The latter contain clasts of buildup facies, skeletal grainstones and exotic black, wackestone clasts. Colonial rugose corals are first recorded at 51.5 m, whilst small heterocorals (Hexaphyllia) are recorded only in the interval between 54.7-57.7 m. Upper bench. The upper 10 m in the 35 m thick middle bench overlap with the lowest strata in the upper bench (see Fig. 4). Nineteen metres of buildup complex (including the 10 m of overlap) is exposed in the SE corner of the upper bench (Loc. 14, Fig. 7) and in the natural scarp east of the quarry. Most of the section is in the same peloidrich wackestone buildup facies as in the lower bench, except here very large cavities are developed (up to a metre in width) filled with laminated, peloidal geopetal sediment and RFC and blocky spar cements. The matrix in the upper buildup facies is noticeably finer and has a volumetrically lower proportion of skeletal debris compared with the lower buildup facies (except for a distinctive thin skeletal grainstone bed at 73 m) (see Fig. 4). Macrofauna is locally

133

conspicuous (Loc. 13, Fig. 7), consisting of large brachiopods (Gigantoproductus maximus), solitary corals (Axophyllum, Dibunophyllum bipartitum and Siphonophyllia siblyi) and scattered colonial rugose corals (Siphonodendron and Lithostrotion). However, above 82 m, fasciculate corals are extremely abundant consisting of large in situ colonies (Figs 10, 11) up to 2 m in diameter, with corallites coated with RFC cements in a lime mudstone matrix. The uppermost beds are locally dolomitized. Approximately 10 m are exposed in natural sections above the upper bench (Locs 15 & 16, Fig. 7). These exposures are mostly in interbuildup facies and contain scattered coquinas rich in bivalves, brachiopods and gastropods, and one oolitic horizon.

Biota, non-skeletal components and fabrics Lower bench. Foraminifers are present in almost all samples, although typically in small numbers, except near the base and at 10m (Fig. 4). Similarly, kamaenids (septate tubular microproblematica), which have been referred to the Chlorophyta (dasycladacean palaeoberesellids) by Mamet & Roux (1974), Skompski (1987) and Deloffre (1988), are present in most samples (Fig. 8d) with peak abundance at 22.5 m. Other algae, notably Koninckopora (dasycladacean green alga) and aoujgaliids (problematical red algae - Rhodophyta of Petryk & Mamet 1972; Mamet & Roux 1977; Chuvashov & Riding 1984), which are otherwise very rare in this part of the buildup, have similar peaks in abundance at 22.5m. Cyanophytes (e.g. Ortonella/Girvanella) are fairly common in lower samples but are absent above 18m. Mirroring the development of cyanophytes are Aphralysia-like domes, which were considered encrusting foraminifers (Belka 1981). Thrombolites (see Aitken 1967) and cyanophyte oncoids are present in the same interval, but absent above 20m. Fragments of fenestrate bryozoans are found in the majority of samples (Fig. 8h), but sponge spicules are very rare, occurring in only three samples, two of which are in reworked clasts (Fig. 8h). In many samples large, well-rounded, micritic intraclasts (Fig. 8a) are conspicuous. They are not bored or encrusted and do not show evidence of a former skeletal fabric that may have been micritized. Middle bench. Foraminifers are most abundant, and red and green algae are more diverse and abundant. Within the interbuildup facies between 48-58m (Fig. 4), Koninckopora (Fig.

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LATE VISEAN BUILDUPS, IRELAND 9a), aoujgaliids (Fig. 9b) and Ungdarella (red algae of Maslov 1956, 1962; Petryk & Mamet 1972; Chuvashov & Riding 1984) occur in most samples, although Ungdarella (Figs 8d, 8e &9c) is not known from lower buildup levels. Cyanophytes (Ortonella/Girvanella) are abundant in the lower part but absent above 51.5 m. Aphralysia-like domes and oncoids occur in this interval, especially at 51.5m, and unusual encrusting foraminifers similar to Calcitornella, encrusting Tetrataxis and Ungdarella (Fig. 9c) occur at 59-61m. Fenestrate bryozoans and sponge spicules are notably absent in most samples.

Upper bench. Foraminifers are relatively sparse in this part of the buildup, apart from the lowest 2 m (65-67 m) and in the beds above the quarry face (Fig. 4). Calcareous algae are much less abundant than in the middle bench, with kamaenids and aoujgaliids in only a few samples. Ungdarella is very rare, whilst Koninckopora is completely absent throughout this upper section. Cyanophytes and cryptalgal fabrics are also very rare except at 73.5 m (Fig. 9d). Near the top of the section (83-85m) however, there is a profusion of phylloid algae resembling the ancestral red coralline alga Archaeolithophyllum (Figs 9e, f). It is associated with Aphralysia, Girvanella, corals, sponge spicules (Fig. 9g) and encrusting bryozoans (Fistulipora and Tabulipora). This crustose and foliaceous phylloid alga has branches and nodes, and formed vertical as well as horizontal sheets which sheltered geopetal cavities, similar to those illustrated by Reid (1986). Unfortunately, the internal structural detail of this alga was obliterated during diagenetic alteration and there is little evidence of cellular structure under cathodoluminescence or ultraviolet light.

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Bedded strata adjacent to the Ardagh buildup Lower bench. The bedded limestones by the crusher (loc. 3) are coarse-grained and thickbedded ( 3 - 5 m thick), well-sorted crinoid-rich intraclastic grainstones with abundant foraminifers and calcareous algae (Fig. 8g). All samples from these beds contain the diagnostic Asbian foraminifer Vissariotaxis which is not recorded elsewhere in the quarry. Rhodophytes are very abundant and include Aoujgalia, Stacheoides (Fig. 8g) and Ungdarella, with rare Solenopora. Chlorophytes are also abundant including Kamaena and Kamaenella (Fig. 9h), but Koninckopora is conspicuously absent from all the samples. Middle bench. A similar intraclastic-skeletal packstone/grainstone lithofacies is present, although here packstones are dominant. On the east face close to the hopper (Locs 5 & 6) are very coarse-grained, intraclastic crinoidal rudites with large crinoid stems (proximal flank facies). As in the lower bench the bedded limestones have abundant Ungdarella and aoujgaliids, but here Koninekopora is common. Cyanophytes are absent. Fenestrate bryozoans are also absent in 14 of 15 samples. Foraminifers are common, especially Eostaffella and bilaminar palaeotextulariids. In the west face (Loc. 9) the limestones are interbedded with a blue-green shale (Fig. 12), which thickens from 10 cm to 4 m over a lateral distance of 10m before terminating against a fault. The shale appears to fill an irregular top of the underlying limestone, which may represent a palaeokarst surface associated with temporary emergence. Palaeokarst surfaces have been described elsewhere from late Asbian platform limestones (see Strogen et al. 1995). At Loc. 10 (west face), limestone beds (up to 10m

Fig. 8. Thin section photomicrographs of lithofacies in Ardagh Quarry. (a) Coarse-grained skeletal packstone (basal coquina) rich in bivalves, brachiopods, ostracodes, crinoids and well-rounded micrite intraclasts (i). Base of roadside scarp, Loc. 1. Scale bar 3 mm. (b) Peloid-rich wackestone buildup facies with irregular spar-filled cavities lined with thin layer of inclusion-rich RFC cements and blocky sparry calcite. Lower bench, 18.5m, Loc. 1. Scale bar 3 ram. (c) Stromatactoid cavities with a thick layer of inclusion-rich, zoned RFC cements (z) and clear, voidfilling blocky cements (b). Buildup facies, lower bench, 16 m Loc. 1. Scale bar 3 mm. (d) Interbuildup skeletal packstone with crinoids and calcareous algae Ungdarella(u) and Kamaena (k). Middle bench, 57.7m, Loc. 12. Scale bar 1.5 ram. (e) Interbuildup skeletal packstone rich in Ungdarella(u). Middle bench, 18.7 m, Loc. 12. Scale bar 1.5 mm. (f) Limestone conglomerate (intraclastic-skeletal rudstone) with clasts of buildup facies, crinoids and Koninckopora (k). Interbuildup facies, middle bench, 48.1 m, Loc. 12. Scale bar 3 mm. (g) Coarse-grained crinoidal-algal-intraclastic grainstone with abundant Stacheoides (s), Aoujgalia (a) and crinoids (c). Proximal bedded facies, lower bench, Loc. 3. Scale bar 3 mm. (h) Hexactinellid sponge spicules (h) within a micrite clast embedded in a mosaic of sparry calcite cement containing fragments of fenestellid bryozoans. Lower bench, 37 m, Loc. 1. Scale bar 1.5 mm.

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Fig. 10. In situ fasciculate colonial rugose coral (Siphonodendron) > 1.25 m high, from the top of the upper bench (Loc. 14). Scale bar 25 cm.

thick) thin towards the massive buildup and have primary depositional dips of 20-30 ° (Fig. 5). These very thick limestone beds are separated by very thin (5-10 cm) green, possibly bentonitic clay bands, the lowest of which is overlain by a thin fenestral micrite. Calcretes and rhizocretions were not recognized in the limestones beneath the clay bands.

Upper bench. Bedded limestones are only present in the N W corner (Loc.17) and appear to form flank beds to the massive buildup (Fig. 5). These fine-grained, well-sorted peloidal packstones contain abundant spar-filled cavities, are poor in foraminifers and devoid of calcareous algae.

Interpretation of buildup complexes Although the base of the > 9 5 m thick Ardagh buildup is not exposed, the lowest 2 m, a coarsegrained skeletal packstone rich in brachiopods, bivalves and gastropods may be interpreted as a possible 'shell bank' or stabilizing skeletal platform on which cyanophytes and calcareous algae formed the pioneer colonizers. As such it could represent the stabilization and colonization stages of Walker & Alberstadt's (1975) ecological succession. This 'basal' coquina may have created a 'local high' on the sea-bed, providing a site preferential for cyanophyte colonization and subsequent vertical accretion of the buildup.

Fig. 9. Thin section photomicrographs of biota from the buildup complex at Ardagh Quarry. (a) Algal grainstone rich in Koninckopora (k). Interbuildup facies, middle bench, 51.5m, Loc. 12. Scale bar 1.5mm. (b) Algal-rich packstone with Ungdarella (u), Koninckopora (k) and Epistacheoides (e). Interbuildup facies, middle bench, 51.5 m, Loc. 12. Scale bar i .5 mm. (c) Wackestone buildup facies with encrusting Tetrataxis (T), Ungdarella (u) and encrusting foraminifer similar to Calcitornella (c). Middle bench, 59 m, Loc. 12. Scale bar 1.5 mm. (d) Coarsegrained, intraclastic-skeletal grainstone with large domiform oncoid with irregular concentric cryptalgal laminae. Interbuildup facies, upper bench, 73.5 m, Loc. 14. Scale bar 1.5 mm. (e) Peloid-rich Archaeolithophyllum boundstone showing horizontal and vertical orientation of phylloid algal sheets. Note branching (b), nodes (n) and cavities (c) beneath horizontal sheets. Buildup facies, upper bench, 83.2 m, Loc. 14. Scale bar 3 mm. (f) Detail of Archaeolithophyllum boundstone showing Archaeolithophyllum (a), intimately associated with Aphralysia (A), Fasciella (F) and Girvanella forming black specks. Buildup facies, upper bench, 83.5 m, Loc. 14. Scale bar 1.5 ram. (g) Wackestone with abundant sponge spicules (mostly in transverse section), with two long spindles in longitudinal section. Note large spar-filled cavity lined with zoned RFC cement. Buildup facies, upper bench, 84m, Loc. 14. Scale bar 3 mm. (h) Algal grainstone rich in fragments of branching Kamaenella (k). Interbuildup facies, middle bench, 41.2m, Loc. 12. Scale bar 3mm.

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Fig. 11. in situ fasciculate colonial rugose coral ('Koninckophyllum') from the top of the upper bench (Loc. 14). Hammer head is 17 cm long. A prolonged period of cyanophyte growth and lime mud entrapment followed, forming a cohesive meshwork of Girvanella and Ortonella thalli with locally developed Solenopora rhodoliths in a peloid-rich matrix, typical of many shallow water Vis6an platform buildups (see Somerville et al. 1992a; Pickard this volume). Buildup growth may have initiated in a moderately deep (>20m) and protected part of

the platform margin, below fair-weather wavebase and at lower light intensities. This is suggested by the abundance of fenestrate bryozoans, the lack of wave and current structures, the absence of mud winnowing and dasycladacean algae (e.g. Koninckopora) in the lower part of the buildup (cf. Horbury & Adams this volume). However, despite the abundance of fenestrate bryozoans, many are fragmentary with

Fig. 12. Wedge-shaped shale unit (outlined by dashed line) thickening rapidly towards the fault plane (F), west face of middle quarry (Loc. 9).

LATE VISEAN BUILDUPS, IRELAND few intact fronds, which may be a result of bioturbation or transportation. This may also explain in part the paucity of large stromatactoid cavities, which are locally abundant in Waulsortian buildups (Phases A and D of Lees & Miller 1985) where fenestellids form support for cavity roofs. Nonetheless, there are abundant small, irregular-shaped, spar-filled cavities developed throughout the buildup. These cavities may be lined with a thin layer of RFC cements similar to those recorded in north Co. Dublin by Somerville et al. (1992a). Although kamaenids occur in the buildup facies, they may not be as reliable a shallow water-depth indicator as suggested by Skompski (1987), because they are also known from the deeper water Waulsortian buildups of the Dublin Basin (Lees & Miller 1985; Strogen et al. 1990; Somerville et al. 1992b). However, the abundance of Kamaena and Kamaenella, often fragmented, in the laterally equivalent bedded limestones and proximal flank facies is indicative of shallow water environments (see Adams et al. 1992). Occasionally, rapid shallowing events occurred as evidenced by skeletal packstones (e.g. at 22m and 29m) rich in kamaenids, Koninckopora and aoujgaliids, and accompanied by abundant foraminifers on the buildup crest. Buildup facies accumulated at Ardagh for much of the lower 37m before there was a gradual change to interbuildup facies (39-43 m and especially 48-58m). At these levels there is a profusion of foraminifers, chlorophytes (especially Koninckopora) and rhodophytes (Ungdarella and stacheiids), suggesting a much shallower water depth, well within the photic zone. Breccia horizons containing clasts of buildup facies, skeletal grainstones derived from locally bedded limestones and exotic, black wackestone clasts occur within the interbuildup facies. The black, fine-grained limestone clasts may represent a similar facies to the Deer Park Formation (a slightly younger platform limestone unit to the NW of Ardagh Quarry; Fig. 3), but they possibly were derived from the coeval quieterwater platform facies, which developed leewards of the buildup and is no longer exposed. The presence of buildup facies clasts in the breccia, which are clearly recognizable by their peloidal wackestone fabric and spar-filled cavities with geopetal sediment, indicates that the lower part of the buildup complex experienced synsedimentary cementation. Following this period of interbuildup facies deposition, cyanophytes were reestablished and

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a second major phase of lime mud accumulation occurred (forming the top of the middle bench and most of the upper bench). However, the contact between the buildup and interbuildup facies is poorly defined at outcrop. The matrix in the upper buildup facies is noticeably finer and has a lower proportion of skeletal debris compared with the lower buildup facies. Conversely, macrofauna is more conspicuous in the upper buildup facies. The absence of Koninckopora in this section may signify a return to greater water depths of buildup development associated with a sea-level rise. However, irrespective of water depth, the absence of Koninckopora may mark its final demise, as the genus became extinct close to the Asbian/ Brigantian boundary (Somerville & Strank 1984; Somerville et al. 1992c; Jones & Somerville this volume). The top of the buildup is Brigantian in age based on its coral faunas (Strogen et al. 1995). A rapid relative sea-level fall occurred near the top of the buildup complex, which exerted a significant control on faunal and floral diversification in the top of the buildup. This is witnessed by the sudden profusion of abundant in situ rugose coral colonies, which acted as baffles, overlain directly by Archaeolithophyllum, encrusting bryozoans and foraminifers (Aphralysia and Tetrataxis). The resulting bindstone fabric is characterized by cavities and corallites lined with RFC cements. This is an ecological community replacement and may represent the diversification stage of Walker & Alberstadt (1975). Unfortunately, the topmost part of the complex is no longer preserved. The ancestral coralline red alga Archaeolithophyllum, which forms crustose and foliaceous sheets several centimetres in length in the buildup, first occurs in the late Vis6an (see Wray 1964, 1977; Belka 1981). Belka (1981) noted that Aphralysia can be found encrusting Archaeolithophyllum and is itself encrusted by Girvanella in algal-foraminiferal mud-mounds. Moreover, of all the Lower Carboniferous calcareous algae, two in particular (the encrusting foliaceous Archaeolithophyllum and the ramose Ungdarella) became cosmopolitan reef formers in the Upper Carboniferous (Wray 1977; Mamet 1991). Recent studies (James et al. 1988) have suggested that Arehaeolithophyllum may be a peyssonelid with an aragonitic wall structure which has been subsequently diagenetically altered. This may explain the lack of cellular detail in the specimens from the Ardagh buildup. Wray (1964, 1977) has suggested that Archaeolithophyllum from Upper Carboniferous

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(Pennsylvanian) bioherms may have formed algal banks similar to those of the Recent Lithothamnium (Bosence 1985), with individual species of Archaeolithophyllum adapted to the more turbulent-water crests of banks. This may be analogous to the Ardagh buildup where Archaeolithophyllum becomes locally abundant just below the top. However, the alga never assumes the volumetrically important role of dominance in the buildup, which forms the culminating stage of the ecological succession model of Walker & Alberstadt (1975). It is inferred that this last stage is associated with a rapid shallowing event. The bedded limestones immediately to the NW of the buildup in the middle bench, although containing clasts of reworked platform limestones, do not contain recognizable buildup derived fragments. The lowest bedded limestones in the middle bench abut the buildup, but higher bedded limestones drape it. Unfortunately there is, as yet, insufficient precision in the foraminiferal/conodont biostratigraphy to differentiate late Asbian and early Brigantian faunas within the peripheral bedded limestones, or between the bedded limestones and the massive buildup facies (see Jones & Somerville this volume). Thus at present it is not possible to demonstrate the nature of the contact, or to determine if there is onlap of an exhumed mound or lateral passage from buildup to bedded flank beds. The buildup complex at Cregg does not show any contact between bedded limestones and massive buildup facies. However, a similar vertical sequence of buildup-interbuildupbuildup facies at Ardagh is recognized in the largest mud-mound at Cregg. At Cregg, as at Ardagh, the top of the buildup complex is not exposed. However, limited field data suggests that the main Cregg buildup was rapidly buried by terrigenous mud with interbedded graded limestone turbidites of the basinal Loughshinny Formation, which crops out immediately to the west. The latter facies is associated with a marine transgression that resulted in the drowning of the southern margin of the Ardagh Platform in the late Asbian-early Brigantian (Fig. 3; Strogen et al. 1995).

Discussion It is generally agreed that in the Lower Carboniferous (Dinantian) there was a major hiatus in reef construction following the collapse of large framework-building metazoans in the late Devonian (James 1984; West 1988). As

Newell (1972) remarked 'this was a time of crisis for reef-building'. The result was that initially in the Dinantian a slow period of recovery was characterized by the development of deeperwater Waulsortian mud-mounds that lack any recognizable organic framework. The prevailing theory on their construction is that they resulted from microbial activity (Lees & Miller 1985, Somerville et al. 1992b). A similar pattern is evident in the role of calcareous algae as reef builders (Chuvashov & Riding 1984). Many algal groups that were present in the Cambrian to Devonian showed a marked decline or became extinct at the end of the Devonian. New algal groups (e.g. KamaenaDonezella group, Archaeolithophyllum group and Ungdarella-Stacheia group) that evolved in the Lower Carboniferous later became important reef-builders in the Middle and Upper Carboniferous (Westphalian-Stephanian) (Chuvashov & Riding 1984; West 1988). Algae are rarely reported as a common reef constituent in the Dinantian, although members of the Renalcis group (a possible cyanophyte; Pratt 1984), which ranges back to the Cambrian (Chuvashov & Riding 1984), occur in some Vis6an bioherms and mud-mounds (Adams 1983; Webb 1989; Somerville et al. 1992a, c). In the early Carboniferous the dominant metazoans in Waulsortian mud-mounds were crinoids, fenestrate bryozoans and sponges (mostly as spicules). Occasionally at the top of these mud-banks and in flank beds, or within interbank facies (in areas of shallower water deposition), chlorophytes (Koninckopora, A tractyliopsis and kamaenids) and cyanophytes (Girvanella) are recorded (Jones et al. 1988; Somerville et al. 1992b). For the most part, calcareous algae are absent from Waulsortian mounds because of their presumed greater (subphotic) water depth (Lees & Miller 1985). Calcareous algae are likely to be present only in mud-banks that have Phase D (shallowest water-depth) components, such as many of the Waulsortian mounds in the Dublin Basin (Somerville et al. 1992b) and slightly younger early Vis~an Waulsortian buildups in NW Ireland (Kelly & Somerville 1992). VisOan bioherm f r a m e w o r k s

In the late Vis6an and Namurian, significant but subtle changes occurred in buildup construction, most notably in the volumetric decline of bryozoans and the rise of calcareous algae and corals in forming frameworks in bioherms and mud-mounds. However, bryozoans (fenestrates,

LATE VISEAN BUILDUPS, IRELAND trepostomes and fistuliporids), and cyanophytes in some cases, still occurred as in situ meshwork or framework builders (Dix & James 1987; Bancroft et al. 1988; Christopher 1990; Lauwers 1992). In the early Vis6an (late Chadian and Arundian), cyanophytes began to emerge as important binders and bafflers of sediment, and formed wave-resistant 'cryptalgal' structures (domal stromatolites and thrombolites) and oncoids. These were constructed by meshworks of Girvanella, Ortonella and Renalcis, assisted by framebuilders such as solenoporids, encrusting foraminifers (Aphralysia and Tetrataxis, interpreted by Cossey & Mundy (1990) as a limpet-like epiphytic encruster), encrusting bryozoans (Fistulipora) and tabulate corals (Adams 1984; Fang & Hou 1987; Somerville et al. 1992a). In the later Vis6an, colonial rugose corals, calcareous algae and non-skeletal microbialite structures are reported (similar to those at Ardagh) (Mundy 1980, 1994; Webb 1989, 1994), and may have been the dominant biotic component in buildups and platform margins bound by syndepositional marine cements (Horbury 1992). The importance of corals in Vis6an bioherms has until recently been unrecognized. Most workers (James 1984; Chuvashov & Riding 1984; West 1988) have commented on the dominance of Waulsortian-type mud-mounds and the absence of well-developed coral reef and algal reef communities, which characterized late Devonian reefs. Vis6an coral reefs referred to by some earlier workers (Johnson 1959; Caldwell & Charlesworth 1962) have subsequently been reinterpreted as coral biostromes. However, Webb (1989, 1990, 1994), in a detailed study of the late Vis6an Lion Creek bioherms from Queensland, Australia, observed that they had a rigid coral-algal framework. The coral-algal boundstone core facies consists of a framework of colonial rugose ('Siphonodendron' and 'Orionastraea') and tabulate (Multithecopora) corals in growth position. The coral colonies are up to a metre in diameter and occur within a microbialite matrix in buildups up to 15m thick. These structures are comparable with the top of the Ardagh buildup. A very large atoll reef over 20km across in southern Japan ranges in age from early Carboniferous to Permian (Ota 1968). In the late Vis6an and Middle Carboniferous period of reef development, colonial rugose (Siphonodendron) and tabulate (multithecoporan) corals formed reef flat communities, assisted by chaetetids (sclerosponges), encrusting bryozoans, encrusting foraminiferans, red algae and cyanophytes (Sugiyama & Nagai 1990, 1994).

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Rodriguez (this volume) has also recognized Upper Vis6an reefs in Spain composed of Siphonodendron and tabulate colonial corals with red algae.

Conclusions (1) Two, massive, late Asbian buildup complexes at Ardagh and Cregg, dominated by cyanophytes and calcareous algae, were developed on a shallow-water carbonate platform close to the margin of a deeper water basin. Buildup relief is difficult to assess but is probably low; nowhere can bedded limestones proximal to the buildup be seen to interfinger with the massive limestones. The bedded limestones at Ardagh appear to thin against, onlap, and drape the buildup, and are coeval with or post-date buildup accumulation. Furthermore, there is no evidence of reef talus in the bedded limestones adjacent to the buildup. However, in the interbuildup facies at Ardagh, one interval of limestone breccia contains a variety of limestone clasts, some of which are of buildup type, whereas others appear to represent more distant sources. (2) Both buildup complexes show vertical variation in lithofacies involving an alternation of fine-grained, peloid-rich algal wackestone/ packstones (buildup facies) and coarse-grained, intraclastic skeletal packstone/grainstones (interbuildup facies). Small, irregular, spar-filled cavities are common in the buildup facies, but particularly near the top of the Ardagh complex, larger cavities contain numerous generations of geopetal sediment. (3) Cyanophyte structures are frequently developed in both buildups and appear to form meshworks and frameworks. However, near the top of the Ardagh buildup, Archaeolithophyllum boundstone facies is associated with thickets of colonial rugose corals, sponges, encrusting bryozoans and foraminifers. For the most part, however, bryozoans and sponges are notably subordinate components and colonial corals are extremely rare. Increasing metazoan diversity towards the top of the Ardagh buildup may represent an ecological community replacement. This is associated with a rapid shallowing event and the development of a wave-resistant skeletal framework of colonial corals and red algae, as the buildup grew into the surf zone. Indication of shoaling is provided by a rare oolitic horizon near the top of the Ardagh buildup. (4) The late Vis6an Ardagh and Cregg buildup complexes contain a complex suite of lithofacies and biofacies with vertical variation in

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textures, microfacies and biota typically related to alternations of buildup and interbuildup facies. This may reflect the vertical stacking or coalescence of smaller discrete mud-banks, and obscure internal geometries. However, the second major phase of buildup facies at Ardagh may be related to a deepening event associated with an early Brigantian transgression, as seen in the abrupt passage from massive bedded, pale grey limestones of the Mullaghfin Formation to dark grey, cherty and argillaceous limestones of the Deer Park Formation. (5) These late Vis6an buildups are more advanced than the earlier Waulsortian (late Tournaisian-early Vis6an) mud-mounds. Cyanophytes are important in providing initial meshworks which acted as baffles, and later under more favourable conditions formed rigid frameworks when associated with other metazoans, principally colonial corals, sponges, encrusting bryozoans, foraminifers, and specialized red algae such as Archaeolithophyllum. These late Vis6an buildups thus appear to represent an important transition in the evolution of Carboniferous reefs: from microbe-dominated Waulsortian mud-mounds in the early Carboniferous, to conspicuous reef-framework buildups of the late Carboniferous. Moreover, it is very significant that some of the algal flora that first emerged in the late Vis6an as contributors to buildup frameworks (e.g. Archaeolithophyllum, Ungdarella and Kamaeneila) later formed the dominant component of Upper Carboniferous bioherms. We would like to thank Roadstone Provinces Ltd, Castleblayney and the quarry manager J. Glynn for permission to visit Ardagh Quarry. We are also indebted to G. Webb and A. Horbury for their invaluable critical reviewing of the paper.

References ADAMS, A. E. 1983. Lower Carboniferous Renalcis from Cumbria. Proceedings of the Yorkshire Geological Society, 44, 327-331. - - 1 9 8 4 . Development of algal-foraminiferal-coral reefs in the Lower Carboniferous of Furness, northwest England. Lethaia, 17, 233-249. - - , HORBURY, A. D. & RAMSAY, A. T. S. 1992. Significance of palaeoberesellids (Chlorophyta) in late Dinantian sedimentation, UK. Lethaia, 25, 375-382. AITKEN, J. D. 1967. Classification and environmental significance of cryptalgal limestones and dolomites with illustrations from the Cambrian and Ordovician of Southwestern Alberta. Journal of Sedimentary Petrology, 37, 1163-1187.

BANCROFT, A. J., SOMERVILLE, I. D. & STRANK, A. R. E. 1988. A bryozoan buildup from the Lower Carboniferous of North Wales. Lethaia, 21, 51-65. BELKA, Z. 1981. The alleged algal genus Aphralysia is a foraminifer. Neues Jahrbuch ffir Geologie und Palaontologie Monatshefte, 1981, 257 266. BOSENCE, D. 1985. Preservation of coralline-algal reef frameworks. Proceedings of the Fifth International Coral Reef Congress, Tahiti, 6, 623-628. BRIDGES, P. H. & CHAPMAN,A. J. 1988. The anatomy of a deep water mud-mound complex to the southwest of the Dinantian platform in Derbyshire, UK. Sedimentology, 35, 139-162. , GUTTERIDGE, P. & PICKARD, N. m. H. 1995. Environmental setting of Carboniferous mudmounds in NW Europe. In: MONTY, C. L. V., BRIDGES, P. H., PRATT, B. & BOSENCE,D. W. J. (eds) Carbonate Mudmounds-Origin and Evolution. Special Publication of International Association of Sedimentologists, 23, 171-190. CALDWELL, W. G. E. & CHARLESWORTH, H. A. K. 1962. Vis6an coral reefs in the Bricklieve Mountains of Ireland. Proceedings of the Geologists' Association, 73, 359-382. CHRISTOPHER, C. C. 1990. Late Mississippian Girvanella-bryozoan mud mounds in southern West Virginia. Palaios, 5, 460-471. CHUVASHOV, B. & RIDING, R. 1984. Principle floras of Palaeozoic marine calcareous algae. Palaeontology, 27, 487 500. COSSEY, P. J. & MUNDY, D. J. C. 1990. Tetrataxis, a loosely attached limpet-like foraminifer from the Upper Palaeozoic. Lethaia, 23, 311-322. COTTER, E. 1965. Waulsortian-type carbonate banks in the Mississippian Lodgepole Formation of central Montana. Journal of Geology, 73, 881-888. DELOFFRE, R. 1988. Nouvelle taxonomie des algues dasycladales. Bulletin de Centres Recherche Exploration Production Elf Aquitaine, 12, 165-217. DIX, G. R. & JAMES, N. P. 1987. Late Mississippian bryozoan/microbial build-ups on a drowned karst terrain: Port au Port Peninsula, western Newfoundland. Sedimentology, 34, 779-793. FANG, S. & HOU, F. 1987. Tatangian (Carboniferous) bryozoan-coral patch reef in Langping of Tianlin, Guangxi. 11th International Congress of the Stratigraphy and the Geology of the Carboniferous, Beijing 1987, Abstracts, 1, 167-168. GEORGE, T. N., JOHNSON, G. A. L., MITCHELL, M., PRENTICE, J. E., RAMSBOTTOM, W. H. C., SEVASTOPULO, G. D. & WILSON, R. B. 1976. A Correlation of Dinantian Rocks in the British Isles. Geological Society of London Special Report No. 7. GUTTERIDGE, P. 1995. Late Dinantian carbonate mud mounds of the Derbyshire carbonate platform. In: MONTY, C. L. V., BRIDGES, P. H., PRATT, B. & BOSENCE, D. W. J. (eds) Carbonate Mudmounds- Origin and Evolution. Special Publication of International Association of Sedimentologists, 23, 289-307.

LATE VISEAN B U I L D U P S , I R E L A N D HORBURY, A. D. 1992. A late Dinantian peloid cement stone-palaeoberesellid buildup from North Lancashire, England. Sedimentary Geology, 79, 117-137. & ADAMS, A. E. 1996. Microfacies associations in Asbian carbonates: an example from the Urswick Limestone Formation of the sothern Lake District, northern England. This volume. JACKSON, J. S. 1955. The Carboniferous Succession of the Kingscourt Outlier with Notes on the PermoTrias. PhD Thesis, University of Dublin. JAMES, N. P. 1984. Reefs. In: WALKER, R. G. (ed.) Facies Models. Geoscience Canada Reprint Series 1 (2nd edn), 229-244. --, WRAY, J. L. & GINSBURG, R. N. 1988. Calcification of encrusting aragonitic algae (Peyssonneliacea): implications for the origin of Late Palaeozoic reefs and cements. Journal of Sedimentary Petrology, 58, 291-303. JOHNSON, G. m. L. 1959. The Carboniferous stratigraphy of the Roman Wall district in western Northumberland. Proceedings of the Yorkshire Geological Society, 32, 83-130. JONES, G. LL., & SOMERVILLE, I. D. 1996. Irish Dinantian biostratigraphy: practical applications. This volume. & STROGEN, P. 1988. The Lower Carboniferous (Dinantian) of the Swords area: sedimentation and tectonics in the Dublin Basin, Ireland. Geological Journal, 23, 221-248. KELLY, J. G. & SOMERVILLE, I. D. 1992. Arundian (Dinantian) carbonate mudbanks in North-West Ireland. Geological Journal, 27, 221-242. LAUWERS, A. 1992. Growth and diagenesis of cryptalgal-bryozoan buildups within a midVis~an (Dinantian) cyclic sequence. Annales de la Soci~t~ G~ologique de Belgique, 115, 187-213. LEES, A. 1964. The structure and origin of the Waulsortian (Lower Carboniferous) 'reefs' of west-central Eire. Philosophical Transactions of the Royal Society London, 247B, 483-531. --1982. The palaeoenvironmental setting and distribution of the Waulsortian facies of Belgium and southern Britain. In: BOLTON, K., LANE, H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. E1 Paso Geological Society and University of Texas at E1 Paso, 1-16. & MILLER, J. 1985: Facies variation in Waulsortian buildups. Part 2: Mid-Dinantian buildups from England and North America. Geological Journal, 20, 159-180. --, HALLET, V. & HIBO, D. 1985. Facies variation in Waulsortian buildups. Part 1: A model from Belgium. Geological Journal, 20, 133-158. MAMET, B. 1991. Carboniferous calcareous algae. In: RIDING, R. (ed.) Calcareous Algae and Stromatolites. Springer-Verlag, Berlin, 370-451. -& Roux, A. 1974. Sur quelques algues tubulaires scalariformes de la Tethys Pal6ozoique. R~vue de Micropaleontologie, 17, 134-156.

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143

& 1977. Algues rouges devoniennes et carbonif~res de la Tethys Occidentales, quatrieme partie. R~vue de Micropaleontologie, 19, 215-266. MASLOV, V. P. 1956. Calcareous algae of the USSR. Transactions of Akadamie de Sciences USSR Geological Institute, 160, 1-301 [in Russian]. - - 1 9 6 2 . Fossil red algae of the USSR. Transactions of Akadam& de Sciences USSR Geological Institute, 53, 1-221 [in Russian]. MILLER, J. 1986. Facies relationships and diagenesis in Waulsortian mudmounds from the Lower Carboniferous of Ireland and Northern England. In: SCHROEDER, J. H. & PURSER, B. H. (eds) Reef Diagenesis. Springer Verlag, Berlin, 311-335. & GRAYSON, R. F. 1972. Origin and structure of the Lower Vis~an "reef" limestones near Clitheroe, Lancashire. Proceedings of the Yorkshire Geological Society, 38, 607 638. & 1982 The regional context of Waulsortian facies in northern England. In: BOLTON, K., LANE H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. El Paso Geological Society and University of Texas at E1 Paso, 17-33. MUNDY, D. J. C. 1980. Aspects of the Palaeoecology of the Craven Reef Belt (Dinantian) of North Yorkshire. PhD Thesis, University of Manchester. --1994. Microbialite-sponge-bryozoan-coral framestones in Lower Carboniferous (Late Vis6an) buildups of northern England (UK). In: BEAUCHAMP, B., EMERY, A. F. & GLASS, D. J. (eds) Pangea." Global Environments and Resources. Canadian Society of Petroleum Geologists Memoir, 17, 713-729. NEVlLL, W. E. 1958. The Carboniferous knoll-reefs of east-central Ireland. Proceedings of the Royal Irish Academy, 59B, 285-303. NEWELL, N. D. 1972. The evolution of reefs. Scientific American, 226, 54-65. ORME, G. R. 1971. The D2-P1 'reefs' and associated limestones of the Pin Dale-Bradwell Moor area of Derbyshire. Comptes Rendues 6th International Congress in Stratigraphy and Geology of the Carboniferous, Sheffield 1967, 3, 1249-1262. OTA, M. 1968. The Akiyoshi Limestone Group: A geosynclinal organic reef complex. Bulletin of the Akiyoshi-dai Science Museum, 5, 1-44. PARKINSON, D. 1957. Lower Carboniferous reefs of northern England. Bulletin of the American Association of Petroleum Geologists, 41, 511-537. PETRYK, A. m. & MAMET, B. 1972. Lower Carboniferous algal flora, southwestern Alberta. Canadian Journal of Earth Sciences, 9, 767-802. PICKARD, N. A. H. 1996. Evidence for microbial influence on the development of Lower Carboniferous buildups. This volume. PRATT, B. R. 1984. Epiphyton and Renalcis: diagenetic microfossils from calcification of blue-green algae. Journal of Sedimentary Petrology, 54, 948-971. -

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RAMSBOTTOM, W. H. C. 1969. Reef distribution in the British Lower Carboniferous. Nature, 222, 765-766. REID, P. R. 1986. Discovery of Triassic phylloid algae: possible links with the Paleozoic. Canadian Journal of Earth Sciences, 23, 2068-2071. RODRIGUEZ, S. 1996. Development of coral reef-facies during the Vis6an at Los Santos de Maimona, SW Spain. This volume. SCHWARZACHER, W. 1961 Petrology and structure of some Lower Carboniferous reefs in Northwestern Ireland. Bulletin of the American Association of Petroleum Geologists, 45,1481-1503. SEVASTOPULO, G. D. 1982. The age and depositional setting of the Waulsortian Limestones in Ireland. In: BOLTON, K., LANE H. R. & LEMONE, D. V. (eds) Symposium on the Environmental Setting and Distribution of the Waulsortian Facies. El Paso Geological Society and University of Texas at El Paso, 65-79. SKOMPSKI, S. 1987. The dasycladacean nature of Late Palaeozoic palaeoberesellid algae. Acta Geologica Polonica, 37, 21-31. SOMERVILLE, I. D. & STRANK, A. R. E. 1984. The recognition of the Asbian/Brigantian boundary fauna and marker horizons in the Dinantian of North Wales. Geological Journal, 19, 227-237. -, PICKARD, N. A. H., STROGEN, P. & JONES, G. LL. 1992a. Early to mid-Visban shallow water platform buildups, north Co. Dublin, Ireland. Geological Journal, 27, 151-172. , STRANK, A. R. E. & WELSH, A. 1989. Chadian faunas and flora from Dyserth: Depositional environments and palaeogeographic setting of Vis6an strata in northeast Wales. Geological Journal, 24, 49-66. , STROGEN, P. & JONES, G. LL. 1992b. MidDinantian Waulsortian buildups in the Dublin Basin, Ireland. Sedimentary Geology, 79, 91-116. , & --1992c. Biostratigraphy of Dinantian limestones and associated volcanic rocks of the East Limerick Syncline. Geological Journal, 27, 201-220. STEVENSON, I. P. & GAUNT, G. D. 1971. Geology of the country around Chapel en le Frith. Memoir of the Geological Survey of Great Britain.

STROGEN, P., JONES, G. LL. & SOMERVILLE,I. D. 1990. Stratigraphy and sedimentology of Lower Carboniferous (Dinantian) boreholes from west Co. Meath, Ireland. Geological Journal, 25, 103-137. - - , SOMERVILLE,I. D., JONES, G. LL. & PICKARD,N. A. H. 1995. The Lower Carboniferous (Dinantian) stratigraphy and structure of the Kingscourt Outlier, Ireland. Geological Journal, 30, 1-23. SUGIYAMA, T. & NAGAI, K. 1990. Growth forms of Auloporidid corals in the Akiyoshi Limestone Group, southwestern Japan: Palaeoecological studies of reef-building organisms in the Akiyoshi organic reef complex. Bulletin of the Akiyoshi-dai Science Museum, 25, 1-25. - & 1994. Reef facies and palaeoecology of reef-building corals in the lower part of the Akiyoshi Limestone Group (Carboniferous) Southwest Japan. Courier Forschungsinstitut Senckenberg, 172, 231-240. WALKER, K. R. & ALBERSTADT, L. P. 1975. Ecological succession as an aspect of structure in fossil communities. Palaeobiology, 1,238-257. WEBB, G. E. 1989. Late Vis~an coral-algal bioherms from the Lion Creek Formation of Queensland, Australia. Comptes Rendu l l th International Congress of the Stratigraphy and the Geology of the Carboniferous, Beijing 1987, 3, 282-295. - - 1 9 9 0 . Lower Carboniferous coral fauna of the Rockhampton Group, east-central Queensland. In: JELL, P. A. (ed.) Devonian and Carboniferous Coral Studies. Association of Australasian Palaeontologists Memoir 10, 1-167. - - 1 9 9 4 . Non-Waulsortian Mississippian bioherms: a comparative analysis. In: BEAUCHAMP, B., EMERY, A. F. & GLASSS, D. J. (eds) Pangea: Global Environments and Resources. Canadian Society of Petroleum Geologists Memoir 17, 701-712. WEST, R. R. 1988. Temporal changes in Carboniferous reef mound communities. Palaios, 3, 152-169. WRAY, J. L. 1964. Archaeolithophyllum, an abundant calcareous alga in limestones of the Lansing Group (Pennsylvanian), southeastern Kansas. Kansas Geological Survey, 170, 1-13. --1977. Calcareous algae. Developments in Palaeontology and Stratigraphy. Elsevier, Amsterdam.

Development of coral reef-facies during the Vis~an at Los Santos de Maimona, SW Spain SERGIO

RODRIGUEZ

Departamento de Paleontologia, Facultad de Ciencias Geoldgicas, Universidad Complutense de Madrid, 28040 Madrid, Spain Abstract: The Siphonodendron Limestone at Los Santos de Maimona Basin is regarded as a reef structure built mainly by rugose corals. This unit is composed of biogenic marls and biostromal limestones containing abundant rugose corals and brachiopods, and frequent calcareous algae, tabulate corals, foraminiferans, bryozoans, echinoderms, ostracodes and molluscs. It is present throughout the basin, which is 12 km long and 3 km wide, but the thickness, development of the framework and distribution of organic components varies from SE (seaward) to NW (landward). It is 40 m thick in the SE, but its thickness reaches only 6 m in the NW. The unit shows a vertical evolution of lithological facies, from biogenic marls at the bottom up to biostromal limestones at the top. The main environmental factors controlling the development of the organic framework are the tidal regime, minor subsidence pulses and periodic storms. Arguments in favour of the reefal origin are structural (distribution, development and relationships between different facies of the Siphonodendron Limestone) and ecological (distribution and relationships of building organisms, environmental indicators, etc.). Objections to the reefal hypothesis (absence of reef crest and talus, low diversity, problematical wave resistance, biostromal nature of the upper beds, storm layers) are discussed. Building structures by corals, brachiopods and calcareous algae are briefly described.

Carboniferous rugose corals are not usually regarded as reef-builders. They were described as the main components of reef mounds or bioherms (Pareyn 1959; Sutherland & Henry 1977; Webb 1989) or as secondary components of reef complexes (Adams 1984), but very rarely as the main builders of reef complexes (Ota 1968). In fact, many authors deny the existence of true ecological reefs (or reef complexes) during the Carboniferous (Heckel 1974; James, 1979, 1983; Copper 1988, 1994) or at least during the Early Carboniferous (West 1988). Nevertheless, Lower Carboniferous true reef complexes have been described from America (Schenk & Hatt 1984), the UK (Adams 1984) and Japan (Sugiyama & Nagai 1994), and at least in the last case, rugose corals are the main component. The Siphonodendron Limestone at Los Santos de Maimona (SW Spain) seems to be an additional example of the building capability of rugose corals during the Carboniferous. The Los Santos de Maimona Basin is located in the Ossa-Morena region, near Zafra (Badajoz, SW Spain, Fig. 1). It is small and elongated in a N W - S E orientation, coinciding with the main Hercynian structures of the region, and limited by tear faults with rotational movement (Odriozola et al. 1983). It comprises Upper Vis~an (Asbian) rocks of varied characteristics, which have been subdivided into eight lithostratigraphic units

numbered from 0 to 7 (Sfinchez et al. 1988, 1991; Falces 1991; Rodriguez et al. 1992). Units 1, 3, 4 and 6 are basically carbonates. The purpose of this contribution is to describe the coral reef-facies of Unit 1, here interpreted as a wave-resistant reef structure developed during the early Asbian. Unit 1 is composed of biostromal and bioelastic marls and limestones containing abundant colonies of the rugose coral genus, and consequently it was named the Siphonodendron Limestone. Large brachiopods, calcareous algae and tabulates participate in the framework of the biostromal levels (Rodriguez et al. 1994). Several sections containing Unit 1 were studied (Albuera River, E1 Almendro, Guadajira River, El Portezuelo, Navafria, E1 Torre6n and Los Santos Hill). Some of these sections illustrate the lateral variations of the Siphonodendron Limestone (Fig. 2). The thickness, development of the framework and distribution of organic components vary largely from SE to NW.

Lithofacies Marly biogenic beds in the lower part of the southern sections (Lower Member; Rodriguez et al. 1992, 1994) are composed of heterogeneous colonies of Siphonodendron and syringoporoids,

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107 pp. 145-152.

146

S. RODRiGUEZ foraminiferans and calcareous algae are abundant in the bioclastic levels. Ooids are also common.

. ......

The hypothesis

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AB

Fig. 1. Location map of the Los Santos de Maimona Basin, with positions of the studied sections. AB, Albuera River; AL, El Almendro; GU, Guadajira River; PT, El Portezuelo; TO, El Torre6n; NA, Navafria; SS, Los Santos Hill. Carboniferous outcrops are indicated by vertical haching in the general map of the Ossa-Morena region.

red and blue-green algae and large gigantoproductids growing in a rigid organic framework (Fig. 2). The dominant microfacies are bafflestones and packstones, but grainstones, bindstones and framestones are not uncommon. Intercolonial spaces are filled mainly with micrite and bioclasts (rugose and tabulate corals, bryozoans, brachiopods, etc.; see Rodriguez et al. 1992). The upper part of the southern sections and almost the whole of the northern sections (Upper Member; Rodriguez et al. 1992, 1994) are composed of biostromal limestones built mainly by corals (bafflestones). Several bioclastic levels (packstones, grainstones), where the Siphonodendron colonies are less developed, occur in the whole area, but more frequently in the northern sections (Rodriguez et al. 1992, 1994). Fragments of corals, brachiopods, bryozoans, bivalves, gastropods, ostracodes, trilobites, echinoderms,

The Siphonodendron Limestone was interpreted by Rodriguez et al. (1994) as the reef-flat of a fringing reef with the open sea towards the south and the land towards the north. The reef was developed on terrigenous sediments that filled the Los Santos Basin during the Middle Vis6an. Carbonate sedimentation occurred only when the water depth was shallow enough to allow the building organisms and carbonate producers to develop at the beginning of the Upper Vis6an (Sfinchez et al. 1988; Rodriguez et al. 1994). The basal, marly levels at the Los Santos Hill and Torre6n sections, which constitute the main reef framework, developed rapidly and arose from the sea bottom to the surface (Fig. 3a, b). The framework began to develop on a very characteristic bed of red algae (Rodriguez & S~nchez-Chico 1995) which provided the corals with a solid substrate on which to grow (Fig. 4). In this part of the reef, coral colonies reach up to 1 m high. When the framework reached the low tide level, vertical growth was substituted by the extensive development of the reef-flat, which is present in most sections (Fig. 3c). Pulses of subsidence in the Los Santos Basin continued, allowing the subsequent development of biostromal beds in the upper part of the sections. After each subsident pulse the growth of the gigantoproductids and the Siphonodendron colonies began again. Periodic storms disturbed the development of the reef so that some beds composed of coral shingle are present in the upper part of the Los Santos Hill, Torre6n and Navafria sections. Fragments are larger and less reworked in the Los Santos Hill section, situated at the south end of the Siphonodendron Limestone outcrops. Northern sections (Guadajira, E1 Almendro and Albuera river) show scarce development of coral colonies, which are replaced by bryozoans and other small colonial organisms.

The basis for recognition as a reef The basis for recognizing the Siphonodendron Limestone as a reef are largely discussed in Rodriguez et al. (1992, 1994), and Rodriguez & Sfinchez-Chico (1995), and will be simply

VISEAN CORAL REEF-FACIES, SW SPAIN

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Fig. 2. Spatial distribution of the studied sections of the SiphonodendronLimestone (Unit 1) at Los Santos de Maimona. The base of the unit is placed in the southern sections at a calcareous bed composed of solenoporacean algae. The boundary between the two members (subunits) is placed at the level where the limestones become dominant. Buildup facies of the lower member are replaced by terrigenous facies in the northern sections. summarized here. The arguments can be divided into two groups; those related to structure, and those related to ecology.

Structure The lower marly, coral member constitutes a lenticular sedimentation unit, 3 km long and up to 20 m thick, which influenced the deposition in adjacent areas (Fig. 2) and that of subsequent units. Also, the thickness of the upper part of the unit decreases to the north (20 to 6 m, Fig. 2), and cycles of gigantoproductid and coral growth are also reduced in thickness.

Ecology The marly, coral member at the lower part of the southern sections (Los Santos Hill, Navafria, Torre6n) is almost totally composed of building organisms surrounded by bioclastic mud (Rodriguez & Sfinchez-Chico 1995). Coral colonies usually grow on ramose solenoporoid red algae (Parachaetetesjohnsoni, Pseudochaetetes, Fig. 4). The most abundant facies in the upper part are biostromal limestones. These biostromes are mainly composed of fasciculate corals which

reach the same growth-level in each bed, sometimes developing typical 'microatoll' structures (Fig. 6), described by Rodriguez & S~nchez-Chico (1995). Surrounding sediments are mainly packstones to wackestones, lacking primary sparry calcite. When corals are firmly established the structure is wave-resistant (but not storm-resistant; Fig. 5). The building organisms decrease to the north, where the proportion of accessory building organisms increases. Vagile organisms are also more abundant to the north (or at least most usually entirely preserved). Storm layers occur throughout the basin at the top of the sections (Fig. 5). The sandy bioclastic debris is proportionally more abundant to the north, but the size of the fragments is larger to the south. Subaerial exposure features occur at the top of many sequences in all studied sections and are regarded as having originated during exceptional low tides or temporary sea-level falls.

The objections There are several objections to the theory that the Siphonodendron Limestone is a true reef complex. Some arise from the accepted concept of a reef, others were posed and explained in Rodriguez et

148

S. RODRiGUEZ

Sea level Stabilitation ~ S l o w

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Sea level

Rapid reef growth to reach file sea level

Slow subsidence

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Fig. 3. Stages of development of the Siphonodendron Limestone. Subsidence is progressivelygreater to the south.

al. (1994); others will be briefly explained here. Further objections were posed by colleagues during the VI Fossil Cnidaria Symposium in Mtinster in 1991, leading the author to revise and study again some aspects of the Siphonodendron Limestone, which are described in Rodriguez & Sfinchez-Chico (1995) and here.

The concept of a reef Detailed discussions on this concept are included in Lowenstam (1950), Heckel (1974), Walker (1974) and Longman (1981) and will not be repeated here. Definitions given by Heckel (1974) include: 'a buildup that displays: (1) evidence of (a) potential wave resistance or (b) growth in turbulent water which implies wave resistance; and (2) evidence of control over the surrounding environment'. The relief, organic buildup and influence on adjacent areas has

already been detailed (Rodriguez et al. 1992, 1994, and here). It only remains to prove that the Siphonodendron Limestone was a waveresistant structure and, in fact, many of the structures present in the upper part of the unit show evidence of this. Some beds at Navafria, Los Santos Hill and El Torre6n sections show clastic material bordering large coral colonies. In addition, some colonies with some parts that are broken away, rest in life position. Nevertheless, some layers are composed of broken colonies mixed with many other bioclasts. They are regarded as storm beds such as those from many present-day reefs (Mather & Bennett, 1984). The lower part of the reef, present in El Torre6n and Los Santos Hill sections, is not required to be a wave-resistant structure because it developed at depth, and the wave-base reached the corals there only sporadically. In any case, the evidence of the development of building organisms growing in turbulent water allow us to infer that the unit was wave-resistant.

VISEAN C O R A L R E E F - F A C I E S , SW SPAIN

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Fig. 5. Stages of buildup development of the upper biostromal beds (1) and storm beds (2, 3) at the top of Los Santos Hill section. Three cycles of unequal development are included in the picture (1 3). The general pattern comprises a bioclastic accumulation level (a), a gigantoproductid-growth level (b), a coral (Siphonodendron, Syringopora, Pleurosiphonella) buildup (e) and a reworking phase (d). Cycle 1 shows the normal sequence. The reworking phase is so strong in Cycle 2 that coral colonies were moved and broken (but not brachiopods, which consequently should be cemented to substratum). Coral colonies are almost intact (but overturned), and brachiopods are completely broken in cycle 3; consequently, the reworking phase was probably very strong but short-lived. Red crusts are present over and below this bed, showing temporary subaerial exposure. No vertical exaggeration.

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Fig. 6. Stages of buildup development in the upper biostromal beds, at E1 Torre6n section. Cycle 1 shows a similar structure to the cycle 2 of Fig. 5. Cycle 2 shows a great horizontal development of the coral colonies (c). The presence of red crusts at the upper surface suggest a temporary subaerial exposure that forced the corals to grow in that way (microatoll growth). The vertical growth later continued (c~), when the relative sea-level rose. No vertical exaggeration.

The reef crest and talus The reef crest and talus are not preserved. In fact, there is no outcrop of the Siphonodendron Limestone southwards from Los Santos Hill, which should be the nearest locality. One possible reason is that the reef crest was totally removed by recent erosion. Another possibility is that a true reef crest never existed, and the reef flat finished with a fore-reef talus. The relief of the bioconstructed beds is small, and such talus should also be small. This possibility agrees with the model described by Cloud (1952) for fringing reefs with slow subsidence. The absence of massive framebuilders may be justified by a moderate level of energy (with the exception of occasional storm episodes). Two facts support this hypothesis: no massive colonies were found in the biostromal layers, but also not in the storm layers, where fragments of the reef crest should be found. Abundant coral debris, possibly belonging to the suggested fore-reef talus, was found in the southern slope of Los Santos Hill. However, this hypothesis is not confirmed because both pieces of evidence are circumstantial.

Bioherms and biostromes Usually, only bioherms are regarded as true reefs. Nevertheless, the upper beds of the

Siphonodendron Limestone are typically biostromal. These beds represent the reef-fiat developed on the buildup and built by corals, algae and gigantoproductid brachiopods. In a very broad sense the Upper Member is also a buildup (6 m thick at Albuera and El Almendro sections, 20 m thick at E1 Torre6n section).

Low diversity One of the most prominent characteristics of coral reefs is a very high diversity of biota (Heckel 1974). In contrast, some cases of bioherms composed of only one or two species have been described (e.g. Rich 1969; Sutherland 1984; M6ndez-Bedia et al. 1994). The diversity of the main builders in the studied example at Los Santos de Maimona is low; only two species of rugose corals (Siphonodendron martini and Siphonodendron irregulate), two tabulate genera (Syringopora and Pleurosiphonella) and one gigantoproductid species (Gigantoproductus aft. semiglobosus) are found. Some red algae (Parachaetetes johnsoni and Pseudochaetetes) play an important role at the beginning of the buildup. Such diversity increases if we take into account secondary builders and associated fauna and flora. At least four species of brachiopods

VISEAN CORAL REEF-FACIES, SW SPAIN (Martinez-Chacdn & Legrand-Blain 1992), 16 foraminiferans (Comas-Rengifo et al. 1992), 12 rugose corals and three tabulate corals (Rodriguez & Falces 1992, 1994), and 13 of calcareous algae (Sfinchez-Chico et al. 1995) have been described. Dental plates and scales of fishes were also described (Soler-Gij6n & Rodriguez 1991). In addition, many species of bryozoans, bivalves, ostracodes, trilobites, echinoderms, and so on, are present in the bioclastic beds. The number of species increases if we consider the components without preservation potential. S t o r m beds

The main objection posed to recognizing the Siphonodendron Limestone as a reef is the presence of storm beds, which are common in the Upper Member. It is claimed that if the framework is wave-resistant, storms should affect little or none of the structure of the reef. However, many actual reefs show periodic disturbance of growth because of storms. Many of them actually show coral shingle beds, and even shingle cays (Mather & Bennett 1984). Periodic storm disturbance does not destroy the reef, but provides a high quantity of material for the reef-flat and back-reef areas.

Conclusions The Siphonodendron Limestone shows many characteristics of a reef (following definitions by Heckel 1974 and Longman 1981) built mainly by corals. It is an exception to the usually accepted idea that rugose corals did not build reefs during the Carboniferous. Some objections remain unanswered, or the explanations are unconvincing (absence of reef crest, low diversity of main builders), but arguments in favour of recognizing it as a reef largely overcome these objections. This study is included in the research project PB910083, supported by the Spanish DGICYT. I am grateful to W. J. Sando, I. D. Somerville, E. Poty and J. R. Nudds for their useful comments on the manuscript.

References ADAMS, A. E. 1984. Development of algal-foraminiferal-coral reefs in the Lower Carboniferous of Furness, northwest England. Lethaia, 17, 233-249. CLOUD, P. E. JR. 1952. Facies relationships of organic reefs. Bulletin of the American Association of Petroleum Geologists, 36, 2125-2149.

151

COMAS-RENGIFO, M. J., RODRiGUEZ, S. & SANCHEZ, J. L. 1992. Foraminiferos. In: RODRiGUEZ, S. (ed.) And~isis Paleontoldgico y Sedimentoldgico de/a cuenca Carbonlfera de Los Santos de Maimona (Badajoz). Coloquios de Paleontologia, 44, 145-157. COPPER, P. 1988. Ecological succession in phanerozoic reef ecosystems: is it real? Pa/aios, 3, 136-152. 1994. Reefs under stress: the fossil record. Courier Forschungsinstitut Senckenberg, 172, 87-94. FALCES, S. 1991. Cartografia y Paleontologla de la Cuenca Carbonifera de Los Santos de Maimona." Corales Solitarios de la Fauna de Cyathaxonia. PhD Thesis, Complutense University of Madrid. HECKEL, P. H. 1974. Carbonate buildups in the geological record: a review. In: LAPORTE, L. F. (ed.) Reefs in Time and Space. Society of Economic Paleontologists and Mineralogists Special Publication, 18, 90-154 JAMES, N. P. 1979. Reefs. In: WALKER, R. G. (ed.) Facies Models. Geoscience Canada Reprint Series, 1, 121-132. 1983. Reef environment. In: SCHOLLE, P. A., BEBOUT, D. G. & MOORE, C. H. (eds.) Carbonate Depositional Environments. Association of American Petroleum Geologists Memoir, 33, 345-440. LONGMAN, M. W. 1981. A process approach to recognizing facies of reef complexes. In: TOOMEY, D. F. (ed.) European Fossil Reef Models. Society of Economic Paleontologists and Mineralogists Special Publication, 30, 9 40. LOWENSTAM, H. A. 1950. Niagaran reefs of the Great Lakes Area. Journal of Geology, 58, 430-487. MARTiNEZ-CHACON, M. L. & LEGRAND-BLAIN, M. 1992. Braquibpodos. In: RODRiGUEZ, S. (ed.) Andlisis Paleontoldgico y Sedimentoldgico de la cuenca Carbonifera de Los Santos de Maimona (Badajoz). Coloquios de Paleontologia, 44, 91-144. MATHER, P. & BENNETT, I. 1984. A Coral Reef Handbook. A Guide to the Fauna, Flora and Geology of Heron Island and Adjacent Reefs and Cays. The Australian Coral Reef Society, Brisbane. MI~NDEZ-BEDIA, I., SOTO, F. & FERN~,NDEZ-MARTiNEZ, E. 1994. Devonian reef types in the Cantabrian Mountains (NW Spain) and their faunal composition. Courier Forschungsinstitut Senckenberg, 172, 161-184. ODRIOZOLA, J. M., PEON, A., VARGAS, I., GARROTE, A. & ARRIOLA, A. 1983. Mapa Geoldgico de Espaffa a escala 1.'50.000. Hoja 854; Zafra. Instituto Geoloegico y Minero de Espafia (2nd edn), Madrid. OTA, M. 1968. The Akiyoshi Limestone Group: a geosynclinal organic reef complex. Bulletin Akiyoshi-dai Science Museum, 5, 1-44. PAREYN, C. 1959. Les rdcifs Carbonifdres du Grand Erg occidental. Bulletin de la Soci~t~ G~ologique de France, 711, 347-364. RICH, M. 1969. Petrographic analysis of Atokan carbonate rocks in central and southern Great Basin. American Association of Petroleum Geologists Bulletin, 53, 340-366.

152

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RODRiGUEZ, S. & FAECES, S. 1992. Corales Rugosos. In: RODRiGUEZ, S. (ed.) Andlisis Paleontol6gico y Sedimentol6gico de la Cuenca Carbonifera de Los Santos de Maimona (Badajoz). Coloquios de Paleontologia, 44, 159-218 & 1994. Coral distribution patterns in the Los Santos de Maimona Lower Carboniferous Basin (Badajoz, SW Spain). Courier Forschungsinstitut Senckenberg, 172, 193-202. -& SJ~NCHEZ-CHICO, F. 1995. Corales rugosos y algas calc~treas de la secci6n de E1 Torre6n (Viseense, Badajoz). Coloquios de Paleontologia, 46, in press. , FAECES, S., ARRIBAS, M. E., DE LA PElqA, ~I. m., COMAS-RENGIFO, M. J. & MORENO-EIRIS, E., 1992. Descripci6n litoestratigrfifica y aspectos sedimentol6gicos de las unidades. In: RODR|GUEZ, S., (ed.) Andlisis Paleontol6gico y Sedimentol6gico de la Cuenca Carbonifera de Los Santos de Maimona (Badajoz). Coloquios de Paleontologia, 44, 49-90. , ARRIBAS, M. E., FAECES, S., MORENO-EIRIS, E. & DE LA PENA, J. 1994. The Siphonodendron Limestone of Los Santos de Maimona Basin: development of an extensive reef-flat during the Vis6an in Ossa Morena, SW Spain. Courier Forschungsinstitut Senckenberg, 172, 203-214 S~,NCHEZ, J. L., COMAS-RENGIFO, M. J. & RODRIGUEZ, S. 1988. Estudio estratigr/lfico de los materiales carbonatados del Carbonifero de Los Santos de Maimona (Badajoz, SO de Espafia). Comunicaciones del segundo Congreso de Geologla de Espa~a, 1, 197-200. & 1991. Foraminiferos del Carbonifero inferior de Los Santos de Maimona (Badajoz, SO de Espafia). Boletin de la Real Sociedad Espa~ola de Historia Natural (secci6n de Geologia), 86, 101-147. S~,NCHEZ-CHICO, F., MAMET, B., MORENO-EIRIS, E. & RODRiGUEZ, S. 1995. Algas calcfireas del Viseense de Los Santos de Maimona (Badajoz, SO de Espafia). Revista Espa~ola de Micropaleontologia, in press. -

-

SCHENK, P. E. & HATT, B. L. 1984. Depositional environment of the Gays River reef, Nova Scotia, Canada. In: GELDSETZER, H. H. J. (ed.) Atlantic Coast Basins. Compte Rendu, Ninth International Congress on Carboniferous Stratigraphy and Geology, 3, 117-130. SOLER-GIJON, R. & RODRiGUEZ, S. 1991. Estudio de un resto de Bradiodonto (clase Chondrictyes) del Viseense de Los Santos de Maimona (Badajoz, SO de Espafia). Coloquios de Paleontologia, 43, 101-114. SUGIYAMA, T. & NAGAI, K. 1994. Reef facies and paleoecology of reef-building corals in the lower part of the Akiyoshi Limestone Group (Carboniferous), Southwest Japan. Courier Forschungsinstitut Senckenberg, 172, 231-240. SUTHERLAND, P. K. 1984. Chaetetes reefs of exceptional size in Marble Falls Limestone (Pennsylvanian), central Texas, Paleontographica Americana, 54, 543-547. - - & HENRY, T. W. 1977. Carbonate platform facies and new stratigraphic nomenclature of the Morrowan series (Lower and Middle Pennsylvanian), northeastern Oklahoma, Geological Society of America Bulletin, 88, 425-440. WALKER, K. R. 1974. Reefs through time: a synoptic review. In: Principles of Benthic Community Analysis; Notes for a Short Course: Sedimenta IV. Comparative Sedimentology Laboratory, University of Miami, 8, 1 20. WEBB, G. E. 1989. Late Vis6an coral-algal bioherms from the Lion Creek Formation of Queensland, Australia. Compte Rendu 11th Congr~s International Congress de Stratigraphie et de Geologie du Carboniferd, 3, 282-295. WEST, R. R. 1988. Temporal changes in Carboniferous reef mound communities. Pataios, 3, 152-169.

Evidence for catastrophism at the Famennian-Dinantian boundary in the Iberian Pyrite Belt CARMEN

MORENO,

SONIA SIERRA & REINALDO

SAEZ

Departamento de Geologia, Universidad de Huelva, 21819 Palos de la Frontera, Huelva, Spain Abstract: Terrigenous shelf sedimentation during the Devonian created a homogeneous basin in the Iberian Pyrite Belt. This shallow south-Iberic basin changed into a mosaic of horsts and graben at the Famennian-Dinantian boundary and subsequent evolutionary history can be explained in terms of locally increased rates of subsidence. Deposits commonly related to highly energetic processes characterize this change (fan-deltas, sediment gravity flows, rapid basin-shallowing). These resulted from convulsive/ catastrophic events related to the Bretonic phase of the Hercynian Orogeny. Marker beds due to convulsive/catastrophic geological events are a valuable correlation tool and a precise key to the analysis of many sedimentary basins. These events can be preserved in the geological record in a variety of ways, generally characterized, by evidence of high transport energy for short periods of time. Identification of 'event beds' is not always easy, but analysis of facies and particularly their accurate interpretation in relation to background sedimentary processes is usually a successful means of identification. Some recent papers dealing with sediment gravity flow deposits, fan-deltas and other coarse-grained deltas are examples of this type of interpretation (Ethridge 1985; Galloway 1985; Kleverlaan 1987; Pilkey 1988; Leigh & Hartley 1992; Mastalerz & Wojewoda 1993). Until recently, the world-famous massive sulphide deposits of our study area, the Iberian Pyrite Belt (IPB) in SW Spain, have absorbed most of the geological research, neglecting other geological features such as those described in this study. The aim of this paper is to show that the sedimentary features of deposits at the Famennian-Dinantian boundary in the IPB are produced by episodic/catastrophic events related to the Bretonic phase of the Hercynian Orogeny, which controlled basin compartmentalization. Knowledge of the morphology of the Dinantian basin of the IPB, which was inherited from older structures, could in turn be a key to understanding the genesis and distribution of Dinantian massive sulphides of the IPB.

Riotinto, Neves-Corvo, Aljustrel, Tharsis and Aznalcollar (Fig. 1). The IPB occupies the southwestern corner of the Iberian Peninsula, extending from Seville, in Spain, to the Atlantic Ocean south of Lisbon, Portugal, in an arcuate belt about 230 km long and 40 km wide. The IPB is a part of the South Portuguese Zone (Julivert et al. 1974) which has been interpreted as a tectonostratigraphic terrane sutured to the Iberian Massif during the middle Carboniferous (Quesada 1991). The sedimentary record of the IPB consists of Devonian and Carboniferous rocks whose conspicuous features include intense Dinantian magmatic activity and the abundance of huge massive sulphide deposits. The accepted stratigraphic sequence for the IPB was established by Schermerhorn (1971) and comprises three main units: the Phyllite-Quartzite (PQ) Group, the Volcanic-Siliceous Complex (VSC) and the Culm Group (Fig. 1).

The Phyllite-Quartzite (PQ) Group This comprises a detrital sequence of shales and sandstones. Limestone lenses near the top of the succession have provided conodonts and other fossils of middle to late Famennian age (Van den Boogaard & Schermerhorn 1975). The extent at depth of the group remains unknown, although estimates of thickness indicate several thousand metres. Sedimentary facies of'the PQ represent marine deposition on a shallow platform.

The Volcanic-Siliceous Complex (VSC) Geological framework The IPB is one of the oldest mining districts in the world. It is characterized by giant and supergiant massive sulphide deposits, including

This complex contains the economically important massive sulphide and manganese deposits. The VSC is late Famennian to Vis~an in age, and consists of a heterogeneous group of rocks

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 153-162.

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Fig. I. Geological map of the South Portuguese Zone of the Iberian Peninsula. After Julivert et al. 1974. Squared area shows location of Fig. 2. displaying rapid lateral and vertical facies changes. Significantly, the thickness of the VSC varies widely, ranging from a few tens to hundreds of metres. Rock sequences consist of both felsic and mafic volcanics interfingering with an envelope of detrital and chemical sediments (Routhier et al. 1980). Volcanic rocks are mainly felsic pyroclastic rocks and mafic flows. Subvolcanic rocks, both felsic and mafic, are ubiquitous, comprising in some places the bulk of the stratigraphic column (Boulter 1993). Sedimentary rocks comprise three main types: volcanic-derived epiclastites whose grain sizes range from conglomerates to silts; black shales rich in organic matter, often associated with massive sulphide bodies; and chemical sedimentary rocks including massive sulphides and manganiferous cherts and jaspers.

The Culm Group

Named the Baixo Alentejo Flysch Group in Portugal (Oliveira et al. 1979), the Culm Group is a thick Upper Carboniferous succession of shales, litharenites and rare conglomerates which overlies the VSC in the IPB. The estimated thickness for this group exceeds several thousands of metres. The Culm Group represents the in fill of a rapidly subsiding basin, mostly by

turbidite sediments coming from at least two source areas: the IPB proper, and the Ossa Morena Zone (Moreno 1993).

Structure and metamorphism

Rocks of the IPB were deformed and regionally metamorphosed in the Asturian phase of the Hercynian Orogeny, from late Vis6an to Westphalian-D times. Three stages of deformation have been recognized in the IPB. The first, D1, generated regional structures and was accompanied by lowgrade regional metamorphism, whereas D2 and D3 only modified the D1 structures. Both deformation and metamorphism seem to increase in intensity from southwest to northeast (Ribeiro & Silva 1983; Munhfi 1990). However, the apparent increase in metamorphic grade could also be related to differential erosion levels from NE to SW. The structure of the IPB has been interpreted in terms of a thin-skinned foreland thrust-andfold belt (Silva et al. 1990; Quesada 1991). Deformation (D1) generated asymmetric folds verging to the SW, which often show transposed bedding on their short limbs, mimicking the structural features of a thrust-belt. Folding was accompanied by development of a penetrative foliation which shows sinistral transection of the

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axial planes. This transection is related to a wrench component of non-coaxial strain during folding (Silva 1983). The largest Devonian outcrops in the IPB lie in the cores of two major antiforms, the Puebla de Guzmfin and Valverde del Camino anticlines. Our study area comprises a part of both structures (Figs 1, 2) and therefore may be considered to be representative of the whole IPB. Two of the main mining districts of the region (Sotiel-Coronada and Tharsis) also lie in the study area.

Sedimentary facies and environments

The Devonian PQ Group is the lowest deposit of regional extent in the IPB. Its sedimentary and petrographical characteristics are similar along the whole IPB, both in Spain and Portugal. It comprises a monotonous detrital sequence of shales and sandstones with a slightly variable lutite/arenite ratio, always >1 except in the upper part of the Group. Thicknesses of single sandstone beds range from 5 to 40 cm. Sedimentary structures such as grading, parallel lamination, planar cross-bedding and bimodal crossbedding are common. Bioturbation is also commonly observed; Skolithos, Nereites and Lophoctenium trails have been identified (Van

den Boogaard 1967). Petrographically the sandstones are quartzarenites and quartzwackes according to the classification of Pettijohn et al. (1972). The total thickness of the group is not known. Rocks of PQ Group were deposited on a shallow marine platform, probably storm-dominated, but a detailed sedimentological study of the complete PQ Group is still to be made. The depositional style near the top of the PQ Group displays a major break from the sedimentary homogeneity described above. Lutite/ arenite ratios show a rapid decrease, and the former Devonian shaly series (with sandy intercalations) changes into a mainly sandy sequence. Moreover, patches of exotic sediments appear among these sandstones, so that the top of the PQ Group exhibits a mosaic of apparently unrelated facies. Four facies associations typify this upper part of the PQ Group: (1) clastic shallow-marine facies association; (2) delta facies association; (3) facies of sediment gravity flow deposits; (4) calcareous shallow-marine facies association. All four facies associations lie at the same stratigraphic level and represent the Famennian Dinantian boundary in the IPB, although in the study area (Fig. 2) calcareous shallowmarine facies are poorly represented. Small, sparsely distributed, calcareous lenses do occur, but the poor outcrop makes them unsuitable for detailed study.

156

C. MORENO ET AL.

Clastic shallow-marine facies Sedimentary rocks of the clastic shallow-marine facies form tabular bodies of sandstones varying in thickness from 3 to 100 m, which commonly form topographic highs. These tabular bodies are built up by interlocking stacked bedforms that vary in size from 1-10 m in length and 0.11 m in height. No internal structures have been observed, but these bedforms are interpreted as megaripples and sandwaves. Overlying these are tabular beds of quartzite (<60cm thick) rich in sedimentary structures such as parallel and bimodal cross-bedding, planar low-angle crosslamination, hummocky cross-stratification, and different types of wave ripple marks (straight, linguoid and rare fiat-topped ripples). Reactivation surfaces are also abundant. The rocks are fine-grained clean quartzarenites made up of a framework of rounded quartz grains and subordinate microquartzite, tourmaline, rutile and zircon. This strong textural and mineralogical maturity points to a sedimentary environment characterized by the vigorous action of tractive currents, which have removed the fine fraction by washing processes (Davis & Ethridge 1975; Odom et al. 1976; Ricci Lucchi 1985), and a mature source area. In addition, the association of sedimentary structures in the sandstone beds suggests a shallow basin dominated by wave and storm current activity (Harms et al. 1975; Hunter & Clifton 1982; Swift et al. 1983; Duke 1985; DeCelles & Cavazza 1992). Both geometric features and stacking forms of these shallow-facies sediments suggest deposition as wave-influenced sand ridges. The lower part of these bars is represented by bedform superposition, probably originating from major storms. The upper part of the sand ridges, corresponding to the tabular thin layers described above, was continously reworked by fair-weather waves. Similar sandbodies are common throughout the geological record from the Precambrian up to the present day (Banks 1973; De Raaf et al. 1977; Dabrio 1982; MacCarthy 1987; Driese et al. 1991; Rine et al. 1991). Sedimentary environments of the same type have been reported previously in the Carboniferous of the IPB (Moreno 1985,1993).

Delta fa cies association Huge sandstone and conglomerate deposits form two large isolated hills at Virgen de la Pefia near Puebla de Guzm~m (Fig. 2). The delta

facies association is composed of quartzarenite, litharenite, conglomerate and shale lithofacies. The bulk of this facies association is made up of quartzites containing litharenite, conglomerate, and shale intercalations. The quartzarenite lithofacies. This lithofacies is characterized by sandstones that are similar in composition, texture, and structure to those in the shallow-marine facies described above. In addition, burrowed levels, in which mainly Palaeophycus, Planolites and other unidentified trace fossils are abundant, and are readliy seen due to the excellent outcrops. The quartzarenite lithofacies constitutes a shallow-marine facies and is interpreted as a series of bars formed in a shallow-marine nearshore environment. The litharenite lithofacies. This unit is characterized by fine- to medium-grained sandstones composed of monocrystalline quartz, chert and subordinate tourmaline, biotite, muscovite, zircon, and opaque minerals, all as monomineralic grains. The lithic components are mainly micaceous quartzites and Fe-Mn oxide fragments. Matrix content of the litharenites ranges from 10-50% (many are technically wackes) including both primary matrix and diagenetic pseudomatrix. The stable components are the same in both lithofacies (quartzites and litharenites) indicating the same provenance; petrographical differences, especially the presence of lithic fragments and abundant matrix, are related to sedimentary reworking processes. At the outcrop scale this lithofacies forms tabular bodies made up of interbedded litharenite and shale sheets. The internal structure of the litharenite beds is characterized by horizontal and undulate lamination, trough crosslamination, hummocky cross-stratification and wave ripples. Strongly bioturbated levels occur. This facies is interpreted as storm-generated sand layers interbedded with fair-weather muds. The conglomerate lithofacies. The conglomerates consist mainly of the clast-supported type. They are polymictic and consist of rounded clasts, generally in contact, set in a sandy matrix. Clasts range from 1 to 25cm in diameter. They are composed of a high percentage of fine- to medium-grained quartzarenites (showing in places an older tectonic foliation), Fe-Mn oxide pebbles derived from destruction of a duricrust (Routhier et al. 1980; Moreno & S~ez 1991), feldspathic arenite pebbles and rare shale clasts.

FAMENNIAN-DINANTIAN CATASTROPHISM Clast-supported conglomerates form channelized bodies with scoured bases, intercalated with the quartzarenites. Four wedge-shaped bodies occur, pinching out to the west (Fig. 3). The thickness of the bodies is variable, but commonly ranges between 1 and 15m. Rare, small isolated lenses also occur. The most prominent internal sedimentary structure is a crude horizontal stratification marked by levels with concentrations of F e - M n oxide pebbles. Roughly graded bedding is common. This facies is interpreted as a channel-fill deposit. Two layers of mud-supported conglomerates also occur in the delta facies association as a subordinate lithofacies. At outcrop, they have tabular geometry with planar, non-erosive bases. The layers are 0.4 m and 3 m thick and show no internal structure. The mud-supported conglomerates are composed of sparse, rounded quartzarenite boulders in a shaly matrix. The occurrence of these mud-supported conglomerate layers seems to have no relationship to the clast-supported conglomerates. They have been interpreted as sporadic mass flows. Unidentified plant debris also appears randomly intercalated among the rocks of the delta facies association.

157

Interpretation. All these facies and facies associations were previously interpreted (Moreno & Sfiez 1990) as wave-reworked coastal bars, in which littoral currents controlled the facies and sediment distribution. A more detailed analysis of facies distributions and the geometric features of the sedimentary bodies led us to re-interpret the whole as part of a wave-dominated coastal delta environment. The major pieces of evidence for this delta system are: (a) the small size of the deposystem; (b) the distribution and geometry of the detrital bodies (Fig. 3); (c) the sedimentary features of each of the individual lithofacies; and (d) the continental provenance of the oxide clasts and plant debris. In this scenario the quartzarenite facies would represent the general, background style of sedimentation; the quartzarenite bar geometry suggests an undulating marine floor with ridges and runnels, where some sediments could be continuously reworked by the waves, whereas others - the litharenite facies remained undisturbed in areas of long-term sediment storage. Predominance of sandy, wave-reworked facies, along with the lack of steeply-dipping foresets, is interpreted as adjustment of this shallow, gently-sloping marine floor. This continuous sedimentation process was

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punctuated by conglomerate deposition, which represents tectonic pulses in the continental source. The clast-supported conglomerates represent the infilling of rapidly prograding channels cutting across the sand bars. Mud-supported conglomerates suggest sporadic deposit of mudflows (probably of flash-flood type). The delta facies association represents a smallsize delta complex with an apical feeder system (type A of Postma 1990). This kind of system is

characterized by radial spreading of sediments and 'commonly by ephemeral, unconfined streams involving mass-flows, [these] are relatively small in radius and form along a basinmargin fault-scarp' (Postma 1990). We suggest that the preserved remains of the whole system represents the submarine portion of the fandelta with a shoal-water profile. The portion corresponding to the 'fan' is absent, and has been either eroded or tectonically displaced, so

FAMENNIAN-DINANTIAN CATASTROPHISM that its behaviour was probably similar to the Mediterranean fan-deltas described by Bardaji et al. (1990).

Facies of gravity flow deposits Rocks of this facies crop out throughout the study area, although the best exposures are near Sotiel village and mine (Fig. 2). Extensive outcrops are more than 3km long and 50m thick. They are formed of large-scale matrixsupported conglomerates different from those found in the delta sediments. These conglomerates are poorly-sorted, with angular monogenic clasts of quartzarenite ranging in size from granules to boulders (0.5-18cm across). Clasts show internal sedimentary structures such as parallel and cross-lamination, and wave ripples; texturally and mineralogically they are similar to quartzites from the shallow-marine facies associations. These clasts are isolated in a shaly matrix which comprises 85-90% of the whole rock. In some cases, bodies of matrix-supported conglomerates are over 50 m thick. At first sight they appear to have no recognizable internal stratigraphy, but the vertical distribution of clasts (Fig. 4) seems to show a crude stratification. Individual layers reach a maximum of 3 m in thickness. Also present, enclosed in a shaly matrix, are large isolated boulders of quartzarenite up to 4m in diameter, and slumped layers of quartzarenite sequences similar to the sandy shallow bar facies described above. This facies is interpreted as the result of sediment gravity-flow deposition, i.e. by dense and highly viscous flows exhibiting plastic behaviour. These are cohesive debris flows (Middleton & Hampton 1976; Lowe 1982; Stow 1986; Pierson & Costa 1987) coming from intrabasin highs and deposited by flow-freezing mechanisms (Fraser & Suttner 1986). Both the physical characteristics of these flows and the thicknesses of the deposits point to the low spreading capability and high competence of the mass flows (Nemec & Steel 1984; Porebski 1984; Nemec 1990), involving a high angle of slope to begin and maintain movement (Campbell 1989; Bjorlykke 1989). These mega-debris flow deposits are thicker than most found in other ancient successions (Leigh & Hartley 1992). The lithological and geometric characteristics of the clasts (granules to boulders, and even slumped beds) suggest an intrabasinal sediment source from positive topographic features on the basin floor. The thickness

159

of mass-flow units indicates that huge volumes of debris were carried down in each single failure event. The failure surfaces in the marine platform had to be deep enough to involve sediments that were at least partly lithified, suggesting a major role for seismic activity in the genesis of these sediment gravity-flow deposits.

Palaeogeographical inferences Terrigenous shelf sedimentation during the Devonian created the homogeneous PQ Group of the IPB. However, this style of sedimentation was interrupted near the Devonian-Carboniferous boundary by highly energetic sedimentary processes, resulting in rapid lateral changes in both facies and thicknesses. These were preserved in the local geological record as facies and facies associations caused by catastrophic events related to large-scale slope failures. This modification of sedimentary style at the top of the Devonian sequence indicates that the broad, continuous and gently-sloping PQ basin changed into a set of sub-basins with different subsidence rates, and intervening horsts. Such a transformation could occur by rifting owing to either a continental extensional tectonic episode or within a strikeslip regime. Either of these would represent the first manifestation of the Hercynian orogenic cycle in the IPB. Several recent papers (Oliveira 1990; Quesada 1991) concerned with the tectonic evolution of the Hercynian Chain suggest that Hercynian movements began in the Lower Carboniferous. However, previous works by L6colle (1977) and Routhier et al. (1980) highlighted evidence for the occurrence of the Hercynian Bretonic phase (Upper Devonian) in the pre-Dinantian palaeogeography of the IPB, with the formation of regional highs and lows as a consequence of 'long radius epirogenic movements'. Our evidence indicates that the lack of uniformity found in the sediments at the Devonian-Carboniferous boundary in the IPB is evidence of the beginning of the Hercynian Orogeny. As long as the higher zones on the already fractured IPB basin remained linked to the emergent continent, shallow-water sedimentation continued. Locally, this littoral sedimentation was interrupted by episodic, catastrophic influx of sediment and the formation of fandeltas. The deepest areas acted as traps where the accumulation of mass flows derived from the shallow-water horsts occurred. These facies and facies associations are characteristic of

C. M O R E N O ET AL.

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convulsive/catastrophic failures related to m a j o r seismic events. Fig. 5 shows an idealized model for these depositional environments in the IPB basin at or a b o u t the D e v o n i a n - C a r b o n i f e r o u s boundary. We wish to thank Inmaculada Gonzalez and M6nica Martin for their company and help during the fieldwork. Critical reviews by F. Barriga, J. M. Leistel and P. Strogen improved the paper and the English text. Financial support for this work was provided by Plan Andaluz de Investigaci6n, PAl, Grupos: 4112 and 4076.

References

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CAMPBELL, C. S. 1989. Self-lubrication on long runout landslides. Journal of Geology, 97, 653-665. DABRIO, C. J. 1982. Sedimentary structures generated on the foreshore by migrating ridge and runnel systems on microtidal and mesotidal coast of Southern Spain. Sedimentary Geology, 32, 141-151. DAVIS, D. K. & ETHRIDGE, F. G. 1975. Sandstone composition and depositional environment. American Association of Petroleum Geologists Bulletin, 59, 239-264. DE RAAF, J. F. M., BOERSMA,J. R. & VAN GELDER, m. 1977. Wave-generated structures and sequences from a shallow marine succession, Lower Carboniferous, County Cork, Ireland. Sedimentology, 24, 451-483. DECELLES, P. G. & CAVAZZA, W. 1992. Constraints on the formation of Pliocene hummocky crossstratification in Calabria (Southern Italy) from consideration of hydraulic and dispersive equivalence, grain-flow theory and suspended load fallout rate. Journal of Sedimentary Petrology, 62, 555-568. DRIESE, S. G., FISCHER, M. W., EASTHOUSE, K. A., MARKS, G. T., GOGOLA, A. R. & SCHONER, A. E. 1991. Model for genesis of shoreface and shelf sandstone sequences, southern Appalachians: palaeoenvironmental reconstruction of an Early Silurian shelf system. In: SwiFT, A. J. P., OERTEL, G. F., TILLMAN, R. W. & THORNE, J. A. (eds) Shelf Sands and Sandstone Bodies. Special Publication of the International Association of Sedimentologists, 14, 309-338.

FAMENNIAN-DINANTIAN DUKE, W. L. 1985. Hummocky cross-stratification, tropical hurricanes, and intense winter storm. Sedimen tology, 32, 167-194. ETHRIDGE, F. G. 1985. Modern alluvial fans and fan deltas. In: FLORES, R. M., ETHRIDGE, F. G., MIALL, A. D., GALLOWAY,W. E. & FOUCH, T. D. (eds) Recognition of Fluvial Depositional Systems and their Resource Potential. SEPM Short Course Notes, 19, 101-126. FRASER, G. S. & SUTTNER, L. 1986. Alluvial Fans and Fan Deltas. A Guide to Exploration for Oil and Gas. International Human Resources Development Corporation, Boston. GALLOWAY, W. E. 1985. Ancient alluvial fans and fans deltas. In: FLORES, R. M., ETHRIDGE, F. G., MIALL, A. D., GALLOWAY,W. E. & FOUCH, T. D. (eds) Recognition of Fluvial Depositional Systems and their Resource Potential. SEPM Short Course Notes, 19, 127-143. HARMS, J. C., SOUTHARD, J. B., SPEARING, D. R. & WALKER, R. G. 1975. Depositional Environments as Interpreted from Primary Sedimentary Structures and Stratification Sequences. SEPM Short Course Notes, 2. HUNTER, R. E. & CLIFTON, H. E. 1982. Cyclic deposits and hummocky cross-stratification of probable storm origin in Upper Cretaceous rocks of the Cape Sebastian area, Southwestern Oregon. Journal of Sedimentary Petrology, 52, 127-143. JULIVERT, M., FONTBOTI~, J. M., RIBEIRO, A. & CONDE, L. E. 1974. Mapa Tect~enico de la Peninsula Ib&ica y Baleares, escala 1 : 1 000000. Instituto GeolSgico y Minero de Espafia, Memoria explicativa, 113 p. KLEVERLAAN, K. 1987. Gordo Megabed: a possible seismite in a Tortonian submarine fan, Tabernas Basin, Province Almeria, Southeast Spain. Sedimentary Geology, 51, 161 180. LI~COLLE, M. 1977. La Ceinture Sud-Ib~rique: un Exemple de Province dt Amas Sulfur& volcanosedimentaries. PhD Thesis, Universit+ Pierre and Marie Curie, Paris. LE1GH, S. & HARTLEY, A. J. 1992. Mega-debris flow deposits from the Oligo-Miocene Pindos foreland basin, western mainland Greece: implications for transport mechanisms in ancient deep marine basins. Sedimentology, 39, 1003-1012. LOWE, D. R. 1982. Sediment gravity flows: II. Depositional models with special reference to the deposits of high-density turbidity currents. Journal of Sedimentary Petrology, 52, 279-297. MACCARTHY, I. A. J. 1987. Transgressive facies in the South Munster Basin, Ireland. Sedimentology, 34, 389-422. MASTALERZ, K. & WOJEWODA, J. 1993. Alluvial fan sedimentation along an active strike-slip fault: Plio-Pleistocene Pre-Kaczawa fan, SW Poland. In: MARZO, M. & PUIGDEFRABREGAS, C. (eds) Alluvial Sedimentation. Special Publication of the International Association of Sedimentologists, 17, 293-304.

-

CATASTROPHISM

161

MIDDLETON, G. V. & HAMPTON, M. A. 1976. Subaqueous sediment transport and deposition by sediment gravity flows. In: STANLEY, D. G. & SWIFT, D. J. (eds) Marine Sediment Transport and Environmental Management. John Wiley, New York, 197-217. MORENO, C. 1985. Shallow Lower Carboniferous "Culm" facies in the Iberian Pyrite Belt (SW Spain). 6th European Regional Meeting of Sedimentology, International Association of Sedimentologists, Lleide, Spain, 309-312. 1993. Postvolcanic Paleozoic of the Iberian Pyrite Belt; an example of basin morphologic control on sediment distribution in a turbidite basin. Journal of Sedimentary Petrology, 63, 1118-1128. & SAEZ, R. 1990. Sedimentaci6n marina somera en el Dev6nico del anticlinorio de Puebla de Guzm/m, Faja Piritica Ib6rica. Geolgaceta, 8, 62-64. & 1991. The paleogeographic significance of shallow marine facies in the PQ Group of the Iberian Pyrite Belt, Spain. International Association of Sedimentolog&ts 12th Regional Meeting, Bergen, Abstracts, 37. MUNHA, J. 1990. Metamorphic evolution of the South Portuguese/Pulo do Lobo Zone. In: DALLMEYER, R. D. & MARTINEZ-GARCIA, E. (eds) PreMesozoic Geology of Iber&. Springer Verlag, Berlin, 363-368. NEMEC, W. 1990. Aspects of sediment movement on steep delta slopes. In: COLELLA,A. & PRIOR, D. B. (eds) Coarse-Grained Deltas. Special Publication of the International Association of Sedimentologists, 10, 29-73. & STEEL, R. J. 1984. Alluvial and coastal conglomerates: their significant features and some comments on gravelly mass-flow deposits. /n: KOSTER, E. H. & STEEL, R. J. (eds) Sedimentology of Gravels and Conglomerates. Memoir of the Canadian Society of Petroleum Geologists, 10, 1-31. ODOM, I. E., DOE, T. W. & DOTT, R. H. JR. 1976. Nature of feldspar grain size relations in some quartz-rich sandstones. Journal of Sedimentary Petrology, 46, 862-870. OLIVEIRA, J. T. 1990. Stratigraphy and synsedimentary tectonism. In: DALLMEYER, R. D. & MARTINEZ-GARCIA, E. (eds) Pre-Mesozoic Geology of lberia. Springer-Verlag, Berlin, 334-347. --, HORN, M. & PAPROTH, E. 1979. Preliminary note on the stratigraphy of the Baixo Alentejo Flysch Group, Carboniferous of Southern Portugal and on the palaeogeographic development, compared to corresponding units in Northwest Germany. Comunica,coes Serci,cos Geol6gicos de Portugal, 65, 151-168. PETTIJOHN, F. J., POTTER, P. E. & SILVER, R. 1972. Sands and Sandstones. Springer-Verlag, Berlin. PIERSON, C. T. & COSTA, J. E. 1987. A rheologic classification of subaerial sediment-water flows. In: COSTA, J. E. & WlECZOREK, G. F. (eds) Debris Flows~Avalanches." Process, Recognition, and Mitigation. Geological Society of America, Reviews in Engineering Geology, VII, 1-12. -

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PILKEY, O. H. 1988. Basin plains; giant sedimentation events. In: CLIFTON, H. E. (ed.) Sedimentologic Consequences of Convulsive Geologic Events. Geological Society of America Special Paper, 229, 93-99. POREBSK1, S. J. 1984. Clast size and bed thick-ness trends in resedimented conglomerates: example from a Devonian fan-delta succesion, southwest Poland. In: KOSTER, E. H. & STEEL, R. J. (eds) Sedimentology of Gravels and Conglomerates. Memoir of the Canadian Society of Petroleum Geologists, 10, 399-411. POSTMA, G. 1990. Depositional architecture and facies of river and fan deltas: a synthesis. In: COLLELA,m. & PRIOR, D. B. (eds) Coarse-Grained Deltas. Special Publication of the International Association of Sedimentologists, 10, 13-27. QUESADA, C. 1991. Geological constraints on the Paleozoic tectonic evolution of tectonostratigraphic terranes in the Iberian Massif. Tectonophysics, 185, 225-245. RIBEIRO, m. & SILVA, J. B. 1983. Structure of South Portuguese Zone. Memorias Servi,cos Geol6gicos do Portugal, 29, 83-90. RICCI-LUCCHI, F. 1985. Influence of transport processes and basin geometry on sand composition. In: ZUFFA, G. G. (ed.) Provenance of Arenites. NATO ASI Series 148, D. Reidel Publishing Company, 19-45. RINE, J. M., TILLMAN, R. W., CULVER, S. J. & SWIFT, m. J. P. 1991. Generation of late Holocene sand ridges on the middle continental shelf of New Jersey, U S A - evidence for formation in a mid-shelf setting based on comparisons with a nearshore ridge. In: SWIFT, D. J. P., OERTEL,G. F., TILLMAN, R. W. & THORNE, J. A. (eds)

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Dinantian depositional environments along the northern margin of the Solway Basin, UK K E L L Y M A G U I R E ], J I L L I A N T H O M P S O N n & S T U A R T G O W L A N D 2

1Geochem Group Limited, Chester Street, Chester CH4 8RD, UK 2 Ichron Limited, 16 Dalby Court, Gadbrook Business Centre, Rudheath, Northwich, Cheshire CW9 7TN, UK Abstract: A sedimentological study has been conducted on the Dinantian strata of the tectonically separated Rerrick and Kirkbean Outliers on the northern margin of the Solway Basin. The Rerrick succession, comprising a sequence of largely non-marine alluvial fan, fluvial and coastal plain clastic rocks, represents a tectonically influenced lateral dispersal system sourced from the northern margin of the Solway Basin. In contrast, the Kirkbean succession is considered to represent the deposits of a tectonically stable shallow-marine shelf subject to both limestone and clastic deposition. Periodically, axial fluvial systems draining westward established delta front conditions on the shelf. The depositional model presented here suggests both lateral and axial dispersal systems were active within the Solway Basin during the Dinantian.

The Solway Basin is an elongate ENE-WSW structure, 25-35km wide, extending along the line of the present-day Solway Firth between northwest England and southwest Scotland

(Fig. 1). This basin contains a Dinantian succession which probably exceeds 5km in thickness at the basin centre (Chadwick et al. 1993; Fig. 2). Along the northern margin of the

2~ location o4 Fig. 2

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From STROGEN, P., SOMERVILLE, I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107 pp. 163-182.

164

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DINANTIAN ENVIRONMENTS, SOLWAY BASIN basin, Lower Carboniferous successions are exposed in five small isolated outliers: Rerrick, Castle Point-Gutcher's Isle, Portling-Portowarren, Southwick Merse and Kirkbean (Fig. 3). The objective of this paper is to demonstrate the wide range of Dinantian depositional environments within the Solway Basin and to describe the evolution of these environments through time. This model is based upon the detailed sedimentological analysis of two of the tectonically separated successions, Rerrick and Kirkbean. The first of these is the Rerrick Outlier, which comprises a succession of largely nonmarine clastic rocks. In contrast, the Kirkbean Outlier exhibits a mixed shallow-marine clasticcarbonate succession. Correlation between these two tectonically separated outliers is hampered by the largely non-marine nature of the Rerrick succession. However, Craig & Nairn (1956) obtained a number of marine fossils including Syringothyris cuspidata and Dictyoclostus teres from the Orroland Limestone Beds of the Rerrick succession (Barlocco Heugh Formation of Deegan 1973). This provides a single tentative biostratigraphical link with the Southerness Beds of the Kirkbean Outlier, where the base of the Syringothyris Limestone was taken as the base of the Chadian by George et al. (1976).

Structural setting T h e present-day Solway Basin outcrop is bounded to the north by the Southern Uplands

165

Massif, comprising Lower Palaeozoic sediments and Devonian granite intrusions, and to the south by the Lake District, also a Lower Palaeozoic Block. To the northeast, the basin is separated from the Northumberland Trough by the north-south trending Bewcastle Anticline (Bewcastle Shelf of Turner et al. 1993). Southwest of here, the Carlisle Basin forms an easterly component of the Solway Basin, being represented by a major half-graben with a westwardthickening hanging wall sequence (Dailly 1990). Although onshore the southern margin of the Solway Basin abuts against the Lake District Block, offshore it is represented by the northern flank of the Ramsey-Whitehaven Ridge (Johnson 1984; Fig. 1). The western margin is taken on the basis of regional gravity data and indicates a marked westward thinning away from Carboniferous/Permo-Triassic depocentres (Bott 1964; Lee 1986; Dailly 1990). The main structural lineaments bounding both the Solway Basin and the Carlisle Basin occur along an ENE trend, and are represented principally by en echelon, syndepositional normal faults which appear to have been most active in early Carboniferous times (Chadwick et al. 1993). Two main fault systems are recognized: (i) North Solway-Water Beck-Gilnockie Fault System to the north; and (ii) the MaryportStublick Fault System in the south (Figs 1, 2). Eastwards these fault systems pass into en echelon systems defining the margins of the Northumberland Trough (Chadwick et al. 1993).

Table 1. Lithostratigraphy of the Rerrick Outlier (amended after Deegan 1973) Group

Formation

Thickness (m)

Depositional environment

Rascarrel Group

Lochenling Rascarrel Burnfoot Castle Muir

60 210 90

Pebbly alluvial fan Braided fluvial channel Braidied fluvial channel

Orroland Group

Black Neuk Scar Heugh Orroland Lodge Barlocco Heugh Dropping Craig Spouty Dennans Hanged Man

100 12 9 10 25 18 90

Pebbly alluvial fan Mixed sand-mud alluvial plain Sandy marine shoreface Mixed sand-mud alluvial plain Mixed sand-mud alluvial plain Mixed sand-mud alluvial plain Pebbly alluvial fan

Wall Hill Sst. Group

Abbey Head

40

Sheep Bught

120

White Port

210

Mixed sand-mud alluvial plain (variably congomeratic) Mixed sand-mud alluvial plain (variably conglomeratic) Mixed sand-mud alluvial plain (variably conglomeratic)

K. MAGUIRE ET AL.

166

Rerrick Outlier

Mixed sand-mud alluvial plain lithofacies

The lithostratigraphy of the Rerrick Outlier is presented in Table 1. Deegan (1973) has divided the Rerrick succession into three groups: the Wall Hill Sandstone Group, the Orroland Group and the Rascarrel Group. Importantly, however, the boundaries of these groups are defined by faults. Three main depositional environments have been recognized within the Orroland Group: mixed sand-mud alluvial plain, sandy marine shoreface, and pebbly alluvial fan. These environments are best demonstrated by the Barlocco Heugh, Orroland Lodge, Scar Heugh and Black Neuk Formations, from each of which a representative section is presented and discussed in the following paragraphs.

Sequences interpreted as having been deposited within fluvial channels and associated interfluves have been identified within the Barlocco Heugh and Scar Heugh Formations. Representative sections through these formations are presented in Figs 4 and 5.

Environment

Barlocco Heugh Formation. Although Deegan (1973) records a thickness of more than 9m for this formation, only the upper 4 m are described here (Fig. 4). This section is dominated by a 3m thick composite sandstone body enclosed in siltstone. It is erosively-based, but with low relief, and appears to comprise vertically stacked and amalgamated beds up to 60 cm thick. Internally, the sandstones are generally poorly-sorted, SEDIMENTOL~L KEY

13m

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"1FORESHORE

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LOWER SHOREFACE

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locally silty, and range in grain size from very fine to very coarse sand grade. Primary sedimentary structures are represented by horizontal lamination, medium-scale trough cross-stratification and, more locally, current ripple cross-lamination. Granule and pebble grade components (volcanic rocks, vein quartz, red-brown sandstones, siltstones and mudstones) occasionally drape foresets and shallow scours, and also occur as a distinctive lag deposit at the base of the sandstone body. The tops of some beds display desiccation cracks and

167

nodular calcite (caliche) resulting from pedogenic processes• Siltstone units developed above and below the sandstone body are around 30 cm thick, grey-brown to red-brown in colour and show some evidence of relict parallel lamination.

Scar Heugh Formation. This overlies the Orroland Lodge Formation and is in turn overlain by the Black Neuk Formation. Although Deegan (1973) records a total thickness of over 27 m for this formation, much of the lower part of the unit comprises sandy marine shoreface deposits more readily assigned to the underlying Orroland Lodge Formation. It is here estimated to be approximately 12 m thick, of which the uppermost 4 m of section is well exposed, comprising pinkish grey and red-brown calcitic sandstones with highly subordinate red-brown siltstones (Fig. 5). The sandstones are moderately-sorted, occur in 30-60 cm thick beds and range in grain size from very fine to coarse sand grade, although occasional pebbly horizons (comprising clasts of volcanic rocks, vein quartz, red-brown sandstone, siltstone and claystone) are also present. Individual beds display sharp, erosive bases and occasionally fining-upward grain-size profiles. Horizontal lamination and trough cross-stratification predominate, although ripple cross-lamination is locally preserved in the finer-grained sandstones. For the most part, primary sedimentary structures have been obliterated by nodular caliche development. The sporadic siltstone developments range in thickness up to 35cm, and horizontal lamination is locally preserved against a mottled "background' attributed to pedogenic calcite development. These sedimentological features suggest deposition of these formations within an alluvial plain setting, subject to frequent sheetflood events during periods of high fluvial discharge. The erosively-based sandstone units are interpreted as the depositional products of initially highvelocity, poorly-confined flows. Deceleration and waning flow conditions are likely to account for the locally observed fining-upward profiles. Intervening siltstone units are attributed to distal floodbasin deposition characterized by very low time-averaged rates of sedimentation. Discrete channel-fills are rarely observed in these sequences, although within the studied Barlocco Heugh interval, the lower half of the sandstone body displays no evidence ofcaliche development and, with its erosive, pebble-strewn base and occasional medium-scale cross-stratification, could conceivably represent a fluvial channel deposit. Unfortunately, the outcrop at this stratigraphic level is not sufficient to enable

168

K. MAGUIRE ET AL.

resolution of channel scale and type; only a fill of dune and small bar-scale structures is evident. Deegan (1973) also recorded discrete finingupward cycles, averaging 2 m in thickness, elsewhere within the Scar Heugh Formation, which could well indicate fluvial channel development.

Sandy marine shoreface lithofacies The Orroland Lodge Formation section studied (Fig. 4) is represented by a composite sandstone body some 9m thick. This displays a sharp, possibly erosive contact with the alluvial plain siltstones of the topmost Barlocco Heugh Formation, although there is no evidence for a marked hiatus (Fig. 4). It largely comprises moderately to poorly sorted, medium- to very coarse-grained sandstones, which are locally supplemented by pebbly horizons comprising a range of lithoclasts: red-brown sandstone, siltstone, claystone, volcanic clasts and pebbles composed of vein quartz and reworked caliche. The lower 4.5m of the sandstone body is characterized by horizontal lamination, mediumscale trough cross-stratification and ripple cross-lamination, although upwards these are increasingly obliterated by intense biogenic reworking. The lowest 1.8m displays a finingupward sequence terminating in 7.5cm of siltstone. Deegan (1973) reported the presence

of abraded coral, bivalve and gastropod remains in the bioturbated section. Some thin beds preserve no primary fabric whatsoever. Skolithos, Diplocraterion and Arenicolites are the main trace fossils present, although Chondrites occurs locally (Fig. 6). The upper 4.5 m of the sandstone body is characterized by very low angle planar cross-stratification reminiscent of beachface accretion surfaces. Horizontal lamination and broad, shallow, scour-and-fill structures are also present. Clastic sequences within this formation are considered to be representative of sandy shoreface deposition. The trace fossil assemblage preserved within the lower part of the sequence is considered representative of the Skolithos ichnofacies (Pemberton et al. 1992). This, together with the sedimentary structures and abraded marine fossils, is suggestive of deposition in a shoreface environment characterized by low net sedimentation rates. An absence of bioturbation within the topmost 4.5m of the studied interval, combined with low-angle stratification, points to relatively high incident wave (or current) strength and the development of beachface accretion surfaces in coarse-grained pebbly sands. The sharp base to the sandstone body and the overlying 1.8m thick finingupward sequence are attributed to initial marine transgression (lower shoreface) and sustained deepening over the underlying alluvial

Fig. 6. Bioturbated lower shoreface deposits from the Orroland Lodge Formation (Orroland Group, Rerrick Outlier, North Solway, grid ref. NX77754630). Crudely flat-bedded sandstones display discrete vertical Skolithos burrows. Hammer is 40 cm long.

DINANTIAN ENVIRONMENTS, SOLWAY BASIN plain. Above the siltstone the shoreface sequence is essentially progradational in nature. However, the prominent pebble and cobble horizon developed 30cm above the siltstone sequence at 6.2m (Fig. 5) is likely to record a major superimposed transgressive event. This horizon is probably a relict transgressive lag deposit and as such constitutes the only record of a former progradational shoreface sequence destroyed through erosional retreat. Both the sandstone bed directly below this horizon and the sandy matrix between the pebbles are extensively bioturbated by Skolithos (trace fossil omission suite).

Pebbly alluvialfan lithofacies The Black Neuk Formation consists of coarsegrained clastic deposits overlying the Scar Heugh Formation in the studied section. Although the contact between these two lithostratigraphical units is sharp, there is no evidence for a marked hiatus. The representative section through the lower part of the Black Neuk Formation, derived from the coastline of the Rerrick Outlier, is 7 m thick (Fig. 5). Above this, the upper part yields stunning cliff faces illustrating lateral facies variations (Fig. 7). This section covers only a small part of the total Black Neuk Formation sequence; Craig & Nairn

169

(1956) recorded over 70m of sediment, whilst Deegan (1973) quoted around 100 m. The studied interval comprises an interbedded assemblage of sandstones and conglomerates. The conglomerates largely occur within deep erosive scours cut into either sandstones or older conglomeratic units, but also display laterally extensive geometries (scale of tens of metres) with little erosional down-cutting (Fig. 5). Both matrix- and clast-supported varieties are present (sometimes within the same channel-fill). Although generally of pebble grade, clasts range up to boulder size, with the larger clasts often well-rounded and the smaller pebble grade clasts typically angular to subangular in form. A variety of clast types are present, including redbrown sandstone, siltstone, claystone, redbrown (diagenetically altered?) volcanic clasts, vein quartz, dark grey dolerite clasts, and undifferentiated purplish-grey igneous clasts with feldspar phenocrysts. The sandstone units, which are poorly-sorted and range in grain size from very fine to very coarse sand, tend to be more laterally persistent than the associated conglomeratic units. However, marked lateral thinning and thickening relationships are present owing to the erosional down-cutting of adjacent conglomeratic channel-fills and also channelling within the sandstones themselves. Generally, channelling is not as evident as within the conglomeratic units,

Fig. 7. General view of pebbly alluvial fan deposits from the Black Neuk Formation (Orroland Group, Rerrick Outlier, North Solway, grid ref. NX78504652). Coarse conglomerates interpreted as chaotic debris flow deposits, infiU a deep erosional scour within coarse gravelly sandstones.

170

K. MAGUIRE ET AL.

being largely confined to broad shallow 'wash outs' on the lateral scale of tens of metres. Primary sedimentary structures are represented by a combination of horizontal lamination, medium-scale trough and planar-tangential cross-stratification in discrete beds up to 1.5m thick. Pebble and granule grade material locally occurs as drapes to foresets and as distinctive lag deposits to small erosional scours. Finingupward motifs occasionally characterize erosively-based units up to 1.8 m thick. This lithofacies is assigned to a pebbly alluvial fan environment. The sandstone units are interpreted as the deposits of laterally extensive sandflats, perhaps primarily deposited by sheetfloods, but subject to more confined (channelized) flows capable of eroding the underlying substratum and transporting sediment in the form of subaqueous dunes and bars. These sandstones appear concentrated in the lower part of the succession where they probably represent a medial fan setting prior to the major avulsion event (channel-switching) inferred to account for the overlying sequence. This event, probably initiated during a period of extremely high discharge, resulted in the emplacement of a major channel complex as part of the lateral migration of the distributary feeder system. Both within and adjacent to this newly emplaced channel system, the overlying conglomeratic units were deposited. The mechanism of deposition may have varied from organized bed-load transport in the form of low amplitude bars and bedforms, to locally disorganized sediment distribution via a range of channelized and non-channelized debris flows.

Kirkbean Outlier The Dinantian Kirkbean Outlier lies on the southeast margin of the Criffel-Dalbeattie granite complex along the northern shore of the Solway Firth (Fig. 3). Craig (1956) divided this mixed clastic-carbonate succession into seven informal intervals (Table 2). His informal nomenclature is retained here. The lowest interval includes red sediments and basaltic lavas, which are arbitrarily assigned to the Old Red Sandstone. The lavas are considered to represent the westernmost extension of the Birrenswark volcanic horizon which has been dated at c. 350 mA (lower-middle Courceyan) by de Souza in George et al. (1976). The remaining six intervals, as defined by Craig (1956), include the Basal Cementstones, Southerness Beds, Gillfoot Beds, Powillimount Beds, Thirlstane

Table 2. Lithostratigraphy of the Kirkbean Outlier (from Craig 1956) Stratigraphic unit

Thickness Depositional (m) environment

>330 Arbigland Group 25 Thirlstane Sandstone 150 Powillimount Beds 13(~200 Gillfoot Beds 150 Southerness Beds >200 Basal Cementstones

Mixed clastic-carbonate marine shelf Braided fluvial channel Mixed clastic-carbonate marine shelf Mixed sand-mud coastal-alluvial plain Mixed clastic-carbonate marine shelf Mixed clastic-carbonate marine shelf

Sandstone and the Arbigland Group (Table 2). The Basal Cementstones are poorly exposed within the area, comprising alternating siltstones, sandstones and limestones, and are not discussed further in this paper. The following section discusses the four main depositional environments (i.e. mixed clastic-carbonate muddy marine shelf, clastic marine shelf, mixed sand-mud alluvial plain, and sandy braided fluvial channel) identified within the remainder of the Kirkbean succession (Table 2).

M i x e d clastic-carbonate muddy marine shelf lithofacies Sequences interpreted to represent muddy marine shelf environments, comprising mixed carbonate and clastic regimes, have been recognized in representative intervals within the lower part of the Southerness Beds, the Powillimount Beds and the upper part of the Arbigland Group (Table 2).

Lower Southerness Beds. The Southerness Beds comprise approximately 150m (Craig 1956) of strata exposed in a broad anticline west of Southerness Point. Representative sections from both the lower and upper parts of the Southerness Beds are illustrated in Figs 8 and 12. Comparison of these sections suggests an upward increase in sandstone lithologies and a concurrent decrease in carbonate lithologies throughout the Southerness Beds as a whole. The lower part of the succession, described here, is attributed to deposition within a mixed clastic-carbonate marine shelf setting, whilst

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Fig. 9. Representative section through the upper part of the Arbigland Group (Kirkbean Outlier, North Solway, grid ref. NX 99755760). Key as in Fig. 4.

the upper part of the Southerness Beds (discussed later) is assigned to a clastic shelf setting. The lower part of the Southerness Beds comprises an interbedded assemblage of grey impure limestones, black shales, 'clean' limestones and rare calcareous sandstones which range in grain-size from very fine to medium sand. Limestone beds range in thickness from

2 c m up to 60 c m a n d c o m m o n l y comprise highly fossiliferous and intensely bioturbated wackestones and packestones. Fossiliferous components are dominated by variably preserved molluscan and brachiopod remains (see Craig 1956 for faunal listings). Bioturbation is generally characterized by non-diagnostic burrow mottling, although well-developed Chondrites and large but non-specific mud-filled horizontal burrows

172

K. MAGUIRE ET AL.

are also identified. The black shales are typically horizontally laminated and range in thickness from < 15 cm to over 1 m. They are generally nonfossiliferous and locally bioturbated, although scattered brachiopod shells are occasionally present. Sandstones within this sequence are confined to just one interval (6-7.75 m, Fig. 8), and exhibit a sheet-like geometry at outcrop scale. They range in grain size from very fine to medium sand grade and are dominated by horizontal lamination, with subordinate planar cross-stratification and current ripple cross-lamination. Upper Arbigland Group. This group was considered by Craig (1956) to comprise at least 330 m of strata, the lowest exposed beds lying in faulted contact (Thirlstane 'Reverse' fault) with the underlying Thirlstane Sandstone. The lower part of the succession is attributed to a clastic shelf facies association (see later), whilst the upper part of the Arbigland Group is assigned to a mixed clastic-carbonate marine shelf. A representative section through the upper part of the succession is displayed in Fig. 9. This sequence, similar in many ways to the lower part of the Southerness Beds, is characterized by interbedded black shales and bioclastic limestones. Shale units range in thickness from only a few cm up to 2 m and generally display horizontal lamination, although in places thin bioturbated horizons are developed (Fig. 9).

Localized concentrations of shelly/crinoidal and coral-rich fossiliferous debris are also present. Conspicuously enveloped within certain shale horizons are excellently preserved hemispherical colonies of Lithostrotion. A variety of both upright in situ (Fig. 10) and overturned forms are present, with diameters commonly up to 75cm. Limestone horizons are pervasively bioturbated, range in thickness from < 2 c m up to a maximum of 75cm and locally exhibit lenticular geometries. A diverse macrofossil assemblage dominated by broken and intact corals, brachiopods and crinoids is present. Petrographically the limestones are poorlysorted wackestones and packestones, often with a terrigenous mud or fine sand component. Powillimount Beds. These beds, which conformably overlie the Gillfoot Beds, were estimated by Craig (1956) to attain a thickness of approximately 150m. A representative 30m thick logged interval from the top of the succession is illustrated in Fig. 11. Although sedimentologically variable, this sequence is similar to those described above, comprising interbedded black shales and grey muddy limestones. Sandstone horizons are more prevalent within the Powillimount Beds and one thin coal horizon (30cm thick), with a rootleted seatearth is also present (Fig. 11). The black shales generally display horizontal lamination,

Fig. 10. An excellently preserved hemispherical colony of Lithostrotion clavaticum enclosed within black shales from the upper part of the Arbigland Group (Kirkbean Outlier, North Solway grid ref. NX99755760). This in situ example is in an upright growth position. Lens cap is 5 cm across.

DINANTIAN ENVIRONMENTS, SOLWAY BASIN although pervasive bioturbation is locally observed. Bed thicknesses range from only a few cm to over 2.5m, with thicker intervals preferentially developed in the upper parts of the Powillimount succession. In shale units below the coal horizon, Craig (1956) recognized a marine fossil assemblage including Spirorbis cf. helicteres, ostracodes, algal fragments and an orthoceratid. The muddy limestones are bioturbated, generally <30cm thick and contain

Environment

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174

K. MAGUIRE ET AL.

The sequences described above broadly reflect deposition in a mixed clastic-carbonate marine shelf environment. Shales represent the relatively continuous background deposition of terrigenous mud below fair-weather wave-base. Carbonate deposition is inferred to have taken place during protracted periods of reduced clastic input. The sandstones are interpreted as distinct event beds attributable to the episodic transportation of sand onto the shelf during storms (Harms et al. 1975; Dott & Bourgeois 1982). These event beds exhibit varying degrees of bioturbation resulting from substrate colonization during periods of quiescence. Their sedimentological/ichnological features are consistent with deposition in a lower shoreface or offshore transition zone setting. Although the environmental setting envisaged for the Powillimount Beds is not too dissimilar to that of the Southerness Beds and the upper part of the Arbigland Group, the presence of a discrete coal horizon, the greater abundance of drifted carbonaceous debris and the more intimate association of limestones and sandstones within the Powillimount Beds suggests deposition in a distal shoreline or shelf region, adjacent to a somewhat embayed and swampy coastal plain. The localized establishment of subaerial conditions is indicated by the development of at least one in situ coal.

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Clastic marine shelf lithofacies Interbedded sandstones and claystones, with inferred marine affinities, have been identified in the upper part of the Southerness Beds and the lower parts of both the Gillfoot Beds and Arbigland Group. These successions display many features in common with the mixed carbonate-clastic regimes described above, but are characterized by a marked abundance of coarse clastic sandstone bodies (Fig. 12).

Upper Southerness Beds. The sandstones within the upper part of the Southerness Beds are enclosed within black shales, and are typically intensely bioturbated, moderately- to poorly-sorted, and very fine- to mediumgrained. The thinner of these sandstone units (<15cm) tend towards the finer end of this grain-size spectrum. Where the fine-grained sandstones are not bioturbated they exhibit current ripple cross-lamination, planar bedding and wave ripple cross-lamination (Fig. 12). The medium-grained sandstones are generally >30cm thick and, in amalgamated form, attain thicknesses of >1 m. These are typified by a

Fig. 12. Representative section through the upper part of the Southerness Beds (Kirkbean Outlier, North Solway, grid ref. NX 98005525). Key as in Fig. 4. variety of well-defined sedimentary structures, minimal bioturbation and distinct fining-upward motifs. Each fining-upward trend is accompanied by changes in bedform configuration. Trough and planar cross-stratified mediumgrained sandstones are succeeded by horizontally stratified and current ripple cross-laminated fine-grained sandstones.

Lower Gillfoot Beds. The Gillfoot Beds conformably overlie the Southerness Beds and comprise between 130-200m of red-purple

DINANTIAN ENVIRONMENTS, SOLWAY BASIN sandstones, conglomerates and shales, with minor limestones (Craig 1956; Fig. 13). Two depositional environments are recognized; the lower, which has marine shelf affinities, is discussed here, whilst alluvial deposits represented within the Gillfoot Beds, are discussed below. Within the basal part of the Gillfoot Beds alternating sandstones and red shales are characterized by non-diagnostic burrow-mottling. One sandstone unit, known locally as the 'Detrital Fossil Band' (Craig 1956; Fig. 13), contains high concentrations of broken fossil debris including an abundant fragmented brachiopod fauna enclosed within a red, poorly sorted, silty sandstone. This horizon also records the lowest occurrence of Lithostrotion within the Kirkbean succession. Elsewhere, the sandstones are thinly-bedded, very fine- to medium-grained, and locally display remnant cross-stratification and rare current ripple cross-lamination, although much of the primary fabric appears to have been destroyed by bioturbation.

Arbigland Group. Sandstones within the lower part of this group, also enclosed within black shales, generally exhibit sheet-like geometries (Fig. 14). They are moderately to poorly sorted and range in grain-size from very fine to medium sand grade. Bed thicknesses range from 15-75cm, although in places stacked bioturbated sandstone units exceed thicknesses of 1.5m. In these cases, gross coarseningupward profiles are developed with a vertical progression from very fine- to medium-grained sandstones. Carbonaceous woody debris is a common component of these sandstones, and shelly material is occasionally concentrated at the bases of sandstone beds. Non-diagnostic burrow mottles are common, although welldeveloped Chondrites and Diplocraterion are locally identified. In addition to the bioturbated sandstones, a range of sharp and erosivelybased current-bedded sandstones are also present within the lower parts of the Arbigland Group. Low angle cross-stratification and hummocky cross-stratification predominate, with subordinate current ripple cross-lamination and rare trough cross-stratification. Although the gross geometry of these sandstone units is generally sheet-like, internal beds are commonly lenticular (Fig. 14). Given the evidence that storm processes were at times operating on the marine shelf, it is possible that a large proportion of the bioturbated sandstones recognized within these sequences represent biogenically reworked sands transferred onto the shelf by relatively high-energy

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K. MAGUIRE ET AL.

176

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limited, of coastal conditions, which contrasts with the generally open-marine shelf conditions inferred for the mixed clastic-carbonate successions detailed earlier. Fining-upward sandstone sequences may be locally indicative of abandoned barforms but, where distinct erosive bases are recorded, it is possible that they are representative of submarine rip-channels excavated on the shelf during storms. Similarly, a tidal channel origin cannot be ruled out, despite the absence of diagnostic tidal bedforms within these rocks.

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Fig. 14. Representative section through the lower part of the Arbigland Group (Kirkbean Outlier, North Solway, grid ref. NX99355700). Key as in Fig. 4. storm-generated currents. The amalgamation of bioturbated sandstone horizons to form compound sequences implies considerable sand influx to the shelf, either as shoreline-attached sheets or more localized low-relief barforms. Locally developed coarsening-upward sequences possibly reflect shoreface progradation and the progressive blanketing-out of lime mud deposition. This is taken as evidence, albeit

Deposits of inferred fluvial channel and alluvial plain origin are represented within the upper part of the Gillfoot Beds, which comprise a mixed assemblage of red shales and interbedded purple sandstones (Fig. 13). The shales are flatlaminated and constitute approximately 50% of the sequence, with individual horizons up to 1.8m in thickness. Sandstones tend to be moderately- to poorly-sorted and very fine- to medium-grained, although they are rarely coarse-grained. Finer grained sandstones occur as thin horizons (only a few centimetres thick) and exhibit current ripple lamination and horizontal lamination. In contrast, the coarser grained sandstones range in thickness from 60 cm to over 2 m and exhibit stacked unidirectional trough and planar cross-sets with palaeocurrent data indicating flow predominantly to the southwest (Fig. 13). The absence of any marine indicators and the stacking of unidirectional structures suggests the accumulation of dune-scale bedforms within an active fluvial channel system. The thinner, finer-grained sandstones reflect deposition by sheetfloods in interchannel areas during periods of peak fluvial discharge. Thus the clastic marine shelf represented by the lower part of the Gillfoot Beds was succeeded by an alluvial plain which was traversed by a fluvial channel system.

Sandy braided fluvial channel lithofacies The Thirlstane Sandstone, which comprises approximately 25m of strata (Craig 1956), is considered to represent deposition within an active braided fluvial channel system. This sandstone interval conformably overlies the Powillimount Beds, and a representative 10m thick section from the base of this sandstone body is illustrated in Fig. 15.

DINANTIAN ENVIRONMENTS, SOLWAY BASIN The basal 75cm of the Thirlstane Sandstone comprises a highly disorganised, poorly-sorted lag of dark purple-red sandstone clasts resting in sharp, erosive contact with the underlying laminated black shales of the Powillimount Beds. Contained within a highly vuggy, medium- to coarse-grained sandstone matrix are pockets of very coarse sand, pebble grade subrounded red mudstone intraclasts and abundant carbonaceous woody fragments (some 15cm in length). Overlying the basal horizon, a continuous sequence of stacked, trough cross-stratified, moderately-sorted, fineto medium-grained sandstones is recognized (Fig. 15). Individual cross-sets generally display gently concave erosive bases and range in thickness from 15-75cm (Fig. 16). The larger sets tend to be preferentially developed in the lower part of the interval. Palaeocurrents measured in these sandstones indicate flow principally in a westerly direction. Most conspicuous within the cross-stratified Thirlstane Sandstone sequence is evidence for variable, but locally intense,

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penecontemporaneous soft-sediment deformation. Within the logged interval, this is typically manifested in the form of oversteepened or overturned foresets and gentle doming of the cross-strata. Although liquefaction is a prerequisite for this type of soft-sediment deformation, it is the shear stress exerted by the overlying sediment-laden currents on the top of the liquified bedform that causes deformation. This is supported by the common truncation of oversteepened/overturned foresets by overlying cross-strata. A more detailed analysis of these soft-sediment deformation features was made by Ord et al. (1988) who, in addition to the above, identified well-defined sand volcanoes and sand dykes, and recognized a large-scale slump structure near the northern limit of the exposure. Most significantly, these authors identified a trend of increasing deformation intensity towards the northern fault termination of the exposure, a phenomenon which they attributed to contemporaneous fault activity on, and tectonic disturbance against, the Thirlstane 'Reverse' fault. In the past, the Thirlstane Sandstone has been variably ascribed to a shallow marine (George et al. 1976; Barrett 1988) or fluvial depositional environment (Ord et al. 1988). In the absence of any positive marine indicators, and in view of the abundance of stacked unidirectional bedforms, the present authors concur with Ordet al. (1988) and attribute the Thirlstane Sandstone to deposition within a braided fluvial channel system. The presence of oversteepened and overturned foresets further precludes a shallow-marine setting, as current velocities and bed shear stress would not be sufficient to produce this deformation within a shallow-marine shelf environment (B. Turner pers. comm. 1995). The stacking of cross-stratified sets is considered to represent the migration and accumulation of transverse dune- and bar-scale bedforms within relatively unconfined low sinuosity fluvial channels. On an outcrop scale, no obvious channel geometries have been observed; as a whole, the Thirlstane Sandstone has a tabular sheet-like geometry, extending along strike for approximately 0.5 km.

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The Rerrick succession (Table 1), comprising largely non-marine clastic rocks, is interpreted as the deposits of an alluvial fan to coastal plain setting influenced by local tectonic activity. The Barlocco Heugh Formation represents a mixed

178

K. M A G U I R E E T AL.

Fig. 16. Stacked trough cross-stratified braided fluvial channel sandstones from the Thirlstane Sandstone (Kirkbean Outlier, North Solway, grid ref. NX 99155645). Cross-strata range in height from <15 cm to around 75 cm and display localized liquefaction features (notably oversteepened/overturned foresets and domed strata) attributed to syndepositional soft-sediment deformation. Palaeoflow directions are principally towards the west. Hammer is 40 cm long.

DINANTIAN ENVIRONMENTS, SOLWAY BASIN sand-mud alluvial plain setting which was subsequently inundated by marine waters resulting in the sandy marine shoreface deposits of the overlying Orroland Lodge Formation. A return to alluvial plain conditions is represented by the Scar Heugh Formation, with the overlying Black Neuk Formation reflecting alluvial fan progradation. The similarities between the semiarid alluvial plain successions of the Barlocco Heugh and Scar Heugh Formations probably reflect the coastal plain environment at periods of relative sea level lowstand, the intervening Orroland Lodge Formation shoreface facies being deposited further out in the basin. The conglomeratic nature of the Black Neuk Formation suggests that, although the coastal plain was traversed by relatively entrenched fluvial systems, it was also subject to episodes of alluvial fan encroachment, possibly triggered by tectonic activity in the hinterland. The presence of pebbles and cobbles up to 15 cm in length in the Orroland Lodge Formation indicates that alluvial fans may occasionally have introduced sediment directly into the marine basin (in the form of fan deltas), or simply that accumulated fan debris was reworked during periods of marine transgression. Palaeocurrent data suggest that the source area for much of the detritus within the Rerrick succession was to the north (Deegan 1973). Furthermore, the immature (feldspathic- and lithic-rich) nature of the sandstones within the Rerrick succession suggests a local source area. This, and the proximity to the North Solway Fault, indicates that the Rerrick Outlier represents a lateral dispersal system which was sourced from the northern margin of the Solway Basin. The Kirkbean succession (Table 2) is considered to represent the deposits of a shallow marine shelf which was episodically inundated by a prograding fluvial system. The Southerness Beds are interpreted as the deposits of a tectonically stable, open-marine shelf setting subject to both limestone and siliciclastic sedimentation. Within the Southerness Beds succession as a whole, limestones progressively give way to sandstones in vertical sequence. This transition reflects increased clastic input which culminated in the inundation of the area by the coastal plain system of the Gillfoot Beds. This low-relief alluvial plain was traversed by fairly entrenched fluvial channels, with the major transport direction from east to west. A return to open marine shelf conditions is reflected by the overlying Powillimount Beds. Clastic material was introduced onto the north Solway shallow marine shelf via an actively prograding

179

fluvial channel system. Fine terrigenous mud was carried in suspension and deposited on the shelf in a low energy setting below fair-weather wave-base. Initially deposited in a more marginal setting, the sand component was subsequently transported onto and across the shelf during storm events. The overall palaeoenvironmental interpretation of the Powillimount Beds is one of a shallow-marine shelf adjacent to a somewhat embayed and swampy coastal plain. The Thirlstane Sandstone represents a major episode of fluvial progradation into the shelf area previously dominated by shallow-marine conditions. The absence of any positive marine indicators and the abundance of stacked unidirectional bedforms suggests deposition within a braided fluvial channel system. Limited palaeocurrent data suggest that the main transport direction was towards the west. Furthermore, the mineralogical and textural maturity of the Thirlstane Sandstone indicates a longer transportation path when compared with the sandstones of the Rerrick succession. A subsequent return to open marine shelf conditions is indicated by the overlying Arbigland Group. This succession exhibits an upward decrease in sandstone lithologies and a concurrent increase in limestones, possibly reflecting a major decrease in the availability of sand grade material as a result of decreasing fluvial influence with time. Correlation of the Rerrick and Kirkbean Outliers with other Dinantian successions in the area would enable the establishment of a depositional model for the northern margin of the Solway Basin during Dinantian times. Figure 17 illustrates the available biostratigraphic data for the Rerrick and Kirkbean Outliers as well as the Langholm and Bewcastle areas. The biostratigraphy of the latter has been revised quite radically since George et al. (1976). However, there is still some debate as to their age assignment. Armstrong & Purnell (1987) concluded that the Lower Border Group of the Bewcastle succession is younger than previously suggested (i.e. Chadian to Holkerian, rather than Courceyan). Subsequently, Purnell (1992) concluded that the base of the Chadian Stage should be placed below the exposed section in Bewcastle, but above the Harden Beds in the Langholm area, confirming the age assignment as illustrated by George et al. (1976, Fig. 10). Most recently, Madhi & Butterworth (1994) have revised the biostratigraphy of the upper part of the Lower Border Group (i.e. below the Harden Beds) in the Langholm area. They disagree with Purnell (1992) and conclude that

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D I N A N T I A N ENVIRONMENTS, SOLWAY BASIN

181

LOW relief coastal plain with elPiSixlic depo~ion largely via tin-ted s h e e t l ~ . O c c a s i o n l l ept~meral lakes (Wayss) and caliche d~'~Wmml, m l ~ easterly derived. llxilll brllldod fluv~lli ~

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the strata below the Harden Beds are of a similar age to the Lower Border G r o u p in Bewcastle, (i.e. Chadian to Holkerian, not Courceyan). However, within the present stratigraphic context, it is suggested that deposition of the Rerrick and Kirkbean successions was broadly contemporaneous. The Rerrick succession is interpreted as a lateral dispersal system, sourced from the northern margin o f the Solway Basin and probably influenced by the synsedimentary m o v e m e n t on the N o r t h Solway Fault. In contrast, the Kirkbean succession is envisaged as representing deposition on a tectonically stable open marine shelf within the Solway marine gulf, with repeated episodes of progradation of a major axial fluvial system along the axis of the basin westwards (Fig. 18). The authors wish to acknowledge the assistance of the following personnel at Geochem Group Limited for their help in preparing this manuscript; to W. Bryan for typing, to P. Horn and S. Hindley for graphics, and to H. Jones for photography. It is also a pleasure to

acknowledge the encouragement of N. J. L. Bailey, Managing Director of Geochem Group Limited.

References

ARMSTRONG, H. A. & PURNELL, M. A. 1987. Dinantian conodont biostratigraphy of the Northumberland Trough. Journal of Micropalaeontology, 6, 97 112. BACON, M. & McQUILLAN, R. 1972. Refraction seismic surveys of the north Irish Sea. Journal of the Geological Society, London, 128, 613 621. BARRETT, P. A. 1988. Early Carboniferous of the Solway Basin: A tectonostratigraphic model and its bearing on hydrocarbon potential. Marine and Petroleum Geology, 5, 271-281. BoTT, M. H. 1964. Gravity measurements in the north-eastern part of the Irish Sea. Quarterly Journal of the Geological Society of London, 120, 369-396. CHADWICK, R. A. & HOLLIDAY, D. W. 1991. Deep crustal structure and Carboniferous basin development within the Iapetus convergence zone, northern England. Journal of the Geological Society, London, 148, 41-53.

182

K. M A G U I R E E T AL.

- - , HOLLOWAY, S. & HULBERT, A. G. 1993. ~I'he evolution and hydrocarbon potential of the Northumberland-Solway Basin. In: PARKER, J. R. (ed.) Petroleum Geology of Northwest Europe." Proceedings of the 4th Conference. The Geological Society, London, 717-726. CRAIG, G. Y. 1956. The Lower Carboniferous Outlier of Kirkbean, Kirkcudbright. Transactions of the Geological Society of Glasgow, 22, 113-132. -& NAIRN, A. E. M. 1956. The Lower Carboniferous outliers of the Colvend and Rerrick shores, Kirkcudbright. Geological Magazine, 93, 249-256. DAILLY, P. 1990. The Late Palaeozoic and Early Mesozoic Structure and Evolution of the Solway and Vale of Eden Basin Complex. PhD Thesis, University of Oxford. DAY, J. B. W. 1970. Geology of the Country around Bewcastle. Memoir of the Geological Survey of Great Britain. DEEGAN, C. E. 1973. Tectonic control of sedimentation at the margin of a Carboniferous depositional basin in Kirkcudbrightshire. Scottish Journal of Geology, 9, 1-28. DOTT, R. H. & BOURGEOIS, J. 1982. Hummocky stratification: significance of its variable bedding sequences. Bulletin of the Geological Society of America, 93, 663-680. FRASER, A. J. & GAWTHORPE, R. L. 1990. Tectonostratigraphic development and hydrocarbon habitat of the Carboniferous in northern England. In: HARDMAN, R. F. P. & BROOKS, J. (eds) Tectonic Events Responsible for Britain's Oil and Gas Reserves. Geological Society, London, Special Publication 55, 49-86. GEORGE, T. N., JOHNSON, G. A. L., MITCHELL, M., PRENTICE, J. E., RAMSBOTTOM, W. H. C., SEVASTOPULO, G. D. & WILSON R. B. 1976. A Correlation of Dinantian Rocks of the British Isles. Geological Society, London, Special Report 7. HARMS, J. C., SOUTHARD, J. B., SPEARING, D. R. & WALKER, R. G. 1975. Depositional Environments as Interpreted from Primary Sedimentary Structures and Stratification Sequences. Society of Economic Palaeontologists and Mineralogists, Dallas, Short Course 2.

JOHNSON, G. A. L. 1984. Subsidence and sedimentation in the Northumberland Trough. Proceedings of the Yorkshire Geological Society, 45, 71-83. LEE, M. K. 1986. A new gravity survey of the Lake District and three-dimensional model of the granite batholith. Journal of the Geological Society, London, 143, 425-435. LUMSDEN, G. I., TULLOCH, W., HOWELLS, M. F. & DAVIES, A. 1967. The Geology of the Neighbourhood of Langholm. Memoir of the Geological Survey, Scotland. MAHDI, S. A. & BUTTERWORTH, M. A. 1994. Palynology of the Dinantian Lower Border Group of the Solway Basin. Proceedings of the Yorkshire Geological Society, 50, 157 171. ORD, D. M., CLEMMEY, H. & LEEDER, M. R. 1988. Interaction between faulting and sedimentation during Dinantian extension of the Solway basin, SW Scotland. Journal of the Geological Society, London, 148, 249-259. PEMBERTON, S. G., MACEACHERN, J. A. & FREY, R. W. 1992. Trace fossil facies models: environmental and allostratigraphic significance. In: WALKER, R. G. & JAMES, N. P. (eds) Facies Models: Response to Sea Level Cbange. Geological Association of Canada, 47-72. PURNELL, M. A. 1992. Conodonts of the Lower Border Group and equivalent strata (Lower Carboniferous) in Northern Cumbria and the Scottish Borders, UK. Royal Ontario Museum, Life Sciences Contributions, 156. SOPER, N. J., WEBB, B. C. & WOODCOCK, N. H. 1987. Late Caledonian (Acadian) transpression in north-west England: timing, geometry and geotectonic significance. Proceedings of the Yorkshire Geological Society, 46, 175-192. TURNER, B. R., YOUNGER, P. L. & FORDHAM, C. E. 1993. Fell Sandstone Group lithostratigraphy south-west of Berwick-upon-Tweed: implications for the regional development of the Fell Sandstone. Proceedings of the Yorkshire Geological Society, 49, 269-281.

Dinantian river systems and coastal zone sedimentation in northwest Ireland JOHN

R. G R A H A M

Department o f Geology, Trinity College, Dublin 2, Ireland Abstract: Northwest Ireland either formed or lay close to the coastal zone and represented the northernmost extent of marine sedimentation for large parts of the late Dinantian. The succession deposited during the initial early Vis~an transgression in Donegal starts with northerly-derived fluvial gravels which pass upwards into coarse-grained fluvial sandstones and associated siltstones. Channel cross-sections are relatively small, suggesting stream widths of a few tens of metres and depths of up to 4 m. These data suggest small rivers with limited drainage basins and are inconsistent with a single drainage basin >95 000 km2, which had been proposed previously. Depositional style in the coastal zone was dominated by laterally migrating tidal creeks, initially containing siliciclastic detrita but with the proportion of carbonate increasing upwards. The succession in North Mayo is similar to that in Donegal but generally lacks the fluvial gravel unit. Widespread deposition of limestones and shales was interrupted in late Arundian times by a southerly progradation of first-cycle feldspathic sandstones, the Mullaghmore Sandstone Formation and its correlatives. In the more proximal environments this is represented by fluvial sediments cutting down into the pre-existing carbonate shelf. This has the basic character of a Type I sequence boundary and represents a major lowering of sea level in northwest Ireland. During this clastic influx, large tidally-influenced fluvial channels developed, with channel widths up to 200m and depths >5 m. This indicates an increase in river size in this area during the Dinantian which may have been due to river capture during the evolution of a relatively low-relief landscape. The regression that led to the southerly progradation of clastics is of similar age to fault movements documented elsewhere in Ireland, and may have a regional tectonic cause. During the Dinantian, Ireland lay near the equator and north of an ocean whose later suture runs from Galicia through Brittany and the Massif Central towards the Vosges. Much of Ireland was the site of shallow-marine sedimentation bounded by land to the north and northwest. This shallow marine area extended eastwards to central Europe and possibly westwards to the maritime provinces of Canada (Sevastopulo 1981; Leeder 1988; Cope et al. 1992). This paper is concerned with rocks in northwest Ireland around Donegal Bay (Fig. 1). This area lay at or close to the coastal zone during most of the Visran (Cope et al. 1992, maps C3 to C5). Although there are significant local variations, the successions of this area show a similar overall sequence. A basal clastic-dominant unit passes upwards into offshore shallow-marine carbonates in the early Vis~an. In late Arundian times there was significant renewed clastic sedimentation before carbonate deposition once more became dominant in the Asbian. The aims of this paper are to describe and interpret some of the non-marine and coastal facies; to assess information on the nature of the fluvial systems supplying sediment to the area; and to comment on the main controls on sedimentation.

Basal clastic rocks On a regional scale there is significant variation in both thickness and facies of the basal clastics, which may be controlled by basement tectonics. The information presented in this paper comes mainly from the two best-exposed sections near the western limits of the exposed Carboniferous sequence in south Donegal and north Mayo (Figs 1 to 4).

Donegal Succession This was subdivided by George & Oswald (1957) into units which, with minor modification, are readily recognizeable in the field (Figs 2, 3), although their nomenclature does not conform to modern usage (Holland et al. 1978). Their older names are given in parentheses when stratigraphic units are first mentioned.

Roelough Conglomerate Formation (Basal Conglomerates). The base of this formation is an unconformity across which the Carboniferous rocks rest on Dalradian rocks (mainly schists). The formation comprises a series of clastsupported, moderately well-sorted conglomerates with subordinate lenses of sandstones and

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 183-206.

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MINNAUN FORMATION

CONGLOMERATE DALRADIAN Fig. 3. Outline stratigraphy for the Donegal Bay area. Clast composition is d o m i n a t e d by psammites and quartzites, with vein quartz c o m m o n in the smaller size ranges. The larger clasts are subrounded to r o u n d e d despite their resistant composition•

rare mudrocks. M a x i m u m particle size (mean of five largest clasts) ranges from 10 to 48 cm, with isolated clasts up to 80 cm in diameter. A l t h o u g h resting unconformably on schists, even the lowest beds are virtually devoid o f schist clasts. oo

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Fig. 4. Geological map of the coastal sections of north Mayo discussed in the text. See Fig. 1 for location.

186

J. R. GRAHAM

Sections of this formation display numerous erosion surfaces although with very limited vertical downcutting. Imbrication is common, but there are also many cross-strata with set thicknesses typically less than 50 cm. Palaeocurrent directions from the cross-strata are consistently from the N N W (Fig. 5). More qualitative information from imbrication, which can only be estimated to an octant of the compass, is in agreement with this. The common imbrication, lack of thick cross-stratal sets, unidirectional palaeocurrents and shallow erosion surfaces suggest deposition in relatively shallow bedload rivers. Isolated Beaconites trace fossils have been found on the tops of some conglomerate beds. Sandstones are thin and laterally discontinuous, and mud, although occasionally preserved as lenses in channel bases, seems to have been mainly flushed through the system. Its localized preservation indicates flows which were very flashy. The previous interpretation as beach or nearshore deposits (George & Oswald 1957) is rejected.

Largysillagh Sandstone Formation (Lower Shalwy Beds). The base of this formation is taken as the lowest thick sandstone (>2m) above which conglomerates become a minor constituent of the succession (Fig. 6). This formation is dominated by coarse sandstones but also contains significant quantities of mudrocks as both relatively tabular bedsets and locally lining erosional hollows. The mudrocks are red, green or grey in colour and sometimes variegated. Most show either desiccation cracks or pedogenic carbonate nodules, thus indicating exposure, often prolonged, of the sediment surface. The sandstones are characterized by numerous erosional surfaces (Figs 7, 8), and internal structures are a mixture of cross-strata and flat or low-angle lamination. Plant debris is common, and Beaconites trace fossils are found on several sandstone surfaces. Only one clear example of inclined heterolithic stratification (IHS; Thomas et al. 1987) has been noted at Shalwy (G646748; Fig. 9). This is interpreted as

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SHALWY POINT

CASTLE PORT Fig. 5. Palaeocurrent data from the basal clastics of southwest Donegal. All data are from cross-strata. Arrows indicate vector means.

D I N A N T I A N R I V E R SYSTEMS, N W I R E L A N D

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Fig. 6. (a) Summary stratigraphical log of the main coastal section at Shalwy (see Fig. 2 for location). Scale in metres. RCF, Roelough Conglomerate Formation; LSF, Largysillagh Sandstone Formation; SF, Shalwy Formation; RPF, Rinn Point Formation. (b) Key for (a) and also for Figs 12, 13, 17, 19, 24 and 25.

188

J. R. GRAHAM

hale eptari, an oncretions Mudstone arbonate glaebules Thinly bedded mudrock •& dandstone Inclined Heterolithic Stratification (IHS) J Ripples / Pari~llel lamination / " ~ Cross-bedding e'eollntraclasts

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Fig. 6. Continued. the deposit of a point bar. However, the exposure on the tidal platform is limited and does not allow accurate estimation of point-bar width. The preserved bar thickness is 2 m and a width of 15m taken from oblique section is only a crude estimate of the latter. Some constraints on the scale of the river channels are also given by the shape and dimensions of the curved erosion surfaces confining the channel fills. These show sandstone-fill thicknesses of 1-3 m, (occasionally up to 4m), and widths of a few metres up to a few tens of

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metres. Palaeocurrents in the sandstones are again northerly (Fig. 5) although showing more variability than those from the Roelough Conglomerate Formation.

Shalwy Formation (Upper Shalwy Beds). The base of the Shalwy Formation is taken at the incoming of intensely bioturbated mudrocks and sandstones, typically with fossils indicating marine or brackish conditions. The basal contact is a relatively abrupt but conformable passage (Fig. 6 at 97m). The lowest beds

.... "+~i

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Fig. 7. Sandstones of the Largysillagh Sandstone Formation at Shalwy (G645748) showing complex intersection of erosion surfaces. The mudrock unit right of centre is 2.2 m thick.

DINANTIAN RIVER SYSTEMS, NW IRELAND

189

Fig. 8. Cut bank of small channel to right of person (circled) at Shalwy (G644748).

typically are plant-rich black mudrocks with a gastropod/ostracode fauna and micrites containing a similar fauna. There are some coarsegrained sandstones lining irregular scours and appearing similar to the sandstones in the underlying Largysillagh Sandstone Formation. However, many sandstones have a more tabular geometry and are extensively bioturbated. These sandstones are characterized by ripple-scale bedforms and flat lamination. Some ripples indicate oscillatory currents (using the criteria of De Raaf et al. 1977) and bipolar currents are also common (Figs 5, 6). Several beds of

sandstone are markedly calcareous and contain considerable amounts of broken shell debris, some grading into sandy limestones. In addition to the lithologies above, a most distinctive feature of this formation is the presence of IHS (Figs 6, 9, 10, 11). Some examples of this indicate palaeocurrents, from ripples in the finer-grained strata to small dunes in the sandier varieties, that are strongly oblique to perpendicular to the dip of the IHS. Some of these are bipolar. The dips of the IHS are low, from a few degrees to a maximum of 20 ° . Marine shell debris is common in these units as,

Fig. 9. Inclined heterolithic strata beneath erosive base of coarser grained sandstones at Shalwy (G646748). Cross-beds at lower left (arrowed) indicate currents strongly oblique to the dip of the heterolithic strata. Rucksack (circled) for scale.

190

J. R. GRAHAM

Fig. 10. Sandstone-dominant inclined heterolithic strata (135-138 m on Fig. 6) at Shalwy (G647746). This is also unit SB1 of Nichols & Jones (1992). Person to right for scale.

locally, is fusain, the latter being interpreted as fossil charcoal (Nichols & Jones 1992). The upper part of this formation passes gradationally upwards into alternating beds of bioclastic limestone and calcareous shale known as the Rinn Point Limestone Formation (Rinn Point Beds) (Fig. 6 at 160m). This gradational passage is characterized by the common presence of quartz granules and pebbles and locally by fine-

grained sandstones showing hummocky crossstratification (HCS). The base of the Rinn Point Formation is taken where tabular beds of limestone and shale first become dominant. This coincides with the disappearance of sandstones. This boundary can be seen on the eastern side of Rinn Point (G672755; Fig. 2), near Shalwy Point (G649745), and at Castle Port, south of Dunkineely (G754741; Fig. 1).

Fig. 11. Muddy inclined heterolithic strata rich in fusain at G648745 (Fusain Unit of Nichols & Jones 1992; 141143m on Figt)6). Person in upper part of unit for scale.

DINANTIAN RIVER SYSTEMS, NW IRELAND Some specific bedsets rich in fusain and the sedimentology of adjacent parts of the sequence were investigated by Nichols & Jones (1992) following earlier palaeobotanical work by Scott & Collinson (1978). They interpreted the gently inclined cross-strata that form these bedsets as representing the lower parts of large sandwave slipfaces. By comparison with some laboratory data on the proportion of bedforms preserved due to random topographic variation (Paola & Borgman 1991), and with dimensions of modern sandwaves, they imply large bedforms, several metres and perhaps up to 15m high, and thus significant water depths. However, they also suggested an inshore estuarine setting in which such water depths are spatially very restricted. The interpretation of these bedsets as sandwave deposits is rejected for the following reasons. (a) In many cases, including some considered by Nichols & Jones (1992), the inclined bedsets (cross-strata) consist of alternations of sand and mud, and some are mud- dominant. This is not consistent with an interpretation as sandwaves. Models of sandwaves based on theoretical considerations allied to field examples by Allen (1980) apply only to cohesionless sediments. (b) The angles of inclination reach a maximum of 20 ° but are typically much less and commonly less than 10° (Figs 10, 11). The area is one of negligible tectonic deformation and thus these represent true values. These are inappropriate for the avalanche faces of sandwaves. Where such low values are predicted for sandwaves by Allen (1980), they are accompanied by downslope migrating dune bedforms; mud does not survive except as intraclasts, and Allen noted that such structures had not been fully described either from modern or ancient sediments. (c) In cases where smaller-scale cross-strata are found within the inclined beds they indicate a flow direction strongly oblique to the inclination of the heterolithic bedsets. This was also noted by Nichols & Jones (1992) who described inclined bedsets dipping NW and smaller-scale bipolar beds indicating E-W tidal currents. (d) The bedsets are found in close proximity to indicators of very shallow water or even exposure (e.g. rootlets at 133.8 m in Fig. 6). (e) The bedsets conform to the description of inclined heterolithic stratification (IHS; Thomas et al. 1987), as noted by Nichols & Jones (1992), for which an origin by lateral migration of point bars was proposed. The true profile view of IHS examples described by Thomas et al. (1987) varies from

191

sigmoidal to straight or slightly concave upwards. They suggested that strongly concave-up IHS sets are typically deposited as abandoned channel-filling sequences. All of these profiles are seen within the Shalwy Formation. Thus an interpretation as tidal point bar deposits for the IHS in the Shalwy Formation seems entirely consistent with the gentle dips, heterolithic character, oblique (in places bipolar) palaeocurrents and marine biota, and is the interpretation favoured here. In a novel approach, Nichols & Jones (1992) also attempted to estimate the size of the drainage basin supplying sediment to this coastal zone by estimating the volume of fusain (fossil charcoal) present, from this the biomass needed to produce it, and hence the areal extent of the drainage basin. They concluded that a drainage basin of >95 000 km 2, a little greater than the present land area of Ireland, supplied a 'Shalwy River System' which reached the sea via an estuarine area in south Donegal. Their calculations of the amount of fusain rely on some conservative estimates of the amount of fusain within their 'Fusain Unit' and also on defining a minimum areal extent of the Fusain Unit. The latter relies on correlating fusain-rich beds from three different coastal sections at Shalwy Point, Rinn Point and Muckros Head (Fig. 2). Of the three exposures described by Nichols & Jones (1992), the fusain unit at Shalwy Point is the thickest and richest in fusain. The exposure at Rinn Point is noted to contain much less fusain and its sedimentological character is also slightly different (cf. Nichols & Jones 1992, figs 6b, 9a). Correlation relies primarily on similar stratigraphic position measuring down from the base of the overlying Rinn Point Formation. Unfortunately, the section at Rinn Point contains a fault across which detailed correlation is not possible. The 'distinctive' black mud with septarian nodules used to correlate the beds at Shalwy Point with those at Muckros Head is not present above the IHS bedset designated as the Fusain Unit at Rinn Point. The section at Muckros Head was recognized by Nichols & Jones (1992) as being in a small fault-bounded block, and correlation was based on the fusain unit having a similar overlying bed to that seen above the fusain unit at Shalwy Point. In fact it is possible to correlate the strata at Shalwy Point with this small fault-bounded block, but not in the way suggested by Nichols & Jones (1992). The lower 6.5 m of this block exposed on the tidal rocks belongs to the upper part of the

192

J. R. GRAHAM

Largysillagh Sandstone Formation and contains blocky desiccated mudrocks, pedogenic carbonate glaebules and Beaconites trace fossils, but lack extensive bioturbation (Fig. 12). The basal beds of the Shalwy Formation are similar to those at Shalwy Point but here are succeeded by an IHS unit containing appreciable amounts of fusain and the first signs of a marine fauna. It is this bedset that Nichols & Jones (1992) have correlated with the fusain unit at Shalwy Point, which, it is suggested here, lies some 40 m higher in the succession (Fig. 6 at 141 m). Thus it appears that the bedsets containing the fusain do not all represent the same stratigraphic level. In this case the area over which the high concentration of fusain exists is only that of each individual 'fusain unit'. These can be seen to be lenticular and of limited extent and are interpreted above as representing lateral accretion of tidal point bars. The gently curved, locally sigmoidal shape of the IHS suggests bar heights of up to 4m. The lateral extent of

top of cliff 1^

SHALWY FORMATION ee)

Gc~

individual beds, interpreted as former bar surfaces, is commonly 15-25 m, suggesting bankfull channel widths of 22-38m (Allen 1965; Ethridge & Schumm 1978). Whilst it could be argued that these tidally influenced channels may not be the largest channels in the system, there is no indication either in the coastal sediments (Shalwy Formation) or the fluvial sediments (Roelough Conglomerate & Largysillagh Sandstone Formations) of any larger channels. The general increase in mud and carbonate content at the expense of sand suggests that much of the siliciclastic material was being trapped in the fluvial system. This is to be expected at times of rising base level. The petrography of the sandstones is both compositionally (20% or more feldspar) and texturally immature as noted by Nichols & Jones (1992), who implied that they were first-cycle sediments from a Dalradian source. This petrography is consistent with a relatively local source but, as noted above, the immediately subjacent schists, which extend for 10 km to the north, are scarcely represented. This, together with the subrounded quartzite clasts in the Roelough Conglomerate Formation, suggests that there was little local erosion at the time of deposition of the Roelough Conglomerate Formation, perhaps due to a generally rising base level promoting aggradation within the fluvial environments. Extensive outcrop of psammites and quartzites presently commences some 10 km to the north, with granitic rocks also extensively exposed north of this. This would be a likely source area.

North Mayo Succession

LARGYSILLAGH SANDSTONE FORMATION -V--

mf m c

Fig. 12. Stratigraphic log of a small fault-bounded section on the western side of Muckros Head (G622742 on Fig. 2 ). The Fusain Unit of Nichols & Jones is at 8 m. See Fig. 6 for legend; scale in metres.

A very similar clastic sequence at the base of the Carboniferous succession is exposed in coastal sections in North Mayo (Figs 1, 3, 4). This also displays a lower coarse-grained fluvial unit passing gradationally upwards into a marginal marine unit which is dominantly clastic, and thence into open marine carbonates. Within the limits of current biostratigraphical resolution, it is of comparable age i.e., post-Courceyan to pre-late Arundian as discussed below. A recently published map by (Geological Survey of Ireland 1992) subdivides the succession into units which approximate to these in terms of their position on the ground, although none of their formations are described or defined in text. Their names are used here with formal definitions to avoid unnecessary replication of nomenclature.

DINANTIAN RIVER SYSTEMS, NW IRELAND

193

25

20

X /A t'3

75 6 -

7o_~ -~_~ 100

65 m' ~c

Fig. 13. Stratigraphic log of the upper part of the Minnaun Formation (MF) and the Downpatrick Formation (DF) on the western side of Downpatrick Head (Fig. 4). See Fig. 6 for legend; scale in metres.

Minnaun Formation. This formation is primarily exposed in cliff sections west of Ballycastle which have only limited direct access, but the formation can also be examined in part on the east side of Bunatrahir Bay (Figs 4, 13), although the base is not seen here. The base is an unconformity across which the Carboniferous rocks rest on Dalradian rocks, mostly psammites. One difference from Donegal is that a basal conglomerate unit comparable to the Roelough Conglomerate Formation is absent, and only a metre or so of pebble conglomerate is present at the base near Port (G023419). Despite limited access, the cliff sections west of Ballycastle are easily examined by binoculars and give good lateral and vertical control. The formation is dominated by medium to coarse-grained sandstones displaying numer-

ous erosion surfaces (Fig. 14). Muds are only locally preserved in the lower part of the formation, as in the Largysillagh Formation in Donegal, but their proportion increases upwards. These are accessible at Minnaun (G039418) and in the sections north of BaUycastle. Colours vary from red to green to grey, and desiccation cracks and calcareous glaebules interpreted as palaeosols are very common. Locally these glaebules coalesce to the extent that they form tabular micrites interpreted as calcrete palaeosols. Some intraclastic conglomerates, produced by reworking of these calcareous palaeosols, are present. The sandstones display abundant trough cross-beds with preserved sets mainly 20-60cm thick, but locally up to 1 m. Palaeocurrents from cross-strata consistently show derivation from

194

J. R. GRAHAM

Fig. 14. Typical cliff section in the lower part of the Minnaun Formation at G033419. Note the localized preservation of mudrocks beneath the curved erosion surfaces at centre left of the cliff. Height of cliff is 30 m.

the NW (Fig. 15). However, there are no large bar structures and the cliff sections do not display any large-scale point bar surfaces despite suitable orientations. Thus, as with the sections in Donegal, the river sizes can only be estimated from the shapes and dimensions of the curved erosion surfaces as being a few tens of metres wide and less than 3 m deep.

Downpatrick Formation. A gradational change into the Downpatrick Formation is exposed SW of Downpatrick Head (Gl18416; Fig. 13). The base is defined by a marked increase in bioturbation, by the common presence of IHS, and by the incoming of scattered marine fossil debris. It is thus comparable to the base of the Shalwy Formation. Dimensions of the IHS are similar to those in the Shalwy Formation, indicating channel depths of up to 4 m and point bar widths of 10--25m. Palaeocurrents from ripples indicate flows strongly oblique to the overall dip of the IHS, and a bipolar pattern of large-scale cross-strata is again in evidence (Fig. 15). Overall micritic palaeosols and evidence for exposure are more common than in the Shalwy Formation in Donegal, and horizons rich in marine shells are less common. In places, micritic palaeosols have been produced by the coalescence of carbonate glaebules and are devoid of fauna. However, most micrites in this formation are tabular, bioturbated, and contain some fauna, mainly ostracodes and gastropods. These are interpreted as lagoonal micrites formed in waters of abnormal salinity. The upper part of the formation (Figs 13, 16) shows a gradational passage into the overlying Moyny Limestone Formation which can be

examined at Downpatrick Head (G123428), west of Ballycastle (G063410, G096403) and east of Downpatrick Head at G144417. The base of the Moyny Limestone Formation is taken to

N

T IINNAUN )RMATION

n = 42

0

10

.....

N PATRIC K IMATION

n = 35

Fig. 15. Palaeocurrent data from the Minnaun and Downpatrick formations. All data are from crossstrata. Arrow represents vector mean.

DINANTIAN RIVER SYSTEMS, NW IRELAND

195

Fig. 16. View of the sea stack of Doonbristy, Downpatrick Head. The lower part of the cliff is formed by the uppermost part of the Downpatrick Formation and shows a muddy channel fill overlain by an IHS set. The upper part is formed by tabular limestones and shales forming the base of the Moyny Limestone Formation. Cliff is 25 m high.

be where tabular beds of limestone and shale first become dominant. The Downpatrick Formation appears to thin westwards, as only c. 15m is present in the cliff sections at Benadereen (G057411) whereas c. 90m is present near Downpatrick Head (Fig. 13).

Early Vis~an carbonates and shales In both Donegal (Rinn Point Formation) and North Mayo (Moyny Limestone Formation), tabular bedded carbonates with subordinate shales succeed the predominantly clastic coastal zone sediments. In Donegal e. 120m of these were deposited before the resumption of significant clastic input, whereas in North Mayo only 50-75 m of the carbonate-dominant sequence is present (Figs 3, 17). Like the Downpatrick Formation beneath it, the Moyny Limestone Formation thins westwards, but the difference in thickness may be partly due to non-deposition and erosion (see below). At the present level of biostratigraphical knowledge it is not possible to assess the temporal correlation of these carbonates with any great precision, but they are clearly post-Courceyan and all appear to belong to the Pu biozone (Higgs 1988). In Donegal these carbonates are overlain by a calcareous shale unit (Doorin Shales) at least

120m thick which is quite different in character to anything lower in the succession. It consists of regularly bedded laminated grey mudrocks, variably fossiliferous, with some horizons showing soft-sediment deformation features affecting several metres of succession. Thin lenticular sandstones and limestones are present within these shales. They have been interpreted as prodelta deposits (George & Oswald 1957), although their sedimentology remains to be investigated in detail. Lithologically this unit can be traced south and east where the terms Coolmore and Bundoran Shales have been used (Oswald 1955; George & Oswald 1957). These shales have also yielded Pu biozone microfloras but are thought to be Arundian in age, at least in part (George et al. 1976). In North Mayo, east of Downpatrick Head, the uppermost 25 m of the carbonate-dominant Moyny Limestone Formation contains more calcareous shales and some fine sandstones, and resembles the upper part of the Downpatrick Formation. Numerous examples of IHS are seen and one example (G162411) abnormally rich in fusain is present (Fig. 18). There is also evidence of exposure in the form of desiccation cracks. These rocks can be interpreted in a similar way to the upper parts of the Downpatrick Formation, as a complex of tidal channels and tidal flats formed in a coastal zone which was a little more carbonate-rich. Much work remains to be

196

J. R. G R A H A M 90_

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Fig. 17. Stratigraphic log of the section on the western side of Bunatrahir Bay (Fig. 4) from the upper part of the Downpatrick Formation to the highest exposed beds of the Mullaghmore Sandstone Formation. Subsidiary columns to the right of the main log indicate major bar surfaces laterally equivalent to the channel sandstones discussed in the text. DF, Downpatrick Formation; MLF, Moyny Limestone Formation; MSF, Mullaghmore Sandstone Formation. See Fig. 6 for legend; scale in metres.

done to clarify lateral relationships at this level of the stratigraphy, where there is clearly significant variability.

Mullaghmore

Sandstone

Formation

Throughout the Donegal Bay area there is a significant coarse clastic intercalation into the otherwise marine, carbonate-dominant Dinantian succession that overlies the basal clastic rocks. This is characterized by coarse-grained, locally pebbly feldspathic sandstones with some indications of exposure (palaeosols, rootlets, etc.). A variety of local names have been u s e d - Mullaghmore Sandstone, Carrowmoran

Sandstone, Donore Sandstone, Kildoney Sandstone, Mountcharles Sandstone - for what appears to be the same unit. The most widely used name, Mullaghmore Sandstone, is preferred here. Biostratigraphical data are still limited and, in particular, correlation between the palynological and foraminiferal biozones requires refinement. TS biozone (Clayton 1985) assemblages have been reported from the Mullaghmore Sandstone at Mullaghmore, Carrowmoran and Mountcharles (Higgs 1988). This biozone appears to be late Arundian to Holkerian in age (Riley 1993). Foraminiferal faunas no older than late Arundian in age have been recovered from the Mullaghmore Sandstone Formation west of Ballycastle (Fig. 17 at

DINANTIAN RIVER SYSTEMS, NW IRELAND

197

Fig. 18. Small-scale inclined heterolithic strata at G 162411 in the upper part of the Moyny Limestone Formation. The set on which the hammer (circled) rests is particularly rich in fusain.

109m) and at Kilcummin Head. Thus the present biostratigraphical data are consistent with the proposed lithological correlation. In the context of this paper it is not possible to describe in detail the large variety of both nonmarine and marine facies in the Mullaghmore Sandstone Formation. Primarily the lower and upper contacts and some of the main fluvial channel sediments will be considered.

tion can be defined by the first major sandstone, there is a significant amount of quartz sand in the upper part of the Moyny Limestone Formation.

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9_ MULLAGHMORE SANDSTONE FORMATION

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Base of the Mullaghmore Sandstone Formation The basal contacts of the Mullaghmore Sandstone Formation are well exposed on coastal sections in North Mayo and Sligo, but on the north side of Donegal Bay the base of the (Mountcharles) sandstones is not well seen. The character of this basal contact varies significantly. In the eastern parts of Donegal Bay there is rapid gradation from highly fossiliferous calcareous shales to calcareous sandstones, locally with granules and small pebbles, displaying evidence of wave action. This type of contact is seen at Portmore (G455344), Rochfort Lodge (G800589) (Fig. 19) and Coolmore (G855663; Fig. 1). Further west, between Downpatrick Head and Kilcummin Head (Fig. 4), there is evidence for shallowing in the upper part of the Moyny Limestone Formation, which resembles parts of the Downpatrick Formation in facies. Although the base of the Mullaghmore Sandstone Forma-

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Fig. 19. Stratigraphic log of the base of the Mullaghmore Sandstone Formation at Rochfo'rt Lodge (G800589) west of Bundoran. See Fig. 6 for legend; scale in metres.

198

J. R. GRAHAM

Fig. 20. Erosional base of the Mullaghmore Sandstone Formation into the Moyny Limestone seen in the upper part of a 30m cliff section at G072412. The sandstone at the top of the cliff is 6.5m thick.

West of Ballycastle, the base of the Mullaghmore Sandstone Formation is spectacularly erosive and cuts down into the tabular carbonates of the Moyny Limestone Formation. On the western limb of an open syncline (G072412) the basal erosion surface has over 4 m of relief (Fig. 20). At G094409 on the eastern limb, only minor relief is visible on the erosion surface (Fig. 17), although fluvial sediments rest directly on shelf carbonates. The basal beds are coarse-grained sandstones with unidirectional cross-strata and are interpreted as fluvial channel deposits. Thus this contact represents a significant local sea-level change, probably at least 8m. In sequence stratigraphy terms this boundary has the characteristics of a Type I unconformity (Van Wagoner et al. 1988). The regional variation in the nature of this boundary suggests that the sea deepened to the south and east. The palaeocurrents from the fluvial facies of the Mullaghmore Sandstone Formation (Fig. 21) indicate transport of coarse sediment from the NNW.

Fluvial channels in the Mullaghmore Sandstone Whilst many of the sandstones in the Mullaghmore Sandstone Formation are clearly marineinfluenced on the basis of bioclastic debris, wave ripples, HCS, and so on, some facies can be interpreted as fluvial. In particular, coarse-grained sandstones with unidirectional cross-strata and

locally associated palaeosols and rootlets appear to be fluvial channel deposits. Coastal exposures in North Mayo and Sligo provide further information on the nature and scale of these channels. Exposures on the western side of Bunatrahir Bay display gently dipping strata cut by an irregular coastline with a few metres of vertical relief. In the lower part of the Mullaghmore Sandstone (Fig. 17 at 70m; G093409), trough cross-bedded sandstones separated by erosion surfaces are well exposed and consistently indicate palaeocurrents from the NNW. These are overlain by a distinctive pale-grey micrite with a buff weathering bioclastic top. This micrite is exposed on both sides of a small headland and then traces around a horseshoeshaped depression in the foreshore where the strata beneath are again exposed. Here they consist of large, gently inclined, heterolithic strata dipping at 10° towards 100 °. Current ripples within the sandy strata indicate palaeocurrents from the N - N W , strongly oblique to these surfaces but roughly parallel to the laterally equivalent trough cross-bedded sandstones. A 30cm Stigmaria is present on one sandstone surface aligned 130-310 ° . These heterolithic strata are interpreted as the products of laterally migrating point bars, and a minimum point bar width perpendicular to the dip of the strata is 60 m. More examples of these structures are exposed further west, particularly at G084409 (Fig. 17,

DINANTIAN RIVER SYSTEMS, NW IRELAND

'CASTLE n = 120

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Fig. 21. Palaeocurrent data from sandstones interpreted as fluvial channel facies from the Mullaghmore Sandstone Formation. All data are from cross-strata; arrows indicate vector means. 102-109m). Here exposures show a continuum from trough cross-beds near the channel centre (Fig. 22) across the gently dipping bar surfaces (Fig. 23). The point bar surfaces dip at 10° towards 180 ° whereas the trough foresets in the channel axis dip at up to 28 ° towards 065 ° . The approximate width of the point bar surface is 75-90 m. Here it is possible to trace the trough cross-bedded sandstones of the axial zone of the

199

channel up to 100 m laterally before encountering a major erosion surface. These field data suggest a channel width of 150-200 m. The top of this sandstone complex is locally draped by thin lenses of limestone containing orthocones and gastropods, and must represent marine flooding of the channel after it was abandoned. A similar style and scale of channel is exposed at Donagh, Co. Sligo (G441353) within the Carrowmoran inlier of the Mullaghmore Sandstone (Fig. 1; Hubbard 1966). Here the point bar surfaces dip at less than 10° to the NNE, whereas well-exposed trough crossbeds in the axial zone of the channel dip consistently towards ESE. Estimated width of the point bar surfaces perpendicular to their dip is 40-50m. Thus the fluvial channel facies of the Mullaghmore Sandstone commonly displays large IHS sets interpreted as former point bar surfaces. The scale of bankfull river width indicated by these structures is a few hundreds of metres rather than the few tens of metres indicated for the channels in the basal clastics. The reasons for this increase in river size are not obvious, and hypotheses are difficult to test. A simple explanation is that river capture increased drainage basin size. Such events may be expected during the evolution of a relatively low relief landscape. However, it is also possible that the increase in drainage basin size during Mullaghmore times was in part due to the exposure of previously shallow-marine areas due to relatively rapid sea-level fall. A full facies analysis of the Mullaghmore Sandstone Formation is currently in progress. Work to date suggests that fluvial channel sands are most common in the most westerly section, west of Ballycastle (Fig. 4), and marine facies are more common further east. However, all sections show limestones with varied marine fauna as well as some unidirectional crossbedded sandstones representing fluvial channels. Thus sedimentation throughout the Mullaghmore Sandstone Formation was in coastal areas, probably in part estuarine. This is consistent with the limestone tops seen on many abandoned fluvial channels (e.g. Fig 17 at 72 m, 100 m, 109 m) and with the presence of bipartite, bipolar sandstones (e.g. Castlenageeha, Figs 4, 24).

Top of the Mullaghmore Sandstone In very general terms the Mullaghmore Sandstone shows an increase in thicker fluvial

200

J. R. GRAHAM

Fig. 22. Large-scale trough cross-strata from the axial channel facies of the sandstone complex at 102-109 m (Fig. 13). Rucksack for scale.

sandstones upsection at Mullaghmore and Carrowmoran (Fig. 1), where most of the formation is seen in one section. In North Mayo only the lower parts of the formation are seen west of Ballycastle, but to the east, at Kilcummin Head (G383209), the upper contact is seen. Here marine sandstones, many with well developed HCS (Figs 25, 26), are abruptly

overlain by bioturbated green mudrocks with common micritic palaeosols (Fig. 27). These palaeosols overlie an erosion surface which has locally cut down into the underlying sandstone (Fig. 28) and which must represent at least a local regressive phase. These micrites and mudrocks are 7 m thick and have at least one horizon of further erosion draped by more

Fig. 23. Inclined heterolithic strata interpreted as point bar surfaces dipping towards the trough cross-bedded sandstones shown in Fig. 22, which lie to the left of this field of view. The map board and rucksack (circled) rest on thin lenses of limestone which represent the flooding of the channel after abandonment.

DINANTIAN RIVER SYSTEMS, NW IRELAND

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micrites and mudrocks. The topmost 70cm of mudrocks are different in character, being extensively bioturbated and displaying trace fossils characteristic of marine horizons seen elsewhere in the formation. Above these rocks is a razor-sharp basal contact of coarse, pebbly bioclastic limestone (Fig. 29) which then passes up into well-bedded limestones and shales with abundant Zoophycos. This contact is taken by the Geological Survey of Ireland (1992) as the base of their Ballina Limestone Formation. Faunas from just above the base of the Ballina Limestone Formation indicate an Asbian age. Some 10km further south, near Killala, this upper contact of the Mullaghmore Sandstone Formation is characterized by extensive development of cross-bedded oolitic limestones, the Killala Oolite Member of the Mullaghmore Sandstone Formation (Geological Survey of Ireland 1992). The upper contact of the Mullaghmore is also seen at Carrowmoran where slightly different aspects are seen on different limbs of an open anticline (Hubbard 1966). To the west

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Fig. 25. Stratigraphic log of the upper part of the Kilcummin Head section showing the upper part of the Mullaghmore Sandstone Formation and the sharp base of the Ballina Limestone Formation. The small column to the right of the main log represents the erosional downcut and subsequent muddy micritic fill which are partly seen in Fig. 28. See Fig. 6 for legend; scale in metres.

(G441353) oolitic limestones are present in the upper part of the Mullaghmore Sandstone just beneath the channel complex described above. To the east (G467344) the topmost sandstone of

202

J. R. GRAHAM

Fig. 26. Hummocky cross-stratified sandstones from the upper part of the Mullaghmore Sandstone Formation at Kilcummin Head. Hammer for scale.

the Mullaghmore is characterized by a pebbly lag and is then abruptly succeeded by relatively clean biomicrites. West of Mullaghmore headland the continuity of exposure is curtailed by sandy beaches. The uppermost beds exposed at Conor's Island (G657525) also show a prominent sandy oolite. Thus the top of the Mullaghmore Sandstone Formation represents a period of winnowing with oolitic and locally pebble lag facies. This is

i ~

typical of many rapid sea level rises seen in the stratigraphical record. In sequence stratigraphy terms such features are termed transgressive surfaces (Van Wagoner et al. 1988). In the more landward parts of the system these transgressive surfaces might be expected to overlie a nondepositional hiatus, possibly including soil zones (Galloway 1989). The sequence at Kilcummin Head (Fig. 25) conforms to this pattern and suggests that some considerable time may be

!

Fig. 27. Mudrocks with micritic nodules, Kilcummin Head. Note the development of micrites as replacements of burrow fills in the left of the photograph. Lens cap (50 mm) for scale.

DINANTIAN RIVER SYSTEMS, NW IRELAND

203

Fig. 28. Erosional downcut (centre) into the prominent tabular calcareous sandstone filled mainly by mudrocks with micritic nodules, Kilcummin Head. Person is standing on the basal bed of the Ballina Limestone Formation. represented by the uppermost 7 m Mullaghmore Sandstone Formation.

of the

Source of the Mullaghmore Sandstone Whilst much work remains to be done on the Mullaghmore Sandstone it is clear that the overall derivation direction of sediment was from the N N W (Fig. 21). The sandstones are relatively coarse-grained, angular and markedly felds-

pathic, and apparently similar in composition to the sandstones from the basal clastic rocks, despite the inferred increase in drainage basin size. A similar interpretation as first-cycle detritus from the Dalradian basement seems reasonable, since this type of basement geology is likely to have extended some considerable distance to the north and west. Notable exceptions to this general transport direction are found in the Killala area. A well

Fig. 29. Basal contact of the BaUina Limestone on the Mullaghmore Sandstone Formation, Kilcummin Head. Hammer head rests on a micritic palaeosol.

204

J. R. GRAHAM

exposed sandstone body at Castlenageeha (G209348; Figs. 4, 24) displays a lower trough cross-bedded quartzose sandstone with abundant drifted plant stems, and is interpreted as fluvial. This is overlain by cross-bedded oolitic shelly sandstones which are interpreted as marine flooding of an abandoned fluvial channel. The fluvial sandstones are southerly derived and the upper oolitic sandstones are northerly derived. Further north at Kilcummin a well exposed section through most of the Mullaghmore Sandstone Formation shows mainly marine facies. However, there is one major fluvial channel sandstone within the sequence which also shows south to north palaeocurrents. The ultimate source for these sandstones is not clear, but they may be evidence for some exposed basement to the south at this time.

Regional extent o f the Mullaghmore Sandstone The Mullaghmore Sandstone Formation is thickest and best developed north and west of the Ox Mountains basement block in the Donegal Bay area. However, it is also recognized in parts of the Lough Allen Basin between the northern parts of the Ox Mountains and the northern parts of the Curlew Mountains, although it does not appear to be present adjacent to the Curlews Block or in the Ballymote Syncline between the southern parts of the Ox Mountains and the Curlews (Fig. 1; Philcox et al. 1992). Important intercalations of sandstone prograding southwards into the marine sequence are known further east in the northern part of Ireland: e.g. the Clonelly Sandstone of the Omagh Syncline, and the Aughnacloy Sandstone of the Clogher ValleySlieve Beagh area, from which Mitchell & Owens (1990) derived a late Arundian age. The recently published maps of the Derrygonnelly and Kesh areas around Lower Lough Erne (Geological Survey of Northern Ireland (GSNI) Sheets 43, 44, & 56) recognize the Mullaghmore Sandstone Formation and show it to be completely Arundian in age, based on foraminiferal assemblages, with the base of the Holkerian recognized in the overlying Benbulben Shale Formation. Further east in Armagh the Drumman More Sandstone Formation (c. 120m) represents deltaic environments with associated thin coal seams, and is dated as late Arundian or Holkerian in age on the basis of fossiliferous limestones which occur above and below the formation (GSNI Sheet 46). It has also yielded

TS biozone microfloras (Higgs et al. 1988). It thus seems as if this progradational event is recognized throughout an area at least 250 x 70 km.

Controls on Dinantian sedimentation in N W Ireland The onset of sedimentation in NW Ireland in the Dinantian was due to the creation of accommodation space in the crust. Ultimately this implies regional subsidence for which large-scale tectonic processes must have been responsible. Reasonable explanations for an overall extensional regime in the northern parts of the British Isles are provided by Leeder (1988). It is clear from the whole of this large area that there were marked variations in subsidence locally, owing to the subsidence being produced by movements on widely spaced, often basement-inherited faults, giving rise to tilt block and basin provinces (Leeder 1988). Marked local variations of this type are documented in the basal clastic sequences in the northern part of Ireland (Sevastopulo 1981; Mitchell & Owens 1990; Graham & Clayton 1994). The main questions which have been asked of these Carboniferous successions is how much they reflect widespread eustatic events, as suggested by Ramsbottom (1973), rather than relatively local tectonic controls. Notwithstanding the local variations, which reflect variation in amounts of extension, the early Vis6an was clearly a time of widespread transgression (see Cope et al. 1992). However, this transgression was gradual (Sevastopulo 1981) and reflected long-term rise of sea level rather than any rapid, short-term change. At this scale the sea level change may well have been eustatic. The widespread progradation of clastic sediment in late Arundian times, exemplified by the Mullaghmore Sandstone Formation, does suggest a change in sea level which was approximately synchronous throughout NW Ireland. The magnitude of the sea level change can only be assessed crudely at present. The fully fluvial sandstones at the base of the Mullaghmore Sandstone Formation west of Ballycastle cut down over 4m into shallow-marine carbonates which show common wave action but a lack of lagoonal micrites, and rare IHS sets. This suggests a minimum sea-level fall of c. 8 m. Gradational basal contacts of the Mullaghmore Sandstone Formation elsewhere (e.g. Fig. 19) are from sediments interpreted as shallow marine, and certainly within the zone of in situ carbonate

D I N A N T I A N R I V E R SYSTEMS, N W I R E L A N D production. Thus a few tens o f metres w o u l d be a m a x i m u m estimate. The top of the M u l l a g h m o r e Sandstone F o r m a t i o n appears to represent a widespread transgression of similar magnitude. In contrast, the alternations between marine and n o n - m a r i n e conditions within the Mullaghm o r e Sandstone F o r m a t i o n c a n n o t be traced over even a few kilometres, a feature previously reported by H u b b a r d (1966). Thus these alternations can be interpreted as due to lateral facies change in coastal sediments. In order to assess the likelihood of a tectonic or eustatic control for the main M u l l a g h m o r e progradational event, it is necessary to look further afield at temporally equivalent successions. In this respect the lack of precise biostratigraphical correlation remains a limiting factor. Nevertheless, there are clear indications elsewhere in Ireland and in Britain of i m p o r t a n t tectonic activity in A r u n d i a n times. This was a time of d e m o n s t r a b l e fault m o v e m e n t s in the Curlew M o u n t a i n s (Philcox et al. 1989), near N a v a n (Philcox 1989), in the D u b l i n Basin ( N o l a n 1989), and in the Craven Basin of N o r t h e r n E n g l a n d ( G a w t h o r p e 1986). Thus it seems likely that the M u l l a g h m o r e Sandstone p r o g r a d a t i o n m a y have been controlled by regional tectonic causes. To test this hypothesis fully requires further biostratigraphical refinement from both this area and elsewhere to determine whether the sea level fall extended b e y o n d areas where it could be linked to tectonic activity. Discussions on the Carboniferous geology of NW Ireland with C. Ni Bhroin, G. Clayton, G. Nichols and G. Sevastopulo have been much appreciated. Thanks to B. Bluck, P. Shannon and, in particular, P. Strogen for their time and effort which helped to improve this paper.

References ALLEN, J. R. L. 1965. The sedimentology and palaeogeography of the Old Red Sandstone of Anglesey, North Wales. Proceedings of the Yorkshire Geological Society, 35, 139-185. - - 1 9 8 0 . Sand waves: a model of origin and internal structure. Sedimentary Geology, 26, 281-328. CLAYTON, G. 1985. Dinantian miospores and intercontinental correlation. Compte Rendu lOme Congres Advancement Etudes Stratigraphie Geologie Carbonifere, Madrid, 4, 9-23. COPE, J. C. W., GUION, P. D., SEVASTOPULO, G. D. & SWAN, A. R. H. 1992. Carboniferous. In: COPE, J. C. W., INGHAM, J. K. & RAWSON, P. F. (eds) Atlas of Palaeogeography and Lithofacies. Geological Society, London, Memoir 13, 67-86.

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DE RAAF, J. F. M., BOERSMA, R. & VAN GELDER, A. 1977. Wave generated structures from a shallow marine succession, Lower Carboniferous, County Cork. Sedimentology, 24, 451 484. ETHRIDGE, F. G. & SCHUMM, S. A. 1978. Reconstructing paleochannel, morphologic and flow characteristics: methodology, limitations, and assessment. In: M1ALL, A. D. (ed.) Fluvial Sedimentology. Canadian Society of Petroleum Geologists Memoir 5, 703-722. GALLOWAY, W. E. 1989. Genetic stratigraphic sequences in basin analysis I: architecture and genesis of flooding-surface bounded depositional units. American Association of Petroleum Geologists Bulletin, 73, 125 142. GAWTHORPE, R. 1986. Tectono-sedimentary evolution of the Bowland Basin, N. England, during the Dinantian. Journal of the Geological Society, London, 144, 59-71. GEOLOGICAL SURVEY OF IRELAND 1992. North Mayo. Bedrock Geology 1 : 100 000 Series. Sheet 6. GEOLOGICAL SURVEY OF NORTHERN IRELAND 1983. CIogher. 1 : 50 000 Series. Sheet 46. - - 1 9 9 1 . Derrygonnelly and Marble Arch. 1:50000 Series. Sheets 44, 56 and 43. - - 1 9 9 4 . Kesh. 1 : 50000 Series. Sheets 32 and 31. GEORGE, T. N. & OSWALD, D. H. 1957. The Carboniferous rocks of the Donegal Syncline. Quarterly Journal of the Geological Society of London, 113, 137-178. - - , JOHNSON, G. A. L., MITCHELL, M., PRENTICE, J. E., RAMSBOTTOM, W. H. C., SEVASTOPULO, G. D. & WILSON, R. B. 1976. A Correlation of Dinantian Rocks in the British Isles. Geological Society, London, Special Report 7. GRAHAM, J. R. & CLAYTON, G. 1994. Late Tournaisian conglomerates from County Donegal, NW Ireland; fault controlled sedimentation and overstep during basin extension. Irish Journal of Earth Sciences, 13, 95-105. HIGGS, K. 1988. Stratigraphic palynology of the Carboniferous rocks in Northwest Ireland. Geological Survey of Ireland Bulletin, 3, 171-201. , MCPHILEMY, B., KEEGAN, J. B. & CLAYTON,G. 1988. New data on palynological boundaries within the Irish Dinantian. Review of Palaeobotany and Palynology, 58, 61-68. HOLLAND, C. H., AUDLEY-CHARLES, M. G., BASSETT, M. G., ET aL. 1978. A Guide to Stratigraphical Procedure. Geological Society, London, Special Report 10. HUBBARD, J. A. E. B. 1966. Facies paterns in the Carrowmoran Sandstone (Vis6an) of western Co. Sligo, Ireland. Proceedings of the Geologists Association, 77, 233-254. LEEDER, M. R. 1988. Recent developments in Carboniferous geology: a critical review with implications for the British Isles and NW Europe. Proceedings of the Geologists Association, 99, 73-100.

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MITCHELL, W. I. & OWENS, B. 1990. The geology of the western part of the Fintona Block, Northern Ireland: evolution of Carboniferous basins. Geological Magazine, 127, 407-426. NICHOLS, G. & JONES, T. M. 1992. Fusain in Carboniferous shallow marine sediments, Donegal, Ireland: the sedimentological effects of wildfire. Sedimentology, 39, 487-502. NOLAN, S. C. 1989. The styling and timing of Dinantian syn-sedimentary tectonics in the eastern part of the Dublin Basin, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication, 6, 83-97. OSWALD, D. H. 1955. The Carboniferous rocks between the Ox Mountains and Donegal Bay. Quarterly Journal of the Geological Society of London, 111, 167-186. PAOLA, C. & BORGMAN, L. 1991. Reconstructing random topography from preserved stratification. Sedimentology, 38, 553-565. PHILCOX, M. E. 1989. The mid-Dinantian unconformity at Navan, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication, 6, 67-81. , BAILY, H., CLAYTON, G. & SEVASTOPULO, G. D. 1992. Evolution of the Carboniferous Lough Allen Basin, Northwest Ireland. In: PARNELL, J. (ed.) Basins on the Atlantic Seaboard." Petroleum Geology, Sedimentology and Basin Evolution. Geological Society, London, Special Publication, 62, 203-215.

--,

SEVASTOPULO, G. D. & MCDERMOTT, C. V. 1989. Intra-Dinantian tectonic activity on the Curlew Fault, north-west Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication, 6, 55 66. RAMSBOTTOM, W. H. C. 1973. Transgressions and regressions in the Dinantian: a new synthesis of British Dinantian stratigraphy. Proceedings of the Yorkshire Geological Society, 39, 567-607. RILEY, N. 1993. Dinantian (Lower Carboniferous) biostratigraphy and chronostratigraphy in the British Isles. Journal of the Geological Society of London, 150, 427-446. SCOTT, A. C. & COLLINSON, M. E. 1978. Organic sedimentary particles. In: W. B. WHALLEY (ed.) Scanning Electron Microscopy in the Study of Sediments. Geo Abstracts, Norwich, 137-167. SEVASTOPULO, G. D. 1981. Lower Carboniferous. In: HOLLAND, C. H (ed.) A Geology of Ireland. Scottish Academic Press, Edinburgh, 147-171. THOMAS, R. G., SMITH, D. G., WOOD, J. M., VISSER, J., CALVERLY-RANGE, E. A. & KOSTER, E. H. 1987. Inclined heterolithic stratification- terminology, description, interpretation, and significance. Sedimentary Geology, 53, 123-179. VAN WAGONER, J. C., POSAMENTIER, H. W., MITCHUM, R. M., VAIL, P. R., SARG, J. F., LOUTIT, T. S. & HARDENBOL, J. 1988. An overview of the fundamentals of sequence stratigraphy and key definitions. In: WILGUS, K. C., HASTINGS, B. S., KENDALL, C. G. ST. C., POSAMENTIER, H. W., ROSS, C. A., & VAN WAGONER, J. C. (eds) Sea Level Changes- an Integrated Approach. SEPM Special Publication, 42, 39-45.

Cyclic emersion surfaces and channels within Dinantian limestones hosting the giant Navan Zn-Pb deposit, Ireland GIANCARLO

RIZZI ! & COLIN

J. R. B R A I T H W A I T E

Department of Geology and Applied Geology, University o f Glasgow, Lilybank Gardens, Glasgow G12 8QQ, UK 1 Present Address." Blackbourn Geoconsulting, 30 Coltbridge Terrace, Edinburgh EHl2 6AE

Abstract: The Carboniferous succession in the Navan Mine area, which lies near the northern margin of the Dublin Basin, rests unconformably on folded Lower Palaeozoic rocks. The informal stratigraphic nomenclature of the mine is used throughout. Sedimentation began in the late Devonian-early Courceyan with Red Beds reflecting deposition in braided streams and alluvial fans. The overlying Laminated Beds consist of shallow-marine barrier sandstones, tidal-flat/lagoonal mudstones, and sabkha evaporites. The Muddy Limestone is a mud-dominated carbonate sequence reflecting deposition in an open elastic-influenced lagoon. The Pale Beds that follow, and on which attention is focused here, contain at least 44 peritidal and shallow-shelf depositional cycles. The overlying Shaley Pale and Argillaceous Bioclastic Limestones reflect deeper water open-sea conditions. This sequence is truncated by a conspicuous erosion surface overlain by Chadian submarine debris-flows and limestone turbidites. The succession reflects a progressive deepening of the waters as subsidence outpaced sediment accumulation, but emersion surfaces capping cycles indicate that neither deposition nor relative sea-level rise were continuous. Cyclicity and emergence are believed to reflect the interaction of regional subsidence, glacio-eustatic sea-level oscillation, and sediment supply. Emersion surfaces occur in the Laminated Beds and throughout the Pale Beds, and include palaeosols, in situ breccias, pinnacled and hummocky surfaces, and karst-modified topography. Deeply incised channels at four intervals point to larger-scale incision. The lithoclast-bearing conglomerates which they contain indicate extensive emergence nearby. The distribution of these features may be related to the margins of outcrops of Old Red Sandstone and Lower Palaeozoic rocks. The Dinantian limestones in this region host several Zn-Pb ore deposits including Europe's largest at Navan. The results of this study suggest a relationship to the margins of former emergent carbonate terrains.

Intraformational emersion surfaces marked by palaeokarst and palaeosols have been widely reported from Dinantian platform limestones in the US (Briskey et al. 1986), Britain (Walkden 1987) and mainland Europe. By contrast, few examples have been described from generally similar limestones in Ireland, perhaps reflecting the paucity of inland surface exposures. However, palaeosols have been reported at Moyvoughly, County Westmeath (Harwood & Sullivan 1991), Andrew & Poustie (1986) described meteoric cements from Tatestown, County Meath, and Pickard et al. (1992) noted both meniscus cements and rhizoliths at Kentstown and Walterstown, in County Meath. Closely spaced drilling around the Navan Z n - P b deposit has generated more than 1000 cores. Over 70 of these have been examined and logged, together with underground exposures. As a result, numerous previously undescribed

emersion surfaces have been identified within the Lower Dinantian Limestones. The aims of this paper are: (1) to present an inventory of these surfaces; (2) to provide an interpretation of them, and finally (3) to consider the implications of this interpretation for the evolution of the Lower Dinantian (Courceyan) ramp on the northern margin of the Dublin Basin.

Summary lithostratigraphy and sedimentology Navan Mine lies about 1 km west of Navan close to the northern margin of the Dublin basin (Fig. I). Detailed descriptions of the lithostratigraphy of the Dinantian succession are available in Andrew & Ashton (1985), Ashton et al. (1986,

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 207-219.

208

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Fig. 1. Solid geology of the Navan area with the location of Navan mine. Modified from Ashton et al. (1992).

1992), Philcox (1989), Anderson (1990), and McNestry & Rees (1992). Rizzi (1992) described the sedimentology summarized below. The Courceyan succession (Fig. 2) rests unconformably on deformed Lower Palaeozoic black shales, turbidites and volcaniclastic rocks and is divided into two groups, a lower Navan Group and an upper Argillaceous Bioclastic Limestone (ABL) Group (cf. Ashton et al. 1992). A formal stratigraphy has been erected for the Dinantian of the region (Rees 1987; Strogen e t al. 1990), but since this paper deals exclusively with the Navan Mine area, the informal mine nomenclature is retained here. The Navan Group is divided into five units: Red Beds, Laminated Beds, Muddy Limestones, Pale Beds and Shaley Pale Limestones. The ABL Group consists of both argillaceous bioclastic limestones s e n s u s t r i c t o and Waulsortian facies limestones. Both groups are truncated by a Chadian 'erosion' surface, and are overlain by the Fingal Group of Nolan (1989), consisting, at the mine, of the Boulder Conglomerate and Upper Dark Limestones. The Red Beds, up to 50m thick, consist of conglomerates, cross-bedded and rippled sandstones, siltstones and mudstones arranged in fining-upwards cycles, some of which are capped by calcretes. They are interpreted as fluvial or alluvial fan deposits (Strogen e t al. 1990; Rizzi 1992).

The Laminated Beds record the initial transgression of the Lower Carboniferous sea (Philcox 1984; Andrew & Ashton 1985) and reflect deposition in a variety of environments (McNestry & Rees 1992; Rizzi 1992). The lower part of the succession is characterized by a transition from silty shales to siliciclastic sandstones, recording the progradation of a shallow-marine barrier-beach complex. Above these, algal-laminated limestones, calcisiltites, fenestral calcite mudstones and black shales are associated with oncolites and a nodular anhydrite (now silicified) and resemble peritidal and lagoonal deposits of the present Trucial coast (cf. Kendall & Skipwith 1968). The succession is capped by thin bioclastic grainstones. The Laminated Beds are cut by a channel sequence about 20 m deep and 100-500 m wide, which isopach data show extends roughly NW-SE (Fig. 3). This contains about 15m of sand to small-pebble grade, burrowed and crossbedded grainstones, forming the Limestone Conglomerate. These are arranged in fining-up cycles about 1.5 m thick with sharp erosive bases. Grains include bioclasts and siliciclastic grains as well as lithoclasts of palaeosols. Cross-sections correlating cycles indicate that there are at least five minor channels stacked vertically within the sequence. The Muddy Limestone consists of about 20 m of shaley, bioclast-rich calcite mudstones and

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Fig. 2. Summary lithostratigraphy of the Lower Dinantian succession at Navan Mine. The stratigraphic nomenclature is the informal one in use at Navan Mine. The prominent marker horizons are shown. UDM, Upper Dark Marker; SDM, Sub-Dark Marker; NOD, Nodular Marker; MC, Microconglomerate; LCG, Limestone Conglomerate; LDM, Lower Dark Marker.

wackestones with burrowed intervals. In situ Syringopora reticulata and an assemblage of brachiopod, coral and other bioclasts reflect deposition in a clastic-influenced marine lagoon. The Muddy Limestone also contains a channel assembly, spatially related to the channel below and locally cutting down into the Laminated Beds. The assembly strikes NE-SW and is over 1200m wide and 20m deep. Individual channels within it are filled with limestone conglomerate resembling that in the Laminated Beds below. The Pale Beds, which host 97% of the Navan ore body, are divided into a lower Micrite Unit and an upper Grainstone Unit (Rizzi 1992). The Micrite Unit, up to 60m thick, consists of approximately 35m-scale, lagoonal/tidal-flat

209

The bases of cycles are typically burrowed mudstones and wackestones, while the tops are of similar lithologies but lack burrows and contain a variety of fenestral structures. A single oolitic grainstone within the Micrite Unit is believed to represent a sand shoal driven onshore. The base of the overlying Grainstone Unit fills in palaeotopography on the surface of the Micrite Unit (Fig. 4 and see below). This is dominated by a NW-SE striking depression in the same geographical position as those in the Laminated Beds and Muddy Limestones and was first recognized by Andrew & Ashton (1985). The Grainstone Unit itself is 150 m thick, and comprises at least six shallow-shelf cycles, each up to 30m thick. Burrowed shaley wackestones form lithostratigraphical markers at the bases of cycles (Fig. 2), and are believed to reflect deposition below fair-weather wave-base. They are capped by oolitic and bioclastic grainstones representing a transition to inner-ramp sandshoals. A third channel assembly is present near the base of this succession, between the two lowest siliciclastic markers. The channel sequence is 30 m thick and at least 2 km wide (Fig. 5). Isopach data show that it occupies a N-S trending depression whose axis lies in the same geographical area as those below. The fill consists of at least 13 individual channels, each up to 1.5 km wide and 8 m deep. Channels are filled with granule- and sand-grade grainstones, collectively referred to as the Microconglomerate. As in the channel-fills below, clasts in these include fragments with meteoric cements (see below). The Shaley Pale Limestones consist of 110m of repeated wackestone, packstone and shale cycles, with a few sandstones and thin graded grainstones, the last with sharp, eroded bases. These were probably deposited offshore, below fair-weather wave-base, with the grainstones representing storm deposits (Philcox 1989; Strogen et al. 1990). The Argillaceous Bioclastic Limestones, which are over 130m thick, consist of shaley wackestones, shales and a few packstones, with the siliciclastic content decreasing vertically. Abundant bioclasts, typically crinoid ossicles, are present throughout, some forming sharp-based beds a few centimetres thick. Generally, these rocks reflect a continuation of deep-water outerramp sedimentation with the thin limestones again representing storm deposits (Strogen et al. 1990). Waulsortian limestones form only isolated knolls in the Navan area (Ashton et al. 1992). Typically, they are interbedded light-grey calcite

210

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mudstones, crinoidal wackestones and packstones and are regarded as reflecting deposition in a deep-water outer-ramp environment (Lees & Miller 1985). A southwards-sloping 'erosion' surface, now interpreted as a submarine slide surface produced during tectonic extension, truncates a substantial part of the Courceyan succession (Philcox 1989; Strogen et al. 1990; Ashton et al. 1992). It is overlain by the Boulder Conglomerate (hosting the remaining 3% of the Navan orebody). There are several conglomerate

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Large-scale channels within the succession have been described above, and there is also a largescale surface limiting the top of the Micrite Unit. However, the Laminated Beds and Pale Beds successions also include a number of surfaces bounding individual depositional cycles. These are of five kinds: (1) brecciated surfaces, (2) hummocky surfaces, (3) pits and dissolution cavities, (4) pinnacled surfaces, and (5) a surface bounding an evaporite.

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Brecciated surfaces Brecciated surfaces are found in the Micrite Unit, capping both lagoonal-tidal flat cycles and the oolitic grainstone (Fig. 6). The breccias have distinctive vertical profiles up to 1 m thick which cut across layering, oncolites and fenestrae within host limestones. Their bases consist of unmodified limestones, but the micrites become progressively more fractured upwards, with fractures forming a rectilinear network of fissures up to l c m wide and at least 10cm deep. Profiles are capped by breccias consisting of angular blocks of limestone up to 10cm in diameter which locally show a jigsaw fit. Some bedding is contiguous between blocks, indicating an in situ origin. Fractures and inter-block cavities are filled with muddy siltstones. Breccia surfaces occur every few metres in hole N975, and in hole N1022 between 555.0m and 547.5m there are three separate breccias.

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These are marked by laterally-linked bowl and funnel-shaped pits separated by blunt pinnacles. They are sharply defined, extending laterally for over 40 m and cutting down locally as much as 3 m into the underlying cycle, truncating individual grains, cross-bedding and other layering. Examples are seen in the Block 6 contour drift and in the 1230 mine level, 126 west slope.

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Pits and dissolution cavities Hummocky surfaces may pass laterally into isolated funnel-shaped pits, or downwards into bedding-parallel dissolution cavities. Funnelshaped depressions are up to 1.5m wide and 3m deep, narrowing downwards. They can be traced across mine passages and are at least 4 m long. Some bedding-parallel cavities up to 30 cm

212

G. RIZZI & C. J. R. BRAITHWAITE

high extend laterally for several metres. They may be either parallel-sided or lenticular and generally extend along interfaces between layering. Pits and cavities truncate sedimentary structures (Fig. 7a). Steeply inclined erosion surfaces recognized within the Micrite Unit (core N1011 at 560.3m depth, and core N629 at 356.3m and 356.5 m depth) probably represent segments of such steep-sided pits.

Pinnacled surfaces A few surfaces are bounded by vertical pinnacles up to 15cm high with pinnacled surfaces extending laterally for at least 30 cm. Individual pinnacles may be 3 cm high with undercut and overhanging side-walls. Terminations may be sharp, blunt or planar. Metre-scale pinnacles truncated obliquely appear as blocks apparently suspended above surfaces. Irregular cavities up to 3 cm in diameter are present in some pinnacles (now filled with blocky calcite); typically these pass vertically into fissures. Pinnacled surfaces have been recognized within the Micrite Unit in core N509 at 56.6 m depth (Fig. 7b) and also in underground exposures in the 1210 mine level, 1902 stope.

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Evaporite surface The Laminated Beds contain a single, 3 m thick cycle, approximately 30m above the base, referred to in the mine as the 'C-D Unit'. This dark grey, delicately laminated calcisiltite contains pseudomorphs of probable evaporite crystals, now replaced by blocky calcite. It is capped by a bed of light grey nodular quartz, the 'Quartz Marker', which extends across the mine area. This may be up to 15 cm thick, consisting either of a massive layer or discrete nodules up to 6cm diameter. The upper surface is sharply truncated with a relief of several centimetres. In thin section the quartz nodules contain numerous inclusions of anhydrite, and in cores EP26 at 834.7m and N864 at 554.5m the bed consists entirely of anhydrite.

Large-scale erosion features In addition to the channels within the Laminated Beds, Muddy Limestone and Grainstone Unit, a large-scale feature forms the upper surface of the Micrite Unit (Fig. 4). There are dramatic variations in the thickness of this unit, which has been measured in over 70 cores.

Fig. 7. Characteristics of limestone surfaces in the Micrite Unit. (a) Stepped surface truncating layering. Core N509, at 45.5 m depth. Microstalactitic cement noted at 's' in overlying rock. Scale bar 6 mm. (b) Former planar surface showing irregular pits and pinnacles. Core N509 at 56.6 m. Scale bar 7 mm. The ornaments separate distinctive limestones.

DINANTIAN EMERSION SURFACES AND CHANNELS Cross-sections constructed using lithostratigraphical markers in the Grainstone Unit, and isopachs based on these data, reveal a striking topography on the upper surface of the Micrite Unit which forms a large-scale hummocky surface. The feature appears to represent real relief and is not an artefact of faulting. The surface is marked by broad highs and lows with a relief of up to 20 m and gradients of as much as 10°. Collectively these depressions form a lineament more than 2 km wide and 50 m deep. The floor of the feature includes smallerscale highs and lows oriented parallel to the strike of the main lineament. Breccias, pinnacles and hummocks extend across the entire upper surface of the unit, including the floor of this depression (Fig. 8), superimposed on the largerscale topography•

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Siliciclastic sediments overlying surfaces Siliciclastic intervals up to 50cm thick overlie brecciated surfaces, pinnacles and other truncation features, forming the bases to the depositional cycles that follow. There are three distinctive lithologies- silty mudstones, sandstones and green mudstones.

Silty mudstones. These dark coloured friable units commonly show centimetre-scale lamination, but this is disrupted locally and some beds are structureless (cf. Fig. 8). Where they overlie brecciated surfaces the mudstones have penetrated between blocks downwards through the breccias. Two types of clast are present: scattered echinoderm grains, and lithoclasts of the underlying rock. The echinoderm grains commonly carry syntaxial overgrowths, but the surfaces of such grains (including overgrowths) are often irregular. Lithoclasts up to 2cm in diameter are locally grain-supported. Sandstones. These are light brown and fine- to medium-grained sands including parallel and lenticular laminae up to 5 mm thick, and crosslamination with foresets dipping at 10-20 ° (core N629 at 356.2 m depth). Some units are graded; they are medium-grained at the base, passing up into silt-grade material. The sandstones contain scattered bioclasts and a few intercalated thin black siltstones, generally with sharp contacts between the two lithologies.

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Green Mudstones. Green illitic mudstones up to 50 cm thick are present in both the Laminated Beds and the Micrite Unit. They are structureless and friable, breaking into granular blocks with waxy lustrous surfaces. The green mudstone in the Laminated Beds lies approximately 22 m above the base, in the 'C-E' Unit of the mine terminology (core N975 at 509.8 m depth). It extends across the mine area and caps a coarsening-upwards sequence consisting of silty shales intercalating with, and passing up into, sandstones. In thin section the mudstone has a distinctive fabric of two sets of discontinuous clay films oriented at 90 ° to each other, which resembles the sepic-plasmic fabric of Brewer (1964). Randomly distributed silt-size siliciclastic grains (the skeletal grains of Brewer 1964) are typically surrounded by circumgranular cracks, with calcite replacing grain margins and penetrating fissures. A few grains form circular clusters which may represent burrow fills; other burrows are filled with pyritic faecal pellets.

214

G. RIZZI & C. J. R. BRAITHWAITE

The green mudstone in the Micrite Unit overlies and fills cavities in a brecciated surface about 10m above the base of the unit and is referred to as the 'Footwall Green Shale'. Its distribution is shown in Fig. 9. Lithoclasts within the mudstone generally have micritic coatings and a few grains are completely micritized. Clay cutans have formed around lithoclasts and line former cavities and bioclast moulds. Cutans consist of clay flakes that form monominerallic coatings, a few microns thick, made up of overlapping discontinuous laminae. Scattered, anastomosing, sub-parallel and planar sparfilled cracks are present in the clay, with circumgranular cracks surrounding individual clasts. The unit can be seen in cores N858 at 527.0 m and N1034 at 656.5 m depth, and also in mine level 1175 stope 274W.

C e m e n t s below surfaces Three non-ferroan calcite cements are present in the limestones below surfaces: microstalactitic cements, syntaxial overgrowths, and granular cements. All three line intergranular and moldic pores, pre-dating grain fracture. Microstalactitic cements are present in the Micrite Unit and are locally coarse enough to be visible to the naked eye. They are radial-fibrous and may include surface-parallel concentric layers of opaque inclusions. Syntaxial overgrowths are large and

inclusion-free and are absent from echinoderm grains only where surfaces were earlier coated or micritized. Granular crystals are equidimensional and, where pore-throats were sufficiently narrow, form meniscate clusters. They are also inclusion-free. Both granular and syntaxial crystals show euhedral terminations. Under cathodoluminescence all three cements are generally non-luminescent, but large crystals may contain 1-5 concentric bright-intermediate orange subzones, up to 10 #m thick.

Interpretation of Surfaces

Brecciated surfaces The jigsaw-fit of clasts and the continuity of bedding in brecciated surfaces points to an in situ fragmentation process, and this, together with their positions at the tops of cycles, suggests that they reflect subaerial exposure. Similar breccias capping shallowing-upwards cycles have been described by James (1984) and Meyers (1988). Meyers (1988) described palaeokarst profiles like those at Navan from Lower Carboniferous limestones in southern New Mexico; these breccias form similar rubble and fissure zones and are also overlain by muddy siliciclastic rocks. They are interpreted as a regolith produced by the dissolution of limestones by vadose meteoric water.

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DINANTIAN EMERSION SURFACES AND CHANNELS

Hummocky surfaces, pits and pinnacles The truncation of sedimentary structures and individual grains by these features indicates that the surfaces bearing them formed postdepositionally, and their sculpted nature points to dissolution as the main agent of erosion. This dissolution was not related to present (surficial) geomorphological processes. They are intraformational and may be overlain by sulphides (Anderson 1990), but none of them are mineralized and there is no basis for regarding them as hydrothermally related (cf. Qing & Mountjoy 1994). They are believed to reflect subaerial dissolution. Dissolution pits and flat-bottomed dissolution pans are common in limestone terrains (cf. Ford & Williams 1989) and similar surfaces have been recorded in the Vis~an of Derbyshire (Walkden 1974). Pinnacles and runnels are common features of present limestone surfaces where water is able to flow freely (cf. Ford & Williams 1989).

Evaporite-related surface The calcisiltite-anhydrite cycle in the Laminated Beds resembles Recent sabkha cycles described by Kendall (1984). The silty composition and delicate lamination are interpreted as reflecting deposition in a lagoon, with evaporite pseudomorphs and the lack of burrows and body fossils suggesting deposition in a stressed, hypersaline environment. The anhydrite nodules at the top indicate a transition to a more arid, peritidal zone. The evaporite is bounded by an irregular erosion surface, similar to surfaces formed in modern sabkha areas by wind deflation (Kendall 1984).

215

The rocks forming the Micrite Unit are interpreted as the deposits of a lagoon/tidalflat system. Anderson (1990) suggested that the channel bounding the upper surface of the Micrite Unit represents a vertically accreting system deposited synchronously with the mudstones of the unit. However, there is a lack of lithological continuity between the two facies, although it is argued that some mudstones near the top of the channel-fill may interfinger with the uppermost beds of the Micrite Unit. However, modern carbonate lagoons and tidal flats have insufficient relief to account for the dramatic variations in thickness of the Micrite Unit as a whole. These variations reflect a palaeotopography with a relief of at least 50m and imply that there is no interfingering between the two units. Similar features have been recognized at both Clogherboy and Tatestown, suggesting that this depression extended N W SE for at least 6km (cf. Andrew & Ashton 1985). The depression is interpreted as a palaeovalley; its surface bears in situ breccias, pinnacles and funnel-shaped pits indicating that the topography was generated by subaerial exposure. It is significant that channels at four stratigraphical levels formed in essentially the same palaeogeographical position, pointing to a wellestablished and long-lived drainage system in the hinterland which was reactivated each time sea level fell. All four channel intervals represent major periods during which the accreting margin was deeply incised. However, both the smallerscale palaeokarst surfaces (which are much more common) and the palaeosols also represent subaerial erosion and substantial breaks in deposition corresponding to relative sea-level falls.

Channels and other large-scale features Channels with distinctive coarse fills are present within the Laminated Beds, the Muddy Limestone, and the Grainstone Unit of the Pale Beds. These are of similar size and shape to coastal channel systems dissecting modern tidal flat, lagoon and barrier island deposits, and extending out into adjacent marine environments (cf. Jindrich 1969). The Limestone Conglomerate and Microconglomerate include fining-upwards cycles and cross-bedding consistent with deposition in such channels. Bioclasts and ooliths were probably derived from adjacent marine environments, while lithoclasts and siliciclastic grains are more likely to have originated from the hinterland.

Siliciclastic beds Siliciclastic rocks overlie many of the surfaces described, including the palaeotopography bounding the upper surface of the Micrite Unit. A similar association has been recognized in the British Dinantian (Walkden & Davies 1983) and elsewhere (James 1984; Meyers 1988). They represent sediment deposited upon previously karstified exposed surfaces (cf. Walkden & Davies 1983; Klau & Mostler 1986; Meyers 1988). Cross-lamination in sandstones indicates that their surfaces were rippled, with graded beds reflecting deposition from decelerating currents flowing over the surface. Comparable sandstones fill depressions in Triassic palaeokarst

216

G. RIZZI & C. J. R. BRAITHWAITE

surfaces in Austria (Klau & Mostler 1986). In Anglesey sandstones filling Dinantian palaeokarst are believed to have been transported over surfaces and along depressions during emergence of the platform (Walkden & Davies 1983). The green mudstones are characterized by clay cutans, a sepic-plasmic fabric, circumgranular cracks, micritized grains, and faecal pellets filling burrows, all features believed to reflect pedogenesis (Brewer 1964). Calcite replacement of siliciclastic grains is common in palaeosols such as those of the Old Red Sandstones of the Welsh Borders (Allen 1986) and are typical of caliche formed in a semi-arid climate. The green mudstones are thus regarded as palaeosols and provide further evidence of emergence. The green mudstone in the Micrite Unit overlies and fills fissures in a brecciated surface, reinforcing the view that breccias formed in a subsoil environment (cf. Wright 1982; Meyers 1988).

Cements The cements found below surfaces at Navan include microstalactitic, granular and syntaxial morphologies showing concentric non-luminescent and bright-intermediate orange luminescent growth-zone couplets. Microstalactitic cements reflect growth in the vadose zone, while granular and syntaxial morphologies may form in either marine or meteoric pore-waters (Meyers 1991; Lavoie & Bourque 1993). Early marine cements are not typically granular and patterns of luminescence resemble those described by Walkden (1987) which were interpreted as of meteoric origin. The cements therefore provide independent support for subaerial exposure. This, together with the timing of cementation (pregrain fracture), indicates that emersion and surface formation took place during the accretion of depositional cycles. Walkden (1987) suggested that similar Lower Carboniferous subaerial erosion events represent periods of up to several thousand years.

Discussion The depositional cycles at Navan vary considerably in thickness but each represents a sedimentological response to a relative sealevel change. The accommodation space required for cycle stacking is believed to have been generated by eustatic sea-level change,

tectonism or regional subsidence (cf. Sami & James 1994; Walkden 1987; Horbury 1989). Harwood & Sullivan (1991) have referred to synsedimentary tectonism at Moyvoughly 50 km to the SW of Navan. Evidence for relative changes in sea level is provided by the emersion surfaces described above but also by channels in the Laminated Beds, the Muddy Limestones and the Pale Beds. These channels were 20-50m deep, suggesting that incision was not the result of tidal action (cf. Ricketts 1994). Channels of different ages are located in the same general geographical area, and the long period over which they were active therefore points to a permanent drainage system on the adjacent land-mass. Reactivation of this drainage basin each time sea level fell accounts for the repeated extension of the channels across the platform. The surfaces described formed complex palaeotopographies with relief ranging from a few centimetres up to the 50m represented by the upper surface of the Micrite Unit, and indicate falls in relative sea-level of equivalent magnitude. Comparable sea-level changes are said to have occurred during accumulation of the Dinantian Urswick Limestone of NW Britain (Horbury 1989). Evidence of subaerial emergence has already been referred to elsewhere in the North Irish Midlands. Harwood & Sullivan (1991) described palaeosols from the Courceyan Moyvoughly Beds at Moyvoughly 50km west of Navan. Andrew & Poustie (1986) referred to meteoric phreatic cements at Tatestown, 3km west of Navan, and meteoric cements are also seen in the Pale Beds at Clogherboy, 3 km SW of Navan (Rizzi 1992). Finally, Pickard et al. (1992) described oolitic grainstones at Kentstown, approximately 15km ESE of Navan, in which meniscus cements and rhizoliths are well developed. All of these features reflect the passage of meteoric waters, and several are also indicators of emergence; these areas were therefore all subject to subaerial exposure. Clasts in the channel fills include palaeosols and fragments with meteoric cements. These suggest derivation from limestones that had already undergone meteoric cementation and were blanketed by palaeosols. The precise location of this emergent terrain is not clear. The channels strike roughly N-S, indicating that drainage may have been from either of these directions. However, since the palaeo-shoreline is believed to have been north of Navan at this time (Phillips & Sevastopolo 1986; Rees 1987), it seems that drainage originated from the north

DINANTIAN EMERSION SURFACES AND CHANNELS with the bed rock in this area including lithified limestones. No evidence of emersion has been found in the Shaley Pales, the Argillaceous Bioclastic Limestones or younger rocks at Navan. Although these may contain signatures of sealevel change, depths were too great to result in emergence.

217

Mines. The support of both these bodies is gratefully acknowledged. Although geologists at Navan Mine were always very busy, they gave freely of their time. We thank, in particular, J. Ashton, P. Powell, A. Black and M. Holdstock. C.J.R.B. thanks the Department of Geology and Applied Geology at Glasgow University for financial support during visits to Navan Mine. We also thank M. Philcox and P. Strogen whose careful reviews significantly improved an earlier draft of the manuscript.

Summary and conclusions (1) Post-Caledonian deposition in the Navan area began during the late Devonian-early Courceyan with Red Beds deposited in alluvial fans and braided streams. These were followed by Laminated Beds, consisting of barrier sandstones, tidal-flat/lagoonal mudstones and sabkha evaporites. The Muddy Limestone succession is dominated by bioclast-rich mudstones and wackestones deposited in an open lagoon. The overlying Pale Beds reflect the onset of shallow open-shelf conditions, but the Shaley Pale Limestones, Argillaceous Bioclastic Limestones and Waulsortian Limestones reflect a progressive deepening. These are all truncated by a major erosion surface, overlain by Chadian submarine debris-flows and limestone turbidites. In general, the succession indicates continuous deepening as subsidence outpaced sediment accumulation. (2) Deposition was not continuous. It was punctuated by relative changes in sea-level, generating sedimentary cycles in the lower part of the succession, causing emersion, and resulting in the formation of a variety of emersion surfaces. These include breccias, palaeosols, and karst-modified palaeotopography believed to have formed subaerially. The wider distribution of these features may be related to the exposed margins of Old Red Sandstone and Lower Palaeozoic rocks, but awaits further investigation. (3) The causes of both emergence and cyclicity are most likely to be found in eustatic sea-level fluctuations superimposed on crustalscale subsidence. (4) The sedimentary cycles and emersion surfaces formed in limestones that host Europe's largest Z n - P b deposit. The influence of these features on the distribution of mineralization here, and elsewhere in Ireland, should now be considered. The work was completed while G. Rizzi was in receipt of a Natural Environmental Research Council CASE award GT4/88/GS/l15 in conjunction with Tara

References ALLEN, J. R. L. 1986. Pedogenic calcrete in the Old Red Sandstones facies (late Silurian-Early Carboniferous) of the Anglo Welsh area, southern Britain. In: WRIGHT, V. P. (ed.) Palaeosols: their Recognition and Interpretation. Blackwell Scientific Publications, Oxford, 58-82. ANDERSON, I. K. 1990. Ore depositional Processes in the Formation of the Navan Zinc~Lead Deposit, County Meath, Ireland. PhD Thesis, University of Strathclyde. ANDREW, C, J. & ASHTON, J. H. 1985. Regional setting, geology and metal distribution patterns of the Navan Orebody. Transactions of the Institution of Mining and Metallurgy, Section B: Applied Earth Science, 94, B66-B93. - - & POUSTIE, A. 1986. Syn-diagenetic or epigenetic mineralization the evidence from Tatestown Zinc-Lead prospect, Co. Meath. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 281-296. ASHTON, J. H., BLACK, A., GERAGHTY, J., HOLDSTOCK, M. & HYLAND, E. 1992. The geological setting and metal distribution patterns of Zn-Pb mineralization in the Navan Boulder Conglomerate. In: BOWDEN, m. A., EARLS, G., O'CONNOR, P. G. & PYNE, J. F. (eds) The Irish Minerals Industry; 1980-1990. Irish Association for Economic Geology, Dublin, 171 210. , DOWNING, D. T. & FINLAY, S. 1986. The geology of the Navan ore body. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and genesis of mineral deposits in Ireland. Irish Association for Economic Geology, Dublin, 635-661. BREWER, R. 1964. Fabric and mineral analysis of soils. Kreiger Publishers, Wiley, New York. BRISKLY, J. A., DINGESS, P. R., SMITH, F., GILBERT, R. C., ARMSTRONG, A. K. & COLE, G. P. 1986. Localisation and source of Mississippi ValleyType zinc deposits in Tennessee, USA, and comparisons with Lower Carboniferous rocks of Ireland. In: ANDREW, C. J., CROWE, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 635-661.

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FORD, D. & WILLIAMS, P. 1989. Karst Geomorphology and Hydrology. Unwin-Hyman, London. HARWOOD, G. M. & SULLIVAN, M. 1991. Sedimentary history of the Moyvoughly area, County Westmeath: evidence for syn-sedimentary fault movements in a mixed carbonate siliciclastic system of Courceyan age. In: LOMANDO,A. & HARRIS, P. M. (eds) Mixed Carbonat~Siliciclastic Sequences. Society of Economic Palaeontologists and Mineralogists, Core Workshop. No. 15, 253-384. HORBURY, A. D. 1989. The relative roles of tectonism and eustacy in the deposition of the Urswick Limestone in south Cumbria and north Lancashire. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society, Occasional Publication, 6, 153 169. JAMES, N. P. 1984. Shallowing upward sequences in carbonates. In: WALKER, R. G. (ed.) Facies Models (2nd edn). Geoscience Canada, Reprint Series, 1, 213-229. JINDRICH, V. 1969. Recent carbonate sedimentation by tidal channels in the lower Florida Keys. Journal of Sedimentary Petrology, 39, 531-553. KENDALL, A. C. 1984. Evaporites. In: WALKER, R. G. (ed.) Facies Models (2nd edn). Geoscience Canada, Reprint Series 1,259-296. KENDALL, C. ST. G. C. & SKIPWITH, P. A. D'E. 1968. Recent algal mats of a Persian Gulf lagoon. Journal of Sedimentary Petrology, 38, 1040-1058. KLAU, W. & MOSTLER, H. 1986. On the formation of Alpine Middle and Upper Triassic Pb-Zn deposits, with some remarks on Irish carbonate hosted base metal deposits. In: ANDREW, C. J., CROWE, R. W. A., F1NLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 663-675.. LAVOIE, D. & BOURQUE, P. m. 1993. Marine, burial and meteoric diagenesis of early Silurian carbonate ramps, Quebec Appalachians, Canada. Journal of Sedimentary Petrology, 63, 233-247. LEES, A. & MILLER, J. 1985. Facies variation in Waulsortian build ups. Part 2: mid Dinantian build-ups from Europe and north America. Geological Journal, 20, 159 180. MCNESTRY, A. & REES, J. G. 1992. Environmental and palynofacies analysis of a Dinantian (Carboniferous) littoral sequence: the basal part of the Navan Group, Navan, County Meath, Ireland. Palaeogeography, Palaeoclimatology, Palaeoecology, 96, 175-193. MEYERS, W. J. 1988. Palaeokarst features in Mississippian limestones, New Mexico. In: JAMES. N. P. & CHOQUETTE, P. W. (eds) Palaeokarst. Springer Verlag, Berlin, 306-328. - - 1 9 9 1 . Calcite cement stratigraphy: An overview. In: BARKER, C. E. & KOPP, O. C. (eds) Luminescence Microscopy and Spectroscopy. Society for Economic Palaeontologists and Mineralogists, Dallas, Texas, Short Course 25, 133-148.

NOLAN, S. C. 1989. The style and timing of Dinantian syn-sedimentary tectonics in the eastern part of the Dublin Basin, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society, Occasional Publication, 6, 83 97. PHILCOX, M. E. 1984. Lower Carboniferous Lithostratigraphy of the Irish Midlands. Special Publication of the Irish Association for Economic Geology, Dublin. --1989. The Mid Dinantian unconformity at Navan, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society, Occasional Publication, 6, 67 81. PHILLIPS, W. E. & SEVASTOPOLO, G. D. 1986. The stratigraphic and structural setting of Irish Mineral Deposits. In: ANDREW, C. J., GROWL, R. W. A., FINLAY, S., PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, Dublin, 1-30. PICKARD, N. A. H., JONES, G. LL., REES, J. G., SOMERVILLE, I. D. & STROGEN, P. 1992. Lower Carboniferous (Dinantian) stratigraphy and structure of the Walterstown-Kentstown area, Co. Meath, Ireland. Geological Journal, 27, 35-58. QING, H. & MOUNTJOY, E. W. 1994. Origin of dissolution vugs, caverns, and breccias in the Middle Devonian Presqu'ile Barrier, host of the Pine Point Mississippi Valley type deposits. Economic Geology, 89, 858-876. REES, J. G. 1987. The Carboniferous Geology of the Boyne Valley Area, Ireland. PhD Thesis, Trinity College, Dublin. RICKETTS, B. D. 1994. Mud flat cycles, incised channels, and relative sea-level changes on a Palaeocene mud-dominated coast, Ellesmere Island, Arctic Canada. Journal of Sedimentary Research, B64, 211-218. RIZZl, G. 1992. The Sedimentology and Petrography of Lower Carboniferous Limestones and Dolomites; Host Rocks to the Navan Zn-Pb Ore Deposit, County Meath, Ireland. PhD Thesis, University of Glasgow. SAMI, T. & JAMES, N. P. 1994. Peritidal carbonate platform growth and cyclicity in an Early Proterozoic foreland basin, upper Pethei Group, northwest Canada. Journal of Sedimentary Research, B64, 111-131. STROGEN, P., JONES, G. LL. & SOMERVILLE, I. D. 1990. The stratigraphy and sedimentology of Lower Carboniferous (Dinantian) boreholes from west County Meath, Ireland. Geological Journal, 25, 103-137. WALKDEN, G. M. 1974. Palaeokarstic surfaces in Upper Visean (Carboniferous) limestones in the Derbyshire Block, England. Journal of Sedimentary Petrology, 44, 1234-1247.

DINANTIAN EMERSION SURFACES AND CHANNELS -1987. Sedimentary and diagenetic styles in late Dinantian carbonates of Britain. In: MILLER, J., ADAMS, A. E. & WRIGHT, V. P. (eds) European Dinantian Environments. John Wiley, Chichester, 131-158. - & DAVIES, J. 1983. Polyphase erosion of subaerial omission surfaces in the Late Dinantian of Anglesey, North Wales. Sedimentology, 30, 861-878.

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WRIGHT, V. P. 1982. The recognition and interpretation of palaeokarsts: two examples from the Lower Carboniferous of South Wales. Journal of Sedimentary Petrology, 52, 83-94.

Microfacies associations in Asbian carbonates: an example from the Urswick Limestone Formation of the southern Lake District, northern England A. D. H O R B U R Y

1 & A. E. A D A M S 2

1 Cambridge Carbonates Ltd., 47 Bedford Road, Ealing, London W13 OSP, and Department of Geology, Royal Holloway College, University o f London, Egham, Surrey TW20 OEX, UK 2 Department of Earth Sciences, University o f Manchester, Oxford Road, Manchester M13 9PL, UK Abstract: Semi-quantitative analysis of allochems from the Urswick Limestone Formation (Asbian) of the southern Lake District area of northern England has revealed a distinctive cyclicity of the microfacies. Cycle-top grainstone microfacies contain an algal flora comprising Koninckopora, Anatolipora and Polymorphocodium, with Girvanellafilaments and Ortonella lumps. Other allochems include intraclasts, large peloids and thick-shelled bivalves and gastropods. The middles of cycles are mostly packstones and micro-grainstones and contain allochems dominated by small peloids and the algae Kamaena,Kamaenellaand Epistacheoides, with the microproblematicum Ungdarellaand relatively high abundances of micritic-walled foraminifera such as endothyrids. Cycle bases contain a diverse algal assemblage including Coelosporella and Stacheoides, with other allochems represented by trilobites, ostracodes, Saccamminopsis, foraminifera such as Archaediscidae, the base late Asbian guide Howchinia, the base early Asbian guides Gigasbiagigas and Vissariotaxis, bivalves, small gastropods, bryozoans, sponge spicules and bored grains. Other allochems are found throughout most cycles, decreasing only in the very shallowest (intertidal) facies, or have an irregular distribution, and include brachiopod debris, crinoid ossicles and coral fragments. There are significant variations in allochem distribution according to palaeogeography. Close to the shelf margin there are higher abundances in the cycle top grainstones of the algae Koninckopora and Anatolipora, and also of the calcified filaments Girvanellaand Ortonella, with thick-shelled gastropods, intraclasts and coarse peloids. At cycle bases, echinoderm arm plates and bryozoans are particularly abundant in packstone-wackestone textures. Platform interior facies are differentiated into a diverse open-marine type, with a high total abundance ofbioclasts in the cycle base pack-wackestones including trilobites, Coelosporella,Stacheoides, Kamaena and bored grains, grainstones are dominated by small peloids, Kamaenella and Ungdarella), and a more restricted cycle type, in which total bioclast abundances are low. Cyclicity on a 2-20 m scale in the late Dinantian (Asbian and Brigantian) carbonates of Europe and North America has been well documented from outcrop macrofacies studies (Somerville 1979a, b,c; Walkden 1987). This cyclicity has been interpreted as a product of fourth order glacio-eustacy (Walkden 1987; Horbury 1989). However, there is relatively little published work on the microfacies aspects of the cyclicity, although the theses of Gray (1981) on North Wales and Horbury (1987) on the southern Lake District cover this in some detail. White (1992) examined thin sections prepared by Horbury, and this thesis contains much information ofpalaeoecological and stratigraphic significance regarding the distribution of foraminiferal genera and species within the microfacies/macrofacies models developed initially by Horbury (1987, 1989). There is abundant palaeontological litera-

ture on microfossils (e.g. Petryk & Mamet 1972; Mamet & Roux 1974; Mamet et al. 1980; Skompski 1984, 1987), but this is mostly related to the problems of classification of the microproblematica which are abundant in the Late Dinantian (e.g. Riding 1977). Relatively little of this literature covers the micropalaeontological aspects of these carbonate systems in a thorough sedimentological sense. The purpose of this paper is therefore to describe and interpret the palaeoenvironmental significance of the microfacies and allochems of Asbian cycle systems by reference to the macrofacies cyclicity. Location The Urswick Limestone Formation crops out along the southern margin of the Lake District

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 221-237.

222

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of northern England (Fig. 1). It is up to 150m thick and both the early and late Asbian are represented (Strank 1981; Horbury 1987). The rocks are strongly cyclic, with approximately 12-20 cycles developed, depending on the locality. The style of cycle development varies throughout the formation; at the base of both the early and late Asbian there are thin grainstone-dominated cycles separated by palaeokarsts and, in the late Asbian, by fenestral limestones. In the upper parts of the early and late Asbian, there are thicker cycles with packstone-wackestone bases and generally poorly developed grainstone cycle tops (Horbury 1989; Fig. 2). These stratigraphic variations are consistently developed throughout the 20 x 30 km study area, and may be comparable to cyclic carbonates described from the Sellafield area 50 km to the northeast (Barclay et al. 1994). In addition, some reference is made to the microfacies variations in contemporaneous outcrop material examined from the Oxwich Head Limestone on the Gower peninsula in South Wales by the second author. This paper is an extension to the work on K a m a e n e l l a presented in Adams et al. (1992), and also acts as a check on the broader significance of observations and conclusions derived from the Urswick Limestone study.

Database Some 3000 acetate peels and 700 thin sections formed the basis of this study. These were selected from 16 outcrop localities (Fig. 1), and are housed in the Geology Department at the Victoria University of Manchester.

Method of study A simple cyclicity (Figs 2, 3) based on four macrofacies (argillaceous packstones-wackestones, grainstones, porcellanous micrites and terrestrial carbonates and clastics) has been defined in the field (Horbury 1989), although it was realized that these facies were often gradational and the actual patterns were more complex (Horbury 1987). Extensive petrography was therefore undertaken because it was found initially that precise field identification of these more subtle sub-macrofacies patterns was unreliable, due to differences in weathering characteristics of shoreline, hillside and quarry exposures, and because of the effects of thorough bioturbation by Thalassinoides and complex diagenesis. Petrography reveals the presence of assemblages of allochems. Although some allochem

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assemblages contained allochems unique to that particular assemblage, most contained the same allochems, but in different proportions. Consequently, specific distinctive microfacies combining allochem assemblages and textural criteria were difficult to identify, since allochem compositions varied gradationally through the cycles. The problem is similar to that encountered with classification of the detailed variation in macrofacies. Position within a cycle and 'microfacies' could best be defined by sequential variations in allochem abundance and Dunham's classification. A sliding scale of 22 'microfacies' were identified, reflecting the gradational nature of allochem assemblages through an idealized cycle, with approximately one 'microfacies' per metre of an idealized cycle (Fig. 4). An attempt at further understanding of allochem distribution resulted in the development of a semi-quantitative method of description. For each microfacies up to 20 thin sections (where available) were selected and the allochems classified on a semi-quantitative logarithmic scale as follows (Table 1). In each case, a magnification of x 30 resulting in a field of view of 2.5 mm was studied; thin sections containing large colonial bioclasts such as Siphonodendron or Chaetetes, for example, were not counted and were considered as part of the macrofacies study (Fig. 3). The average number of each allochem per field of view was assessed, and a value per microfacies of each allochem could then be calculated. Given the sequential microfacies succession within an idealized shallowing-up depositional cycle, trends of allochem abundance could then be plotted (Fig. 4). From this figure, the association of allochems with particular positions within depositional cycles can be easily demonstrated visually. This made interpretation of depositional environments simpler,

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Argillaceous packstones and wackestones

Dibunophyllum Clisiophyllum Palaeosmilia Caninia Rotiphyllum Siphonodendron (junceum, sociale, martini, pauciradiale) Lithostrotion araneum Syringopora orthocones trilobites Staparollus turretellid gastropods ?Wilkingia productoids chonetids Thalassinoides Teichichnus

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Fig. 3. Idealized Asbian cycle showing macrofacies.

and helped to constrain the depositional significance of some allochems which had not previously been studied in a palaeoenvironmental sense. Point counting is unsatisfactory for description of carbonate microfacies, since important palaeoenvironmental criteria typically depend on the abundance of individuals, rather than being area-related. The semi-quantitative method of characterization allows consistent assessment of individual allochem abundances in each specimen that can be compared to texturally different specimens. It is also both faster and avoids problems of closure inherent in point-count data collection and interpretation.

Allochem distribution in an idealized cycle Cycles in the Urswick Limestone Formation vary between 2 and 20m in thickness (Horbury 1989). Allochem composition and trends in

Dunham's classification of the 'idealized' Urswick Limestone cycle in Figs 3 and 4 typically show a shallowing-upwards succession. However, at the base of some cycles the macrofacies and microfacies may indicate an upwardsdeepening trend, but these form only a minor part of the formation, are typically thin (< 1 m) and are considered later in the paper.

Description Cycle-base allochems. It is clear from Fig. 4 that some allochems are present only at the base of cycles in wackestone-packstone fabrics, for example, the algae Coelosporella, Saccamminopsis, Stacheoides, and other bioclasts such as trilobites, bryozoans and sponge spicules (Fig. 4, Figs 5a--c). Some small fragments of Coelosporella could be Kulikia (Mamet et al. 1980; Skompski 1986), although differentiation of these two when fragmented is difficult. Observations in

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Fig. 4. Distribution of 'bioclasts' through an 'ideal' cycle.

Mamet et al. (1980) suggested that Kulikia is part of a similar microfacies association, for example with Earlandia. Other bioclasts found in mmor quantities at cycle bases include Stacheoides and oncoids with coats of Girvanella. In addition, foraminifera show a great generic diversity at the cycle base (White 1992), with Archaediscidae, the base late Asbian guide foraminifera Howchinia, the base early Asbian guide foraminifera Gigasbia gigas and Vissariotaxis (Horbury 1987), and the families Lasiodiscidae and Tetrataxidae (White 1992) being notably common. Some bioclasts, such as the alga Coelosporella, foraminifera (family Tetrataxidae) and Saccamminopsis, are particularly notable at the very base of cycles, whilst other bioclasts such as the foraminifera family Lasiodiscidae only appear above the cycle base (Horbury 1987; White 1992).

Table 1.

Allochems per field of view

Value

Over 10 5-9 3-4 1-2

10 6 3 2

Ostracodes, bivalves, small gastropods and bored grains are developed throughout the idealized cycle, but are dominant in cycle-base to cycle-middle packstone textures (Figs 5d, e). At the base of cycles, bivalves are generally thinshelled and gastropods turreted. In the cycle base to cycle middle microfacies, a fine peloidal texture is locally preserved, such that cycle base microfacies are typically microsparry (Figs 5a-e) or micropeloidal (Fig. 5f) with much the same allochem abundances in both types of sediment.

Cycle-middle allochems. The middle of cycles mostly comprise clotted peloidal packstones and micro-grainstones and contain allochems dominated by small peloids, algae such as Epistacheoides, and the important palaeoberesellid algae Kamaena and Kamaenella with the microproblematicum Ungdarella. Kamaenella is particularly abundant in the middles of cycles, and may constitute up to 95% of all bioclasts present (Figs 5h, 6a). Adams et al. (1992) reported that at the locality of Ilston in the Gower area of South Wales, 67% of samples examined contain this bioclast as the most abundant, in 37% of the samples it comprised over half the bioclasts by volume, and in 12.8% of samples it comprised over 80% of the total bioclasts by volume. There are also relatively high abundances but often low

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ASBIAN MICROFACIES, S. LAKE DISTRICT generic diversity of foraminifera families, which are typically not well developed in the cycle-base lithologies, such as Endothyridae (Horbury 1987), Ozawainellidae and Loeblichiidae with occasional Tournayellidae (White 1992).

Allochems in cycle-top grainstones. Allochems found in grainstones at the top of cycles comprise Anatolipora, Girvanella, Ortonella, Koninckopora and Polymorphocodium, intraclasts, large peloids, ooids, thick-shelled bivalves and gastropods (Figs 6b-h, 7a). Commonly there are early marine isopachous cements (Figs 6g, h) and the coarse micritized grains often possess a matrix sediment of fine peloids which post-dates early marine cementation (Fig. 6g). Allochems in cycle-top lime mudstones and wackestones. In the uppermost porcellanous mudstones and wackestones, abundances of calcispheres, ostracodes, thin-shelled bivalves and gastropods, intraclasts and oncoids increase (Figs 7b-d). Sediment fabric often appears to have been originally a compacted micropeloidal grainstone, and may also have abundant microbial-thrombolitic textures preserved where relatively uncompacted. At the top of cycles, pedogenic features such as alveolar textures, laminar calcrete and rhizocretions are also common (Fig. 7e). Allochems that show irregular or no regular trends through cycles. Some allochems as identified in thin section are found throughout most cycles, decreasing only at the very top or base of cycles, or have an irregular distribution. They are all macrofossil fragments and include brachiopods, crinoid ossicles and corals. The development of

227

these macrofossils is often clearer from field logs, due to the difficulty of correctly identifying genera or species from fragments in thin sections; these relationships are discussed later.

Other trends and microfabrics. Echinoderms, ostracodes and foraminifera may locally be dominant in some samples, or absent from samples in which they are otherwise typical (Horbury 1987). At the base and middle of cycles, ostracodes are locally abundant, and often occur as nests or refuges within coral colonies such as Siphonodendron (Fig. 7f). Micritized omission surfaces are also commonly developed in the middle and particularly at the top of cycles (Fig. 7g).

Interpretation Water depths. Most Asbian bioclasts do not have modern analogues; however, their water depth ranges can be determined approximately by restoration of typical cycle thicknesses, allowing for sediment compaction. The lower limit of significant micritization and peloid development is taken at 10m water-depth, which is also the approximate depth of fair-weather wave-base and significant grainstone development. This also corresponds to the top of significant sponge spicule/bryozoan development and is the optimum depth of Kamaenella development. It is likely by extrapolation that the deepest water sediments were deposited at or about 15-20m water-depth. The shallowest water sediments are relatively rare, and represent high-energy intertidal/shallow subtidal environments (ooid shoals) and intertidal-supratidal low-energy restricted sediments (porcellanous micrites).

Fig. 5. Photomicrographs of important litho- and biofacies. Scale bar 200#m in all cases. (a) Cycle-base bioclastic wackestone, dominated by sponge spicules (now calcite casts, arrowed), and earlandiid foraminifera (e). Early Asbian, Farleton Fell. (b) Cycle-base bioclastic packstone. Diverse biota with fragment of Coelosporella (c), also earlandiid (a) and endothyrid (e) foraminifera, ostracode valve (o), bivalve (b), sponge spicule (s) and echinoderm ossicle (x). Early Asbian, Stainton Quarry. (c) Cycle-base bioclastic packstone with diverse biota including well-preserved Coelosporella thallus (c), earlandiid foraminifera (e), bryozoan (b), ostracode (o) and bivalve (v) fragments. Early Asbian, Stainton Quarry. (d) Cycle-base bioclastic packstone dominated by benthic foraminifera, mainly earlandiids (e) and endothyrids (d), also ostracod valves (arrowed) and Kamaenella in a microsparry matrix. Late Asbian, Sandside Quarry. (e) Cycle-base bioclastic packstone with a diverse biota including archaediscid (a) and endothyrid (e) foraminifera as well as earlandiids (r), ostracode valves (o), bivalve fragments (b) and echinoderm plates (x). Early Asbian, Farleton Fell. (f) Cycle-base bioclastic peloidal packstone-grainstone with small peloids in a sparite cement. Bioclasts include bivalves (b), ostracode valves (arrowed) and Epistacheoides (e). Early Asbian, Warton Crag. (g) Probable cycle-base bioclastic packstone, dominated by well-preserved Kamaenella showing branching (arrowed), from an intraclast in a cycle-top grainstone unit. Late Asbian, Stainton Quarry. (h) Cycle-middle grainstone dominated by comminuted Kamaenella (pale grey tubes and circular sections), with endothyrid foraminifera (arrowed); a generally lowdiversity biota. Early Asbian, Warton Crag.

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depositional setting. The bioclast assemblages indicate a typically open marine environment, with evidence of slight restriction and faunal impoverishment, mainly where the sediment was deposited in very shallow water environments. Additionally, totals of bioclast values for each microfacies show that the most diverse assemblages of bioclasts occurred close to the base of cycles (Fig. 8), suggesting that this was a highly productive depositional setting. The slight decrease in abundance of bioclasts at the cycle base (about 22 m beneath the idealized cycle top) reflects an increase in the proportion of lime mud and the development of more wackestone facies; work by Davies (1984) on slightly thicker Brigantian cycles in North Wales suggested that the lithofacies at 20 m or more beneath cycle tops comprises mostly mudstones and wackestones. This was probably the depth at which lowered pO2 and light penetration may have become critical factors in carbonate sediment production on these shelves. However, at the base of cycles the presence of bored grains, and the encrusting of grains by bryozoans and Girvanella, all indicate the presence of depositional hiatuses, probably during the lag time between rise in relative sea-level and the beginning of significant carbonate production (cf. Schlager 1981). The bioclasts that show little or no dependence on position within the cycle, as observed in thin section, are typically large robust organisms which were probably readily transported and did not break-up easily in the higher energy shallow-water environments. Most brachiopod, coral and echinoderm genera probably lived in the sub-fairweather wave-base environments as indicated by macrofacies, but would

229

have been transported into shallower water during storms. The presence of foraminifera as abraded fragments, in packstone intraclasts or, with micritic chamber fills, in grainstones, indicates that there was much reworking of material into shallow water (White 1992).

Palaeoenvironmental indicators. Algal genera have proved to be the most useful markers, since different algae tend to occur throughout each cycle, and enable a zonation to be built up. Algal families, (e.g. Dasycladaceae) are, however, environmentally wide-ranging (for example, containing Coelosporella, Kamaena, Kamaenella, Anatolipora and ?Koninckopora). Foraminifera are useful environment guide-fossils for the basal and middle part of cycles, but are not particularly abundant in the upper parts of cycles (White 1992). However, they are environmentally diagnostic at a family as well as genus level. Significance of Kamaenella.

Distribution of bioclasts through the depositional environments represented by the microfacies succession may have been controlled by the presence of abundant Kamaenella. For example, the distribution of foraminifera genera and palaeoberesellids is typically mutually exclusive through cycles in the Urswick Limestone Formation (White 1992). The Kamaenella-dominated middle part of the cycle separates the basal cycle allochem assemblages from the cycle-top allochem assemblages. White (1992) argued that " . . . t h e environments which allowed the palaeoberesellids to become established or resulted from their successful colonisation may have been restricted in some way which was

Fig. 6. Photomicrographs of important litho- and biofacies. Scale-bar 200/zm in all cases. (a) Cycle-middle bioclastic grainstone with a biota dominated by Kamaenella (arrowed) and occasional Kamaena (k), endothyrid foraminifera (e) and echinoderm plate (o), with well-developed inclusion-rich marine cements between bioclasts. Late Asbian, Sandside Quarry. (b) Cycle-middle bioclastic peloidal grainstone dominated by micritized unidentifiable debris and recognizable endothyrid foraminifera (e) and Kamaenella (arrowed) with traces of isopachous cement. Late Asbian, Back Lane Quarry. (c) Cycle-middle bioclastic peloidal grainstone, dominated by fine (?faecal) pellets, together with Ungdarella (u) and Kamaenella (arrowed). Late Asbian, Stainton Quarry. (d) Cycle-top bioclastic peloidal grainstone dominated by Ungdarella (u) and peloids formed from heavily micritized bioclasts (dark areas), also Kamaena (k). Late Asbian, Stainton Quarry. (e) Cycle-top bioclastic peloidal grainstone with large Koninckopora (k) in a matrix of comminuted Kamaenella (grey, inclusion-rich areas), foraminifera (f), echinoderm plates (e) and peioids (p). Early Asbian, Middlebarrow Quarry. (f) Cycle-top bioclastic peloidal grainstone with a diverse biota including Anatolipora (a), Ungdarella (u), bivalve (b) and ostracode (o) fragments, micritized foraminifera (f) and peloids formed from heavily micritized bioclasts (dark). Late Asbian, Back Lane Quarry. (g) Cycle-top intraclastic peloidal grainstone; some intraclasts are grapestones with superficial ooid coatings (g), with infiltrated fine peloidal sediment (p) and local isopachous cement. Late Asbian, Dunald Mill Quarry. (h) Cycle-top bioclastic peloidal grainstone with most peloids probably being heavily micritized Anatolipora and foraminifera (f). Isopachous cements have locally spalled off grains (arrowed) before precipitation of pore-filling cement. Late Asbian, Warton Crag.

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unfavourable to many foraminifera with the exception of Endothyridae and Archaediscidae". Similar exclusion of bioclasts by beresellid algae has been suggested for Donezella in buildups in the Namurian of Spain (Bowman 1979). The development of abundant Kamaenella thickets at or about fair-weather wavebase is typical of Asbian platform facies (Adams et al. 1992). Kamaenella may thus have served to prevent, by baffling or scarcity of growth sites, the colonization of habitats by other larger calcareous organisms. The high rates of production of Kamaenella sediment may also have swamped production of other bioclasts and allochems in this zone. It was noted during study of the Kamaenella cementstone reef at Dunald Mill Quarry (Horbury 1992) that the reef core facies contained few bioclasts other than the foraminifera family Endothyridae and the genus Earlandia, thickwalled ostracodes and the microproblematicum

231

Ungdarella. The few other bioclasts in the core facies were transported in, despite being present in abundance in flanking facies (e.g. echinoderms and bryozoans). Horbury (1992) concluded that the observed density of preserved Kamaenella in the reef core would probably not have been sufficient to have prevented other organisms from colonization of this substrate. However, the actual live reef may have been a site of denser Kamaenella encrustation. Adams et al. (1992) demonstrated that the Kamaenella grainstones were mostly reworked during storms, such that the Kamaenella thalli growing on the reef were periodically broken off, and were only rarely cemented into the reef framework. The exclusion of foraminifera from Kamaenella-dominated sediment has been attributed by White (1992) as being a function of substrate or local microenvironment control determined by algal colonization, with the few genera of foraminifera able to survive being present in high abundances (e.g. Endothyracea and Ozawainellidae and large Archaediscidae). These robust foraminifera also tend to be spherical to subspherical or spindle-shaped, and may have been adapted to cope with the higher-energy upper part of cycles, or were more resistant to abrasion (White 1992). Foraminifera from the Lasiodiscidae family and fragile members of Endothyraceae and Ozawainellidae are completely absent from higher parts of the idealized cycle, indicating the difficulty in distinguishing substrate versus preservation control on foraminiferal abundance in the geological record (White 1992). However, the notable changes in allochem composition at about the level of the fair-weather wave-base indicates that there was a significant environmental break at this level which was exploited and partly controlled by the abundant development of Kamaenella.

Fig. 7. Photomicrographs of important litho- and biofacies. Scale bar 200 #m in all cases. (a) Cycle-top bioclastic oolitic grainstone. Larger bioclasts (foraminifera, f) and large peloids have two or three superficial coatings, in a matrix of finer peloids and Kamaenalla (arrowed). Late Asbian, Stainton Quarry. (b) Cycle-top mudstone showing fenestral cavities (c), desiccation cracks (arrowed) and articulated ostracodes (o). Biota are of very low diversity and abundance. Late Asbian, Stainton Quarry. (c) Cycle-top microdetrital grainstone dominated by finely-comminuted Kamaenella (arrowed) and small peioids; this is probably a depositional grainstone that was later compacted. Late Asbian, Warton Crag. (d) Cycle-top bioclastic wackestone dominated by ostracode valves (arrowed) and with a few Kamaenalla fragments (k). Late Asbian, Farleton Fell. (e) Cycle-top palaeosol, showing alveolar textures lining/filling intergranular cavities (arrowed) and micritic bridges between grains (m). Grains visible include Ungdarella (u) and peloids (p). Late Asbian, Warton Crag. (f) Intra-Siphonodendron biota, typified by articulated ostracodes (o) and earlandiids (e) preserved between the corallites of S. junceum (s). Early Asbian, Warton Crag. (g) Micritized surface at cycle top. Micritization is of a fine earlandiid Kamaenella peloidal grainstone (g), and overlying sediment is a coarse grainstone with Ungdarella, foraminifera (f) and peloids (p). Late Asbian, Stainton Quarry. (h) Cycle-base biota, typified by Vissariotaxis (arrowed) in sponge spiculedominated wackestone-packstone. Early Asbian, Stainton Quarry.

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A. D. HORBURY & A. E. ADAMS

Allochem variation through the Urswick Limestone Formation Similar broad microfacies are developed throughout the formation. However, some allochems are stratigraphically controlled, in that they are not always present in otherwise similar microfacies throughout the formation. In the early Asbian there is an abundance of Coelosporella in cycle-base packestones and wackestones which is not mirrored in the late Asbian. Microfossils present in the late Asbian but absent in the early Asbian include Saccamminopsis in cycle-base packstones and wackestones; in grainstones of the cycle middle and top Ungdarella appears, and in cycle-top grainstones there are Anatolipora and Polymorphocodium. Microfossils commonly developed in the late Asbian but poorly developed in the early Asbian include Girvanella and Ortonella in cycle-top grainstones. These differences mirror established macrofossil variation in the local stratigraphy, for example, the presence of Davidsonina (Cyrtina) septosa in cycle-top grainstones in the late Asbian but not in the early Asbian (Rose & Dunham 1977; Horbury 1987), and the presence of Lithostrotion araneum in the cycle-base wacke-packstones of the late Asbian but not the early Asbian. Because these differences are typically of abundance rather than absolute presence or absence, caution must be applied in the use of these microfossils for precise stratigraphic subdivision, besides the obvious problems with the strong facies control on many of these microfossils. For example, Ungdarella in South Wales is only well developed in the latest Late Asbian where it appears to replace Kamaenella in abundance (Adams et al. 1992), whereas it is developed in the earliest late Asbian in the Urswick Limestone Formation. Walkden (1982) suggested a high energy depositional setting for Koninckopora, as in this study, whilst Schofield (1982, p. 33), working on the Holkerian of Derbyshire, noted the reverse, in that Koninckopora was indicative of low energy environments and the red alga Stacheia was indicative of high energy deposition. Reasons for this apparent change in habitat between the Holkerian and Asbian are unclear.

Microfacies variation across the platform There are significant variations in allochem distribution according to palaeogeography, which deviate from the 'idealized cycle' concept. The late Asbian shelf margin near Lancaster

contains higher than average abundances of the algae Koninkcopora and Anatolipora in cycle-top grainstones, and also of the calcified filaments Girvanella and Ortonella, with thick-shelled gastropods, intraclasts and coarse peloids. In cycle-base packstones and wackestones, which are generally thin in this setting (Horbury 1989; White 1992), echinoderm arm plates and bryozoans are commoner, particularly in the lee of Kamaenella-dominated reefal facies (Horbury 1992). White (1992) noted that the greatest diversity of foraminiferal families, typically Endothyridae and Archaediscidae, occurred within the medium-grained grainstones rather than in the crinoidal-bryozoan flanking facies to the Kamaenella-cementstone reef. Also, the foraminifera of the family Lasiodiscidae are restricted to shelf margin localities (White 1992). Platform interior facies are typically diverse open marine types, with a high total abundance of bioclasts in the cycle-base packstones and wackestones including notable trilobites, Coelosporella, Stacheoides, Kamaena and bored grains; cycle-top grainstones are dominated by small peloids, Kamaenella and Ungdarella. Storm-influenced platform interior microfacies mainly occur in the early Asbian in the Warton Crag area, and are typically defined by the presence of mid-cycle microfacies dominated by Kamaenella occurring in the basal part of cycles, indicating a localized lowering of wave-base. In South Wales, Kamaenella is similarly abundant in the "[agoonal' facies as developed in the northern exposures of the Gower outcrop (Adams et al. 1992). There are notable variations to the typical platform allochem assemblages and microfacies in the Farleton Fell/Holme Park Quarry/Hutton Roof area when compared with the platform interior to the east. Sponge spicules are common in the cycle-base wackestones and packstones of the early Asbian, but are rarely found in the early Asbian elsewhere. In the late Asbian, total bioclast abundances in cycle-base wackestones and packstones are low, and cycle-base wackestones are relatively porcellanous. The microproblematicum Ungdarella is found throughout the early Asbian, whilst elsewhere it is only found in the uppermost 20m (Horbury 1987). In the five depositional cycles comprising the early Asbian in the study area (Horbury 1989), the percentage of coarse peloidal grainstones present in the cycle-top grainstones decreases upwards through the formation as follows: from 40%-28%-28%-1%-3% (Horbury 1987). This reflects the coastal onlapping and development of the nearshore sediment variant of the

ASBIAN MICROFACIES, S. LAKE DISTRICT idealized cycle in the basal three cycles in the study area (Horbury 1989). In the condensed platform-interior grainstone section of late Asbian age in Stainton Quarry, there is a high abundance of ooids, intraclasts, thick-shelled gastropods and Koninckopora. This assemblage is similar to the platform margin microfacies assemblage, but also contains evidence of coastal onlap associated facies (see later section on sequence stratigraphy). Similar thin, condensed oolitic-intraclast dominated platform interior-facies, with only scarce Kamaenella, occur in the northern outcrop of the Asbian succession in South Wales (Adams et al. 1992), and may also indicate the presence of an onlap margin.

Sequence stratigraphic controls on microfacies development Most of the marine sediment in the Asbian platform stratigraphy was deposited during sealevel highstands, at eustatic maxima (Horbury 1989). The idealized microfacies 'cycle' reflects, therefore, highstand deposition, characterized by a bioclastic-peloidal composition. This is similar to the specialized shelf margin sediment composition of the late Asbian and early Brigantian of northern Derbyshire noted by Gawthorpe & Gutteridge (1990). In any cyclic stratigraphy, there must be a record of the lowstand and transgressive systems tract (Vail et al. 1977). In the case of the Asbian of the study area, at Leaper's Wood Quarry there is a 26 m deep palaeovalley fill, recording karstification, lowstand and transgressive systems tract deposition which is significantly different to the typical platform-interior highstand facies. There are also lower relief features at other localities that have similar lowstand/ transgressive systems tract facies, allowing a more detailed picture to be built up for facies/ microfacies changes through a complete cycle (Fig. 9). Lowstand deposits are reflected by a variety of palaeokarst and local palaeosol features, followed by fluvial-estuarine-fan delta clastic carbonates. These have a scarce but completely different biota in growth position compared with the biota of the highstand carbonates, with a dominance of pectinid and other bivalves (Horbury 1987). All other bioclasts present show evidence of significant abrasion and transport, and were probably derived from karstification of highstand carbonate. Transgressive systems tract sediments are typified by an abundance of oolitic-peloidal-

233

microbioclastic grainstones, which are locally associated with plant debris, stromatolites, peritidal micrites and oncolitic limestones where the preserved transgressive systems tract is a thin unit (for example, at Trowbarrow Quarry where an oolite/stromatolite unit marks the initial flooding of a late Asbian palaeokarst). These sediments are faunally impoverished, with only occasional Chaetetes macrofauna observed in growth position and no significant in situ microbioclastic components. The differences in macrofacies/microfacies of the lowstand, transgressive systems tract and highstand sediments is indicative of a freshwater/brackish lowstand system, very shallow marine transgressive systems tract sediments and the fully open marine highstand system. During the lowest sea levels, the connection of the lowstand rift basin seas with open oceanic circulation was probably very poor, and fresh water may have significantly affected at least the surface waters of these basins. For this reason, brackish faunas are developed close to the basin margin. During transgression, the thick units of oolite developed at Leaper's Wood Quarry in the palaeovalley fill reflect deposition in a coastal onlap setting, with aggradation of highenergy shallow-water environments which later spilled out onto the karstified platform top. Only following the maximum flood event did the platform carbonates begin to accumulate, following a hiatus at which time grains and intraclasts were bored, and transgressive systems tract allochems were bioturbated into the overlying units. Highstand deposition commenced with water depths of up to 20 m and fully open marine conditions.

Relationship of microfacies to macrofacies The microfacies broadly correspond to the macrofacies of Horbury (1989) as noted in Fig. 4. Macrofossil elements show a similar style of variation in abundance through the idealized cycle as do the microfauna and flora (Fig. 3). Combination of both macro- and micro-facies/ faunal/floral data have helped to constrain four marine-environment groups (Horbury 1987). These were used to interpret the anatomy of the platform interior environments, and comprise: (a) sub-fairweather wave-base mudstones to packstones; (b) storm-influenced sub-fairweather wave-base packstones and grainstones; (c) above fair-weather wave-base shoal grainstones; and (d) intertidal-supratidal flat low-energy mudstones and wackestones.

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ASBIAN MICROFACIES, S. LAKE DISTRICT From these environment groups, deviation of cycles from the 'idealized cycle' are better accounted for. Horbury (1987) defined a total of seven sedimentary systems derived from an understanding of these local microfacies distributions. There are three variants in the early Asbian (nearshore cycle, storm cycle and normal shelf cycle types) and four in the late Asbian (normal-marine and restricted-marine cycle bases, normal shoaling and platform-edge shoaling cycle tops). Platform margin facies of early Asbian age were not encountered in the study. From these geographical variations in microfacies and macrofacies data, a schematic chronostratigraphy of a shelf cycle can be constructed (Fig. 9). This must be used in conjunction with the 'idealized' cycle in order to interpret fully the depositional environments across the study area.

Discussion The basal parts of the cycles contain many bioclasts which have been used to define formation and stratigraphic tops, e.g. Saccamminopsis, Girvanella, Howchinia and Vissariotaxis (Fig. 7h). These 'diagnostic' bioclasts are only present where the style of cyclicity is such that the basal facies of cycles are developed. Since development of cycle-base facies was typically a function of local subsidence style (Horbury 1989), the stratigraphic precision of these bioclasts may be called into question. The main formation/chronostratigraphic boundaries in the southern Lake District all coincide with major sequence-stratigraphic boundaries and changes in style of stacking of sedimentary cycles. Where major lithology changes occur, the 'facies boundaries' typically reflect missing gradational section, and changes in relative sea-level may possibly be inferred to control such surfaces, both at the top and base of cycles and within cycles. The generally flat-bedded character of the Urswick Limestone in outcrop, together with the general scarcity of progradational features, suggests that the platform top mostly comprised very low angle ramps and banks over which gradual environmental variation took place. Rapid lateral microfacies changes on the platform top are generally scarce, with evidence of extensive coastal onlap and storm activity controlling variations in the earliest Asbian. During the late Asbian, coastal onlap was important only in the west at Stainton Quarry. Vertical variations from cycle to cycle reflect the longevity of the coastal onlap

235

influence, and also reflect variations in cycle stacking patterns and possibly the development of the platform margin. Compositions of highstand, lowstand and transgressive systems tract sediments are important in understanding the effects of eustacy on the half-graben basin system. It is obvious from this study that the main production of sediment as loose grains was during the highstand, whilst the subtidal deposition during the lowstand was too brackish and during the transgressive systems tract was too restricted to produce a significant amount of loose sediment. The only sediment shed at this time was probably meteorically stabilized fragments of karstified limestone (Horbury & Gawthorpe 1993). This has implications for the timing of platform progradation and the relationship of turbidite type to sequence stratigraphy in half-graben basins. Palaeogeographies should illustrate these three extremely different types of depositional setting by more than one map reconstruction.

Conclusions Microfacies indicate strong differences between highstand, transgressive and lowstand systems tract deposition. The highstand systems tract sediments typically comprise open marine platform-top carbonates deposited in water depths of 0-20m. Lowstand systems tracts were brackish and coastal (where they are preserved), whilst the transgressive systems tracts were dominated by high-energy shallow subtidal-intertidal settings typical of coastal onlap margins. The most sensitive environmental indicators through an 'ideal' cycle are calcareous algae, both in terms of water depth indication and position on the platform. Calcareous algae also form the major bioclast component in most of the platform sediment. Foraminiferal genera and families are also indicative of depositional environments, but only in the middle and base of cycles. Other types of bioclasts and allochems are generally restricted to the top or base of cycles (e.g. ooids and bryozoans), or are developed throughout cycles and have little diagnostic use as microfacies or environmental markers (e.g. brachiopods and echinoderm debris). The development of abundant Kamaenella, which grew in thickets at or about wave-base, probably controlled the division of cycles into three broad biological zones.

236

A. D. HORBURY & A. E. ADAMS

Many aUochems used for stratigraphic purposes in the Late Dinantian are only found in cycle-base lithologies. Their presence or absence is mostly dependant on cycle stacking patterns and was ultimately controlled by local tectonics rather than by biological evolution. Cycle bases were the most productive environments in terms of diversity. Microfacies can be correlated with, and help constrain, interpretation of field macrofacies. Both macro- and microfacies are intergradational, with no 'sharp' compositional boundaries due to bioturbational mixing of sediment and habitat overlap. Where major lithology changes occur, 'facies boundaries" typically reflect missing gradational section, and changes in relative sea-level may possibly be inferred to control such surfaces. Sediment accreted vertically as a system of very low angle ramps and banks within each cycle. There is little or no evidence of significant progradational infiU of the platform-top accommodation space. Lateral variations in microfacies were probably due to coastal onlap and storm activity; vertical variations from cycle to cycle mainly reflect the longevity of coastal onlap systems, and cycle stacking patterns. Important microfacies variations from the ideal platform interior cycle comprise shelf margin, storm-dominated, restricted platforminterior and nearshore coastal onlap cycle variants.

References ADAMS, A. E., HORBURY, A. D. & RAMSAY, A. T. 1992. Significance of palaeoberesellids (Chlorophyta) in Dinantian sedimentation, UK. Lethaia, 25, 375-382. BARCLAY, W. J., RILEY, N. J. & STRONG, G. E. 1994. The Dinantian rocks of the Sellafield area, west Cumbria. Proceedings of the Yorkshire Geological Society, 50, 37-49. BOWMAN, M. B. J. 1979. The depositional environments of a limestone unit from the San Emiliano Formation (Namurian/Westphalian), Cantabrian Mts., NW Spain. Sedimentary Geology, 24, 2543. DAVIES, J. R. 1984. Sedimentary cyclicity in late Asbian and early Brigantian (Dinantian) limestones of the Anglesey and Llandudno districts, North Wales. Proceedings of the Geologists Association, 95, 392-393. GAWTHORPE, R. G. & GUTTERIDGE, P. 1990. Geometry and evolution of platform-margin bioclastic shoals, late Dinantian (Mississippian), Derbyshire, UK In: TUCKER, M. E. (ed.) Carbonate Platforms. Special Publication of the International Association of Sedimentologists, 9, 39 54.

GRAY, D. I. 1981. Lower Carboniferous Shelf Palaeoenvironments in North Wales. PhD Thesis, University of Newcastle-upon-Tyne. HORBURY, A. D. 1987. Sedimentology of the Urswick Limestone in South Cumbria and North Lancashire. PhD Thesis, University of Manchester. - - 1 9 8 9 . The relative roles of tectonism and eustacy in the depositon of the Urswick Limestone Formation in south Cumbria and north Lancashire. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication, 6, 153-169. - - 1 9 9 2 . A Late Dinantian peloid cementstonepalaeoberesellid buildup from North Lancashire, England. Sedimentary Geology, 79, 117-137. -& GAWTHORPE, R. L. 1993. Sequence Stratigraphic Controls on Highstand vs Lowstand Shedding: a Diagenetic Perspective. Abstracts of BSRG workshop, Manchester University, December 1993. MAMET, B. L. & ROUX, A. 1974. Sur quelques algues tubulaires scalariformes de la Tethys Palaeozoique. Revue de Micropaleontologie, 17, 134-156. --, DEJONGHE, L. & Roux, A. 1980. Sur la presence de Kulikia (Dasycladac~e) dans la Vis~en des Grandes Malades (Jambes). Soci~te Belgique de G+ologie, Bulletin, 89, 291-295. PETRYK, A. A. & MAMET, B. L. 1972. Lower Carboniferous algal flora, southwestern Alberta. Canadian Journal of Earth Sciences, 9, 767-802. RIDING, R. 1977. Problems of affinity in Palaeozoic calcareous algae. In: FLUGEL, E. (ed.) Fossil Algae, Recent Results and Developments. SpringerVerlag, Berlin, 202-211. ROSE, W. C. C. & DUNHAM, K. C. 1977. Geology and Hematite Deposits of South Cumbria. Memoir of the Geological Survey of Great Britain. SCHLAGER, W. 1981. The paradox of drowned reefs and carbonate plaforms. Geological Society of America Bulletin, 92, 197-211. SCHOFIELD, K. 1982. Sedimentology of the Woo Dale Limestone Formation of Derbyshire. PhD Thesis, University of Manchester. SKOMPSKI, S. 1984. The functional morphology of the Carboniferous dasycladacean genus Kulikia. Neues Jahrbuch fur Geologie & Palontologie Monatshefte, 7, 427 436. - - 1 9 8 7 . The dasycladacean nature of Late Palaeozoic palaeoberesellid algae. Acta Geologica Polonica, 37, 21-31. SOMERVILLE, I. D. 1979a. A sedimentary cyclicity in early Asbian (Lower DI) limestones in the Llangollen district of North Wales. Proceedings of the Yorkshire Geological Society, 42, 397-404. 1979b. Minor sedimentary cyclicity in late Asbian (Upper D1) limestones in the Llangollen district of North Wales. Proceedings of the Yorkshire Geological Society, 42, 317-341.

ASBIAN M I C R O F A C I E S , S. L A K E D I S T R I C T - - 1 9 7 9 c . A cyclicity in early Brigantian (D2) limestones east of the Clwydian Range, North Wales, and its use in correlation. Geological Journal, 14, 69-86. STRANK, A. R. E. 1981. Foraminiferal biostratigraphy of the Holkerian, Asbian and Brigantian stages of the British Lower Carboniferous. PhD Thesis, University of Manchester. VAIL, P. R., MITCHUM, R. M. & THOMPSON, S. III. 1977. Seismic stratigraphy and global changes of sea level, Part 4: Global cycles of relative changes of sea level. In: PAYTON, C. A. (ed.) Seismic Stratigraphy. Applications to Hydrocarbon Exploration. American Association of Petroleum Geologists Memoir 26, 83-97.

237

WALKDEN, G. M. 1982. Field Guide to the Lower Carboniferous Rocks of the South-east Margin of the Derbyshire Block." Wirksworth to Grangemill. Department of Geology and Mineralogy, Unversity of Aberdeen. 1987. Sedimentary and diagenetic styles in late Dinantian carbonates of Britain. In: MILLER, J., ADAMS, A. E. & WRIGHT, V. P. (eds) European Dinantian Environments. Wiley, Chichester, 131-155. WHITE, F. M. 1992. Aspects of the Palaeoecology and Stratigraphic Significance of Late Dinantian (Early Carboniferous) Foraminifera of Northern England. PhD Thesis, University of Manchester.

The stratigraphy and cyclicity of the late Dinantian platform carbonates in parts of southern and western Ireland STEPHEN

J. G A L L A G H E R

Geology Department, University College, Belfield, Dublin 4, Ireland Present address." School o f Earth Sciences, University o f Melbourne, Parkville, Victoria 3052, Australia Abstract: The late Dinantian platform carbonate successions of the Burren, Buttevant and Callan areas in Ireland have been correlated using detailed litho- and biostratigraphy. Several subdivisions or lithofacies associations (LA) of the Asbian to Brigantian part of the succession have been recognized in each area (lithofacies associations 1-5). Holkerian(?) to early Asbian ramp carbonates of LA 1 underlie the late Asbian successions in both the Burren and Buttevant areas. The influx of the bilaminar palaeotextulariid Cribrostomum lecomptei coincides with the onset of late Asbian shallow-marine cyclic platform sedimentation (LA 2) in all areas. The sediments of LA 2 have abundant Kamaenella and foraminifera, and are characterized by palaeokarstic surfaces and shales that cap a minimum of nine shallowing-upward minor cycles in the Burren and Buttevant areas. Two new biostratigraphic subdivisions (Cf671 and Cf672) of the late Asbian/Cf67 subzone are described for the first time in this part of the succession. Brigantian sedimentation (LA 3) is typified by amalgamated beds of crinoidal limestone and abundant Asteroarchaediscidae. Thin peloidal limestones with coral thickets cap shallowing-upward cycles within this succession. These cycles, which are thinner and less numerous than those of the underlying late Asbian, rarely terminated in emergence, but most reached shallow subtidal depths prior to the next transgression. The change in cyclicity style across the Asbian/Brigantian boundary may be related to the sedimentation rates of the crinoidal limestones, due to increases in cyclic oscillation in the Brigantian. Brigantian LA4 is characterized by deep subtidal cherty limestones, with abundant algal-coated wackestone intraclasts. Fasciellaand Howchiniabradyana are typical microfossils. No cyclicity is observed, probably owing to the deep subtidal nature of this unit. LA4 is overlain by another unit of cyclic crinoidal limestones (LA 5) in the Burren, which has no correlatives in the other areas studied. The succession in all areas is unconformably overlain by Namurian siliciclastic rocks. The nature and number of minor cycles in the late Dinantian of Ireland is similar to those of platform successions of the same age elsewhere in the British Isles, suggesting that eustatic changes were one of the major controls on cyclicity during the late Dinantian.

The late Dinantian palaeogeography of Ireland (Fig. 1) was dominated by a wide (c. 500 kin) shallow-marine carbonate platform bounded by land to the north and by the South Munster Basin to the south, with at least two intervening fault-controlled intraplatform basins, the Dublin Basin and the Shannon Trough (Fig. 1; Sevastopulo 1981; Somerville & Strogen 1992; Pickard et al. 1994). During Asbian time mud-mound development was locally important, especially near the margins of the Dublin and Shannon basins. Apart from relatively deep-water limestone-shale successions in northwest Ireland (Schwarzacher 1964; Sevastopulo 1981) and the Dublin Basin (Jones et al. 1988; Strogen et al. 1990), most of Ireland was covered by shallowmarine fossiliferous carbonates during the Asbian. A similar palaeogeography persisted into Brigantian times, although the first Yoredale cyclothems appear in the Sligo/Leitrim area

during this time (Sevastopulo 1981). Late Dinantian platform carbonate sedimentation ceased with uplift and the unconformable deposition of Namurian deltaic clastic rocks. Cyclicity in the late Dinantian platform carbonates of Ireland has been described by Schwarzacher (1964, 1975, 1989, 1993) from the Sligo region in northwest Ireland and from the Burren area, Co. Clare (referred to in Walkden 1987). Although the nature of the cyclicity is well known from the mixed carbonate-clastic Sligo successions (Schwarzacher 1993), knowledge of the nature and extent of cyclicity in the carbonate successions of this age in Ireland is scarce. Recent work by the author (Gallagher 1992) has revealed the presence of well-developed cyclicity in the late Asbian to Brigantian carbonates of the Buttevant, Burren and Callan areas, lying about 100 km apart in southern and western Ireland (Fig. 1). These cycles are not

From STROGEN,P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 239-251.

240

S. J. G A L L A G H E R --[t =acei

~

Shales and sandstones.

~ Cyclic crinoidal I ¢ * " * 1 limestones.

~ ,~,

Bedded chert, free limestones.

c=°

~

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-

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,-

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.

200

C1 Localities A= The Callan area in Co.s "Hpperary/Kilkenny. B= The Buttevant area in Co. Cork. C= The Burren area Co. Clare. C1= The Ballard Bridge Section. C2= The Black Head Section.

:-~ . I1,1.

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abbreviations

D.B.= Dublin Basin. S.T.= Shannon Trough. S.M.B.= South Munster Basin. S.I.P.= Southern Irish Platform

S.M.B.

Fig. 1. A stratigraphic correlation of the late Dinantian platform carbonates from southern and western Ireland. The vertical scales in columns A, B and C are in metres. The inset map shows the approximate position of the major intraplatform basins and the northern shoreline.

only similar throughout the area studied, but they are also comparable to cycles reported elsewhere in the British Isles (Walkden 1987). This study combines detailed lithostratigraphic analysis with a well-constrained biostratigraphy to provide a framework in which to compare and correlate cycle style across the Southern Irish Platform (Fig. 1). Using this combined

approach, the late Asbian to Brigantian succession of the three areas studied can be subdivided into five divisions. Since formal lithostratigraphies have been independently erected for these three areas (Fig. 1), it is not possible to give these divisions formal status and they are therefore referred to informally as lithofacies associations (LA). They are defined as follows:

LATE DINANTIAN PLATFORM CARBONATES, IRELAND LA 1: ?Holkerian to early Asbian non-cyclic ramp carbonates LA 2: a late Asbian cyclic unit LA 3: an early Brigantian cyclic unit LA 4: a late Brigantian, cherty, non-cyclic unit LA 5: a late Brigantian cyclic unit.

241

cf. cornucopiae, and the brachiopods Septarinia leuchtenbergensis, Latiproductus latissirnus, Gigantoproductus cf. giganteus, and Semiplanus

fragiZis. Microfossils

Late Asbian to Brigantian biostratigraphy A significant part of this study involved detailed biostratigraphy using macrofauna (corals and brachiopods) and microfossils (foraminifera and calcareous algae). The focus of this paper concerns the late Asbian to Brigantian cyclicity and stratigraphy. Details of the ?Holkerian to early Asbian biostratigraphy of the successions studied will be described elsewhere. It was possible to erect two new biostratigraphic units in the late Asbian stage/Cf6~/subzone that can be correlated with other late Dinantian platform successions in Co. Limerick (Somerville et al. 1992), in southern Co. Cork (Gallagher 1992) and Belgium (Conil et al. 1990). The first section of this paper will therefore include a brief review of the proposed late Asbian biostratigraphic scheme (Fig. 2).

Macrofauna The macrofaunal zonation schemes outlined by George et al. (1976) and Mitchell (1989) for Great Britain, and by Conil et al. (1990) for France/ Belgium, were used for this work. The first significant macrofaunal biostratigraphic markers Siphonodendron junceum and Dibunophyllum bipartitum appear at the base of LA 2 and indicate an Asbian age. The late Asbian brachiopod Davidsonina septosa, abundant in all areas, first occurs in the middle of LA2. Other diagnostic macrofauna which first occur in the uppermost part of LA2 are Gigantoproductus cf. edelbergensis, Axophyllum vaughani, Clisiophyllum garwoodi, Lithostrotion maccoyanum, Siphonodendron sociale, Diphyphyllum lateseptatum and Gigantoproductus edelbergensis. It was possible using macrofauna to define the Asbian/ Brigantian stage boundary, which coincides with the LA2/LA3 boundary in the Burren, using the first appearance of the diagnostic Brigantian coral Palastraea regia. It was more difficult to do so in the other areas due to poor exposure. Other rare Brigantian macrofauna are confined to strata above the base of the Brigantian; these include the corals Lonsdaleia duplicata, Actinocyathus floriformis and Caninia

All three successions were sampled at regular intervals at an average spacing of 1-2 m, and over 800 thin sections were examined. The ranges of important fossils are illustrated in Fig. 2. The microfossil zonation scheme used incorporates those of George et al. (1976), Conil et al. (1980), Fewtrell et al. (1981 ), Somerville & Strank (1984), Conil et al. (1990) and Somerville et al. (1992). The base of LA 2 coincides with the base of the Cf6"), Neoarchaediscus subzone (---early/late Asbian boundary). The base of the subzone is defined by the first appearance of bilaminar palaeotextulariids e.g. Cribrostomum lecomptei, the calcareous red alga Ungdareila, and Koskinobigenerina. The microfossil assemblage compares with that recorded from the laterally equivalent Dromkeen Limestone Formation in Co. Limerick, (Somerville et al. 1992). Assemblages indicative of the Cf6-~ subzone range throughout LA2. Variations in the microfossil assemblages allow the lower part of the subzone (Cf6"~l) to be distinguished from the upper part (Cf6"72); these appear to correspond with the unnamed subdivisions of the subzone reported by Conil et al. (1990, fig. 2). The formal definitions of the new subdivisions are as follows: (i) Zone Cf6 (Asperodiscus (= Neoarchaediscus) zone) (=Asbian-Brigantian). Definition (Conil et al. 1977, 1990; Paproth et al. 1983) based on the first appearance of Neoarchaediscus (Asperodiscus), Neoarchaediscus (Neoarchaediscus) and archaediscids at angulatus stage. Vissariotaxis compressa, which in Belgium occurs at the base of this zone, was found in this work to appear first at some distance above the base. 'Bilaminar' palaeotextulariids appear in Belgium at the base of Cf6. (ii) Subzone Cf6"~ (= late Asbian). Definition (Conil et al. 1977, 1990; Paproth et al. 1983) based on the first appearance of bilaminar palaeotextulariids (in this work), Cribrostomum, Palaeotextularia ex. gr. longiseptata and the monolaminar palaeotextulariid Koskinobigenerina. In Belgium the base is also defined on the first occurrence of Howchinia bradyana and Bradyina rotula, but in the areas studied here these taxa do not appear at the base of this

242

S. J. G A L L A G H E R

.Lithofacies

... K:a..m..a..e~tLa. (8 .L..G..A) .........................................

Tetrataxis conica .... /~.0q..~t:~'k..o..qq...~...te. nu!ra...mo...~ .(A. L.G..A.) ............. Paraarchaediscus with nodes

Paraarchaediscus @ involutus Paraarchaediscus @ concavus "M'Sih~ii~ihiJii~irlS~ii~i~e'~Liiiii~id

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

...A.sterosphaera Plectogyranopsis ...Kon!nck°~ra.m/nuta ( A L G.(~) ...................... Pseudolituotubn

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

. .A..r~a.ediscus

e..a~g.~La.t~

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

...Paraarc!?f!edisqus.O anf}tdatus .................. Globoendothyra globulus

.Bog... ,She!!.a.. z.ig.a~ns!.s ..................................... ...Pa/a~te.~u!at~a..e.x:..g.r:...conso.b...n.'n.a ............. .. ,~eoa~h.a..~LSco. s..(..N~...r~....a~./s~.s)...sp;.. Neoarchaediscus (Aspq~iscus) SO. D.~fania. b.ilqba (.PFo..B..L!E!~!AT!CI~!M) ........ Koninckopora inflata (ALGA)

Koskinotextularia Forschia prisca Millerella Planoarchaediscus concinnus Gigas~ia gigas l .... ii II..K..o...n/nckP ..p~..~ m0. ttq~m..a.n .s..i.(AL..G...A.). .............

ii

Pseudolituotubella

=1

Scaleb/ina Valvuline/la Exvotarisella (ALGA) . . . . . Asteroarchaediscus .... V!s~ ~tax~i-~ .qom,om.~.a ................................ .....K.u~ki.a (A.t:G,5) .................................................

Endostaffella ,,. B!!am!.nar .p.a_!a._e~...textuladid Ungdarella (ALGA) . . . . . .

...P.. a/a.~te.xt.u./, arfa. ex:.gr:../on, g!.s..e.t?t..a.ta. .......... .... K~k.,'nqU~.~.,'.n..a

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

Cribrostomum lecomptei Condrustrella Biorbis ...H.a.P./.o~hraeme!/..a

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

Cnbrospira ... B!t~.~.d.xa

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

Nevillea ...H..o.~h~n~a.q.~a.na

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

...KP.n?nck~.oPO.~.SP:.B.(ALGA).......................... Fasciella kizilia (ALGA) Neoarchaediscus stellatus Bradyina rotula ....... ..Sa~amm/nqo.S.is ............................................... Neoarchaediscus incertus .. L.qe.b!ich/a p...a.~am~S. .......................... ...C...a~f...o/./u..m... ( .A..L.G.~). ........................................... .... ..c.~./.o. ~//a.(.A.LGA.)~ ....................................... Howchinia trans. Monotaxinoides

Archaediscus @ tenuis Brenckleina

British stages ~ m subzones New Subzonal divisions SERIES

Late Asbian Cf65 I

VISEAN (part.)

Fig. 2. The biostratigraphic ranges of important foraminiferal and algal microfossils in the late Dinantian of parts of southern and western Ireland. The thicker bar denotes particular intervals in which an abundance of a certain microfossil type occurs. L/fac., lithofacies.

LATE DINANTIAN PLATFORM CARBONATES, IRELAND subzone. They first occur some distance above the base, in Cf672 (see below). (iii) New subzonal division Cf6"71 (= earlier part of the late Asbian). Defined by the appearance of bilaminar palaeotextulariids, Ungdarella, Koskinobigenerina and Haplophragmella (rare). The important bilaminar palaeotextulariid taxa appearing near the base of this division are Palaeotextularia ex. gr. longiseptata and Cribrostomum lecomptei. Vissariotaxis compressa disappears at the end of Cf671. Although Ungdarella first appears at the base, it occurs only in relatively low amounts, but by the end of Cf6-~l it is a very important bioclastic constituent in all the platform successions studied (Fig. 2). In the Dromkeen Formation of the Limerick Syncline, Somerville et al. (1992) reported the first appearance of Koskinobigenerina breviseprata, Cribrostomum lecomptei and Palaeotextularia ex. gr. longiseptata at the base of the Cf6-y subzone. In Belgium (Conil et al. 1990) Koskinobigenerina and Cribrostomum lecomptei appear at the base. Cribrospira, a distinctive endothyrid with a cribrate aperture, appears towards the middle of this subzone; this taxon seems to be confined to the late Asbian, from mid-Cf671 to Cf672. It is relatively common in the sections studied. Bibradya also first occurs, with Cribrospira, in the middle part of the subzone, although this taxon is quite rare. Nevillella (Nevillea) sp. first appears in the upper part of this subzone in the areas studied. In the Limerick Syncline (Somerville et al. 1992) Cribrospira sp. first occurs in the 'middle' of the Cf67 subzone, although Jones & Somerville (this volume) do not recognize the widespread utility of this taxon. Somerville et al. (1992) also reported the occurrence of Nevillea dytica from the base of Cf6"7. In Belgium Conil et al. (1990) discovered the first occurrences of Cribrospira panderi and Bibradya grandis in the middle unnamed division of Cf63, subzone, so the possibility exists that an additional subzonal division within Cf67 may be proven with further data collection. Strank (1981, 1983) reported the occurrence of these two genera from the early Asbian of Great Britain. (iv) New subzonal division Cf672 (=later part of the late Asbian). The guides are the first appearance of the foraminifera Howchinia bradyana (rare), Bradyina rotula, Neoarchaedis-

cus (Asperodiscus) stellatus, Neoarchaediscus (Neoarchaediscus) incertus and Saccamminopsis sp., along with the calcareous alga (problematicum) Koninckopora sp. B and Fasciella kizilia (rare).

243

Koninckopora sp. B (Gallagher 1992; Somerville et al. 1992) is a monolaminar calcareous alga, similar to Koninckopora sp. A (Marchant), except that the cells of the thallus are about twice as large (i.e. in sp. A these cells are less than 50 #m in diameter, in sp. B they are over 100 #m). Koninekopora sp. B seems to be confined to this division. The two Neoarehaediscus species and Saccamminopsis first occur in Cf672 in small amounts; these taxa are most common in the Brigantian. Howchinia bradyana and Fasciella kizilia are extremely rare in this division; these components are most abundant in the middle Brigantian/Cf66 successions studied. In the Limerick Syncline (Somerville et al. 1992), all the guide taxa listed above, including the taxon Koninckopora sp. B, first occur in the upper part of the Cf6-'/ subzone. In Belgium (Conil et al. 1990) Howchinia bradyana and Bradyina rotula appear at the base of the Cf67 subzone. (v) Subzone Cf66 (=Brigantian). Although macrofaunal evidence (see above) indicates that the base of LA3 equates with the Asbian/ Brigantian boundary, Brigantian foraminifera diagnostic of the Cf66 subzone, such as Warnantella, Loeblichia and Janichewskina, were not found in beds immediately overlying the base. However, in the absence of the rare macrofaunal indicators, the base of the Brigantian has been located using the following microfossil evidence: (a) the extreme rarity or absence of Koninckopora species (a calcareous alga taxa most common in pre-Brigantian times); (b) the relative increase in the abundance and diversity of the stellate archaediscid faunas; (c) the occurrence of abundant Saccamminopsis, which invariably occur in beds overlying the highest palaeokarst in the late Asbian (this work). L A 4 has a distinctive Brigantian microfossil assemblage that is characterized by Fasciella-coated wackestone intraclasts or bioclasts, plus common

Howchinia bradyana.

Lithostratigraphy Lithofacies association 1 The Holkerian(?) to early Asbian successions constituting LA 1 in the Burren and Cork areas show no cyclicity and exhibit the following macroscopic features and microfacies. Rocks of LA 1 typically consist of a variety of limestone units, which include bedded chert-free grey limestones, dark cherty limestones and mudbank limestones. Two contrasting sections through

244

S. J. GALLAGHER

LA 1 occur in the lower part of the Burren Formation, Co. Clare (C1 and C2, Fig. 1). The Black Head Section (C2) consists predominantly of skeletal peloidal limestones which are rich in the dasycladacean alga Koninckopora. In contrast, the Ballard Bridge Section (C 1) is typified by skeletal limestones that are bryozoan-rich. Both sections (which are 25 km apart) are overlain by the Maumcaha Member and underlain by the Tubber Formation; they are lateral equivalents. Other limestone units that comprise LA 1 include the cherty, thin-bedded and bryozoanrich limestones of the Cecilstown Member, which faunally are the lateral correlatives of the mudbank limestones of the Cracoan 'Reef' in North Co. Cork, and the massive Koninckopora-rich limestones of the Maumcaha Member in the Burren (Fig. 1).

Interpretation. The Koninckopora-rich limestones of LA 1 were deposited in a subtidal, open marine palaeoenvironment, above normal wave-base. In contrast, the bryozoan-rich cherty limestones were probably deposited in an open marine palaeoenvironment below normal wavebase. The lack of cyclicity in LA 1, together with the presence of laterally correlative units of mudbanks (which were the prevalent late Dinantian lithofacies in southern Ireland at this t i m e - see George et al. 1976), deep-subtidal cherty and shallow-subtidal chert-free limestones, suggests that the palaeogeography of southern Ireland during the Holkerian(?) to early Asbian may have been dominated by a roughly southerly-dipping carbonate ramp. Lithofacies association 2 All the late Asbian successions of LA 2 studied have similar lithologies. Essentially, limestones at this stratigraphic level form cyclic units (minor cycles, referred to below as cycles for brevity) of bedded to poorly-bedded steely-blue limestones. In the Burren succession these form characteristic terraces in the topography. The late Asbian LA 2 comprises up to a maximum of 12 cycles in the Burren, although only nine terraces, referred to as T1 to T9, occur; three thin cycles do not form individual terraces. In north Cork there are nine cycles (C1~C9), none of which form terraces. The macroscopic features and microfacies of LA 2 can be summarized as follows: (a)

the late Asbian cycles in the areas examined typically comprise thickly-bedded to unbedded dark limestones;

(b)

the top of individual cycles display palaeokarstic features such as scalloped surfaces and palaeosols; cycles are typically 10-20 m thick, although they may be as thin as 2 m. (c) macrofauna-rich horizons are commonly found at the top (rarely at the base and middle) of each cycle; (d) where bedding thickness can be measured it usually increases upwards within a cycle; (e) the predominant microfacies is a fine grained Kamaenella, foraminifera and calcisphere-rich peloidal packstone to grainstone, although in the lower four cycles in north Cork and T2 of the Burren, packstone/wackestone textures prevail; (f) in the Burren (T4 upward), Kilkenny/ Tipperary and north Cork (C3 upward) successions, Ungdarella and Kamaenellarich packstones to grainstones are common in the uppermost sample and slightly less common in the lowermost sample in each cycle (Figs 3a, b); this microfacies is usually absent or rare in the middle of each cycle; (g) small amounts of the dasycladacean alga Koninckopora occurs near the top and sometimes at the base of cycles, although it is usually rare in the middle (except in T7 of the Burren, Fig. 3b); (h) fenestrate bryozoan hash is rare in all the cycles. It is absent at the top of cycles T1-TS, T7 and T9 of the Burren succession, and is also missing in the upper part of Cycles 1, 3, 5 and 9 in the north Cork Dromdowney Member. In the Callan area it is also found to be rare at the top of cycles. Where fenestrate bryozoan hash occurs, it is most commonly found in the middle of cycles; (i) occasionally, crinoidal packstones to grainstones are found towards the centre of some cycles; crinoids, although present in moderate amounts, are usually rare at the tops of cycles; (j) apart from occasional well-washed grainstones at the tops of cycles, oolitic and oncolitic microfacies are not present in the Burren, and are extremely rare in north Cork. In contrast, cross-laminated, oolitic and peloidal facies, along with oncolitic horizons, can be found at the top of cycles in the Callan area; (k) no microfacies indicative of intertidal/ supratidal deposition were found in the Burren or north Cork areas; however, fenestral micrites were found beneath one palaeokarst in the Callan area.

LATE D I N A N T I A N P L A T F O R M C A R B O N A T E S , I R E L A N D

245

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Fig. 3. Summary lithologs of two Asbian and Brigantian minor cycles from the Buttevant area, north Co. Cork and the Burren area, Co. Clare. (a) Minor cycle 9, Ballyclogh Limestone Formation, Ballyclogh Quarry, north Co. Cork. (b) Minor cycle T7, Burren Formation, northeast Aillwee section, Co. Clare. (c) A section through the Templemary Member of the Liscarrol Limestone Formation, north Co. Cork. (d) A log of part of the Lower Crinoidal Limestones of the Slievenaglasha Formation, Co. Clare.

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S. J. GALLAGHER

Interpretation.

The arrangement of macroscopic features and microfacies suggest that most cycles show a shallowing-upwards trend. The presence of the calcareous algae Koninckopora and Ungdarella at the top and sometimes at the base of cycles and their conspicuous absence toward the middle, where crinoids and fenestrate bryozoans are most common, is significant. It indicates local preservation of shallow subtidal facies at the base of cycles, followed by deeper, quiet subtidal conditions which allowed fenestrate bryozoan thickets to develop. Towards the top of cycles, shallower, in some cases slightly higher energy, subtidal conditions recurred, culminating in emergence for most (if not all) of the cycles. No two cycles sampled were found to be exactly alike. To help demonstrate the variability of the sedimentary cyclicity more clearly, the distribution of macroscopic features and microfacies of two closely sampled cycles in north Cork and the Burren are illustrated (Figs 3a, b). These typical cycles show most of the characteristics mentioned above, and have similarities with those described by Somerville (1977, 1979a, b, c). The dominant bioclastic constituent in the late Asbian limestones is Kamaenella. Thickets and banks of this calcareous green alga were extensive on the southern and western parts of the South Irish Platform (Fig. 1) during much of the late Asbian. In the latest Asbian, Ungdarella (probably a calcareous red alga) seems to have replaced much of the palaeoberesellids. These two bioclastic constituents must have flourished in an environment very similar to the Halimeda banks accumulating off the coast of Florida today (Tucker 1985). Adams et al. (1992), in a review of the significance of the palaeoberesellids in Dinantian sedimentation in the UK, observed that they are abundant in late Dinantian (Asbian) sediments. They concluded that these probable chlorophytes were the most important carbonate-producing organisms in low to moderate energy platform environments in the Asbian of the UK. The rapid accumulation of vast quantities of these two algal types and the volume of sediment created by them was thought by these authors to have been a possible contributory factor in the progradation of Asbian cyclic shallowmarine successions.

Lithofacies association 3 In southern Ireland, the palaeoberesellid and ungdarellid-dominated shallow subtidal facies of

the late Asbian were replaced by laterally extensive, shallow to deeper subtidal crinoid and bryozoan limestones in the Brigantian. In the lower section of all the Brigantian successions studied there are at least five cycles. These are typically 3-10 m in thickness, but some may reach 20 m. Included in this style of cyclicity are the 'crinoid sands' of the Templemary Member of north Cork, the Ballydonnell Member of Kilkenny/Tipperary, and the lower cyclic crinoidal limestones of the Burren (Fig. 1). All these units have the following macroscopic features and microfacies (Figs 3c, d): (a)

(b)

(c)

(d)

(e)

the base of each cycle is marked by a thick bed or an amalgamated thickly-bedded series of coarse-grained crinoidal packstones to grainstones, in which fenestrate bryozoan hash may be abundant, but foraminifera are rare; the uppermost bed of each cycle is a fine- to medium-grained crinoid-poor peloidal limestone. In the Burren these are peloidal grainstones; in north Cork peloidal wackestones to packstones are also present. Moderate amounts of foraminifera and Kamaenella often occur in this part of the minor cycle. Fenestrate bryozoan hash may be absent or extremely rare; macrofaunal bands, in the form of fasciculate rugose coral thickets of Siphonodendron or scattered cerioid corals (Lithostrotion) and brachiopod coquinas, are typically found at or near the tops of cycles; only the uppermost beds of two cycles show pedogenic features (pedotubules); these occur in the north Cork Templemary Member; very rarely, the top of a cycle may be marked by a horizon of stromatolites, as in the Lismalin Member in Callan.

Interpretation. These cycles are predominantly subtidal in nature, like those of the underlying late Asbian succession. Evidence for shaUowingupwards trends in the cycles are pedotubules, stromatolites and the presence of Koninckopora only at their tops. However, shoaling in the majority of Brigantian minor cycles appears to have terminated in a shallow subtidal environment, since the majority of the cycles exhibit no evidence of palaeokarstic or pedogenic phenomena at their tops. Transgressive events established deeper subtidal conditions and prompted flourishing crinoid and bryozoan thicket growth. These crinoid 'gardens' or 'meadows' must have been extensive on the platform during earliest

LATE DINANTIAN PLATFORM CARBONATES, IRELAND

247

Brigantian times. Reworking by currents, storms and bioturbation would have produced laterally continuous sheets of crinoids (Cain 1968; Aigner 1985) in an analogous environment to that described for the Kamaenella-rich facies of the 'terraced' succession. The switch in limestonebuilding allochem types was also accompanied by a change in the nature of the microbiota and sedimentation patterns. Typical cycles from the Templemary (north Cork) and lower crinoidal units (the Burren) have been selected to show the nature of the cyclicity and distribution of the microbiotic constituents (Figs 3c, d).

lower 'crinoid sands' of the basal Brigantian succession of the Burren (LA3, Fig. 1). The major differences are that (a) the base of the cycles in LA 5 are cherty crinoidal limestones, characterized by packstone and rare wackestone textures, whereas in LA3 these are grainstones, and (b) the limestone beds that cap the cycles in LA 5 are similar to those in LA4 (i.e. they are rich in Fasciella and intraclasts).The uppermost metre of LA5 is characterized by a well-washed grainstone bed with micritized bioclasts and pedotubules, which is disconformably overlain by Namurian sandstones and shales.

Lithofacies association 4

Interpretation.

A well developed non-cyclic cherty unit overlies LA3 in all three areas studied; it is very distinctive and has the following characteristics. The rocks are well-bedded, dark grey to black bioturbated packstones to wackestones with chert bands. No cyclicity is apparent. In addition to crinoids and fenestrate bryozoans, Faseiella (a problematical alga) and trepostome bryozoans are the dominant bioclastic constituents. Intraclastic packstones to wackestones are very common, with individual wackestone intraclasts or bioclasts often partially coated (due to partial burial in the substrate) with Fasciella. This interval lacks shallow-water indicators such as micritized grains and dasycladacean alga, and possesses a very distinctive microbiota (discussed above).

Interpretation.

It is suggested that most (if not all) of these limestones were deposited below normal wave-base but above storm wave-base in a subtidal palaeoenvironment. The presence of wackestone intraclasts and bioclasts suggest reworking by a combination of storm events and bioturbation. The deeper subtidal aspect and apparent lack of cyclicity in this succession reflects the onset of transgressive conditions, possibly due to increased subsidence rates, that persisted into the upper half of the Brigantian in the Burren succession.

Lithofacies association 5 and Namurian overstep succession In the Burren succession there is a second cyclic crinoidal limestone unit (LA5) above LA4, which has no correlative in the other areas studied. Superficially LA5 is similar to the

The cyclic crinoidal limestones of LA 5 formed in a deep subtidal palaeoenvironment, where crinoid thickets were continually reworked by storm action below normal wavebase. This is followed by a change to lower energy facies at the top of cycles. The lithological evidence points to continued deepening from the base to the upper part of the Brigantian, underlining the predominantly transgressive nature of the Brigantian successions studied. However, there is evidence for emergence (pedotubules) and shallow subtidal facies in the uppermost bed of LA 5, suggesting that a rapid regressive event and subaerial exposure took place prior to the deposition of deltaic Namurian clastic rocks. In the other areas studied, Namurian sediments overlie the Dinantian successions at an angular unconformity.

Asbian to Brigantian bioclast changes There is a relative shift in abundance of skeletal ailochems from the early Asbian to Brigantian stages in all the successions examined in this work. These changes are summarized in Table 1.

Changes in sedimentation and style of cyclicity across the Asbian]Brigantian boundary The change from a shallow to a deeper subtidal sedimentation pattern across the Asbian/Brigantian boundary as demonstrated by LA 2 to 3 is not unique to the areas in this study, but has similarities to that recorded from other platform areas in the British Isles (Walkden 1987). However, compared with the late Asbian/ Brigantian successions in other UK localities, where average minor-cycle thicknesses increase

248

S. J. GALLAGHER

Table 1. Main contributors to lime sediment in southern Ireland during the early Asbian to late Brigantian

Interval

Limestone-building bioclasts

Earliest Asbian

Koninckopora or fen. bryozoans Koninckopora* and Kamaenella

Early Asbian to l o w e r part of the late Asbian Lower to middle parts of the late Asbian Middle to upper parts of the late Asbian Upper part of the late Asbian Brigantian

Kamaenella Kamaenella* and Ungdarella Ungdarella* and Kamaenella Crinoids and bryozoans

* Signifies the more abundant component where there are two listed.

T e c t o n i c effects

In Ireland, fault activity controlled the development and subsidence of the intraplatform basins such as the Dublin Basin throughout the Vis6an (Pickard et al. 1994). Platform carbonates deposited at the margins of these basins were controlled by faulting, but this fault activity was of relatively local importance; elsewhere no major faults active in the late Asbian to Brigantian of Ireland have been reported (but see Strogen et al. this volume ). It is therefore unlikely that these faults controlled episodic subsidence and uplift sufficient to generate extremely similar minor-cycle styles and facies across a wide platform (at least 100km in width). However, it is likely that rates of subsidence due to 'sag' influenced overall platform development.

E u s t a t i c effects

above the Asbian/Brigantian boundary (Walkden 1987), the successions studied in this work (especially well-exposed successions in the Burren and North Co. Cork) show the opposite pattern. The apparent deepening above the Asbian to Brigantian boundary has been attributed to faster subsidence rates in the Brigantian, or to an increase in the amplitude of the cyclic oscillations (Walkden 1987). If the latter is true, then the reduction in the thicknesses of the minor cycles observed across this boundary in southern Ireland might be related to low sedimentation rates of the crinoidal limestones here relative to subsidence rates. Hence, the build-up of crinoidal material at the base of the Brigantian minor cycles may not have been sufficiently rapid to create subaerial emergence at the tops of minor cycles. It also may explain why fewer minor cycles are recognized in the Brigantian of southern and western Ireland compared with the underlying late Asbian successions.

Causes of late Dinantian cyclicity The origin of the minor cyclicity in late Dinantian platform carbonate successions has been the subject of much debate in the last two decades. Discussions are provided in Bott & Johnson (1967), Ramsbottom (1973, 1979), Somerville (1979a, b,e), Bridges (1982), Tucker (1985), Walkden (1987), Leeder (1988) and Horbury (1989). The two factors below are considered to be important in controlling cyclicity of platform carbonates in the areas studied.

The transgressions and regressions that produced the cyclical stratigraphy in the three areas may have been eustatically controlled. The style of sedimentation and number of minor cycles observed between the areas studied is very similar, and all the late Asbian and earlier Brigantian members that have been described can be correlated across southern and western Ireland. This suggests that whatever process operated to produce the inherent cyclicity, it must have been operating across this platform. Hence transgressions and regressions may well have been almost instantaneous platform-wide events. A eustatic origin is suggested to account for this pattern. The number of cycles generated in the late Asbian and Brigantian Irish successions is of the same order as those reported by Somerville (1979b) and Walkden (1987) in the rest of the British Isles. This is further evidence to support possible eustatically-controlled synchronous cyclic events on a platform-wide scale. The duration of these events was calculated by Leeder (1988) to be in the order of 190000 to 240000 years (if the Asbian to Brigantian stages lasted 9 Ma), or 319 000 to 405 000 years, if the duration of these stages was 15Ma, as suggested by Walkden (1987). Harland et al. (1990) suggested that the Asbian to Brigantian stages lasted only 6.6Ma; the cycles would then be of the order of 140 000 to 176 000 years. Thus, whatever values are used to calculate cycle durations, they are within the range of values for Milankovitch climatic cycles (Schwarzacher 1993). Sea-level changes of this (fourth) order are probably

LATE DINANTIAN PLATFORM CARBONATES, IRELAND caused by glacio-eustatic effects (Donovan & Jones 1979), and it is significant that the first Dinantian records of Gondwanan glaciation (Crowell & Frakes 1970; Crowell 1978) were reported from rocks of latest Vis6an age in the southern hemisphere, just at the time when cyclicity is best developed in the British Isles.

Summary and conclusions Two new biostratigraphic divisions are recognized in the late Asbian/Cf6~ subzone, denoted Cf671 and Cf672. The definition of Cf671 is based on the first occurrence of the bilaminar palaeotextulariid Cribrostomum lecomptei and the base of the Cf672 division by the first occurrence of Howchinia bradyana, Bradyina

rotula, Neoarchaediscus stellatus, Saccamminopsis and Neoarchaediscus incertus. A review of the criteria defining the base of the overlying Brigantian/Cf66 subzone suggests that although the base of this stage can be located on macrofaunal grounds, diagnostic Brigantian foram taxa are absent from beds immediately above the Asbian/Brigantian boundary at any locality. In the absence of diagnostic microfossil components, the base of the Brigantian/Cf66 subzone was located based on an increase in stellate archaediscids, the relative rarity of Koninckopora, on lithostratigraphic criteria and the presence of Saccamminopsis-rich horizons. The late Dinantian succession in all three areas can be subdivided into distinctive lithofacies units, denoted LA 1 to 5, which can be correlated litho- and biostratigraphically. In the Burren succession LA 1 consists of early Asbian shallow to deeper subtidal carbonates, whereas in North Co. Cork it is characterized by early Asbian 'reefal' and deep subtidal limestones. It is suggested that the palaeogeography of southern Ireland during the early Asbian may have been dominated by a southerly-dipping carbonate ramp. The deposition of late Asbian LA 2 marked the onset of relatively stable platform conditions in all areas. This succession is cyclic, and characterized by at least nine minor cycles of thickly bedded, shallow subtidal, palaeoberesellid-rich limestones. Each shallowing-upwards minor cycle is capped by a palaeokarst surface, on which evidence of pedogenesis can be found, and a clay. Brigantian LA 3 is also characterized by a series of thin minor cycles, each composed of basal amalgamated subtidal crinoidal limestones capped by a shallower subtidal peloidal

249

crinoid-poor limestone. Brigantian LA 4 is noncyclic and comprises thinly-bedded, cherty deepsubtidal limestones with a distinctive intraclastic wackestone microfacies. A second subtidal cyclic crinoidal unit (LA 5) overlies LA 4 in the Burren area, and Namurian siliciclastic rocks unconformably overlie the Dinantian succession in each area. A distinctive change in rock-forming bioclast types and calcareous algae is observed from the early Asbian to Brigantian. In the earliest Asbian rocks the dominant bioclasts are Koninckopora or fenestrate bryozoans. Palaeoberesellids become dominant from the upper part of the early Asbian to the late Asbian. In the latest Asbian Ungdarella replaces much of the palaeoberesellids. The onset of Brigantian sedimentation is marked by a switch to crinoid and bryozoan-rich facies with few palaeoberesellids and ungdarellids. The palaeoenvironment in which the late Asbian carbonates were deposited consisted of laterally continuous thickets of calcareous alga very similar to the Halimeda banks accumulating off the coast of Florida today. Similar crinoid 'gardens' or 'meadows' characterized Brigantian sedimentation. A major change in sedimentation cycle style is seen across the Asbian/Brigantian boundary in all areas. Cycles in the late Asbian are capped by palaeokarsts and have a thickness of 10-20m, averaging 12m. In contrast, minor cycles in the Brigantian rarely terminated in emergence and attain thicknesses of 5-10m, averaging only 6m. This pattern may be accounted for by the low accumulation rates of the crinoidal limestones in the Brigantian compared with to the high accumulation rates of the algal limestones in the Asbian. In general, Brigantian sedimentation was deeper subtidal in nature compared with the underlying Asbian. The change in sedimentation pattern is not unique; it is reported in the UK by Walkden (1987), where the pattern is attributed to faster subsidence rates or an increase in the amplitude of cyclic oscillations in the Brigantian compared with the Asbian. The research for this work was partly funded by an Eolas Basic Research Grant in Science. The author wishes to thank I. D. Somervilleof University College, Dublin, for his ongoing advice and for reviewing an earlier draft of the manuscript. In addition the author would like to thank G. Jones and C. MacDermot for allowing access to unpublished data on the Callan and Burren areas. The text was greatly improved by the detailed comments of A. E. Adams and J. Davies who reviewed the manuscript.

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phy of the Dinantian Limestones in the Llangollen Area and East of the Clwydian Range, North Wales.

of Holkerian, Asbian and Brigantian Stages of the British Lower Carboniferous. PhD Thesis, Uni-

PhD Thesis, Queen's University of Belfast. - - 1 9 7 9 a . A sedimentary cyclicity in early Asbian (lower D1) limestones in the Llangollen district of North Wales. Proceedings of the Yorkshire Geological Society, 42, 397-404. --1979b. Minor sedimentary cyclicity in late Asbian (upper D1) limestones in the Llangollen district of North Wales. Proceedings of the Yorkshire Geological Society, 42, 317-341. - - 1 9 7 9 c . A cyclicity in the early Brigantian (D2) limestones east of the Clwydian Range, North Wales and its use in correlation. GeologicalJournal, 14, 69-86. & STRANK, m. R. E. 1984. The recognition of the Asbian/Brigantian boundary fauna and marker horizons in the Dinantian of North Wales. Geological Journal, 19, 227-237. & STROGEN, P. 1992. Ramp sedimentation in the Dinantian limestones of the Shannon Trough, Co. Limerick. Sedimentary Geology, 79, 59-75. , & JONES, G. LL. 1992. The biostratigraphy and sedimentology of Dinantian limestones in the Limerick Syncline, Ireland. Geological Journal, 27, 201-220.

versity of Manchester. 1983. New stratigraphically significant foraminifera from the Dinantian of Great Britain. Palaeontology, 26, 435-442. STROGEN, P., JONES, G. EL. & SOMERVILLE, I. D. 1990. Stratigraphy and sedimentology of Lower Carboniferous (Dinantian) boreholes from west Co. Meath, Ireland. Geological Journal, 25, 103-137. --, P., SOMERVILLE, I. D., PICKARD, N.A.H., JONES, C. LL. & FLEMING, M. 1996. Controls on ramp, platform and basinal sedimentation in the Dinantian or the Dublin Basin and Shannon Trough, Ireland. This volume. TUCKER, M. E. 1985. Shallow-marine carbonate facies and facies models. In: BRENCHLEY, P. J. & WILLIAMS, B. P. J. (eds) Sedimentology." Recent Developments and Applied Aspects. Geological Society, London, Special Publication, 18, 147-169. WALKDEN, G. M. 1987. Sedimentary and diagenetic styles in late Dinantian carbonates of Britain. In: MILLER, J., ADAMS, A. E. & WRIGHT, V. P. (eds) European Dinantian Environments. Geological Journal Special Issue 12, 131-156

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Initiation, growth and decline of a tectonically controlled Asbian carbonate ramp: Cuilcagh Mountain area, NW Ireland JOHN

G. K E L L Y

Department of Geology, University College, BelfieId, Dublin 4, Ireland Present address: CSA Ltd, Parkview House, Beech Hill, Clonskeagh, Dublin 4, Ireland

Abstract: The Lough MacNean Valley and the surrounding Cuilcagh and Belmore Mountains expose carbonates and subordinate shales of late Arundian to Asbian age. The Benbulben Shale, of late Arundian to Asbian age, is composed of fossiliferous shales with limestone bands. This is succeeded by the Glencar Limestone, composed predominantly of argillaceous micritic wackestones and packstones, and fossiliferous shales. This is overlain by the Asbian Dartry Limestone, which exhibits extreme lateral and vertical facies variations; it is quite different on east, north and west Cuilcagh, which are described in turn. Lateral variations in facies and thickness in the Dartry Limestone are attributed to carbonate ramp development on a differentially subsiding block, initiated by sinistral strikeslip movement on the Castle Archdale-Belhavel and Curlew~logher Valley fault systems. The sediments exposed on east Cuilcagh represent the proximal portion of the ramp, whilst the Knockmore and Cloghan Hill Members represent the mid and distal ramp. The chertyfacies limestones mark widespread open shelf conditions at the end of subsidence. Inversion is marked by the shallow-water Cloghany Limestone, the Carn Limestone and the topDartry disconformity prior to deposition of the Meenymore Formation. The Cuilcagh-Lough MacNean area forms part of the Lough Allen Basin (Fig. 1). It was originally covered as part of Griffith's mapping project of the 1830s and remapped by Wilkinson and Cruise of the Geological Survey of Ireland (GSI) in the late 19th century. Sheridan (1972) discussed the area using the original GSI stratigraphy. Brunton & Mason (1979) reclassified the Dinantian succession in the West Fermanagh Highlands using Oswald's (1954) nomenclature up to and including the Dartry Limestone, and Brandon's (1972) nomenclature for the succession above. This lithostratigraphy has been generally adopted by subsequent workers. The Cuilcagh area was first remapped on a systematic basis by the Geological Survey of Northern Ireland (GSNI) in 1979-83 (GSNI 1:50000 sheet 44, 56 & 43, Derrygonnelly and Marble Arch, 1991) who established five members within the Dartry Limestone. Kelly (1989) subdivided the Dartry Limestone into four facies assemblages (A-D), attributed to proximal-, middle-, distal- and post-ramp positions. This paper describes the facies and thickness variations seen across the Cuilcagh-MacNean area and attributes these to carbonate ramp development on a differentially subsiding fault block. The GSNI stratigraphy is presented in Fig. 2, and a correlation between this stratigraphy and the facies assemblages of Kelly (1989) is presented in Table 1. A correlation along the section line A - A of Fig. 1 is presented in Fig. 3.

Structure The Cuilcagh Mountain area lies between the Castle Archdale-Belhavel Fault Zone and the Clogher Valley-Curlew Mountain Fault Zone (Fig. 1). Tectonic activity of Asbian age is well documented from the Northwest Basin of Ireland (Kelly 1989; Mitchell & Owens 1990; Mitchell 1992) and has had a considerable controlling effect on sedimentation within this area. Structural analysis indicates that N-S directed compression in Asbian times induced dextral strike-slip movement on these fault zones, leading to the development of transtensional areas in the Lough Allen-Lough MacNean and Slieve Rushen areas. This transtensional movement led, as indicated by the evidence presented below, to the development of differentially subsiding fault blocks, with greater subsidence in the Cuilcagh-MacNean area to the northwest.

Stratigraphy and biostratigraphy A generalized stratigraphy for the Cuilcagh area is shown in Fig. 2. The biostratigraphy of the rocks below the Glenade sandstone are described below. George et al. (1976) erected six stages for the Dinantian rocks in Britain and Ireland, and Riley (1993) reviewed all published biostratigraphical data for the Dinantian and established

From STROGEN,P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 253-262.

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faunal and floral ranges for significant taxa. Correlation of the lithostratigraphic divisions of Legg et al. (1995) with the stages of George et al. (1976) is presented in Fig. 2.

Benbulben Shale Formation The base of the Benbulben Shale Formation has yielded a microfauna of Arundian age (Legg et al. 1995). At the top of the formation, small, highly fossiliferous outcrops of the Benbulben Shale in the Cladagh River (H127357) have yielded a diverse coral fauna which includes Clisiophyllum

delicatum nanum, Koninckophyllum vaughani, Siphonodendron martini, S. scaleberense, S. sociale and Siphonophyllia benburbensis. A foraminiferal fauna including Archcediscus krestovnikovi, A. reditus and A. stilus at both angulatus and concavus stages has also been obtained from the limestone bands. The presence of arch~ediscids at both concavus and angulatus stages in the Cladagh River outcrops indicates that the Benbulben Shale at this locality is of late Holkerian or early Asbian age, as this is the interval in which these two arch~ediscid stages overlap (Riley 1993). The presence of Koninckophyllum vaughani and

ASBIAN RAMP, IRELAND

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Siphonophyllia benburbensis, however, indicate a definite early Asbian age as they are first recorded in the Asbian (Mitchell 1989). The Holkerian must therefore be restricted to the middle of the Benbulben Shale Formation and does not include the upper Bundoran Shale and Mullaghmore (Derrygonnelly) Sandstone as proposed by George et al. (1976) and Brunton & Mason (1979).

for micropalaeontological study proved to be barren. The top of the underlying Benbulben Shale, however, has been shown above to be lowermost Asbian in age, and definite Asbian faunas have been collected from the overlying Dartry Limestone. Dibunophyllum sp. and goniatites of the Beyrichoceras micronotum group have been recorded in the type locality (Oswald 1954), indicating an Asbian age.

Glencar Limestone Formation

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No macrofauna was collected from the Glencar Limestone in the study area. Samples collected

Considerable numbers of corals were collected from localities in the Cuilcagh region. Thin

256

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KELLY

Table 1. Comparisonof the facies assemblages of Kelly

benburbensis. Foraminifera recovered include Archcediscus karreri, A. krestovnikovi, A. varsanofievae and A. stilus at the angulatus stage, Gigasbia gigas and Vissariotaxis sp.

(1989) with the lithostratigraphy of the Geological Survey of Northern Ireland (1991) Kelly (1989)

GSNI (1991)

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Carn Limestone Member

Facies C Facies B Facies A

Knockmore Limestone Member. The Knockmore Limestone Member produced no useful taxa.

Cloghany Limestone Member Carrickmacsparrow LimestoneMember Dartry Limestone Formation Cloghan Hill Limestone Member Knockmore Limestone Member Dartry Limestone Formation

Cloghan Hill Limestone Member. At Cloghan Hill (H089360) this member yielded a large number of specimens including Axophyllum sp., Clisiophyllum keyserlingi, Dibunophyllum sp., Koninckophyllum vaughani, Pseudozaphrentoides juddi, Siphonodendron intermedium, S. martini, and S. pauciradiale.

sections also yielded significant numbers of foraminifera from near the top of the formation. Various significant taxa described below were identified in the various members of the Dartry Limestone.

Cloghany Limestone Member. The Cloghany Limestone Member contains archaediscids at angulatus stage and Howchinia bradyana.

Dartry Limestone cherty facies.

Pseudozaphrentoides juddi, Siphonodendron pauciradiale, Koninckophyllum vaughani, Siphonophyllia benburbensis and Clisiophyllum keyserlingi) first appear in the Asbian (Mitchell

A considerable number of the corals identified in the Dartry Limestone (Dibunophyllum sp.,

Cherty limestones exposed at Legacurragh Gap (H158308) include Haplolasma densum, Siphonodendron

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ASBIAN RAMP, IRELAND Arch~ediscids at the angulatus stage indicate an Asbian age for these rocks (Riley 1993, Jones & Somerville, this volume) and the presence of Howchinia bradyana in the Cloghany Limestone indicates a late Asbian age for this member (Riley 1993; Jones & Somerville this volume), suggesting that the boundary between the early and late Asbian lies within, or at the base of, this stratigraphic horizon.

Meenymore Formation The Meenymore Formation lies above the Dartry Limestone of Asbian age. There is, as a result of the restricted environments indicated by the sediments of this formation, a paucity of coral/ brachiopod data with which to date this unit. No previous micropalaeontological collecting has been reported. However, Brandon & Hodson (1984) reported an Asbian (B2) goniatite fauna from this formation. Foraminifera recovered from the formation in the Polliniska stream section (H 164290) include arch~ediscids with well-developed angulatus sutures (Archaediscus karreri, A. krestovnikovi, A. reditus, A. stilus and A. of the group

varsanofievae/mellitus), Endostaffella fucoides, Eostaffella parastruvei, Gigasbia gigas, Howchinia bradyana, Koskinobigenerina sp., Nodasperodiscus sp., Nodosarcheediscus sp., Mediocris mediocris, Valvulinella sp., and Vissariotaxis sp. The presence of arch~ediscids with well developed angulatus sutures, Nodasperodiscus sp., Nodosarchcediscus sp., Howchinia bradyana and Koskinobigenerina sp. indicate an upper Asbian age for this section.

Sedimentology The sedimentology is described in turn from the three principal areas of outcrop in east, north and west Cuilcagh, and finally from two deep boreholes in the area.

East Cuilcagh The base of the Dartry Limestone Formation is not exposed on east Cuilcagh, and total thickness is estima~;ed at 120 m. In a number of quarries and outcrops in Gortalughany Townland (H 170304) over 30m of shallow-water carbonates are exposed. These consist of tabular-bedded, black, fine-grained organic-rich carbonate mudstones with a restricted fauna mainly comprised of calcispheres, ostracodes and thin-shelled brachiopods. The tops and bottoms of these beds have a

257

laminated appearance, probably due to an increasing argillaceous content, although algal laminations may be present. The carbonate mudstones are interbedded with thin shales, and thin black cherts are developed along bedding planes. Interbedded with these dark well-bedded limestones are a number (at least six) of coarse, pale crinoidal grainstone beds up to 2.6m thick, which are commonly dolomitized. Of particular note within these grainstone beds is the presence of intraclastic material derived from the black mudstones, of glauconite, cross-bedding, and beds with irregular tops and vertical fissures. The fissures are filled with bioclastic sediment different from the grainstones and the overlying carbonate mudstone. This sequence of black carbonate mudstones, cherts, thin shales and grainstones was termed 'Facies A' by Kelly (1989) and is informally referred to here as the Gortalughany Beds. These may be, in part, laterally equivalent to a 4 m thick grainstone, the Carrickmacsparrow Limestone Member, identified by Legg et al. (1995) on Belmore Mountain some l0 km to the north. The Gortalughany Beds are interpreted as a shallow-water package of restricted, hypersaline lagoonal limestones and coarse carbonate sandbar deposits. The infilled fissures are interpreted as evidence for emersion. This sequence is believed to have been deposited on the proximal portion of a carbonate ramp. The Gortalughany beds are overlain by approximately 50m of carbonate mudstones, packstones and cherts, similar to the Dartry Limestone described from Co. Sligo by Oswald (1954). Bedding is less regular, becoming wavy in appearance, and cherts tend to form more irregular nodules and bands in comparison with the well-defined bedding and tabular cherts of the Gortalughany beds. The top of the cherty facies of the Dartry Limestone on east Cuilcagh includes a distinctive horizon containing abundant solitary and colonial rugose corals near the top (Kelly 1989; Legg et al. 1995). The coral horizon is particularly well exposed at Legacurragh Gap, a large dry glacial spillway. Cherts are abundant and silicification of fossils is very common. Features interpreted as due to synsedimentary slumping have been identified at Legacurragh (Legg et al. 1995). These coralline limestones are succeeded by the Cloghany Limestone Member, which is up to 1.3m thick (Geological Survey of Northern Ireland 1991). This member contains a distinctive chert-free compact limestone composed of

258

J.G.

medium-grained packstones and grainstones. These contain an abundant and diverse fauna, predominantly of crinoid ossicles and fenestellid, trepostome and encrusting bryozoans, but also including foraminifera, echinoid fragments, brachiopod shells and spines, sponge spicules, ostracods, trilobites, algosponges, calcareous algae and calcispheres. Many bioclasts are micritized or coated in micrite, much of the matrix in the packstones is micritic and glauconite is common. The micritization, micrite coatings and presence ofglauconite indicate shallow-water deposition for this member. The Cloghany Limestone Member here is overlain by unfossiliferous to poorlyfossiliferous shales and micritic bioclastic wackestones and packstones, similar to the Carn Limestone of Brunton & Mason (1979) and formally referred to as the Carn Limestone Member. The thickness of the Carn Limestone on east Cuilcagh is variable; this is interpreted as the result of subsequent emersion and erosion prior to deposition of the Meenymore Formation (Legg et al. 1995).

North Cuilcagh

In northern Cuilcagh the Glencar Limestone Formation is overlain by the Knockmore Limestone Member of the Dartry Limestone Formation, which is composed almost entirely of carbonate mudbanks and minor associated argillaceous bands. The contact may be erosional or may display lateral facies variations between Glencar Limestone facies and the basal part of the Knockmore Limestone Member. Erosional contacts are common; lateral interfingering of Glencar Limestone facies and the Knockmore Limestone is well exposed at Hanging Rock (Legg et al. 1995) and in Cascades Rising Cave (H 12283498). The carbonate mudbanks of the Knockmore Limestone are developed in two distinct forms, tabular sheets and knoll-shaped banks which 'stack up' vertically and laterally. Large knollshaped bank accumulations form a number of the distinctive limestone hills seen in the Marlbank area. Argillaceous crinoidal limestones, which represent a flank facies to the mudbanks, are common in the Knockmore Limestone in the Tullybrack uplands but are not present on Cuilcagh, possibly due to erosion. The mudbank limestones are mainly formed of polymicrites in which an extensive Stromatactis cavity network has developed. These

KELLY Stromatactis cavities are infilled with micrite muds, multiple cryptofibrous spars and blocky spars. Extensive portions of the mudbanks are composed solely of intact fenestellid bryozoans encrusted with polyphase cryptofibrous calcite (CFC) and small residual cavities infilled with polymicrites and/or blocky spars. This texture is similar to the 'veines bleues' facies described from Waulsortian mudbanks in Belgium. Bioclastic material generally forms only a small proportion of the total sediment within the banks (5-15%) and is mainly restricted to intact fenestellid bryozoans, crinoids, brachiopods, sponge spicules and gastropods. In a small number of areas bioclast-rich pockets occur, bioclastic material comprising up to 60% of the total volume. Here the fauna is composed predominantly of brachiopods and crinoid debris, although gastropods, trilobites, ostracodes, echinoid fragments, sponge spicules and rugose corals including Siphonodendron sp. and zaphrentids have been recovered. Brachiopods are generally articulated and the corals are in life position. Component analysis of the carbonate mudbanks of the Knockmore Limestone Member and comparison with the Waulsortian depth zonation proposed by Lees & Miller (1985) suggests a Phase 'C' type assemblage, indicating deposition below the photic zone and storm wave-base. The banks contain few shale horizons and are poorly bedded or massive, the bank limestones being composed of up to 99% CaCO3. Preferential dolomitization of the micrite sediment is common adjacent to faults in the area, and total dolomitization also occurs locally. The Knockmore Limestone Member is overlain to the south by cherty packstones of the Dartry Limestone, and the top of the Dartry Limestone Formation, the Cloghany Limestone Member, has been identified in the Owenbrean River by Legg et al. (1995). Further west, at Tullyard (H209332) on the Fermanagh/Cavan border, some 30 m of cherty packstones containing a rugose coral band 3 m thick directly overlie the Knockmore Limestone. This is dramatically exposed in Polltullyard (H209332) a pothole over 45 m deep developed at the contact between the cherty packstones and the carbonate mudbanks. The coral horizon exposed here is at a similar stratigraphic level to that seen in east Cuilcagh. Extensive coral horizons dominated by Siphonodendron sp. extend over several kilometres, and are present in Dartry Limestone-equivalents in the Carrick-on-Shannon Syncline (Caldwell 1959) and the Bricklieve Mountains (Caldwell &

ASBIAN RAMP, IRELAND Charlesworth 1961; Dixon 1972). The total thickness of the Dartry Limestone in north Cuilcagh is approximately 240m, of which 210 m is composed of carbonate mudbank limestones, although in outcrop the Knockmore Limestone thickens from east to west.

West Cuilcagh In west Cuilcagh, approximately 220m of the Knockmore Limestone Member are overlain by a sequence of argillaceous bioclastic limestones, micritic boundstones and isolated carbonate mudbanks, collectively termed Facies 'C' by Kelly (1989) and formally named the Cloghan Hill Limestone Member. At Pollnagossan (H064353) and Barran Townland (H0435), in Co. Cavan, the Cloghan Hill Member is overlain by cherty packstones, although the distinctive coral horizon has not been observed in this area. The Cloghan Hill Limestone Member consists of a laterally and vertically variable package of limestones. The basal beds, directly overlying the Knockmore Limestone Member, comprise thinly-bedded, argillaceous crinoidal packstones which weather easily. The removal of this easilyeroded lithology has exposed the top surface of the more resistant Knockmore Limestone Member over much of west Cuilcagh. The rocks of the Cloghan Hill Limestone ]Vlember above this basal horizon are composed of two main lithotypes. The most striking rocks are carbonate mudbanks, which form isolated hills resting on the fiat-lying Knockmore Limestone Member beneath. The banks are predominantly interbedded with argillaceous bioclastic limestones, in which solitary and colonial rugose corals and small brachiopods are abundant. The corals are commonly crushed and not in life position. The colonial corals are mostly Siphonodendron sp. and are rarely found intact. The Cloghan Hill Limestone Member mudbanks differ from the Knockmore Limestone Member banks in that they form isolated 'patch banks' as opposed to the laterally and vertically extensive intergrown banks of the Knockmore Limestone Member. Within these patch banks, three separate facies can be identified: a massive, unbedded core facies; a well- to poorly-bedded flank facies; and a bedded cover facies. The basal mound flank sediments are wellbedded bioclastic packstones, distal from the mound core. From the flank towards the bank core, these sediments progressively change in character. Individual beds thicken dramatically

259

as bedding becomes indistinct and finally bedding disappears into the massive bank core facies. Micrite content increases towards the bank core, initially as distinct micrite nodules which coalesce to form micrite-dominant beds, and the bioclastic proportion decreases, becoming dominated by crinoids and fenestellid bryozoans. Stromatactis cavities are developed close to the core. Similar flank-to-core transitions have been described from Waulsortian banks (Miller 1986) and Arundian mudbanks (Kelly & Somerville 1992). The bank core is a wackestone consisting of bioclasts set in a matrix of several generations of micrite and pelleted micrite. The pelleted form is dominant over homogeneous micrites in much of the bank. Within the bank, the fauna is relatively abundant, frequently comprising up to, or even over, 40% of the total sediment (excluding calcite cements). Although dominated by crinoid ossicles and fenestellids (both intact and fragmented), fauna is diverse and includes trepostome bryozoans, encrusting bryozoans (such as Fistulipora incrustans), sponge spicules, echinoid spines, brachiopods, trilobites, gastropods, simple and plurilocular foraminifera, algospongia (Aoujgaliia sp. and Stacheoides sp.), heterocorals, calcispheres and rugose corals. The rugose corals are invariably intact and composed only of solitary forms, such as Axophyllum sp., Zaphrentids and Clisiophyllum keyserlingi. The corals are commonly bored, the boring being restricted to areas which do not have a coat of CFC, indicating that the presence of an early cement has prevented boring. Crinoid ossicles are rarely bored, but are frequently micritized. Within the banks stromatactis-type cavities are common. The roofs of these cavities are commonly controlled by bryozoan fronds and brachiopod shells, but may be irregular in form. Geopetal micrite mud-fills are common, the residual cavities being filled with CFCs and blocky spars. Shelter cav.ities formed beneath bioclasts and filled with blocky spar are also encountered, but these can be distinguished from true stromatactis cavities because they lack geopetal muds and CFCs. The bank top sediments are composed of micritic packstones with crinoid ossicles, fragmented brachiopods, bryozoans and rugose corals, more abundant than in the bank sediments, set in a fine carbonate mud. The sediment contains intraclasts, believed to be derived from the bank, and may appear rubbly in places due to a combination of stylolitization and the presence of a high proportion of intraclasts.

260

J.G.

The extent of fragmentation and the inclusion of a large proportion of intraclastic material suggests that the sediments have been affected by storm activity, bank development having been terminated by upward growth of the bank to storm wave-base. Comparison of the Cloghan Hill bank components with the depth zonation proposed for the Waulsortian by Lees & Miller (1985) shows a Phase 'D' type assemblage, indicating deposition within the photic zone. Boreholes

Two deep boreholes have been put down in the area, MacNean No. 1 and Big Dog No. 1 boreholes, and a third, Owengarr No. 1, to the SW of the Clogher Valley Fault (Figs 1, 3). MacNean No. 1 borehole was drilled by Marathon in the early 1960s and collared to the SE of the Castle Archdale-Belhavel Fault Zone (Fig. 1). Borehole logs (Fig. 3) show the Dartry Limestone to be approximately 320m thick and it has been interpreted (Sheridan 1972) as being composed entirely of rocks of carbonate mudbank facies. Reexamination of surface outcrop in the area immediately south of the MacNean drillhole site, however, indicates the presence of at least 60m of mudbanks and argillaceous limestones of the Cloghan Hill Member, overlain by cherty packstones of the Dartry Limestone. The underlying Glencar Limestone Formation in the MacNean Borehole appears to be 45m thick. Big Dog No. 1 borehole was also drilled by Marathon in the early 1960s and collared to the NW of the Castle Archdale-Belhavel Fault Zone (Fig. 1). Borehole logs (Fig. 3) indicate that the Dartry Limestone is entirely composed of cherty packstones totalling 140m in thickness (Sheridan 1972). The Glencar Limestone Formation appears to be 30 m thick. Owengarr No. 1 borehole lies SW of the Clogher Valley Fault (Fig. 1) and penetrated almost 700 m of Dartry Limestone Formation rocks.

Comparison between the Asbian carbonate mudbanks, Waulsortian and Arundian carbonate mudbanks The carbonate mudbanks of the Knockmore and Cloghan Hill Limestone members display a range of features which are similar to many

KELLY features present in Waulsortian mudbanks (carbonate mudbanks of Courceyan-Chadian age according to Lees 1988) and mudbanks of Arundian age in NW Ireland (Kelly & Somerville 1992). Gutteridge (1995) describes mudbanks of Brigantian age from Derbyshire which also have similarities with Asbian mudbanks from NW Ireland. The main features of importance are: (a)

both Asbian and Waulsortian mudbanks are primarily composed of micrite muds produced in situ; (b) all the mudbanks referred to above lack framework-building organisms; (c) the Asbian mudbanks display both tabular and mound form and may coalesce to form thick, laterally extensive complexes, similar to forms described from the Waulsortian by Lees (1964); (d) the Asbian mudbanks have distinct core, flank and bank top facies, similar to those described from Waulsortian, and from Arundian and Brigantian mudbanks; (e) Stromatactis cavity networks, infilled by internal geopetal micrites, single or polyphase CFC and coarse blocky spars are common in Waulsortian, Arundian, Asbian and Brigantian mudbanks; (f) polyphase CFC coatings on fenestellid bryozoan sheets (veines bleues) are common in the Knockmore Limestone and in Waulsortian mudbanks; (g) sedimentological variations from mudbank flank to core in the Cloghan Hill Member are similar to those described by Miller (1986) from Waulsortian mudbanks in the Craven Basin; (h) macrofauna is sparse in both Asbian and Waulsortian mudbanks, although faunal diversity is highly variable, but can be correlated with a crude depth zonation, diversity increasing with shallower depositional depth. In summary, therefore, it appears that there are a number of features common to Dinantian carbonate mudbanks of Courceyan-Chadian (Waulsortian), Arundian, Asbian and Brigantian age. These similarities suggest a common origin for the carbonate mud sediment comprising these deposits and related depositional mechanisms within them. Variations between Dinantian carbonate mudbank deposits in different areas may reflect differences in depositional environment (such as tectonic regime or water depth) rather than fundamental differences in sediment origin.

ASBIAN RAMP, IRELAND Conclusions

Activation of preexisting structures of Caledonide trend in the Asbian caused dextral strikeslip motion, leading to the development of a transtensional structural regime in the Lough Allen Basin. Differential subsidence of a major fault block was initiated, possibly during deposition of the Glencar Limestone Formation, and led to the development of a carbonate ramp during Dartry Limestone deposition. Shallow water, proximal ramp facies, including restricted lagoonal limestones and carbonate sand-bars, are exposed in east Cuilcagh. Midramp facies are represented by the deeper water carbonate mudbanks of the Knockmore Limestone Member, and the distal ramp is represented by the Knockmore Limestone Member and the argillaceous limestones and patch carbonate mudbanks of the Cloghan Hill Limestone Member. Cessation of differential subsidence led to widespread deposition of the cherty, open-shelf facies and the distinctive coral horizon in the Dartry Limestone. The shallow-water Cloghany Limestone Member, the Carn Limestone Member and the end-Dartry disconformity indicate a period of shallowing and subsequent emersion, possibly due to inversion. Coral and foraminiferal faunas recovered from the sequence indicate that ramp development began in the early Asbian and ended before the deposition of the late Asbian Cloghany Limestone Member. I would like to thank G. Jones, I. Somerville, I. Mitchell and G. Millar for discussion and information concerning the geology of the Cuilcagh area. I would also like to thank the referees, who made many valuable comments which have significantly improved the initial manuscript. Research was partly funded by the Department of Education (NI). References

BRANDON, A. 1972. The Upper Vis6an and Namurian Shales of the Doagh outlier County Fermanagh, Northern Ireland. Irish Naturalist's Journal, 17, 159-170. & HODSON, F. 1984. The stratigraphy and palaeontology of the late Vis+an and early Namurian Rocks of Northeast Connaught. Geological Survey of Ireland Special Paper, 6, 1-54.

BRUNTON, C. H. C. & MASON, T. R. 1979. Palaeoenvironments and correlations of the Carboniferous rocks in west Fermanagh, Ireland. Bulletin of the British Museum of Natural History, 32, 91-108.

261

CALDWELL, W. G. E. 1959. The Lower Carboniferous rocks of the Carrick Syncline. Quarterly Journal of the Geological Society, London, 115, 163-187. CHARLESWORTH, H. A. K. 1961. Vis6an coral reefs in the Bricklieve Mountains of Ireland. Proceedings of the Geologists' Association, 73, 359-382. DIXON, O. A. 1972. Lower Carboniferous rocks between the Ox and Curlew Mountains, Northwest Ireland. Journal of the Geological Society, London, 128, 71-101. GEOLOGICALSURVEYOF NORTHERN IRELAND 1991. 1 : 50 000 Sheets 44, 56 and 43. Derrygonnelly and Marble Arch. GEORGE, T. N., JOHNSON, G. A. L., MITCHELL, M., PRENTICE, J. E., RAMSBOTTOM, W. H. C., SEVASTOPULO, G. D. & WILSON, R. B. 1976. A correlation of Dinantian rocks in the British Isles. Geological Society of London Special Report, 7, 1-87.

GUTTERIDGE, P. 1995. Late Dinantian carbonate mud mounds of the Derbyshire carbonate platform. In: MONTY, C. L. V., BRIDGES, P. H., PRATT, B. & BOSENCE, D. (eds) Carbonate Mud Mounds: Origins and Evolution. International Association of Sedimentologists Special Publication, in press. JONES, G. LL. & SOMERVILLE, I. D. 1996. Irish Dinantian biostratigraphy: practical applications. This volume. KELLY, J. G. 1989. The Late Chadian to Brigantian Geology of the Carrick-on-Shannon and Lough Allen Basins, North West Ireland. PhD Thesis, National University of Ireland. -t~ SOMERVILLE, I. D. 1992. Arundian (Dinantian) carbonate mudbanks in north-west Ireland. Geological Journal, 27, 221 243. LEES, A. 1964. The structure and origin of the Waulsortian (Lower Carboniferous) 'reefs' of west central Eire. Philosophical Transactions of the Royal Society of London, 247B, 483-531. - - 1 9 8 8 . Waulsortian "reefs" - the history of a concept. MOmoires de la Institute gOologique de la UniversitO de Louvain, 34, 43-55. & MILLER, J. 1985. Facies variation in Waulsortian buildups, Part 2" Mid-Dinantian buildups from Europe and America. Geological Journal, 20, 159-180. LEGG, I. C., JOHNSTON, T. P., MITCHELL, W. I. & SMITH, R. A. 1995. Geology of the Country around Derrygonnelly and Marble Arch. Memoir of the Geological Survey of Northern Ireland, in press. MILLER, J. 1986. Facies relationships and diagenesis in Waulsortian mud mounds from the Lower Carboniferous of Ireland and Northern England. In: SCHROEDER,J. H. &; PURSER, B. H. (eds) Reef Diagenesis. Springer Verlag, Berlin, 311-335. MITCHELL, M. 1989. Biostratigraphy of Visban (Dinantian) rugose coral faunas from Britain. Proceedings of the Yorkshire Geological Society, 47, 233-247.

262

J. G.

MITCHELL, W. I. 1992. The origin of Upper Palaeozoic sedimentary basins in Northern Ireland and relationships with the Canadian Maritime Provinces. In: PARNELL, J. (ed.) Basins on the Atlantic Seaboard. Petroleum Geology, Sedimentology and Basin Evolution. Geological Society, London, Special Publication, 62, 191-202. -& OWENS, B. 1990. The geology of the western part of the Fintona Block, Northern Ireland: evolution of Carboniferous basins. Geological Magazine, 127, 407-426.

KELLY OSWALD, D. H. 1954. The Carboniferous rocks between the Ox Mountains and Donegal Bay. Quarterly Journal of the Geological Society of London, 111, 167-182. RILEY, N. J. 1993. Dinantian (Lower Carboniferous) biostratigraphy and chronostratigraphy in the British Isles. Journal of the Geological Society, London, 150, 427-446. SHERIDAN, D. 1972. The Upper Old Red Sandstone and Lower Carboniferous of the Slieve Beagh Syncline and its setting in the Northwest Carboniferous Basin of Ireland. Geological Survey of Ireland Special Paper, 2, 1-129.

Controls on ramp, platform and basinal sedimentation in the Dinantian of the Dublin Basin and Shannon Trough, Ireland P. S T R O G E N 1, I. D. S O M E R V I L L E 1, N. A. H. P I C K A R D 2, G. L L . J O N E S 1 & M. F L E M I N G t

l Department o f Geology, University College Dublin, Belfield, Dublin 4, Ireland 2Institute of Biology and Geology, University o f Tromso, Dramsveien 201, N-9037 Tromso, Norway Abstract: In the Dublin Basin a Courceyan ramp phase of sedimentation was followed in

the Chadian by tectonic break-up of the basin into distinct shallow-water platforms, on which production of carbonate sediments continued in considerable volume, and 'deep' basinal areas in which it ceased. Progradation of the platforms across these basinal areas was limited, and mainly confined to the dip-slope of hanging wall blocks; progradation across fault scarps was rare. In the Shannon Trough basement-fault control was evident in the distribution and migration patterns of volcanic centres in the Chadian to Arundian, but despite this, ramp sedimentation occurred throughout the Dinantian, evolving into a purely constructional large platform by late Dinantian time. There was no break-up of the basin as in the case of the Dublin Basin. The reason for the contrasting behaviour of the two basins is related to the rate of upwards movement of extensional faults relative to sedimentation rates. In the Dublin Basin these faults penetrated to the palaeosurface to form scarps by the late Chadian, and this topography survived into the Brigantian. In the Shannon Trough these faults failed to surface, but deep basement structures controlled the distribution of Dinantian volcanic centres which lie on a series of ENE-trending lineaments. These lineaments, which parallel the axis of the Shannon Trough, almost certainly mark the traces of active down-to-basin faults that controlled its half-graben structure. The basement rocks of the two basins are clearly of a different nature; the Dublin Basin is floored by basement of a much more heterogeneous nature than the Shannon Trough, the former lying south and the latter north of the putative Iapetus Suture line.

The Dublin Basin sensu lato (Pickard et al. 1994) and the Shannon Trough are two of a series of basins of Dinantian age in Ireland which are connected by shallow shelf areas (Fig. 1). The maximum thickness of Lower Carboniferous sediments is approximately the same (2-3 km) in both basins; net subsidence rates are closely comparable, but despite this the styles of sediment-fill in the two basins are quite different (Fig. 2). The Shannon Trough behaved as a ramp system throughout the Dinantian (Somerville & Strogen 1992; Fig. 3), whereas the Dublin Basin evolved from an early ramp into discrete platforms separated by deep-water basins by late Chadian time (Fig. 4; Pickard et al. 1994). This paper summarizes recent work on the two basins and attempts to explain the differences in their behaviour in terms of basement structural control. Both basins are now inverted, but the level of tectonic deformation is low. In the Dublin Basin, which has many similarities with the Craven Basin in England, the horst blocks that formed the Visran platforms have suffered greatest uplift and a good deal of the fiat-lying shallow-

water carbonates have been eroded; the basinal areas are gently deformed by folding but their stratigraphy is largely intact. The Shannon Trough has suffered only gentle folding (less than 10% shortening) with minor overthrusting. Even now its basin-margin faults are covert. The geophysical signature of the basement to the Dublin Basin is complex on both gravity (Murphy 1981) and magnetic maps (Morris & Max 1995), and a plethora of sub-Carboniferous faults have been detected (Andrew 1992, Fig. 4). On the other hand the basement signature of the Shannon Trough is s i m p l e r - it is magnetically quiet, apart from shallow anomalies caused by Carboniferous volcanic rocks, and it coincides with a broad gravity high (Murphy 1981). Basement control on sedimentation has been widely invoked in the Dublin Basin (MacDermot & Sevastopulo 1972; Strogen & Somerville 1984; Nolan 1986, 1989; Rees 1987; Jones et al. 1988; Philcox 1989; Andrew 1992; Pickard et al. 1992, 1994; Somerville et al. 1992a), in which there has been extensive drilling programmes. Much less has been published on the sedimentology of the

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 263-279.

264

P. STROGEN E T A L . N

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=========================

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Fig. 1. Map of Irish Dinantian sedimentary basins in late Chadian to Arundian times showing the Dublin Basin, Shannon Trough, South Munster Basin, and the Northwest Basin. Amended after Strogen et al. (1990). The trace of the Iapetus Suture is taken from Morris & Max (1995). The letters A-E indicate the general areas of the stratigraphic successions shown in Fig. 2. The locations of Figs 5 & 11 are also indicated. Shannon Trough (Strogen 1988; Somerville and Strogen 1992; Somerville et al. 1992b) and even less on its deep structure. Such deep structure is of great significance to, inter alia, basin analysts and mineral explorationists. Many major mineral deposits must remain to be detected by the recognition of subtle and hidden basement structures from a study of the Dinantian sedimentary record. The Irish Dinantian is host to many zinc-lead-silver-barytes deposits of enormous economic importance, such as at Navan (one of the world's largest zinc orebodies), worked-out mines such as Tynagh and Silvermines, and mines just coming on-stream such as Lisheen and Galmoy (for details see Johnston et al. this volume). All of these have been found close to surface; the next generation of mines will almost certainly be deeper and less readily detected by surface geophysics. More subtle modelling of fluid-flow pathways and the basement structures that control them will be required.

The geological account of the Dublin Basin draws heavily on recent work by the authors on deep (> 1 km) cored boreholes, and on areas such as Navan and Walterstown where large numbers of shallower (200-700m) boreholes have been drilled. Analysis of the Shannon Trough relies heavily on work on the Limerick volcanics (Strogen 1983, 1988) augmented by fieldwork on the enclosing limestones, and updated biostratigraphic control by Somerville et al. (1992b). Limited numbers of boreholes have also become available in the Shannon Trough (see below).

Dublin Basin Two phases of sedimentation can be recognized in the Dinantian of the Dublin Basin: ramp sedimentation in the Courceyan to early Chadian (late Tournaisian), and platform and basin sedimentation from the late Chadian to Brigantian (Vis~an).

D U B L I N BASIN A N D S H A N N O N T R O U G H

Ramp phase

265

f o u n d in Jones et al. (1988), Strogen et al. (1990) and Pickard et al. (1992, 1994). The n o r t h w a r d s advancing transgressive phase in the n o r t h of the Dublin Basin overlies a thin and discontinuous Old Red Sandstone fluvial facies, with clasts of

Figure 5 is a m a p o f the D u b l i n Basin, and the formal lithostratigraphy of key areas is shown in Figs 2 and 4. Details of the lithofacies can be ~ ,l,~ A

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266

P. STROGEN E T AL. RATHKEALE

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Fig. 3. Geological successions at the three points of the dog-leg section shown on Fig. 11. The axis of the Shannon Trough lies along the estuary, in the NW corner of the map and westwards. Note (i) the greatly thickened Limerick Limestone Formation to the northwest and west, and (ii) the lateral passage from shallowwater grainstones in the Limerick Syncline through mid-ramp nodular wackestones into the deep ramp shales and storm beds of the Rathkeale Beds. The two westerly successions both display shoaling sequences.

local origin. This is overlain by peritidal heterolithic mudstones, sandstones and limestones of the Liscartan Formation and the laterally variable micrites and grainstones (inner ramp) of the barrier-shoal complex of the Meath Formation. Deeper water and relatively uniform stenohaline conditions across the mid-ramp are represented by the overlying Moathill and Slane Castle Formations in the north of the basin, and the equivalent upper part of the Malahide Limestone Formation in the south. Deeperwater Waulsortian mud-mounds on the outer ramp are recorded by the overlying widespread Feltrim Limestone Formation. The Malahide Limestone, a nodular argillaceous packstonewackestone sequence displaying weak cyclicity (Strogen & Somerville 1984; Lipiec et al. 1994) shows a simple isopach pattern in the southern part of the basin (Jones et al. 1988; Fig. 6a). This pattern reflects increased accommodation towards the East Midlands Depocentre, but during deposition there seems to have been little difference in depositional conditions, especially

water depth, across the ramp; lithofacies in outcrop and boreholes show little systematic variation. In the northern part o f the basin around Navan there is, over the same time interval, a definite trend to more muddy facies from north to south and a distinctly mud-prone, thinly bedded interval, the Woodtown Member of the Slane Castle Formation, can be recognized, notably in the Woodtown Borehole (Figs 2, 5). This trend culminates in a shale-prone area around the Trim Borehole, but on a basinal scale these lateral changes are very gradual. The Waulsortian-facies Feltrim Limestone Formation (Fig. 6b) behaves in the same manner; no systematic changes in depth-related flora and fauna (cf. Lees & Miller 1985) can be found across the basin (Somerville et al. 1992c) except in the very marginal areas where it is replaced by slightly argillaceous, shallow-water bioclastic limestones (e.g. the Kilbride Limestone Formation in the Kingscourt Outlier (Strogen et al. 1995) and the Lane Limestone on the Balbriggan Block).

D U B L I N BASIN A N D S H A N N O N T R O U G H

North Dublin Basin

STAGES NAMURIAN

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Fig. 5. Simplified geological map of the Dublin Basin showing the localities and boreholes referred to in the text. Boreholes: A, Athboy; E, EP25; FT, Flemingstown; FH, Fairyhouse; W, Woodtown.

268

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banks

Fig. 6. Isopach maps for the sub-Waulsortian succession (a) and the Waulsortian Feltrim Limestone Formation (b) in the southern part of the Dublin Basin. Note the basin-filling geometry and smooth gradients for both successions. Modified with additional data from Jones et al. (1988) and Somerville et al. (1992c).

DUBLIN BASIN AND SHANNON TROUGH The early, rapid, lateral variation in shallowwater facies, followed by monotonously uniform or very gradually changing deeper-water facies, is indicative of sedimentation on a shallow homoclinal ramp during a major relative sealevel rise. Although the Balbriggan Block clearly acted as a positive area in the east of the basin (Figs 5 and 6), the area of the basin between it and the Leinster Massif simply behaved as a double-sided ramp. Thickness variations reflect the slow differential subsidence which created accommodation space, in turn matched by sediment accumulation, and at no time did any appreciable relief exist across the basin. The northern part of the basin seems to have been the windward side, since the wavedominated barrier-beach and lagoon system of the Meath Formation is by far the thickest on that side of the basin and is represented on the southern side mainly by argillaceous limestones of the lower part of the Malahide Formation. A landmass clearly existed on the northern side of the basin, which was a mixed carbonatesiliciclastic system since much of the Meath Formation is sandy, whereas the Leinster Massif appears to have been submerged, if only to shallow depths, at this time. Andrew (1992) emphasized the existence of old Caledonide fractures in the sub-Tournaisian floor, and inferred that there is no evidence (from regional isopachs and lithofacies) that these were significantly active during deposition of the Navan Group (Fig. 4). During deposition of the overlying Cruicetown Group (Figs 2, 4) the thickness variation appears to indicate the influence of these structures, since a number of depocentres are outlined on Andrew's maps by these basement fractures. Two points can be made concerning these faults. Firstly, recent detailed logging of a large number of closely spaced boreholes in the south of the Kingscourt Outlier north of Kells (Fig. 5) has shown that there is a local depocentre in the Liscartan and lowermost Meath Formations, which coincides in position with the proximal hanging wall of the later Kingscourt Fault (Strogen et al. 1995, Fig. 8). The Walterstown Fault (Pickard et al. 1994) was also slightly active at this time, and some soft-sediment deformation occurred in the hanging-wall rocks. The influence of these faults is observed in subtle changes, indicating only minor fault-controlled relief prior to the late Asbian. Secondly, the multiple depocentres were almost certainly created by extensional tectonics (e.g. see Gibbs 1987, Fig. 5), but these faults did not penetrate through the sediment pile to the surface.

269

P l a t f o r m a n d basin p h a s e

This simple situation changed abruptly in the late Chadian when the ramp rapidly broke up into distinct platforms and basins sensu stricto. The dividing faults had surfaced, and erosion occurred on many of the footwall blocks such as the Kentstown Block (Pickard et al. 1994). On the platforms, limestones of the Milverton Group were then deposited. These are clean, shale-free, thickly bedded grainstones and packstones with scarce carbonate buildups (Fig. 4; Somerville et al. 1992a; this volume). In contrast, mudstones of the Tober Colleen Formation were deposited in the basins and form a major basin-wide event, and were followed by shale-rich successions of calciturbidites, debris flows and other mass-flow deposits of the Rush Conglomerate and Lucan Formations. These rocks contain blocks of lithified carbonates and pre-Carboniferous clasts, particularly in the Rush Conglomerate Formation, but otherwise are dominated by platform-shed shallow-water fauna and flora. The majority of beds are graded and finely laminated. They become finer-grained away from the platform areas and are almost barren calcisilts in areas distal from platforms such as the Flemingstown Borehole (Fig. 5; Pickard et al. 1992, 1994). No autochthonous carbonate is present, and it is clear that the basin floors had subsided below the level of benthic carbonate production. From regional mapping and abundant borehole control it is clear that the platforms were the footwall blocks from the late Chadian till the Brigantian, whilst the basins were the hanging-wall blocks of active fault systems. Uplift and erosion of the footwall blocks due to isostatic readjustment to faulting (Barr 1987) was widespread at this time (Pickard et al. 1992, 1994), and in consequence Vis6an limestones lie directly on Lower Palaeozoic rocks around the margins of the Balbriggan Block (Figs 2, 4). The onset of this faulting event was sudden and apparently due to a phase of regional tectonic activity, as suggested for the northern and central England basins at this time by Gawthorpe et al. (1989) and Ebdon et al. (1990), but fault movement persisted intermittently and less dramatically until the late Dinantian (Brigantian) when a second major tectonic event occurred. During this time interval the margins of the platforms seemed unable, in general, to prograde out into the basins. Such progadation across the fault margins of the blocks depends on:

270 (a) (b)

(c)

(d)

P. STROGEN E T A L . the rate and extent of relative displacement along the bounding faults; the rate of carbonate production on the footwall b l o c k s - a function largely of water depth and total production, which is a function of platform area; the rate of accommodation provided on the platforms themselves by subsidence of the footwall blocks; other physical controls on sediment e x p o r t - wave and storm intensity, directions of drift, tidal scour etc.

For example, near Walterstown (Fig. 5; Pickard et al. 1992) a thin Milverton Group platform succession is known from a series of 31 Boreholes. However, only 400m away to the SW across the Walterstown Fault, the Skreen borehole penetrates only basinal facies of the Lucan Formation. This borehole terminated at 817 m, still in Holkerian calciturbidites (Pickard et al. 1992, fig. 8). The thickness of basinal carbonate rocks here is far in excess of that remaining on the platforms, but still the deposystem was incapable of prograding across the active Walterstown Fault, despite the obvious lack of accommodation space on the platform. Since carbonate production was obviously copious, it appears that the controlling factor in this situation was a high rate of movement on the Walterstown Fault. Since there is no sign of shoaling within the basinal sequence, it is possible that movement, probably on a smaller scale, occurred on the Skreen Fault continuously or intermittently throughout the Vis6an, rather than as a single event in the late Chadian as proposed by earlier workers (Nolan 1989; Pickard et al. 1992). It seems improbable that the enormous amount of new accommodat i o n - well over 800 m - could be created in a single tectonic event, followed by purely passive carbonate export down the steep fault scarp. Northwards progradation down the gentler slopes of the tilted footwall block did take place. At Beauparc (Fig. 5) the transition from basin to platform successions has been recorded (Fig. 7), with distal slope (turbidites of the Lucan Formation) passing up into proximal slope (massive limestone breccias of the Fennor Formation) and platform margin (skeletal wackestones and packstones of the Mullaghfin Formation). Progradation of the platform across the tilted footwall of the Walterstown Block can be traced through a series of boreholes (Fig. 8). The thin succession remaining on the Waiterstown Block must simply reflect the low rate of

25 -

20 _

Asbian

Asbian

T

PLATFORM MARGIN (MULLAGHFIN FM.)

Asbian

15

PROXI ~L SLOPE (FENNOR FM.) 10 ,Ikerian to early Asbian

DISTALSLOPE - BASIN (LUCAN FORMATION)

0

~1

Skeletalwaeke/packstones Massive limestone breccias Calcareous shales & rare micrites 1 Graded, laminated calciturbidites

Fig. 7. Stratigraphic log of the succession in Beauparc Quarry (Figs 5, 8) showing the rapid transition from basinal to platform facies in the late Holkerian to early Asbian.

net subsidence of this particular block. Other footwall blocks accumulated far greater thicknesses of platform carbonates: for example, at Kingscourt the extensive Ardagh Platform succession is 900 m thick compared with 300 m at Walterstown (Pickard et al. 1992; Strogen et al. 1995). The Kingscourt Block is one of the few platform areas in which progradation occurred across the bounding fault. Outcrops

DUBLIN BASIN AND SHANNON TROUGH

271

Brigantian Asbian Holkerian Arundian Late Chadian

~

LowerPalaeozoic ~ &earlyCourceyan rocks

CruftyFm. ~

Holmpatrick& ~ MullaghfinFms.

Basinal ~ facies

Basinslope facies

MILVERTON GROUP

Fig. 8. Model for progradation of the platform facies of the upper Holmpatrick Formation and Mullaghfin Formation across the tilted footwall of the Kentstown Block after the initial transgression of the Crufty and lower Holmpatrick Formations. Modified with additional data points from Pickard et al. (1992).

of the Ardagh Platform carbonates including buildups (Somerville et al. this volume), occur several kilometres south of the known position of the basin bounding fault at Nobber (Fig. 5): i.e. they overlie basinal calciturbidites for several kilometres south of the earlier platform margin. There is some suspicion that this huge mass of platform carbonate may be allochthonous (see discussion in Strogen et al. 1995), but this in itself could be the result of oversteepening of the platform margin, as demonstrated by the failure of the Florida platform edge described by Mullins et al. (1986), or Cambrian examples described by Stewart et al. (1993). Clearly the Ardagh Platform was a far greater exporter of carbonate sediment, probably because of its greater area of production compared to the Walterstown Block. In Co. Longford, in the northwest of the Basin, three deep boreholes (2780/1, 3363/1 and 666/1) only a few kilometres apart (Fig. 5) also demonstrate limited progradation into the basin (Fig. 9). The pre-Tober Colleen Formation succession is lithologically uniform in the three boreholes, although it thickens down to basin (boreholes 3363/1 and 666/1). Break-up into platform and basin is indicated by the mud blanket of the Tober Colleen Formation. In the vicinity of borehole 2780/1 the Tober Colleen Formation is overlain by platform carbonates of the Oakport Limestone Formation. The latter is a succession of shale-free, thickly bedded grainstones and packstones, with thin micrites near the base. It is equivalent to the lower part

of the Milverton Group to the east (Fig. 2). In contrast, in boreholes 666/1 and 3363/1 a very thick Tober Colleen Formation is succeeded by basinal turbidites of the Lucan Formation. It is notable that the platform succession here is much thicker than the adjacent contemporaneous basin succession: i.e. accommodation space was created here far more quickly than on the Walterstown Block, thereby reducing the rate of export of carbonate sediment. By Arundian time the truly basinal Lucan Formation in 3363/1 had been superseded by the Ballymore Formation, a slope facies of bioturbated nodular pack-wackestones, although true platform carbonates are not seen in this borehole. There is a strong suggestion here that in this part of the basin the marginal fault was still active, but that the massive carbonate production from this platform (which is known to extend over several thousand square kilometres of Counties Longford, Roscommon and Leitrim to the W and NW) was slowly causing progradation by the late Arundian. By late Asbian time the upstanding blocks in the northern part of the Dublin Basin began to be overwhelmed, and the shale-rich basinal Loughshinny Formation extended rapidly across the platforms, passing up into the black shales of the late Brigantian to Namurian Donore Formation (Fig. 4; Rees 1987; Pickard et al. 1994). The cause of this event is uncertain, but carbonate production was terminated by a massive influx of terrigenous muds onto the platform areas. The Drumanagh Member

272

P. STROGEN E T A L .

3363/1 23.5

O13= Overburden 2780/1 [ ~ ....

r',.",.";r.~n~,-b~d,~o.~u~r phase). Basement faults were subtly active in the :".."..":~a~ t~mesto~es ~ Massive pale ..... ~aJo-f,o~J=~s,....

first phase, highly active in the second when they created surface scarps (giving rise to block and r~lnterbeddodcalcibasin topography), and dormant during the •7-'7-'.4' turbiditesand shales E~Terrigenousmud_ Namurian. They were inverted during late ~f os~°s Argillaceous r e'~n mrozlular m~r'°~ a t Hercy i oniann de- _

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Palaeozoic

Volcanic r o c k s in the D u b l i n Basin

Volcanic rocks are far less voluminous in the Dublin Basin than their counterparts in Limerick, and their exact range in age is uncertain. Thin, very fine-grained tuff bands are known in boreholes from the late Courceyan Slane Castle Formation, both in Counties Meath and Westmeath (Strogen 1995) and in Co. Longford (Fig. 9). Further south they also occur at this level at Lisheen (Hitzman et al. 1992; Shearley et al. this volume). The exact nature of these ash bands is uncertain. Some may be intrusive tuffisites, which commonly follow bedding for considerable distances, as observed in closely spaced boreholes from the Gortdrum area of SE Limerick. Only at Croghan Hill, Co. Offaly (Fig. 5), are there substantial outcrops of volcanic rocks. These are coarse vitric tufts and ankaramite lavas, over a hundred metres thick, which have also been encountered in boreholes towards Edenderry and near Lough Ennel (Strogen

Fig. 9. Comparative logsofboreholes2780/1, 3363/1 and 666/1 from the Longford area (Fig. 5). Note the obvious correlation of the lower part of the succession and the breakdown of this simple scheme in the late Chadian. Borehole 3363/1 contains elements of both the basin and platform stratigraphies for this part of the basin (see Fig. 2). V, tuff band within the late Courceyan rocks.

Shannon Trough

(Fig. 4) of the Loughshinny Formation contains many beds of limestone breccias, indicating a major tectonic event. This appears to be of regional extent since it has been recorded in the northern England basins at this same time by Ebdon et al. (1990) and Adams et al. (1990). Clearly a regional tectonic cause seems most likely, although a thermal subsidence phase has been suggested (Leeder 1982) for this period of Carboniferous basin evolution, In simplistic terms the evolution of the Dublin Basin seems to have progressed through three phases (Fig. 10). The first was one of slow stretching (ramp phase), culminating in a phase of active rifting (platform and basin phase), followed by thermal subsidence or tectonic uplift of terrigenous source areas (platform demise

There has been far less deep drilling in the Shannon Trough, although scattered boreholes such as the Pallaskenry (Li/68/10), Kilteeely (O1/ 81/3). Cappagh (TU-76/4) and Roche's Castle (Li/69/1) boreholes (Fig. 11; Somerville & Jones 1985; Somerville et al. 1992b) have now become available, and drilling activity is increasing. Despite less detailed knowledge, however, the broad stratigraphy and thickness variations are clear (Strogen 1988). Since the total subsidence in this basin was comparable with that in the Dublin Basin (Fig. 10), one might expect a comparable series of phases here along the Shannon Estuary. In fact this did not happen. The Courceyan history of the Shannon Trough is, however, closely comparable with that of the Dublin Basin (Figs 2, 3). An early marine transgression gave rise to the sandy Mellon House and Ringmoylan

1995). They appear to be of Arundian age and are similar in chemistry to lavas of the Asbian Knockseefin Formation in Limerick.

DUBLIN BASIN AND SHANNON TROUGH Dublin ramp phase Dublinrifting phase Dublinlhermal [ [ sag phase

o

500

-qj,~j~,

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if_

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3000

DUBLIN ,N

Subsidence~1"

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Fig. 10. Cumulative subsidence and sediment fill curves for the Dublin Basin and Shannon Trough. Sediment thickness data from Shephard-Thorn (1963), Jones et aL (1988), Strogen (1988), Nolan (1989), Strogen et al. (1990), Pickard et aL (1992) and Somerville et aL (1992a, c). Timescale from Harland et al. (1989). The shaded area broadly indicates the water depths in each basin at any given time.

Shale Formations, the latter rich in brachiopods, followed by richly fossiliferous argillaceous packstones and wackestones of the Ballymartin and Ballysteen Formations, which are comparable in facies with, and timeequivalent to, the Malahide Formation in south Dublin. This is followed, as in Dublin, by the deeper water Waulsortian facies of the Limerick Limestone Formation. The Shannon Trough in the late Courceyan appeared to undergo rapid deepening (Fig. 10), comparable with the earlier (Strunian) subsidence phase of the South Munster basin to the south (Strogen 1988; Naylor et al. 1987). This Courceyan to early Chadian succession in the Shannon Trough is very similar in facies (Fig. 2) and thickness (c. 900-1100m Fig. 10) to that in the south of the Dublin Basin. Following this stage, break-up of the Dublin Basin began, but in the Shannon area, despite widespread volcanism at this time (Strogen 1974, 1988), ramp conditions prevailed (Somerville & Strogen 1992). Despite the vagaries of outcrop and lack of deep drilling, it is possible to establish that in the Limerick Syncline area (Fig. 3), clean shallow-water grainstones (inner

273

ramp) accumulated throughout the Chadian to Asbian, whereas on the Shannon Estuary, late Chadian nodular argillaceous limestones of the Cooperhill facies (mid-ramp) shoaled upwards into cleaner grainstones of the Arundian Mungret Limestone Formation (Fig. 3), some parts of which contain cross-sets up to 13m high. This succession can be traced westwards into storm-dominated packstones and grainstones, and further west still into the cleaved mudstones and thin storm-beds of the Rathkeale Formation (outer ramp). The latter in turn shoals upwards into the nodular to clean grainstones of the Durnish Limestones (Shephard-Thorn 1963) and into the shallow-water micrites of the Shanagolden Limestones and Parsonage Beds, with goniatites and a Cyathaxonia-phase coral fauna. Graded beds are rare and confined to thin storm-beds within the Rathkeale Beds. Throughout this succession there is an absence of the abrupt lateral facies changes so characteristic of the Dublin Basin at this time, and everywhere there is evidence of slow but continuous shoaling. The picture, therefore, is one of a ramp system initiated in the Courceyan, sloping from E to W, which evolved very slowly into a shelf-like geometry by Asbian time (Somerville & Strogen 1992). The subsidence history of the Shannon Trough is far less complex than that of the Dublin Basin (Fig. 10) and is shown schematically in Somerville & Strogen (1992, Fig. 8). There are parallels between events in the two basins in periods of sudden deepening (e.g. during deposition of the Waulsortian facies of the Feltrim and Limerick Limestone Formations), but complete contrasts in behaviour after the Chadian. There is an apparent lack of direct basement-fault control of sedimentation in the Shannon Trough, where steady and progressive subsidence was matched by carbonate sedimentation (Fig. 10). No overt down-to-basin faults can be detected. However, a closer look at the volcanic rocks indicates that these faults are present. Basement control is subtle, but it is covert, and best identified from the distribution of contemporaneous volcanic rocks.

Volcanic r o c k s in the L i m e r i c k S y n c l i n e

In the Limerick Syncline volcanic rocks range in age from late Chadian to Asbian (Strogen 1988; Somerville et al. 1992b), extending into the Brigantian at Tulla, Co. Clare (Schultz & Sevastopulo 1963). Two volcanic formations, the Knockroe and Knockseefin Formations,

274

P. STROGEN E T AL.

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Fig. ! 1. Geological map of County Limerick showing the dog-leg section of Fig. 3, and the lineaments defined by

the volcanic centres within the Knockroe Volcanic Formation. Boreholes: P, Pallaskenry (Li-68-10); C, Cappagh (Tu 76-4); K, Kilteely (OI 81-3). Also shown is the Roche's Castle borehole (Li-69-1) and the outline of Fig. 12. occur in the Limerick Syncline (Strogen 1988; Fig. 12). The Knockroe Formation (Chadian to Arundian) consists of a bimodal suite of basalts and trachytes. The basalts are mildly alkaline and olivine-normative, with high TiO2, KzO and other large ion lithophile (LIL) element contents. This geochemistry is consistent with a very low degree of melting of an undepleted withinplate mantle. The trachytes result from highlevel, low-pressure fractionation of olivine, clinopyroxene and plagioclase and are leucocratic, commonly potash-rich rocks. The Asbian-aged Knockseefin Formation is made up of primitive ankaramites, unaccompanied by trachytes. These ankaramites are similar to those found as intrusives in the former Gortdrum Mine to the east of the Syncline and in the Arundian volcanics at Croghan Hill, Co. Offaly. Arundian basic volcanics also occur to the south at Subulter and Ballygiblin in Co. Cork (Clipstone 1992). Six volcanic centres have been identified within the Knockroe Formation. At each there was a parallel evolution in volcanic phase (Table 1). The early phases of eruption at any volcanic centre were Surtseyan in style - highly explosive basaltic eruptions caused by the interaction of

basaltic magma with seawater pouring into the subaqueous vent, as at Surtsey in 1964 (Thorarinson 1967). This activity built up lowangle vitric tuff-rings which were later buried and preserved, once the vents became watertight, by quiet subaerial emission of voluminous basaltic laves; these flowed non-explosively into the sea (Strombolian to Hawaiian phases). The individual tuff-rings, which range from 6 to 8 km in diameter, can be mapped out; overlapping rings can often be distinguished by their different crystal contents, reflecting the phenocryst content of the rising basaltic magma. Groups of basaltic lava flows can be followed for distances of up to 10 km. Magma residence time in the upper crust was prolonged, and many of the laves are highly porphyritic. Fractional crystallization gave rise to hawaiites and trachytes. The latter were emplaced mainly as intrusives in dykes, sills and vents, and as crater fills and short stubby flows. They are confined to areas close to major volcanic centres. Later phases of eruption from the same vents gave rise to early phases of Vulcanian vent-clearing, producing proximal lithic tufts rich in trachyte fragments. Several phases of trachyte intrusion occurred. Epiclastic

DUBLIN BASIN AND SHANNON T R O U G H

275

:~;~;~;!;~ I

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Knockroe Volc. Fro.

of basaltic breccJas

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Limerick Lst. Fm.

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Lineaments

Fig. 12. Map of the Limerick Syncline emphasizing the distribution of intrusive trachytes and syenites, the thinning and fining directions of tufts, and the flow directions of lavas within the volcanic piles. The three lineaments defined by the volcanic facies of the Knockroe Formation are shown.

volcanic sandstones are found interbedded with the volcanic rocks at all levels in the succession. Analysis of volcanic facies distribution, flow directions in lavas, and of fining and thinning directions in tufts (Fig. 12) has identified six volcanic centres during the Chadian to Arun-

dian interval (Strogen 1983). These centres lie on two ENE-trending lineaments (2 & 3 in Figs 11, 12), and the volcanic centres migrated from WSW to ENE with time as shown by the sense of overlap of the volcanic piles. These two major lineaments also pass, to the SW, through the

Table 1. Sequence of volcanic phases in a Dinantian volcanic centre in the Limerick area

Phase of eruption

Environment

Eruptive mechanism

Products

Geometry

Epiclastic

Subaerial to marine

Aeolian, fluvial and shallow marine erosion

Volcanic sandstones Marginal and interdigitating

Vulcanian

Subaerial

Vent clearing

Lithic tufts

Steep ash cones

Trachyte emplacement

Subaerial to intrusive

High level fractionation

Trachytes

Tholoids, stubby flows, dykes and sills

Hawaiian

Subaerial to submarine

Quiet lava emission

Basaltic lavas

Extensive lava sheets

Surtseyan

Marine

Phreatomagmatic explosions

Vitric basaltic tufts

Low tuff rings

276

P. STROGEN E T AL.

volcanic centres at Knockfeerina and Ballinleeny, which cut rocks of Old Red Sandstone facies. Roughly equidistant between these two lineaments, the volcanic pile around the Limerick Syncline thins to only a few metres thickness. These lineaments have little surface expression, but their control on the distribution of the volcanic rocks is clearly related to fundamental basement structures; other such lineaments exist. Trachytes do not flow far from their parent vents or intrude great horizontal distances. The line of trachyte intrusions south of Kilteely (Figs 11 & 12) is phacolithic in geometry and cannot have been emplaced from any of the six volcanic centres, but mark the trace of a third lineament (1 in Fig. 12), which extends eastwards through the trachyte plug at Cullen, tuff-breccia bodies in railway cuttings near Oola (Fig. 11) and NW of Cashel. Further trachyte intrusive bodies exist at Maddyboy (4 in Fig. 11), and almost certainly mark the site of a fourth lineament, and the volcanic rocks at Carrigogunnel (Strogen 1974), Roche's Castle (Penaroya, Li-69-1) and Shanagolden (Fig. 11) mark a fifth ENE-trending lineament. The chemistry of these volcanic rocks (Strogen 1983) indicates that volcanism was due to a low degree of mantle melting and the slow, passive rise of relatively small amounts of basalt magma towards the surface. This is indicative of only a moderate degree of crustal extension and low /3-values. The volcanism is in many ways similar to basaltic volcanism in the North Sea during Jurassic extension, especially in its mild alkalinity, low overall volume and patchy distribution. Clearly basement structures were active in controlling the distribution and migration of volcanic centres across the Shannon Trough during much of the Vis6an, yet direct evidence of their influence on sedimentation is elusive. Nevertheless, these lineaments almost certainly mark the traces of active down-to-basin faults as suggested by Strogen (1988), which controlled the half-graben structure of the Shannon Trough. All five lineaments trend approximately ENE-WSW parallel to the axis of the Shannon Trough, regional isopach trends (e.g. for the Waulsortian facies Limerick Limestone Formation), and the suggested trend of the Iapetus Suture in this area (Morris & Max 1995).

Discussion The question must be asked as to why these lineaments in the Shannon Trough had no

surface expression and did not influence sedimentation directly as in the Dublin Basin. In any basin undergoing extension, basement faults must penetrate upwards through the sediment carapace (Gibbs 1987), and such faults existed in both basins. In both the Dublin Basin and Shannon Trough the initial ramp phase was due to tectonically generated sags. In the Dublin Basin the extensional basement faults propagated up to the sedimen~water interface and initiated a phase of block-and-basin sedimentation. Storm waves, tidal currents and eustatic sea-level changes may have contributed to the export of sediment from the platforms, but tectonics determined the amount of accommodation space created on the platforms and in the receptor basins alongside them. The size of the platforms and the accommodation space provided on them was the most important single factor in the attempted progradation of the carbonate platforms into the basins. In the Shannon Trough tectonic control was more subtle. Major basement faults were present but never surfaced. As a result, export of carbonate sediment from the shallow carbonate-producing areas was less dramatic, and the early ramp slowly evolved closer and closer to a true constructional carbonate platform. Fault topography never appeared and purely sedimentary processes dominated. Despite this, tectonics was the dominant control on the evolution of both basins. The question now becomes: why was the behaviour of the basement faults different in the two areas, one set penetrating to the surface, the other not? This must surely be related to the nature of the basement in the two areas. The behaviour of the Dublin Basin is closely comparable with many of the Dinantian basins in northern England, such as the Craven Basin (Gawthorpe et al. 1989). There, many of the carbonate platforms lie on tectonic blocks cored by granites. It is suggested that the same is true in the Dublin Basin, although evidence for widespread buried granites beneath the southern part of the Longford-Down Massif (Fig. 5) is not strong, but known or suspected beneath the Balbriggan Block (see Pickard et al. 1992) and the Leinster Massif. Recent geophysical analysis of the basement in central Ireland (Morris & Max 1995) indicates that the Iapetus Suture passes south of the Limerick Syncline and north of Navan. Thus the crust beneath the Dublin Basin is largely derived from south of the suture, and beneath Limerick from north of it. It is suggested that the old Caledonian crust to the north of the suture is much more homogeneous than that to

D U B L I N BASIN A N D S H A N N O N T R O U G H

the south, so that there is little tendency for buoyant blocks to rise at the expense of their denser neighbours. However, in the Dublin Basin blocks of contrasting density existed, the lighter ones becoming uplifted to form the platform areas. It is surely significant that in the South Munster Basin, which also lies south of the suture, synsedimentary fault scarps were widely developed (Price 1989). The Northwest Basin (Fig. 1), which also lies north of the suture but at a distance of over 200 km away from the Shannon Trough and over 80kin north of the suture, exhibited block-and-basin sedimentation with some similarities to that in the Dublin Basin (Kelly & Somerville 1992; Philcox et al. 1989, 1992). In contrast, in its lack o f basement heterogeneity the S h a n n o n T r o u g h seems to be a very unusual D i n a n t i a n basin. Regardless o f their f u n d a m e n t a l origin, it remains to be emphasized that during the Irish D i n a n t i a n , tectonic events had a far greater control on the gross pattern of c a r b o n a t e facies distribution than either autocyclic or eustatic events. We wish to thank D. Downing and B. O'Donovan of Tara Prospecting for access to countless boreholes from the Dublin Basin, and P. Powell of Tara Mines for access to selected boreholes from the Navan area. The Geological Survey of Ireland are thanked for permission to examine exploration boreholes from Co. Limerick now stored by them. D. Hunt and A. Sleeman are thanked for their constructive comments on earlier drafts of the manuscript.

References ADAMS, A. E., HORBURY, A. D. & ABDEL AZIZ, A. A. 1990. Controls on Dinantian sedimentation in south Cumbria and the surrounding areas of northwest England. Proceedings of the Geologists" Association, 101, 19-30. ANDREW, C. J. 1992. Basin development chronology of the lowermost Carboniferous strata in the Irish north-Central Midlands. In: BOWDEN, m. m., EARLS, G., O'CONNOR, P. G. & PYNE, J. F. (eds) The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin, 143-169. BARR, D. 1987. Lithospheric stretching, detached normal faulting and footwall uplift. In: COWARD, M. P., DEWEY, J. F. & HANCOCK, P. L. (eds) Continental Extensional Tectonics. Geological Society of London Special Publication, 28, 75-94 CLIPSTONE, D, 1992. Biostratigraphy and Lithostratigraphy of the Dinantian of the Kilmaclenine Area, North County Cork. PhD Thesis, National University of Ireland.

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EBDON, A. C., FRASER, A. J., HIGGINS, A. C., MITCHENER, B. C. & STRANK, A. R. E. 1990. The Dinantian stratigraphy of the East Midlands: a seismostratigraphic approach. Journal of the Geological Society, London, 147, 519-536. GAWTHORPE, R. L., GUTTERIDGE, P. & LEEDER, M. R. 1989. Late Devonian and Dinantian basin evolution in northern England and North Wales. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication 6, 1-23. GIBBS, A. D. 1987. Development of extension and mixed-mode sedimentary basins. In: COWARD, M. P., DEWEY, J. F. & HANCOCK, P. L. (eds) Continental Extensional Tectonics. Geological Society of London Special Publication, 28, 19-33. HARLAND, W. B., ARMSTRONG, R. L., COX, A. V., CRAIG, L. E., SMITH, A. G. & SMITH, D. G. 1989. A Geologic Time Scale 1989. Cambridge University Press, Cambridge. HITZMAN, M., O'CONNOR, P., SHEARLEY, E., SCHAFFALITZKY, C., BEATY, O. W., ALLAN, J. R. & THOMPSON, T. 1992. Discovery and origin of the Lisheen Zn-Pb-Ag prospect, Rathdowney Trend, Ireland. In: BOWDEN, A. A., EARLS, G., O'CONNOR, P. G. PYNE, J. F. (eds) The Irish Minerals Industry 1980-1990. Irish Association for Economic Geology, Dublin 211-226. JOHNSTON, J. D., COLLER, D., ~MILLAR, G. & CRITCHLEY, M. F. 1996. Basement structural controls of Carboniferous-hosted base metal mineral deposits in Ireland. This volume. JONES, G. LL., SOMERVILLE, I. D. & STROGEN, P. 1988. The Lower Carboniferous (Dinantian) of the Swords area: sedimentation and tectonics in the Dublin Basin, Ireland. Geological Journal, 23, 221-248. KELLY, J. G. & SOMERVILLE, I. D. 1992. Arundian (Dinantian) carbonate mudbanks in North-west Ireland. Geological Journal, 27, 221-242. LEEDER, M. R. 1982. Upper Palaeozoic basins of the British Isles- Caledonide inheritance versus Hercynian plate margin processes. Journal of the Geological Society, London, 139, 479-491. LEES, A. & MILLER, J. 1985. Facies variation in Waulsortian buildups, Part 2. Mid-Dinantian buildups from Europe and North America. Geological Journal, 20, 159-180. LIPIEC, M., STROGEN, P. & SOMERVILLE, I. O. 1994. Shoaling sequences in the Courceyan of the Malahide coastal section, County Dublin, Ireland. In: European Dinantian Environments II. University College Dublin, 6-8th, September 1994, Poster Abstracts, 49 50. MACDERMOT, C. V. & SEVASTOPULO, G. D. 1972. Upper Devonian and Lower Carboniferous setting of Irish mineralisation. Geological Survey of Ireland Bulletin, 1, 267-280.

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MORRIS, P. & MAX, M. D. 1995. Magnetic crustal character in Central Ireland. Geological Journal, 30, 49-67. MULLINS, H. T., GARDULSKI, A. F. & HINE, A. C. 1986. Catastrophic collapse of the west Florida carbonate platform margin. Geology, 14, 167-170. MURPHY, T. 1981. Geophysical evidence. In HOLLAND, C. H. (ed.) A Geology of Ireland. Scottish Academic Press, Edinburgh, 225-229. NAYLOR, D., SEVASTOPULO,G. D. & SLEEMAN, A.G. 1987. Subsidence history of the South Munster basin, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication 6, 99-109. NOLAN, S. C. 1986. The Carboniferous Geology of the Dublin Area. PhD Thesis, University of Dublin. - - 1 9 8 9 . The style and timing of Dinantian synsedimentary tectonics in the eastern part of the Dublin Basin, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication 6, 83-97. PHILCOX, M. E. 1989. The mid-Dinantian unconformity at Navan, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds). The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication 6, 67-81. , BAILY, H., CLAYTON, G. & SEVASTOPULO, G. D. 1992. Evolution of the Carboniferous Lough Allen Basin, northwest Ireland. In: PARNELL, J. (ed.) Basins on the Atlantic Seaboard. Geological Society Special Publication, 62, 203-216. , SEVASTOPULO, G. D., & MACDERMOT, C. V. 1989. Intra-Dinantian tectonic activity on the Curlew Fault, northwest Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication 6, 55-66. PICKARD, N. A. H., JONES, G. LL., REES, J. G., SOMERVILLE, I. D. & STROGEN, P. 1992. Lower Carboniferous (Dinantian) stratigraphy and structure of the Walterstown-Kentstown area, Co. Meath. Geological Journal, 27, 35-58. , REES, J. R., STROGEN, P., SOMERVILLE, I. D. JONES, G. LL. 1994. Controls on the evolution and demise of Lower Carboniferous carbonate platforms: northern margin of the Dublin Basin, Ireland. Geological Journal, 29, 93-117. PRICE, C. A. 1989. Some thoughts on the subsidence and evolution of the Munster Basin, southern Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation in the British Isles. Yorkshire Geological Society Occasional Publication 6, 111 122.

REES, J. C. 1987. The Carboniferous Geology of the Boyne Valley Area, Ireland. PhD Thesis, University of Dublin. SCHULTZ, R. W. & SEVASTOPULO,G. D. 1965. Lower Carboniferous volcanic rocks near Tulla, Co, Clare, Ireland. Scientific Proceedings of the Royal Dublin Society, A2, 153-162. SHEPHARD-THORN, E. R. 1963. The Carboniferous limestone succession in north-west County Limerick, Ireland. Proceedings of the Royal Irish Academy, 62, 267-294. SOMERVILLE, I. D. & JONES, G. LL. 1985. The Courceyan stratigraphy of the Pallaskenry Borehole, County Limerick, Ireland. Geological Journal, 20, 377-400. & STROGEN, P. 1992. Ramp sedimentation in the Dinantian limestones of the Shannon Trough, Co. Limerick, Ireland. Sedimentary Geology, 79, 59-75. - - , PICKARD, N. A. H., STROGEN, P. & JONES, G. LL. 1992a. Early to mid-Vis6an shaUow water platform buildups, north Co. Dublin, Ireland. Geological Journal, 27, 151-172. --, STROGEN, P. & JONES, G. LL. 1992b. The biostratigraphy of Dinantian limestones and associated volcanic rocks in the Limerick Syncline, Ireland. Geological Journal, 27, 201-220. & JONES, G. LL. 1992c. Mid-Dinantian Waulsortian buildups in the Dublin Basin, Ireland. Sedimentary Geology, 79, 91-116. & SOMERVILLE, H. E. m. 1996. Late Vis6an buildups at Kingscourt, Ireland: possible precursors for Upper Carboniferous bioherms. This volume. STEWART, W. D., DIXON, A. O. & RUST, B. R. 1993. Middle Cambrian carbonate-platform collapse, southeastern Canadian Rocky Mountains. Geology, 21, 687-690. STROGEN, P. 1974. The volcanic rocks of the Carrigogunnel area, County Limerick. Scientific Proceedings of the Royal Dublin Society, 5, 1-26. 1983. The Geology of the Volcanic Rocks of Southeast County Limerick. PhD Thesis, National University of Ireland. - - 1 9 8 8 . The Carboniferous lithostratigraphy of SE County Limerick, Ireland, and the origin of the Shannon Trough. Geological Journal, 23, 121-137. 1995. Lower Carboniferous volcanic rocks of the Limerick Syncline. In: ANDERSON,K., ASHTON, J., EARLS, G., HITZMAN, M. & TEAR, S. (eds) Irish Carbonate-hosted Zn-Pb Deposits. Irish Association for Economic Geology and Society of Economic Geologists Field Guide, May 1995, 75-80. & SOMERVILLE, I. D. 1984. The stratigraphy of Upper Palaeozoic rocks in the Lyons Hill area, County Kildare. Irish Journal of Earth Sciences, 6, 155-173. , JONES, G. LL. & SOMERVILLE, I. D. 1990. Stratigraphy and sedimentology of Lower Carboniferous (Dinantian) boreholes from West Co. Meath, Ireland. Geological Journal, 25, 103-137.

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D U B L I N BASIN A N D S H A N N O N T R O U G H -, SOMERVILLE, I. D., PICKARD, N. A. H. & JONES, G. LL 1995. Lower Carboniferous (Dinantian) stratigraphy and structure in the Kingscourt Outlier, Ireland. Geological Journal, 30, 1-23.

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THORARINSON, S. 1967. The Surtsey eruption and related scientific work. Polar Record, 13, 571-578.

The influence of climatic change on exposure surface development: a case study from the Late Dinantian of England and Wales SIMON VANSTONE

Postgraduate Research Institute for Sedimentology, University of Reading, Whiteknights, Reading, Berkshire RG6 2AB, UK Present address." Integrated Exploration and Development Services, Enterprise House, Cirencester Road, Ilsom, Tetbury, Gloucestershire GL8 8RX, UK Abstract: Exposure surfaces represent an integral part of Asbian-Brigantian cyclothemic platform carbonates in England and Wales. These are characterized by the association of clay palaeosols, calcrete and palaeokarst and in most instances would appear to have been polygenetic. Alternating calcrete-karst stratigraphies associated with individual exposure surfaces indicate that the climate changed from semi-arid to humid to semi-arid conditions during each sea-level fall/rise cycle. Lowstand intervals were humid and resulted in karstification of the cyclothem-top sediments and the formation of a mineral soil. In contrast, regressive/transgressive phases were semi-arid and resulted in calcretization of the emergent platform carbonates. The influence that climatic cyclicity had upon exposure surface development was modulated by variations in platform bathymetry, subsidence and spatial climatic variation, and platforms exhibit their own individual record of what was essentially an idealized sequence of events. As with the sea-level oscillations responsible for cyclothemic sedimentation, the climatic cyclicity is thought to be the product of orbital forcing and probably reflects either eccentricity-driven shifts in the locus of monsoonal precipitation, or precession-driven variations in monsoonal intensity. If precessional in origin, exposure surface development represents a single minimum to minimum excursion, some 20 ka in duration, whereas if eccentricity-driven this may have been appreciably longer. Nevertheless, the immature nature of the exposure surfaces suggests that emergence was probably only of the order of a few tens of thousands of years.

Sedimentary cyclicity is a prominent feature of late Dinantian (Asbian-Brigantian) carbonate platforms, not only in Britain (Somerville 1979a, b; Berry 1984; Davies 1984; Horbury 1987; Ramsay 1991), but also in many other parts of the world (Ramsbottom 1973; Walkden 1987). Individual cycles are characterized by an upward-shoaling sequence of carbonate lithofacies and are bound at their base and top by exposure surfaces (Gray 1981; Berry 1984; Davies 1984; Horbury 1987). The majority of cyclothems display an incomplete shoalingupward sequence, and subaerial phenomena frequently modify subtidal carbonates. Cycle tops are marked by a distinctive suite of features interpreted as having formed through penecontemporaneous subaerial exposure of the platform carbonates. These include palaeokarstic surfaces, pyritized 'terra rossa'-type insoluble residues, laminar, rhizocretionary and mottle calcretes and bentonitic clay palaeosols (Walkden 1972, 1974; Somerville 1977; Walkden & Davies 1983; Berry 1984; Horbury 1987; Davies 1991; Wright & Vanstone 1991; Vanstone 1993).

Whereas many workers have discussed individual aspects of exposure surface development (Walkden 1972, 1974; Somerville 1977; Berry 1984; Horbury 1987), few have presented integrated models linking palaeosol, palaeokarst and calcrete formation. Two important exceptions are the unpublished work of Gray (1981 ) and a model published by Davies (1991). These two models differ radically in their interpretation of the controls on exposure surface development. Gray's (1981) model isolates three factors that are considered to have influenced subaerial modification of the emergent North Wales Shelf: proximity to the shoreline prior to emergence (essentially bathymetry), which influenced the duration of exposure; climatic change associated with each emergence event, 'which brought about the transition from calcretization to karstification of the cyclothem top; and the accumulation of a mineral soil, which compounded the effects of climatic change. In contrast, Davies (1991) preferred a process-response model for exposure surface development, where cementation of the cyclothem cap terminated intrastratal calcretization and initiated the formation of crustal

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advancesin Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 281-301.

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S. VANSTONE

veneers, and the accumulation of a mineral soil promoted karstification. Furthermore, topographic division of the exposure surfaces into broad, shallow (<3 m) depressions and intervening highs was considered to account for variability in profile form, more mature profiles that formed within depressions being lateral equivalents of simple profiles that developed upon adjacent highs. Although both models are attractive in the sense that they offer explanations of important aspects of exposure surface development, neither model adequately explains the relationship between calcretization and karstification of the emergent platform carbonates and regional variations in exposure-surface character. Analysis here of a large number of exposure surfaces from a variety of different carbonate platforms, indicates, for example, that both autocyclic and allocyclic controls were important. The purpose of this paper is to offer a detailed model for the penecontemporaneous subaerial modification of late Dinantian platform carbonates. It is not the

D D

and areas Fluviodeltaic (yoredale)

intention of the paper to provide detailed descriptions of individual exposure phenomena, as these have been described extensively elsewhere (Walkden 1972, 1974; Berry 1984; Horbury 1987; Davies 1991; Vanstone 1993).

Study areas

Natural exposures are generally unsatisfactory for the study of late Dinantian exposure surfaces, associated palaeosols either being densely vegetated or lying beneath scree slopes at the base of limestone crags. As such their study has almost entirely been confined to quarries, which widely exploit late Dinantiari:i limestones throughout England and Wales. An exception to this general rule has been the analysis of coastal sections in Anglesey and Cumbria, where exhumed exposure surfaces have been studied. With the exception of the Alston and Askrigg Blocks, exposure surfaces have been studied

Shallow marine platformal carbonates Basinal shales / turbidites

Fig. 1. Asbian-Brigantian palaeogeography of England and Wales (adapted from Horbury 1991). Sedimentation rates were profoundly influenced by the underlying tectonic framework and the definition of platform and basinal areas incorporates the sediment isopach data of Fraser et al. (1990). Study areas (.) are also indicated: 1, Gower Peninsula, South Wales Shelf; 2, Mold district and 3, Anglesey, North Wales Shelf; 4, Derbyshire Platform; 5, south Cumbria/north Lancashire Shelf; 6, west Cumbria Shelf.

LATE DINANTIAN CLIMATE

to mixed layer kaolinite-smectite and kaolinite, in a small, but significant number of the palaeosols (Vanstone 1993), indicates, for example, that the smectite was unstable under the prevailing climatic conditions. In general, smectite is only known to form a significant proportion of the total clay content of mature soils in areas of low annual rainfall (Barshad 1966; Birkeland 1984). An additional factor contributing to the smectitic nature of the soils is likely to have been poor internal drainage, high base saturation greatly reducing the rate at which smectite degradation proceeds (Birkeland 1984; Weaver 1989). Thin basal pyritic crusts veneering the underlying palaeokarstic surface (Fig. 2) are interpreted as pyritized 'terra rossa'-type insoluble residues associated with karstification of the cyclothem top, while pyrite concretions within the palaeosols are thought to pseudomorph pedogenic iron oxy-hydroxide precursors (Vanstone 1993). Vertic features, particularly pseudoanticlines, are observed in a small proportion of the palaeosols. Examples include two Asbian palaeosols in Derbyshire, one in Tunstead Quarry (BNGR: SK 097740) and the other in Hillhead East Quarry (SK 079692), and an early Brigantian example in Pant Quarry, North Wales (SJ 198702).

from most of the platform sequences of England and Wales (Fig. 1). These include the Asbian around Buxton and Wirksworth in Derbyshire, the late Asbian and early Brigantian of the Mold district in North Wales, the Asbian of the Gower Peninsula in South Wales, and Asbian and early Brigantian sections around Cockermouth, Egremont, Carnforth and Shap in Cumbria. In total, some 187 exposure surfaces (152 Asbian and 35 Brigantian) have been studied. These represent approximately 50 different stratigraphic levels (up to 40 in the Asbian and 9 in the early Brigantian); however, correlation between quarry sections is in many cases uncertain.

Exposure

283

surfaces

Clay palaeosols Bentonitic clays are a ubiquitous feature of the exposure surfaces (Fig. 2) and represent clay palaeosols formed through the degradation of basaltic volcanic ash (Walkden 1972). Prior to diagenetic modification, the soils were probably almost entirely smectitic in composition (Vanstone 1993), a feature attributed primarily to their immaturity. Degradation of the smectite

Smectitic clay palaeosol. Generally up to 2m thick. Smectite locally degraded to kaolinite and interstratified kaolinite-smectite

Scattered pyrite euhedra

\ Pyrite crust. Venee: cyclothem top sediments. Interpret~ as a terrarossa-type insoluble residue resulting from karstification

~yriteconcretions, many of which probably ~seudomorph pedogenic iron oxy-hydroxide concretions

Laminar calcrete. . Occurs both as veneers of non-karsted and karstified cyclothem top sediments. 1ram 0.5m thick Rhizocretions. ' / Occur within the uppermost 2m of the cyclothem top. Frequently display a common depth of termination

Mamillated

palaeokarstic surface. Gently undulating to ~otholed morphology. Up to 5m relief

D• • ~ a •

• • • • • • • ~

• •

•• •

~:

•

e

•

•

•

•

•

•

•

•

•

6 • ~ •

Calcrete mottle horizons.

"~

Generally bedding parallel. Either in contact with or detached from cyclothem top. Occur up to 5m below the exposure surface

Fig. 2. Diagram illustrating the range of pedogenic and early diagenetic features associated with late Dinantian exposure surfaces of England and Wales. Although some exceptionally well developed exposure surfaces display most of these features, others usually display only a small selection of them.

284

S. VANSTONE

Calcretes A number of macromorphological forms of calcrete are commonly associated with the exposure surfaces. These include rhizocretions, laminar calcretes and calcrete mottle profiles (Walkden 1974; Berry 1984; Horbury 1987; Davies 1991; Vanstone 1993). Rhizocretions are an abundant and almost ubiquitous feature of exposure surfaces in Derbyshire and North Wales. However, they are relatively infrequently associated with exposure surfaces in Cumbria and are notably absent in the Gower sections of South Wales. Rhizoconcretions are usually small (generally <2 cm in diameter), occur within the uppermost 1-2m of the cyclothems and exhibit a strong vertical component (Fig. 2). Laminar calcretes represent another important form of pedogenic carbonate. These are similar in their geographical distribution to rhizocretions and are particularly abundant in the Mold area of North Wales (Somerville 1977, 1979b) where they are present in about a third of the exposure surfaces studied. Laminar calcretes primarily occur as crusts veneering the upper surface of the cyclothem top, although also occur infilling cracks and as anastomozing stringers within the uppermost metre of limestone. Crusts vary in thickness from 1 mm to 0.5m and include calcified rootmats and micritic laminar calcretes. Calcrete mottle profiles within the cyclothem-top sediments exhibit a somewhat curious distribution, being abundant associated with Cumbrian exposure surfaces, but extremely rare in Derbyshire and North Wales, an inverse relationship to that shown by rhizocretions. Rhizocretions are, however, associated with some Asbian calcrete mottle profiles in Cumbria. In all, two distinct types of calcrete mottle profile are distinguished: those in contact with the overlying exposure surface and those detached from it. These have been likened to pedogenic and groundwater profiles respectively (Horbury 1987; Vanstone 1993), being comparable to those documented by Semeniuk & Searle (1985). Brecciation is generally an unimportant feature of the calcrete profiles, although it is common in Anglesey (Davies 1991).

Palaeokarst Palaeokarst is a ubiquitous feature of the exposure surfaces (Fig. 2). This exhibits a smooth, mamillated appearance, a feature

attributed to its formation beneath a soil cover (Walkden 1974). Its spatial distribution is, however, such that lateral expanses of karstified limestones are interspersed with non-karstified surfaces. These non-karstified surfaces form a datum below which the karstified surface extends to form palaeokarstic depressions (Vanstone 1993). These are divisible into individual palaeokarst pits that represent the building blocks of the karstified surfaces. Immature pits typically exhibit a narrow, cylindrical form (Walkden 1974, fig. 5) and are considered to have formed through the concentration of acidified waters by stemflow drainage beneath trees making up the vegetation cover (Vanstone 1993). Palaeokarst pits frequently exhibit a common depth of termination below the cyclothem top, as do rhizocretions, and either indicate the presence of a shallow watertable during formation, or a cementation effect linked to the former presence of the watertable (Vanstone 1993). Enlargement of the palaeokarst pits was also by lateral expansion rather than vertical deepening, suggesting that downward dissolution was either much slower owing to the more tightly cemented nature of the carbonates, or that dissolution was inhibited by saturated groundwaters. Possible evidence for a cementation effect comes from Solomon & Walkden's (1985) study of a late Asbian exposure surface in North Wales, in which porosity occlusion by vadose cements was shown to increase downwards, reaching a maximum of 30% at a depth of 1 m.

Vadose cements Other vadose cements include probable speleothem deposits associated with a late Asbian exposure surface in Eskett Quarry (BNGR: NY 053169), west Cumbria. These include botryoidal growths of radiating columnar calcite crystals that veneer the palaeokarst pits and stellate nodules within the palaeosol fill (Fig. 3.50 in Vanstone 1993) and are similar to 'cave popcorn' speleothem deposits documented by Kahle (1988) and Webb (1994).

Palaeokarst/calcrete relationship Central to an understanding of exposure surface development is the relationship between the calcrete and palaeokarst. These would appear to represent distinct phases of development. For example, thick developments of laminar calcrete crust are commonly observed to both veneer and

LATE DINANTIAN CLIMATE

Fig. 3. Stage 1 laminar calcretes (c) veneering nonkarsted cyclothem-top sediments are illustrated in (a) and (b). Such features are frequently laterally extensive over many tens of metres, as in (a). However, they are also commonly observed to have been truncated by palaeokarst (P), as in (b). This indicates that the laminar calcrete formed during the initial phase ofcalcretization. A similar relationship is also shown by the surface illustrated in (a) and this is represented diagramatically in Fig. 5. The calcrete mottle profile illustrated in (c) also formed during the initial phase ofcalcretization (stage 1), subsequent dissolution having preferentially removed the intermottle carbonate to produce a rubbly cyclothem top (R). This passes gradationally downwards into cemented mottle carbonate (m). (a) Late Asbian, Clintz Quarry, Moota, west Cumbria (BNGR: NY 160357). (b) Late Asbian, Grange Mill Quarry, Derbyshire (SK 242574). (c) Late Asbian, Barland Quarry, South Wales (SS 576896). Metre rule for scale in (a); scale bars in (b) and (c) are graduated in centimetres.

285

Fig. 4. Palaeokarst veneered by stage 3 laminar calcretes. (a) A laterally extensive laminar rootmat calcrete crust (c). This veneers a gently undulatory palaeokarst, and the laminar rootmat thickens into lows on the surface (arrows). Early Brigantian, Graig Quarry, North Wales (BNGR: SJ 205565). (b) This shows a palaeokarst pit (arrow) in which the entire surface has been veneered by a later laminar rootmat calcrete (c). Early Asbian, Tunstead Quarry, Derbyshire (SK 097740). Scale bars are graduated in centimetres. be truncated by palaeokarst. Similar relationships have also been noted by previous workers (Walkden 1974; Somerville 1977; Gray 1981; Berry 1984; Horbury 1987; Davies 1991) and suggest that calcretization or karstification prevailed over prolonged periods of time largely to the exclusion of the other. Systematic evaluation of the relationship between the calcrete and palaeokarst suggests that each phase of platform emergence was characterized by two distinct phases of calcretization separated by a phase of karstification. An initial phase of calcretization (stage 1)

286

S. VANSTONE

resulted in the formation of rhizocretions and mottle calcrete within the unconsolidated cyclothem top and produced laminar calcrete veneers of their non-karsted upper surfaces (Fig. 3a-b). These features are frequently observed to have been modified by subsequent karstification (stage 2). Stage 1 laminar calcretes are, for example, frequently observed to have been truncated by palaeokarst pits (Fig. 3b), whilst solution modified calcrete mottle horizons are characterized by a poorly-cemented rubble of solution-rounded mottles in which the intermottle carbonate has been partially or completely removed (Fig. 3c). Palaeokarst pits that dissect mottle horizons also typically exhibit irregular, rubbly walls, testifying to solutional modification of an earlier-formed calcrete mottle horizon. Rhizocretions are frequently restricted to upstanding palaeokarst hummocks and probably indicate that they formed prior to karstification. In contrast, laminar calcrete crusts that veneer the palaeokarstic surfaces (Fig. 4a-b), faithfully coating their surface irregularities, indicate a second phase of calcretization subsequent to karstification (stage 3). Some key exposure surfaces illustrating the three stages of development are described in Table 1.

0

Distinction between stage 1 and stage 3 calcretes is also particularly well illustrated by an exposure surface in Clintz Quarry, Moota, west Cumbria (Fig. 5). Stage 1 calcretes are widespread and include a thick (0.25-0.50m) laminar calcrete veneer of a non-karsted cyclothem top (Fig. 3a), abundant rhizocretions that extend to a depth of 0.9 m, and a surfacedetached calcrete mottle horizon, between 0.9 and 1.9m sub-surface. Low-density palaeokarst pits (stage 2) truncate both the stage 1 laminar calcrete veneer and the calcrete mottle horizon (Fig. 5) and extend to a depth of some 3m. Where plugged by sandstone, the palaeokarst pits exhibit abundant evidence for a second phase of calcretization (stage 3). This consists of laminar calcrete stringers, pervasive throughout the sandstone fill, and laminar calcrete crusts which partially or completely veneer the palaeokarst pits (Fig. 5). These vary from 2 cm thick associated with the vertical walls, to 20 cm thick around their bases, and are typically meniscate where they veneer the base of a pit. The change from calcretization to karstification and vice versa cannot purely be attributed to a process-response model, such as that offered by Davies (1991). Rather, as has previously been suggested by Gray (1981) and

3m

15m

l

0

Stage1laminarcalcreteveneer :."2.~,

StageI calcretemottles

W S

Stage 3 laminar calcrete veneer Stage 3 laminar calcrete stringers

Stage2 palaeokarstpits withsand or day palaeosolfill Fig. 5. Late Asbian exposure surface in Clintz Quarry, Moota, west Cumbria (BNGR: NY 160357) illustrating three distinct stages of development. An initial phase of calcretization (stage 1) resulted in rhizocretions, a calcrete mottle profile and a thick laminar calcrete veneer of the non-karsted cyclothem top. Within the eastern part of the quarry, deep palaeokarst pits (stage 2) truncate both the laminar calcrete and calcrete mottle profile and indicate that these calcretes formed prior to karstification. A second phase of calcretization (stage 3) followed karstification, resulting in laminar calcrete veneers of the palaeokarst pits and laminar calcrete stringers within their sandstone fills. The insert represents a composite section through the Clintz Quarries with exposure surfaces marked and the horizon in question arrowed.

Bee Low Limestone (late Asbian)

Oxwich Head Limestone (late Asbian)

Cefn Mawr Limestone (early Brigantian)

4

1

7

2

8

1

Grange Mill Quarry, Derbyshire (SK 242574)

Barland Quarry, South Wales (SS 576896)

Graig Quarry, North Wales (SJ 205565)

Hendre Quarry, North Wales (SJ 194679)

Shap Fell Quarry, east Cumbria (NY 587138)

Holme Park Quarry, south Cumbria (SD 535788)

Horizon numbers refer to those of Vanstone (1993).

Urswick Limestone (late Asbian)

Knipe Scar Limestone (Asbian)

Cefn Mawr Limestone (early Brigantian)

Bee Low Limestone (early Asbian)

14

Tunstead Quarry, Derbyshire (SK 097740)

Age

Horizon

Quarry

Occasional rhizocretions occur to a depth of 1.6m (stage 1). Surface-attached calcrete mottle horizon, up to 1.6 m subsurface (stage 1). Potholed palaeokarstic surface, with a relief of 2 m, truncates calcrete mottle profile.

Occasional rhizocretions occur to a depth of 0.2 m (stage 1). Surface-attached calcrete mottle profile, up to 3 m subsurface (stage 1). Potholed palaeokarstic surface, with a relief of 0.9 m, truncates calcrete mottle horizon (stage 2). 0.1 m thick micritic laminar calcrete crust veneers many of the palaeokarst pits (stage 3).

Rhizocretions occur to a depth of 0.4m (stage 1). Surface-detached calcrete mottle profile at 1.4m subsurface (stage 1). Well-developed palaeokarstic depression, some 8.5 m across and 1.2 m deep (stage 2). Palaeokarst relief infilled by palaeosol. Thick laminar rootmat calcrete crust, up to 0.15 m thick, veneers palaeokarst surface and is laterally persistent over several tens of metres (stage 3).

Abundant rhizocretions occur to a depth of 0.55 m (stage 1). A thin clay palaeosol, up to 0.1 m thick, infills the relief on a gently undulatory palaeokarstic surface (stage 2). A thick laminar rootmat calcrete crust, up to 0.1 m thick, veneers the palaeokarst (Fig. 4a, stage 3). This is laterally persistent over intervals of up to 40 m.

Solution-modified (stage 2) surface-attached calcrete mottle horizon (stage 1, Fig. 3c). Prominent 1.5 m thick clay palaeosol with geothite nodules and upper coaly horizon (stage 2).

Rhizocretions occur to a depth of 0.8 m (stage 1). Micritic laminar calcrete crust, up to 6 c m thick, veneers non-karsted cyclothem top (stage 1). Palaeokarstic surface, with 0.15 m relief, truncates calcrete crust (stage 2, Fig. 3b).

A b u n d a n t rhizocretions occur to a depth of 1.2 m (stage 1). Potholed palaeokarstic surface, with up to 1 m relief, infilled by a clay palaeosol with goethite nodules (stage 2). Laminar rootmat calcrete crust, up to 0.1 m thick, veneers some palaeokarst pits (stage 3, Fig. 4b).

Exposure surface character

Table 1. Key features shown by exposure surfaces illustrating well-developed calcrete-palaeokarst relationships

288

S. VANSTONE

Berry (1984), the changes are thought to reflect fluctuations in climate. Holocene calcretes can develop under a wide variety of rainfall regimes (Goudie 1983). The general lack of associated evaporites and the strong biogenic affinity of the calcretes described here suggests, however, that they formed under a semi-arid to subtropical climate. This inference is supported by their similarity to calcretes documented from these climatic regimes (e.g. Multer & Hoffmeister 1968; Arakel 1982; Semeniuk & Searle 1985). In these instances, rainfall is usually confined to the winter months and is interspersed by prolonged dry seasons during which intense evaporation leads to a net annual moisture deficit and the precipitation of calcrete. In contrast, karstification generally characterizes more humid climatic conditions (Wright 1988, 1994). For example, denudation rate has been shown to be directly proportional to total rainfall received, a tenfold increase in precipitation resulting in an order of magnitude increase in the rate of limestone dissolution (Smith & Atkinson 1976; White 1984). Soil PCO2 also profoundly influences the rate of karstification (Trudgill 1985). This reflects the level of biological activity and is itself a function of rainfall regime and temperature. It is therefore possible to recognize three climatically distinct stages of exposure surface development: an initial phase of calcretization (stage 1), karstification (stage 2) and a second phase of calcretization (stage 3). These not only form the basis of the model outlined below, but also represent a means for evaluation of other controls, such as bathymetry and subsidence, influential during their development.

Model for exposure surface development Stage 1 - Initial phase of calcretization (Fig. 6a) Calcretization of the newly emergent platform was in most instances responsible for the transition from an unconsolidated to partially or completely case-hardened cyclothem top. Poor water-retention would have characterized the newly exposed sediments, a function of both diffuse flow through a relatively homogeneously porous substrate, and evaporation associated with the prevailing climatic conditions. As a result, vegetation colonizing the land surface is likely to have experienced at least seasonal water-shortage. The dominantly vertical inclination of the rhizocretions coupled with their

common depth of termination suggests, for example, that the plants were phreatophytes, sinking root systems to tap a shallow watertable some 0.5-2m sub-surface. Carbon isotopic signatures of calcified rootmarks also suggest that the plants were C4 (Hatch Slack cycle) or CAM (crassulacean acid metabolism) photosynthesisers (Wright & Vanstone 1991). Such photosynthetic techniques are widely employed by plants living in water-stressed environments of the present day and represent physiological adaptations that allow more conservative use of water (Moore 1983, 1989). Early cementation of the cyclothem top also frequently resulted in the formation of calcrete mottles. These probably reflect the incorporation of impurities, such as organic matter, iron and clays derived from the overlying soil profile, within early meteoric cements (Horbury 1987). Surface-attached mottle profiles probably formed in response to a gravitational supply of impurities, whereas surface-detached mottle profiles probably precipitated within the capillary-rise zone, impurities being derived from the watertable. Given sufficient time, cementation of the emergent platform sediments would have resulted in the formation of an indurated and impermeable substrate, equivalent to the plugged horizon of Gile et al. (J966). Plant roots impinging upon such indurated surfaces were deflected laterally and became interwoven to form rootmats, the calcification of which produced rhizolite laminar calcretes (sensu Klappa 1980). Whereas these probably developed beneath an ash cover, evaporative precipitation associated with bare limestone surfaces probably resulted in micritic laminar calcretes. Formation of laminar calcrete veneers during this initial phase of development was, however, very much a function of time, evaporative climatic conditions rarely being sufficiently prolonged to enable both casehardening of the substrate and subsequent formation of laminar calcrete. As such, earlyformed crustal veneers of laminar calcrete are only observed in some 5% of the exposure surfaces. In addition, brecciation and early joint formation, features that characterize mature calcrete profiles (Arakel 1982; Machette 1985), only occasionally accompanied cementation of the cyclothem top. Throughout this phase of development, windblown volcanic ash accumulated over the newlyemergent land surface, a process that was probably partially assisted by vegetative baffling. The prevalence of evaporative conditions

(a)

/" - Rooting depth governed by initial depth of the watertable (0.5-2m below surface)

Sparse vegetation cover ' ~ _N~

Laminar calcrete veneers formed where the sediment became case hardened

Soil/ash-cover was thin or absent and afforded little water retention,

Newly exposed carbonate sediment acted as a soil for colonising vegetation

- -

-4

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'-.

~

WT

/ Calcrete mottles formed through early cementation of the carbonate substrate

Deep root systems tapped pelicular and ground water sources

Dense vegetation cover

Relief of karst determined by the depth of a shallow watertable

I

Evolution in karst depression form resulted from the lateral amalgamation of component pits

1 •

Karst pits initiated by stemflow drainage

.....

Volcanic ash, accumulating at the sediment surface, degraded to form a smectitic soil

Terra rossa t y p e insoluble residue accumulated as a thin crust veneering the karstified cyclothem top sediments

(e) Micritic laminar calcretes probably developed through evaporation of carbonate-rich fluids flowing over exhumed karstic surfaces [

. uf.JL

.I,

Sparse vegetation cover Thin soil-cover afforded little water retention

Case-hardened sediment surface prevented significant root penetration, roots being deflected laterally (plantpot effect) to form woven mats of rootlets. Where calcretised, these formed laminar rootmat calcretes,

~/,~.

Where a thick soil-cover remained laminar calcrete veneers frequently did not develop probably due to the surfaces greater water retention

Fig. 6. Model for exposure surface development. (a) An initial phase o f calcretization (stage 1) resulted in the f o r m a t i o n o f rhizocretions and calcrete mottle profiles within the newly exposed p l a t f o r m sediments a n d laminar calcrete veneers o f the n o n - k a r s t e d cyclothem top. (b) Karstification (stage 2) o f the cyclothem top modified these earlier calcretes. Palaeokarst pits were initiated at sites o f stemflow drainage from trees m a k i n g up the vegetation cover. These evolved into clay-filled depressions t h r o u g h the lateral a m a l g a m a t i o n o f the c o m p o n e n t pits. (e) A second phase o f calcretization (stage 3) resulted in the f o r m a t i o n o f laminar calcretes, these veneering the karstified c y c l o t h e m - t o p sediments.

290

S. VANSTONE

suggests, however, that its degradation would have been slow and true mineral soils are therefore unlikely to have been a characteristic feature of the land surface at this time.

Stage 2 - karstification (Fig. 6b) As the climate ameliorated, probably through progressive increase in seasonal rainfall, calcretization was gradually arrested and the calcretized land surface became subject to karstification. This systematically removed and modified the stage 1 calcretes. Rhizocretions were extensively removed from much of the cyclothem top, frequently only being preserved within palaeokarst hummocks. In addition, many larger rhizocretions were exhumed upon the palaeokarstic surfaces, a reflection of the greater resistance to dissolution offered by their micritic envelopes. Calcrete mottle horizons also became solution-modified and were truncated, as were laminar calcrete veneers of the cyclothem top. As a direct consequence of climatic amelioration a number of other changes are also envisaged to have taken place. Vegetation density, would have almost certainly increased, whilst the taxonomic forms present may also have changed and/or increased in diversity. In addition, a prominent soil-cover, often many metres thick, formed during this stage of development through the chemical degradation of volcanic ash. This was probably partially inherited from stage 1, although with increased adhesion and vegetative baffling under these wetter climatic conditions, it would have accumulated at a faster rate. Despite being of considerable thickness, however, the soils were generally mineralogically immature, a feature at least partially attributable to the short-lived duration of this pluvial stage, clay mineralogical transformations having had insufficient time to equilibrate with the prevailing climatic conditions. In addition, poor internal and external drainage, coupled with high base saturation, may also have influenced the rate at which these transformations took place. During this stage of development the land surface became differentiated into vegetated soilfilled karstic depressions and intervening, relatively bare inter-karstic areas. Depressions are thought to have been initiated around areas of vertical solution pipes that formed as a result of stemflow drainage beneath trees capable of acidifying and concentrating rain-water at specific sites within the carbonate substrate (Vanstone 1993). Spatial compartmentation of

these trees was probably inherited from stage 1. This possibly reflects the preferential colonization of slightly lower areas, a strategy that may have conferred advantages in water conservation. Alternatively, it may reflect reproduction in plants with a poorly developed dispersal mechanism, the daughter plants growing in close proximity to the parent plants. Once initiated, evolution of the depressions would have been self-perpetuating. Improved water retention resulting from the accumulation of a soil cover and increased vegetation density, coupled with greater acidification of the soil waters due to elevated PCO2, would, for example, have promoted karstification at these sites. Whereas the depressions evolved primarily by lateral amalgamation of component karst pits (Vanstone 1993), the interpalaeokarstic areas are likely to have evolved by vertical denudation because of the absence of a vegetation cover. As such they were probably characterized by serrated surfaces typical of subaerial karstification (Sweeting 1972); however, subsequent stylolitization has largely destroyed these features. Where such surfaces later became mantled by a soil cover, these were modified to form mamillated karst. Accumulation of 'terra rossa'-type insoluble residue accompanied karstification, and this formed a thin veneer over the karst surface. In addition, concentrations of iron also occasionally developed within the mineral soil, as a result of its downward translocation through the solum. However, such concentrations are rare, and, together with the clay mineralogy of the palaeosol, indicate that the soils were immature. Other pedogenic features that developed during this stage of development include rare vertic features, such as pseudoanticlines. These features are diagnostic of smectitic soils, termed Vertisols, that form under seasonal climatic regimes (Birkeland 1984). Coals present in many of the soil profiles may also have developed during this time interval, although a more likely scenario is that they formed in association with the hydromorphic event that would have accompanied platform drowning. Vadose cements probably also formed during this stage of development. Widespread cementation of the cyclothem tops (e.g. Solomon & Walkden 1985), particularly within a few metres of the sediment surface, could, for example, have slowed downward penetration of the palaeokarst pits such that their growth was primarily by lateral amalgamation rather than vertical deepening. Rare 'cave popcorn' speleothem deposits are also intimately associated with the

LATE DINANTIAN CLIMATE palaeokarst and probably represent localized redistribution of calcium carbonate. Speleothemic vadose cements also occur as fissure-fill deposits in Brigantian mud-mounds of the Derbyshire Platform (Gutteridge 1995). These are post-dated by a phase of calcretization and may represent stages 2 and 3 of the model presented here.

Stage 3 - second phase of calcretization (Fig. 6c) This stage of development marked the gradual return to a semi-arid climate and resulted in a second phase of calcretization. Whereas stage 1 calcretes were an almost ubiquitous feature, those formed during this stage are only associated with some 5% of the exposure surfaces. These consist of laminar calcrete veneers of the karstified cyclothem top, and are notably absent from many palaeokarstic surfaces exhibiting a thick palaeosol cover. Such a distribution probably reflects the greater water retention afforded by the soil cover, hydrological conditions that would have been detrimental to the formation of laminar calcrete veneers. Reflecting the return to evaporative climatic conditions, the land surface is likely to have been characterized by a sparse vegetation cover and, coupled with drier soil conditions, would probably have resulted in periodic deflation of the inherited soil cover. Drier soil conditions would also have largely terminated clay mineralogical transformations. Associated with bare limestone surfaces, brecciation is occasionally observed to have modified the cyclothem-top sediments (e.g. in Anglesey).

Overriding controls influencing exposure surface development Being responsible for platform emergence-submergence, each sea-level fall/rise cycle was therefore accompanied by changing climatic conditions. Each regressive interval was characterized by the transition from a semi-arid to humid climate, a trend which was reversed during the subsequent transgression. A similar relationship has also been documented by Webb (1994) from an exposure surface at the MississippianPennsylvanian boundary in northwest Arkansas. Here, karstification was also followed by a phase of calcretization immediately prior to

291

marine transgression and provides further evidence for the model envisaged here. While climatic change was a characteristic feature of each sea-level oscillation, the effect it had upon exposure surface development was dependent upon platform bathymetry, subsidence and spatial climatic variation, these overriding factors determining precisely how each individual platform evolved. As such each platform typically exhibits its own individual record (Table 2) of what is essentially an idealized sequence of events.

The influence of bathymetry Platform bathymetry was a particularly important control on exposure surface development, determining the point in the climatic cycle when the platform both became emergent and was subsequently drowned. At its most extreme, shallow platforms that became emergent early during the regressive phase and remained so until late in the transgressive phase may exhibit all three stages of development, whilst relatively deep platforms that became emergent only for a short period of time toward lowstand may only have undergone karstification (Fig. 7). Hence spatial and temporal variation in bathymetry possibly accounts for much of the observed variation in exposure surface form. Representing the first and last stages of exposure surface development, calcretization was probably particularly sensitive to variation in platform bathymetry, its interaction with ameliorating climatic conditions during stage 1 probably accounting for many of the observed differences in calcrete profile form between platforms. Profiles dominated by rhizocretions typify exposure surfaces in the Asbian of Derbyshire and the late Asbian-early Brigantian of North Wales (Table 2) and constitute one end of the spectrum, whilst calcrete mottles, the principal calcrete type in Asbian and early Brigantian exposure surfaces of Cumbria (Table 2), represent the other spectral extreme. Semeniuk & Searle (1985) noted similar, although less extreme, variation in profile form along a climatic transect in southwestern Australia. Here, rhizocretions and calcrete mottles, components of mixed profiles under more humid climatic conditions, decreased in abundance at a disproportionate rate as climatic conditions became drier and more evaporitic. Under these later conditions, rhizocretions are dominant, although would appear to be more sparsely distributed than in the late Dinantian

292

S. V A N S T O N E

Table 2. Exposure surface characteristics in the principal areas of study. Area of study

Exposure surface characteristics

South Wales (Gower Peninsula)

Asbian exposure surfaces of the Oxwich Head Limestone are noted for the absence of calcretes. Rhizocretions are, however, present in the Oxwich Head Limestone of nearby Porthcawl sections. Palaeokarstic surfaces are subdued to moderaely mature. Exposure surfaces are absent in the Brigantian (Oxwich Head Limestone and Oystermouth Beds).

North Wales (Mold district)

Stage 1 calcretes are a uniquitous feature of late Asbian (Loggerheads Limestone) and early Brigantian (Cefn Mawr Limestone) exposure surfaces. These consist primarily of rhizocretions together with laminar calcretes and occasional calcrete mottle profiles. Palaeokarstic surfaces are typically subdued, although are usually present. Stage 3 laminar calcretes are an important feature of many surfaces.

North Wales (Anglesey)

Exposure surfaces in the Moelfre Limestone (late Asbian), Traeth Bychan Limestone (early Brigantian) and Red Wharf Cherty Limestone (late Brigantian) are frequently more mature than observed elsewhere. Rhizocretions (stage 1) are abundant. Laminar calcretes veneer both non-karsted (stage 1) and karsted surfaces (stage 3). Brecciation is also a characteristic feature of many exposure surfaces. Palaeokarstic surfaces are frequently well developed and pits commonly exhibit a sandstone fill.

Derbyshire

Exposure surfaces throughout the Asbian Bee Low Limestone exhibit a relatively uniform signature. Stage 1 calcretes include rhizocretions and occasional laminar calcretes and are an almost ubiquitous feature. Palaeokarstic surfaces are generally well developed. Stage 3 laminar calcretes locally veneer palaeokarstic surfaces.

West Cumbria

Exposure surfaces are well developed in both the late Asbian (White and 5th Limestones) and early Brigantian (Rough, Spotted, Potholes, Saccamina and Junceum Limestones). Stage 1 calcrete mottle profiles are a common although not ubiquitous feature. Rhizocretions and laminar calcretes are associated with some late Asbian surfaces. Palaeokarstic surfaces are generally present and are frequently well developed. Stage 3 laminar calcretes are occasionally present, although are restricted to Asbian exposure surfaces.

South Cumbria/ north Lancashire

Exposure surfaces are best developed in the upper Urswick Limestone (late Asbian). Stage 1 calcrete mottle profiles are a common, although not ubiquitous feature. Rhizocretions and very rare laminar calcretes are occasionally observed. Palaeokarstic surfaces are generally present and are frequently well developed. Stage 3 laminar calcretes locally veneer the palaeokarstic surfaces.

East Cumbria (Shap district)

Exposure surfaces are a prominent feature of the Asbian Knipe Scar Limestone. Stage 1 calcretes, generally in the form of calcrete mottle profiles, are widespread. Rhizocretions are locally present. Palaeokarstic surfaces are usually well developed. Stage 3 laminar calcretes locally veneer the palaeokarstic surfaces.

profiles. Following this line of reasoning, profiles d o m i n a t e d by rhizocretions are speculated to represent calcretes f o r m e d u n d e r m o r e arid climatic conditions, whilst those d o m i n a t e d by calcrete mottles are suggested to have formed u n d e r a slightly m o r e h u m i d climate. As such, early emergence of shallower platforms might

have favoured the formation of rhizocretions, whilst the emergence of slightly deeper platforms, subsequent to partial climatic amelioration, might have resulted in the f o r m a t i o n of calcrete mottles (Fig. 8). Such a model is supported by the phreatophytic nature of the vegetation cover, seasonal

LATE DINANTIAN CLIMATE ~ arid

1

/

i

High 5 stand /~ /\Sea-level . . . . . . \curve

Humid Low 3 Shallow Platform

stand

Deep platform

1 Stage 1 calcretisation

2 Karstification

Karstification

Stage 3 ¢alcretisation

5 Inferred climate = cyclic

Inferred climate = humid

Fig. 7. Diagram illustrating the influence of platform bathymetry on exposure surface development. Shallow platforms that became emergent early during the regressive phase and remained so until late in the transgressive phase are likely to exhibit all three stages of development, whereas relatively deep platforms that only became emergent for a short period of time toward lowstand are only likely to have undergone karstification. While exposure-surface records of shallow platforms may therefore suggest cyclical fluctuations in climatic regime, the signature produced in deeper platforms is more likely to suggest a uniformly humid climate. Such differences possibly account for much of the observed variation in exposure surface characteristics.

moisture deficiency causing the plants to sink deep root systems to tap the watertable. In contrast, emergence under a slightly more humid climate would have fulfilled two necessary requirements for mottle genesis: a medium for the transport of impurities into the underlying sediments, and conditions conducive to the release of finely particulate organic detritus

293

from decaying vegetation. In addition, cementation of the cyclothem top may also have been slower, enabling mottle development to continue as the watertable became more deeply entrenched within the sediment column in response to continued sea-level fall. As such, the generally deeper and more variable values for watertable depth obtained from the mottle horizons might be explained. Mixed calcrete profiles probably reflect the overprint of an early-formed rhizocretionary profile by calcrete mottles, these having developed in response to an increase in seasonal rainfall as the climate ameliorated and in association with a more depressed watertable. Such an origin is supported by the similarity in rooting depth exhibited by rhizocretions in mixed and rhizocretionary profiles, and by the calcrete mottles overlapping, although generally deeper, distribution. Mixed profiles probably therefore reflect the emergence of platforms intermediate in water-depth between that required for the formation of rhizocretionary and calcrete mottle profiles (Fig. 8). Cyclothem tops in which palaeokarst is present, but calcretes are absent, represent a further variation on this theme, emergence probably having occurred once the climate had ameliorated to such a degree that karstification was the first subaerial process (Fig. 8). From the change in style of cyclothemic sedimentation patterns (transition from peritidal cycles in the early Asbian to subtidal cycles in the late Asbian and sub-wavebase cycles in the early Brigantian), water depths are considered to have increased through the late Dinantian (Walkden 1987). As such, exposure surface character might also be expected to evolve through time. One possible example of this is seen in Cumbria. Here, rhizocretions and laminar calcretes, associated with many of the late Asbian exposure surfaces, are absent from early Brigantian surfaces, calcretes being solely of the mottle variety (Table 2). Whereas such a change could result from ameliorating climatic conditions, the associated increase in cycle thickness and reduction in cycle numbers indicates that a bathymetric control is more likely. Furthermore, associated palaeokarst is equally well developed in both the late Asbian and early Brigantian. Another good example is seen in the Gower Peninsula of South Wales. Here, there is a transition from sub-wavebase cycles with karstified tops in the Asbian, to an entirely marine succession in the Brigantian (Ramsay 1991). The Asbian succession probably records deposition on a relatively deep

294

S. VANSTONE \

" \N

I I

\\

Rhizocretionary profiles ".. *r q ~ x .i become case-hardened ~" *l~.),,~, "~ • N .(f~,~,4 through tune and capped by \ . p" I N N . a laminar calcrete veneer x \ "<'~J-" I \ \

".. b-

~)~.

Rhizocretionaryprofiles overprinted by calcrete mottle profiles as climate ameliorates

~."t]~ ~t'~,~, K,_~ < "P"-.,~ "~"5.

Rhizocretions :::~; Calcretemottles Lammar calcrete

~bf"~. ,.~,~ ,~ff~,,~ - ,¢ ~ r , ~

PROFILES LATER MODIFIED BY KARSTIFICATION

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Calcrete mottle profiles P = pedogenic and G = groundwater

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xx C~ % .

\ \

Late emergence of the platform leads to karstdication without prior calcretisation

.d

RHIZOCRETIONS SEMI-ARID

i

CALCRETE MOTTLES RAINFALL REGIME

KARSTIFICATION SUB-HUMID

Fig. 8. Model for catcrete development illustrating the possible interaction between evolving climatic conditions and the timing of platform emergence. Early emergence under more evaporitive climatic conditions is suggested to have been conducive to the formation of rhizocretions and laminar calcretes. In contrast, emergence under slightly more humid climatic conditions may have resulted in calcrete mottle profiles. Intermediate scenarios resulted in rhizocretionary profiles overprinted by calcrete mottles as the climate ameliorated. Late emergence resulted in karstification of the cyclothem-top sediments without prior calcretization. platform, late emergence accounting for the almost complete absence of calcretes and predominantly subdued palaeokarst. In contrast, the platform apparently became isolated from subsequent platform emergence events during the Brigantian due to increased water depth. The influence of bathymetry is also well illustrated in Anglesey, where exposure surfaces are amongst the most mature of any to be found in the British Isles. As well as exhibiting the full range of exposure surface phenomena found elsewhere, brecciation is also a common feature of the cyclothem tops (Davies 1991). Many surfaces also possibly exhibit evidence for continued development through successive cycles, palaeokarstic surfaces being stacked one on top of the other (Walkden & Davies 1983). Such complexity reflects the area's location toward the landward margin of the North Wales Shelf, the platform becoming emergent and remaining while other areas were covered by the sea.

The influence of subsidence As with platform bathymetry, subsidence may also have profoundly influenced the effect that climatic change had upon exposure surface development. Recognition of its signature within the exposure surface record is, however, somewhat more uncertain. For example, omission of stages of development, particularly karstification and the last stage of calcretization, owing to forced resubmergence would mirror the effects that might be observed due to lateral variation in exposure surface form. Considering the importance of tectonism in platform evolution (Berry 1984; Horbury 1987, 1989; Ramsay 1991), however, subsidence is likely to have been an important control on exposure surface development. Walkden (1987) also suggested that the progressive increase in bathymetry indicated by changes in the style of cyclothemic sedimentation might have resulted from an increased rate of subsidence through the

LATE DINANTIAN CLIMATE late Dinantian. As such, subsidence may have increased in importance in the Brigantian. During periods of tectonic quiescence, exposure-surface character was primarily controlled by the bathymetric configuration of the platform. In contrast, during intervals of heightened tectonic activity in which subsidence periodically outpaced eustatic sea-level change, exposuresurface character would have been modified by the omission of certain stages of development from the record owing to the premature submergence of the platform. In such instances, emergence of the platforms would be dependent upon whether regression or subsidence occurred at the faster rate. If the rate of subsidence were faster, emergence would not have occurred. In contrast, where sea-level fall outpaced subsidence, the platform would have become emergent, and the style of subaerial modification would have been dependent upon the timing of emergence. As the rate of sea-level fall tailed off toward lowstand, however, the relative increase in subsidence would have led to platform drowning, terminating exposuresurface development. As such, exposure surfaces influenced by subsidence-related drowning are likely to be characterized by immature to mature stage 1 calcretes, by subdued palaeokarst and by the absence of stage 3 calcretes.

The influence o f spatial climatic variation The influence of spatial climatic variation is particularly evident in exposure surfaces of the North Wales Shelf (Mold area), where welldeveloped calcretes indicate prolonged periods of emergence, yet relatively poorly developed palaeokarst suggests that the pluvial phase was less intense or of a shorter duration than elsewhere. Possible spatial climatic effects are also evident from comparison of exposure surfaces of a similar age from different platforms (Table 2). For example, the formation of calcrete mottle profiles in Cumbria and rhizocretion-dominated profiles in North Wales and Derbyshire might, in addition to possible bathymetric effects (Fig. 8), also reflect slightly more humid climatic conditions over the Cumbrian Platform. Spatial climatic variation is also suggested by the occurrence of tidal-flat evaporites in the late Asbian Meenymore Formation of County Leitrim, northwest Ireland (formerly the Aghagrannia Formation; West et al. 1968). These would not appear to have been subject to periodic karstification, and support a climate that tended towards aridity.

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Evaporitic peritidal carbonates have also been documented in the Brigantian of Derbyshire; however, early meteoric dissolution of the evaporites indicates that they probably formed within a semi-arid climate (Gutteridge 1989). Whereas the absence of calcretes within the Gower region of the South Wales Shelf could also be attributed to a relatively humid climate, the generally subdued nature of the associated palaeokarst, coupled with the presence of rhizocretions associated with exposure surfaces in the nearby Porthcawl area (Wilson et al. 1990), indicates that a bathymetric control is more likely. Spatial climatic variation, present over such a small geographical area, could only have arisen if precipitation and evaporation were near isovolumetric, slight regional variation in rainfall regime causing the finely balanced system to be weighted differentially toward calcretization or karstification. Such a situation may have arisen where platforms were positioned leeward of orographic barriers or laterally extensive expanses of land, predominantly dry winds producing below-threshold rainfall density.

Underlying controls of the palaeoclimatic cyclicity There is a growing wealth of evidence linking cyclothemic sedimentation in the late Dinantian with glacioeustatic sea-level oscillations. Not least of these is the identification of probable late Dinantian ice centres in western Argentina (Gonzalez 1990) and eastern Australia (Roberts 1981), evidence for southern hemisphere ice build-ups and a driving mechanism for cyclical oscillations in sea level. Such a mechanism would also best explain the widespread development of a cyclothemic pattern of sedimentation during this time interval (Ramsbottom 1973; Walkden 1987). The tectonic history of individual platforms was, however, also clearly influential in the signature that resulted (Horbury 1989; Ramsay 1991). Because orbital forcing is recognized as the pacemaker of Pleistocene glaciation and sealevel record (Imbrie & Imbrie 1980; Chappell & Shackleton 1986), evidence linking Milankovitch orbital perturbations with Carboniferous sea-level oscillations has been sought by many workers (Schwarzacher & Fischer 1982; Heckel 1990; Collier et al. 1990; Maynard & Leeder 1992). Studies here primarily focused on cyclothem periodicity, and Walsh power spectral analyses indicate periodicities of between 93ka and 147ka for late Carboniferous

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cyclothems in Britain and North America (Maynard & Leeder 1992). Such periodicities are comparable to the 100 ka eccentricity rhythm responsible for growth and retreat of the Pleistocene ice sheet, and provide further evidence for Milankovitch-induced glacioeustatic sea-level oscillations during the late Dinantian. Considering the likelihood of such a control and the apparent regularity with which the envisaged climatic fluctuations are associated with each sea-level oscillation, it seems likely that the climate was also driven by orbital perturbations in solar insolation. Such a link between climate and sedimentation was first suggested by Gray (quoted by Tucker 1985) who postulated that 5th order climatic cyclicity might be superimposed upon 4th order climate cycles responsible for cyclothemic sedimentation. An extrinsic control of the climate is also supported by the failure of an autocyclic response model to explain the observed fluctuations. Lowering of sea level would, for example, have greatly reduced the areal extent of the intervening seaways, and whereas these intervals were characterized by an increase in humidity, would more likely have resulted in increased desiccation of the emergent land surface. As such the climatic fluctuations are thought to reflect orbitally-forced perturbations in solar insolation with an equal or shorter periodicity to that forcing platform emergence. These include the 100ka eccentricity, 31ka obliquity and 17ka/20ka precession rhythms, periodicities for the latter two orbital parameters having been corrected for a shorter earth-moon distance during the Carboniferous (Collier et al. 1990). Each of the orbital parameters exerts its effect upon climate by producing changes in the intensity of the seasonal cycle (Berger 1988). Obliquity, although it causes significant changes in the receipt of solar radiation at high latitudes, has little effect within tropical and equatorial regions (Bradley 1985; Crowley & North 1991; Dawson 1992). Considering Britain's location close to the equator during the AsbianBrigantian (Scotese & McKerrow 1990), the palaeoclimatic cyclicity documented here is therefore unlikely to reflect variation in solar insolation due to changes in obliquity. In contrast, orbital eccentricity and precession of the equinoxes primarily exert their effect upon seasonality within low-latitude areas (Bradley 1985; Crowley & North 1991; Dawson 1992) and represent a more likely proposition. Variation in orbital eccentricity produces cyclical variation in

the relative intensity of the seasons (Bradley 1985; Dawson 1992). Seasonal contrast is least pronounced during periods of eccentricity minima; however, during periods of maximum eccentricity, the amount of solar radiation received between aphelion and perihelion may vary by as much as 30%, and these time intervals exhibit marked seasonality. Precession of the equinoxes also influences seasonality by determining whether maximum solar insolation, associated with perihelion, coincides with summer or winter. Perihelial summers and aphelial winters produce the greatest seasonal contrast (Bradley 1985; Dawson 1992). This dual control of the climate within lowlatitude areas is such that the precessional effect is modulated by variations in orbital eccentricity (Bradley 1985; Fischer & Bottjer 1991). During periods of eccentricity minima the seasonal timing of perihelion is inconsequential and the precessional effect is minimal. In contrast, toward eccentricity maxima seasonal timing is crucial and precession exerts its maximum effect. Such a relationship is particularly well illustrated by the African-Asian monsoonal record over the past 150 ka (Prell & Kutzbach 1987). Using a variety of palaeoclimatic records considered to be sensitive to fluctuations in monsoonal intensity, these authors showed that the interglacial record (eccentricity maxima) exhibits a strong correlation between precessional increases in solar radiation and intervals of monsoonal intensification, whereas the intervening glacial record displays a far less coherent history of fluctuations in monsoonal intensity. The formation of the exposure surfaces during periods of glacial lowstand, intervals that equate with eccentricity minima (Fischer & Bottjer 1991), suggests that precession would have been less effective in forcing climate than if they had formed during interglacials. The extent to which the precessional effect was modulated, however, depends upon the degree of ellipticity of the orbit during times of eccentricity minima and the timing of emergence of the platforms, eccentricity varying in a quasiperiodic fashion with 100ka and 400ka periodicities. As such, the precessional effect may have been sufficient to affect the climate during some periods of emergence, although not in others. Schwarzacher & Fischer's (1982) study of the Lower Asbian Glencar Limestone in northwest Ireland indicates, however, that the precessional effect might have been sufficiently coherent to exert an influence upon climate, even during intervals of eccentricity

LATE DINANTIAN CLIMATE minima. In this study, bedding was demonstrated to be cyclic, each cycle consisting on average of five (:t:2) component beds, an arrangement attributed to 100ka eccentricity rhythms with superimposed precessional effects. As such, precession-driven climatic fluctuations cannot be ruled out as the mechanism responsible for the cyclic record documented here. Alternatively, the observed cyclicity may reflect eccentricity-driven climatic change, with interglacial intervals, characterized by an evaporative climatic regime, alternating with glacial intervals during which the climate was distinctly humid. As with precession, the 100 ka eccentricity rhythm is also modulated by a longer frequency cycle, in this case the 400ka eccentricity perturbation, and as such the strength of the 100 ka climatic oscillations, associated with consecutive periods of emergence would have varied in intensity. For example, 100 ka fluctuations associated with 400ka troughs may have been insufficient to produce a significant shift in climatic regime and may account for some of the more mature palaeokarstic surfaces, humidity persisting throughout the lowstand interval. The interglacial-glacial climatic signature within equatorial regions during the Pleistocene was, however, of opposite polarity to that indicated for the Dinantian here, glacial intervals being characterized by a shift toward increased aridity (Dawson 1992). This is considered in part to reflect an increase in global surface albedo, which, coupled with lower average temperatures, reduced the overall amount of evaporation such that cloud buildup was greatly reduced (Manabe & Hahn 1977). In addition, equatorward shift of polar highpressure areas resulted in displacement of the trade-wind belts. This resulted in a corresponding displacement of monsoonal rainfall patterns in both Africa and South America, and precipitation occurred over much more restricted land areas (Dawson 1992). Whereas the general effect was to increase aridity within low-latitude areas, equatorial North Africa experienced a significant increase in precipitation regime (Prell & Kutzbach 1987; Rognon 1987; Dawson 1992), resulting in the formation of a series of pluvial lakes in the Middle East (Rognon 1976). These changes are believed to reflect a shift in locus of monsoonal precipitation. Wetter climatic conditions have also been suggested for some tropical areas during glacial intervals of the Neogene (Valentine 1984; Baron 1985). As such, regional shifts in monsoonal circulation resulting from displacement of the

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climatic belts during glacial intervals could provide an explanation for the climatic cyclicity in the late Dinantian. Monsoonal circulation has been speculated to have been a fundamental feature of the climate within low-latitude areas during the late Carboniferous for some time (Rowley et al. 1985; Besly 1987). However, the precise timing of its onset is somewhat uncertain. Rowley et al. (1985) have speculated that monsoonal circulation may have been an important feature of the northern margin of Gondwanaland from the Vis~an onwards, whilst Wright (1990a) has suggested that the existence of monsoonal climatic conditions could provide an explanation for pronounced equatorial aridity necessary for the formation of calcretes and evaporites throughout the mid-Dinantian (late Tournaisian-Arundian). Wright (1990a) and Wright et al. (1991) have also conjectured that periodic fluctuation in the pedologic signature of exposure surfaces throughout this time period (calcretes/evaporites v. palaeokarst) might reflect shifts in the seasonal moisture budget resulting from changes in monsoonal circulation. Palaeogeographical reconstructions of Scotese & McKerrow (1990) also indicate that the continental configuration during the Dinantian was conducive to monsoonal circulation (Valdes pers. comm. 1993). As such, both eccentricity-driven shifts in the locus of monsoonal precipitation and precession-driven variations in monsoonal intensity could provide explanations for the observed palaeoclimatic cyclicity. A precessional origin is possibly supported by evidence for multiple fluctuations in climate recorded at exposure surfaces in Anglesey (Walkden & Davies 1983). These fluctuations were invoked to explain the polyphase development of channel sandstones and associated palaeokarsts, wetter intervals, during which the sands were deposited and during which karstification of the underlying limestones took place, alternating with drier periods during which the sands underwent lithification. These wet-dry cycles, which in the Trwyn Dwlban palaeokarst number as many as four, are recorded at individual exposure surfaces and provide at least cursory evidence for the precessional origin of the cyclicity documented here.

Other evidence for palaeoclimatic cyclicity Although fluctuations in climatic regime during the Carboniferous are becoming increasingly

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widely documented, most of those exhibiting well-defined cyclicity would appear to be of a shorter duration than those documented here, e.g. Loftus (1985) and Elrick et al. (1991). Alternations between ferruginous and calcrete palaeosols within the Etruria, Newcastle and Keete Formations (Westphatian C/D-Stephanian) of central England (Besly 1987) would, however, appear to reflect more prolonged fluctuations in climate and may be of a similar periodicity to that documented here. Similar climate cycles are also apparent from the stratigraphic distribution of climatically-sensitive rocks within the mid to late Pennsylvanian succession of the Appalachian Basin (Cecil 1990). Here, coals/leached palaeosols, considered to represent deposition under a humid climate, alternate with lacustrine/pedogenic carbonates (Vertisols/Aridisols) formed under semi-arid climatic conditions. Comparable shifts in pedological signature have also been observed within the mid-Dinantian successions of South Wales (Wright 1988, 1990a; Wright et al. 1991) and southwest England (Vanstone 1991): palaeokarst coupled with rare podsolic soils formed during humid climatic intervals, alternating with calcretes and evaporites that developed under semi-arid and arid climatic regimes. Van der Zwan et al. (1985) also illustrated fluctuations in the composition of climaticallysensitive microfloral assemblages within the late Holkerian-Asbian of northwest Canada, the Brigantian of northwest Europe and the early Namurian of Oklahoma in the US, and these may represent an incomplete record of those fluctuations documented here. As such, climatic cyclicity would appear to have persisted throughout much of the Dinantian and Late Carboniferous time interval.

Duration of emergence Although the immature nature of the calcrete, palaeosols and palaeokarst indicate that platform emergence was short-lived, uncertainty concerning the driving force of the climatic fluctuations means that no definite figure can be suggested at present. Were precession the underlying control, the three stages of exposure surface development would represent a single minimum to minimum excursion in precessionforced insolation, and their cumulative duration would amount to some 20ka. Such precise quantification is not possible were eccentricity the driving mechanism, the proportion of the 100ka cycle represented by emergence (as

opposed to carbonate deposition) being difficult to quantify. Estimates based on comparison of palaeokarst relief/calcrete crust thickness and the rate of formation of such features in the Holocene and Pleistocene are considered to be fraught with difficulties (Wright 1990b) and likely to give erroneous results. Using such an approach, it would also be difficult to quantify the likely time lags between a change in climate and discernible changes in exposure-surface character. Nevertheless, it is the author's opinion that the average duration of emergence is unlikely to have been more than a few tens of thousands of years. Exceptions to this general rule include instances where platforms remained emergent through, more than one cycle. This is likely to have been the case in the more proximal areas of land-attached platforms. The Anglesey area of the North Wales Shelf, being characterized by more mature calcrete and stacked palaeokarst, is probably one such example.

Conclusions Three distinct stages of exposure surface development are recognized. An initial phase of calcretization (stage 1) led to the formation of rhizocretions and calcrete mottles within the newly emergent sediments, and to laminar calcrete veneers following cementation of the cyclothem top. These were systematically removed and modified by the subsequent phase of karstification (stage 2), a process that was enhanced by the accumulation of a mineral soil. A second phase of calcretization (stage 3), resulting in the formation of laminar calcrete veneers of the karstified cyclothem top, was generally much less widespread than stage 1. These three stages of development represent a response to cyclical shifts in climatic regime, the transition from semi-arid to humid to semi-arid climatic conditions accompanying each sea-level fall/rise cycle. The influence that this climatic cyclicity had upon exposure surface development was, however, very much dependent upon a number of overriding controls, and each platform exhibits its own individual record of what is essentially an idealized sequence of events. Particularly influential was platform bathymetry, which modulated the climatic effect by determining the point in the climate cycle when the platform both became emergent and was drowned. Subsidence is also likely to have been important, intervals of heightened tectonic activity resulting in omission of the later stages of development due to the premature

LATE D I N A N T I A N C L I M A T E resubmergence of the platform. A further factor influencing the style of subaerial modification was spatial climatic variation. The climatic cyclicity is considered to reflect either eccentricity-driven shifts in the locus of m o n s o o n a l precipitation or precession-driven variations in m o n s o o n a l intensity. If precessional in origin, the three stages w o u l d represent a single m i n i m u m to m i n i m u m excursion some 20 ka in duration. Such precise quantification is not possible were eccentricity the driving force; however, the i m m a t u r e nature of the calcretes, palaeosols and palaeokarst indicates that the d u r a t i o n of emergence is unlikely to have been m o r e than a few tens of thousands of years. This research formed part of my PhD undertaken at the Postgraduate Research Institute for Sedimentology, University of Reading. I am particularly grateful to my supervisor, P. Wright, for his help and enthusiastic support throughout. I am also indebted to R. Hill for her patience and encouragement. The various quarrying concerns are thanked for their cooperation in providing access to quarry exposures. J. Davies (BGS, Aberystwyth) and A. Horbury (Cambridge Carbonates) are thanked for many helpful discussions. The author is also grateful to P. Gutteridge (Cambridge Carbonates) and I. Somerville (University College, Dublin) for constructive criticism at review. The Natural Environmental Research Council and BP are acknowledged for their financial support. Thanks are also extended to Integrated Exploration and Development Services for use of their computer facilities.

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LATE D I N A N T I A N C L I M A T E - - 1 9 7 9 b . A cyclicity in the early Brigantian (lower D2) limestones east of the Clwydian Range, North Wales and its use in correlation. Geological Journal, 14, 69-87. SWEETING, M. M. 1972. Karst Landforms. Macmillan Press, London. TRUDGILL, S. T. 1985. Limestone Geomorphology. Longman, London. TUCKER, M. E. 1985. Shallow-marine carbonate facies and facies models. In: BRENCHLEY, P. J. & WILLIAMS, B. P. J. (eds) Sedimentology: Recent Developments and Applied Aspects. Geological Society of London Special Publication 18, 139-16 I. VALENTINE, J. W. 1984. Neogene marine climate trends: implications for biogeography and evolution of the shallow sea biota. Geology, 12, 647-650. VAN DER ZWAN, C. J., BOULTER, M. C. & HUBBARD, R. N. L. B. 1985. Climatic change during the Lower Carboniferous in Euramerica, based on multivariate statistical analysis of palynological data. Palaeogeography, Palaeoclimatology, Palaeoecology, 52, 1-20. VANSTONE, S. D. 1991. Early Carboniferous (Mississippian) palaeosols from southwest Britain: influence of climatic change on soil development. Journal of Sedimentary Petrology, 61,445-457. - - 1 9 9 3 . Soil Development and the Use of Palaeosols in the Assessment of Palaeoclimate: a Case Study from the Late Dinantian of Britain and Newfoundland. PhD Thesis, University of Reading. WALKDEN, G. M. 1972. The mineralogy and origin of interbedded clay wayboards in the Lower Carboniferous of the Derbyshire Dome. Geological Journal, 8, 143-160. 1974. Palaeokarstic surfaces in Upper Vis6an (Carboniferous) limestones of the Derbyshire Block, England. Journal of Sedimentary Petrology, 44, 1232-1247. - - 1 9 8 7 . Sedimentary and diagenetic styles in late Dinantian carbonates of Britain. In: MILLER, J., ADAMS, A. E. & WRIGHT, V. P. (eds) European Dinantian Environments. John Willey, Chichester, 131-155. & DAVIES, J. R. 1983. Polyphase erosion of subaerial omission surfaces in the late Dinantian of Anglesey, North Wales. Sedimentology, 30, 861-878.

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WEAVER, C. H. 1989. Clays, Muds and Shales. Developments in Sedimentology No. 44. Elsevier, Amsterdam, 103-188. WEBB, G. E. 1994. Palaeokarst, palaeosol and rocky shore deposits of the MississippianPennsylvanian unconformity, northwestern Arkansas. Geological Society of America Bulletin, 106, 634-648. WEST, I. M., BRANDON, A. & SMITH, M. 1968. A tidal flat evaporitic facies in the Vis6an of Ireland. Journal of Sedimentary Petrology, 38, 1079-1093. WHITE, W. B. 1984. Rate processes: chemical kinetics and karst landform development. In: LAFLEUR, R. G. (ed.) Groundwater as a Geomorphic Agent. Allen and Unwin, London, 227-248. WILSON, D., DAVIES, J. R., FLETCHER, C. J. N. & SMITH, M. 1990. Geology of the South Wales Coalfield, Part VL the Country around Bridgend. Memoir of the British Geological Survey, Sheets 261 and 262. WRIGHT, V. P. 1988. Palaeokarsts and palaeosols as indicators of palaeoclimate and porosity evolution: a case study from the Carboniferous of South Wales. In: JAMES, N. P. & CHOQUETTE, P. W. (eds) Palaeokarst. Springer-Verlag, New York, 329-341. 1990a. Equatorial aridity and climatic oscillations during the early Carboniferous, southern Britain. Journal of the Geological Society, London, 147, 359-363. 1990b. Estimating rates of calcrete formation and sediment accretion in ancient alluvial deposits. Geological Magazine, 127, 273-276. 1994. Palaeosols in shallow marine carbonate sequences. Earth Science Reviews, 35, 367-395. - & VANSTONE, S. D. 1991. Assessing the carbon dioxide content of ancient atmospheres using palaeocalcretes: theoretical and empirical constraints. Journal of the Geological Society, London, 148, 945-947. & ROBINSON, D. R. 1991. Ferrolysis in Arundian alluvial palaeosols: evidence of a shift in the early Carboniferous monsoonal system. Journal of the Geological Society, London, 148, 9-12.

Reconstruction of a lost carbonate platform on the shelf of Fennosarmatia: evidence from Vis~an polymictic debrites, Holy Cross Mountains, Poland ZDZISLAW

B E L K A 1, S T A N I S L A W

JANINA

SKOMPSKI 2 &

SOBON-PODGORSKA

3

1Geologisch-Paldiontologisches Institut, Universitdit Tfibingen, Sigwartstrasse 10, 72076 Tfibingen, Germany 2 Institute of Geology, Warsaw University, Al. Zwirki i Wigury 93, 02-089 Warszawa, Poland 3 Polish Geological Institute, Upper Silesian Branch, ul. Krolowej Jadwigi 1, 41-200 Sosnowiec, Poland

Abstract: The sedimentary sequence of a totally eroded carbonate platform on the shelf of Fennosarmatia in Poland has been reconstructed based on the analysis of detrital material derived from the platform and deposited in an adjacent basin. The database was taken from a polymictic debrite unit intercalated in the Vis6an basinal succession exposed in the southwestern Holy Cross Mountains. The unit represents a gravity-flow deposit and contains carbonate clasts ranging from Frasnian to Vis6an in age. They provide evidence for a Frasnian carbonate platform located south of the Holy Cross area that drowned during Famennian and Tournaisian times but subsequently, during the early Vis6an, started to recover. This reversed trend is interpreted to have resulted from the combined effect of eustatic sea level fall and tectonic uplift. The geographical extent of the inferred carbonate platform system, named here as the Nida Platform, cannot be precisely outlined, but most probably corresponded with the Jedrzejow High, an elevated fragment of Precambrian basement.

A tropical epicontinental sea covered much of southwestern Fennosarmatia during Middle Devonian to Early Carboniferous times. The central part of Poland, which lies outside the Variscan orogenic belt, constituted an inner fragment of this shelf area. During the Middle Devonian the shelf was the site of a vast, shallow-water carbonate platform with a very uniform facies pattern. During the Frasnian the platform disintegrated into a mosaic of blocks affected by differential subsidence. Although each of these fragments underwent different depositional evolution, a general stepwise drowning can be observed. It created a very pronounced shelf bathymetry with isolated, relatively small carbonate platforms and seamounts separated by small intracratonic basins. In the area of the Holy Cross Mountains all elevated blocks had already subsided before the latest Famennian (Szulczewski 1977; Szulczewski et al. 1996). In the southern part of Poland, however, the platform-to-basin topography controlled the depositional evolution of the shelf until the late Vis~an (Belka 1987; Paszkowski 1988). Detailed palaeogeographical reconstruction of the shelf topography is difficult

because of syn- and post-Variscan erosion which removed much of the Devonian and Lower Carboniferous succession, and especially that of the elevated carbonate platforms. The only record of the history of such platforms are platform-derived sediment gravity-flows deposited in the basins. This study is an attempt to reconstruct the sedimentary sequence of a totally eroded carbonate platform based on the analysis of detrital material which was derived from this platform and transported into the adjacent basin. The database comes from the polymictic debrites intercalated in the Lower Carboniferous basinal succession exposed in the Ostrowka quarry near Galezice, in the southwestern Holy Cross Mountains (Fig. 1).

Geological background Lower Carboniferous deposits terminate the Variscan succession of the Holy Cross Mountains. Owing to post-Variscan erosion they are preserved only locally in the axial parts of synclines. Their cumulative thickness reaches

From STROGEN, P., SOMERVILLE, I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 315-329.

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Z. BELKA E T A L .

Fig. 1. Simplified geological map of the Galezice area with location of the Ostrowka quarry (modified from Zakowa 1976). Inset shows the outcrops of Palaeozoic rocks of the Holy Cross Mountains and the location of the study area. Arrows on the main figure show the studied sections where the debrite unit is exposed. over 600 m in the central part of the Holy Cross Mountains area, but decreases to about 200 m in the southwestern part of the Palaeozoic exposures (Fig. 1). This difference, however, is not a secondary erosional feature but appears to have resulted from a higher subsidence rate in the central segment of the Holy Cross Mountains. The Lower Carboniferous succession comprises predominantly argillaceous and other clastic sediments (for review see Zakowa 1980, 1981), deposited in a basinal realm that developed in the Holy Cross area, after drowning of the Devonian carbonate platform (Szulczewski et al. 1996). The Tournaisian succession, which is locally separated from the Devonian by a significant stratigraphic gap, usually starts with pelagic clays interbedded with thin mudstones to wackestones and tephra horizons. Overlying these are black siliceous shales, known as the Zareby Beds, that represent a typical starvedbasin facies accumulated below the level of benthic carbonate production. They commonly contain phosphatic nodules and extremely abundant radiolarian fauna. Shelly calcareous fauna and evidence of bioturbation are totally absent. The onset of this organic-rich facies in the Holy Cross Mountains is clearly diachronous over a distance of several kilometres (Zakowa 1980; Zakowa & Paszkowski 1989) owing to differential subsidence of the blocks forming the basin floor. In the Galezice area, the Zareby Beds are of early to middle Vis6an age

(Szulczewski et al. 1996). During late Vis6an time, a radical change occurred and clastic input began to affect the sedimentary regime in the basin. A thick series of clastic rocks, known as the Lechowek Beds, was deposited, which is dominated by sandy shales with sandstone and siltstone intercalations. According to Zakowa (1981), the Lechowek Beds show an upward regressive trend indicated by the gradual disappearance of marine fauna and records of land plant remains at the top. The only significant exception to the facies development described above occurs in the succession exposed near Galezice village (Fig. 1). Here, shallow-water platform carbonates of Frasnian age are overlain at an angular unconformity by Fammenian pelagic cephalopod limestones (Units A, B in Fig.2), in turn overlain by pelagic carbonates of Tournaisian age and the radiolarian shales of the Zareby Beds (Units C, D). Most significantly, these radiolarian shales are here separated from the Lechowek Beds (Unit F) by a coarse-grained carbonate unit (Unit E). This carbonate unit forms three lenticular bodies, dipping to the north, several hundred metres long and a maximum of 25 m in thickness. At the base is a polymictic breccia (Figs 3, 4) overlain by a package of poorly-sorted, coarse-grained crinoidal packstone to rudstone layers, each 15-100 cm thick (Fig. 5). All these layers contain extremely abundant and diverse shallow-water benthic fauna (for review see Belka & Skompski 1988)

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Fig. 2. Cross section through the ledge of the Ostrowka quarry (modified from Belka & Skompski 1988). The stratigraphic information is simplified; for detailed conodont-based stratigraphy see Szulczewski et al. (1995). A,,peritidal carbonates (Kowala Formation); B, condensed crinoidai-cephalopod limestone; C, clays with limestone intercalations; D, black radiolarian shales with cherts (Zareby Beds); E, breccia and crinoidal limestones (carbonate gravity flows); F, shales with intercalations of siitstones and sandstones (Lechowek Beds).

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Z. BELKA E T AL. derived from different ecological niches. In the upper portion of the unit, two thin cephalopod wackestone layers are sandwiched between the coarse-grained packstones. The shallow-water character of the abundant biota seems to be in conflict with the position of the carbonate unit within a basinal sequence. Detailed sedimentological investigations (Belka & Skompski 1988), however, have shown that the unit represents a range of gravity-flow deposits forming three distinct lobes of a submarine deep-water fan. The detrital material, along with the shallowwater benthic fauna, was transported from an adjacent carbonate platform and deposited in the deeper, lower-slope environment. This is indicated by the mixing of fauna originating from different ecological niches, and by many sedimentary features, such as the anatomy of the carbonate lenses, grain-supported texture, chaotic clast arrangement, preferred orientation of elongated bioclasts (rugose corals, crinoid stems), and the presence of reworked fragments from the substrate. The thin interbeds of cephalopod wackestone (Fig. 5), however, appear to represent basin plain deposits.

Fig. 3. Lower part of the Upper Vis6an carbonate gravity flows exposed near Jazwiny Hill. The basal debrite lens from which the clasts were sampled for petrographical examination is arrowed.

Debrite characteristics The debrite investigated in this study is a polymictic limestone breccia thought to consti-

Fig. 4. Polished slab of the studied debrite showing the polymictic clast composition and the chaotic clast arrangement. The clasts are predominantly Frasnian and Vis6an carbonates associated with fragments of phosphatic nodules and rugose corals. Scale is 2 cm long.

CARBONATE PLATFORM RECONSTRUCTION, POLAND

[Todowa Hill I

exposed (Figs 3, 5). The base of the breccia is erosive and sharply defined. In many places injection structures resulting from loading of the underlying radiolarian shales are observed. The top of the bed is relatively planar, with only rare large projecting clasts giving a slight hummocky appearance. This clast-supported breccia is poorly-sorted and exhibits a chaotic clast arrangement (Fig. 4). Clasts are mostly angular, and range in size from gravel to metre-sized blocks. Lithoclasts of various exotic limestones are dominant, mixed together with coarsegrained Vis6an skeletal debris. The largest blocks are represented by colonies of Vis6an tabulate and rugose corals, such as Syringopora and Lithostrotion (Fig. 5). Associated with these are slope-derived fragments of radiolarian shales, usually squashed between limestone clasts, and phosphatic nodules up to 5 cm in diameter. The granular matrix, composed of fine-grained carbonate particles, comprises generally from 10 to 20% by volume of the deposit. Clast size distribution has not been quantitatively studied, but comparison of the breccia fabric in the Todowa Hill and Jazwiny Hill sections revealed a higher abundance of larger blocks in the former outcrop (Fig. 5), suggesting that it lies in the more axial part of the lobe.

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Fig. 5. Measured sections through the carbonate gravity flow deposits at Todowa and Jazwiny Hills (modified from Belka & Skompski 1988). a, black radiolarian shales (at the base of the sections) and sandy shales (at the top); b, carbonate breccia (debrite); c, large colonies of rugose (Lithostrotion) and tabulate (Syringopora) corals; d, coarse-grained crinoidal packstone/rudstone (mean grain size < 10 mm); e, coarse-grained crinoidal packstone/ rudstone (mean grain size >10 mm); f, graded bioclastic limestone; g, cephalopod wackestone. There is an exposure gap of c. 3 m near the base of the Todowa Hill section. tute part of the western lobe of a submarine fan. At present, it is best studied in the middle part of the face of Ostrowka quarry, between Todowa and Jazwiny Hills, where the lobe of the fan is cut and displaced by a fault for a distance of about 200m (Fig. 1). This lenticular, massive layer attains about 5 m in thickness near Todowa Hill but wedges out laterally. In the section at Jazwiny Hill its marginal wedge-shaped fragment is

In order to reconstruct the sedimentary sequence of the source area from which the exotic clasts were derived, we have examined 110 lithoclasts in terms of biostratigraphy and microfacies. The clasts were systematically sampled in the Jazwiny Hill section. Only clasts larger than 5 cm in diameter were sampled because of the requirement for large samples for conodont studies. Conodont zonation applied in this paper follows schemes proposed for the Upper Devonian by Ziegler & Sandberg (1984), and for the Tournaisian by Sandberg et al. (1978) and Belka (1985). Foraminifers were also used to determine the age of the clasts.The age assignment of the foraminiferal assemblages is based on the zonation of Conil et al. (1990). Petrographic studies involved the examination of clasts in thin section and as polished slabs. The carbonate classification used in this study is that of Dunham (1962), expanded by Embry & Klovan (1971).

Results Stratigraphical and petrographical examination of clasts reveals a polymictic composition for the

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carbonate material. The clasts range from Frasnian to Vis6an in age, being surprisingly similar to the age range of the section exposed in the Ostrowka quarry (Fig. 2). Only four samples did not contain any microfossils. Conodonts were found in 28 clasts and foraminifers in 76. Twelve samples provided only problematical calcareous microfossils (e.g. calcisphaerids), which are usually referred to calcareous algae or foraminifers and are not of stratigraphic importance. We interpret these clasts as being of Frasnian age based on their lithological features. Determination of the precise age of the Vis6an foraminiferal assemblages is hindered in many cases because of their isolated character and, therefore, they cannot be placed within a continuous biotic succession. The stratigraphic composition of the investigated clast material is shown in Fig. 6. The clasts represent a wide spectrum of lithologies, dominated primarily by shallow-water platform carbonates. The range ofclasts is described below in stratigraphic order.

Frasnian clasts The Frasnian clasts (15 samples) constitute a very uniform group, clearly recognizable due to the creamy coloration of the limestones and their very fine-grained lithology. These are peloidal, calcisphaerid wackestones to packstones with numerous foraminifers (Fig. 7a), mostly representatives of Diplosphaerina, Radiosphaera and Pachysphaerina. The rocks typically show fenestral fabric and are interbedded with algal laminites. Sporadic micritized fragments of stromatoporoids occur, floating in the peloidcalcisphaerid groundmass. This lithotype is characteristic of Devonian shallow-water environments and is usually interpreted as having been deposited in restricted lagoons, between stromatoporoid-coral mounds (Larsen et al. 1988). It corresponds to type microfacies 9 in the scheme of Preat & Mamet (1989). In the Holy Cross area, calcisphaerid limestones are widespread in the Givetian/Frasnian Kowala Formation (e.g. Racki 1993; Racki & SobonPodgorska 1993). The Frasnian age of clasts is documented by the occurrence of multilocular foraminifers Eogeinitzina devoniea rara (Fig. 7b), Eonodosaria rausereae, Tikhinella measpis, Tikhinella fringae and Nanicella ex gr. galloway. According to Kalvoda (1986) and Zadorozhnyj (1987), the assemblage is indicative of the middle-upper Frasnian. This is supported by the presence of the ostracode Bairdiocypris samsonowiczi, which

Tot

2%) lian (14%) unknown (4%)

Vis6an (66%) Fig. 6. Diagram showing the average stratigraphic composition of the clast material within the debrites. is known to occur in the upper Frasnian of the Holy Cross Mountains (e.g. Malec & Racki 1993). However, stratigraphically important forms have been found in three samples only. As mentioned above, remaining clasts are attributed to the Frasnian, but it is important to note that the same facies had already started to develop in the Holy Cross Mountains area during Middle Devonian time; thus the Givetian age of these clasts cannot be excluded.

Famennian clasts Only two Famennian clasts were recognized in the material investigated. Both limestones are bioclastic wackestone rich in comminuted skeletal debris (Fig. 7c). The biotic components include primarily fragments of trilobites, crinoids, ostracodes and juvenile goniatites. The conodont fauna is very rich. While in the first sample crinoids and trilobites are present in equal proportions, in the second one trilobite fragments clearly predominate and are additionally associated with some well-preserved bryozoans. The matrix is micritic or microsparitic. Both samples resemble pelagic limestones widely developed in the Devonian of Variscan Europe (Tucker 1974; Franke & Walliser 1983). The characteristic fossil content, the micritic matrix, and the lack of calcareous algae and micritization indicate a depositional environment below the photic zone and storm wave-base. The high content of conodonts suggests a relatively low rate of sedimentation. Conodont fauna recovered from the first sample is diagnostic for the Middle expansa Zone. The second sample contained conodonts represented by long-ranging taxa that fall within the interval from the Middle expansa Zone to the Lower praesulcata Zone.

C A R B O N A T E PLATFORM RECONSTRUCTION, P O L A N D

321

Fig. 7. (a) Thin section of typical Frasnian calcisphaerid-peloidal wackestone. (b) Frasnian foraminifer Eogeinitzina devonica rata Lipina. (e) Photomicrograph showing characteristic biotic components of the Famennian pelagic limestones - trilobites, crinoids and ostracodes. (d) Infilling of Tournaisian neptunian dyke with numerous broken calcite crystals (white). (e) Example of Vis+an crinoidal packstone with micritic coatings on bioclasts. Scale bars are 1 ram, except on (e) where bar is 0.1 mm.

Tournaisian clasts

Two groups of clasts are recognized a m o n g the Tournaisian samples. These are mudstones (six samples) and fragments of limestone breccia (nine samples).

Mudstones. These are very distinctive yellow to buff weathering rocks. A l t h o u g h particular clasts may represent different stratigraphic ages, the mudstones display a uniform, m o n o tonous lithology with poor fossil content. Only single shells of e n t o m o z o i d ostracodes and rare

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fragments of crinoids and trilobites are observed. Conodonts are also not very common. The biota and the micritic lithology indicate a deep marine, pelagic depositional environment, certainly deeper than the conditions under which the upper Famennian wackestones accumulated. Conodont data show that the deposition of mudstones started not later than delicatus Zone time, possibly earlier, during the sandbergi Zone, and persisted up to the anchoralis Zone.

Limestone breccias. Most of the samples belonging to this group are composed of angular clasts coated with and set in a peculiar carbonate matrix, which includes a large quantity of dispersed calcite crystals (Fig. 7d). The calcite crystals are broken, poorly-sorted, and usually sharp-edged, but in some cases are diagenetically corroded. Single broken crinoids and small micritic intraclasts occur as accessory components. Conodonts are very common, and can often be observed within the matrix in thin sections. All samples provided mixed conodont faunas. The presence of conodonts and crinoids points to marine conditions. There is no doubt that the samples represent the infill of neptunian dykes within the Frasnian host rocks. Of crucial importance for the understanding of the origin of these dykes are the dispersed calcite crystals. Certainly, they are neither syngenetic nor diagenetic. It is probable that the dykes developed along tectonic fissures previously affected by calcite mineralization. The next episode of tectonic activity must have caused re-opening of these fissures. This brecciated the host-rock and caused subsequent infilling of the fissures with lime mud mixed together with broken crystals from the calcite veins. The youngest conodont elements found in the dykes are diagnostic for the anchoralis Zone. This is evidently the age of the majority, if not all of the dykes. Only in one sample were the youngest conodonts recovered indicative of the interval from the crenulata Zone to the delicatus Zone. The stratigraphic admixture in the dykes displays a characteristic pattern. Older conodonts extracted from the clasts are of Famennian (from the Lower marginifera Zone up to Middle praesulcata Zone) and middle Tournaisian (from the sandbergi Zone up to the cuneiformis Zone) age. Moreover, based on lithology and foraminiferal fauna we have identified some clasts of Frasnian limestones. There is, however, no evidence of early Tournaisian clast material in the fissures.

VisOan clasts The Vis6an clasts constitute two-thirds of all samples investigated (Fig. 6). They can be classified into three lithological groups: biolithites, peloidal limestones, and bioclastic limestones. All Vis6an limestone clasts are characterized by the presence of diverse foraminiferal fauna dominated by representatives of the genera

Endothyra, Howchinia, Valvulinella, Archaediscus and Tetrataxis. The stratigraphic examination of the foraminiferal assemblages revealed that the three lithofacies were deposited during the Cf5 and Cf6 foraminiferal zones of the middle and late Vis6an. The youngest samples bear foraminifers diagnostic of the Cf6-,/Zone. The abundance of foraminifers was the reason why only a few clasts were dissolved in search for conodonts, which offer relatively poor stratigraphic resolution within the Vis6an. The conodont data obtained are not very rigorous, but they confirm the middle and late Vis6an age of the clasts. Conodonts are indicative either of the bilineatus Zone or of the pre-bilineatus interval.

Biolithites. This group (27 samples) includes carbonates with textures from bafflestone to floatstone. They are characterized by a relatively low diversity of biotic components. Moreover, the fauna is sparse and primarily represented by rugose corals (Lithostrotionidae) and auloporid tabulates (e.g. Sinopora polonica, Multithecopora sp.; Figs 8d, e). Other fossils such as brachiopods, bryozoans, heterocorals and crinoids do occur, but they are never volumetrically important. Generally the bioclasts are surrounded by cryptalgal coatings with encrusting foraminifers (Figs 8f, g). Common features within the rock are spongiform matrix (Pratt 1982) and numerous, irregular cavities showing geopetal structures. The cavities are usually lined by slightly yellowish, isopachous cloudy cement and filled by micritic or pelmicritic internal sediment, often neomorphosed, with blocky calcite in the top part. Sporadically, bioclastic material is present as cavity infill. In some neomorphic pseudospar areas, 'ghosts' of cryptalgal wrinkled lamination can be observed. Peloidal limestones. These rocks (33 samples) display mostly wackestone texture and are somewhat similar to the biolithites because of the peloidal character of the matrix. A highly varied spectrum of bioclasts, dominated by fenestrate bryozoan hash and crinoid debris, is present. It also includes fragments of gastropods, brachiopods and trilobites. A number of grains are

CARBONATE PLATFORM RECONSTRUCTION, POLAND micritized and encrusted. Foraminifers occur very commonly, especially encrustations of Pseudolituotuba gravata (Fig. 8c). Characteristic is a high abundance of girvanellids. In addition, single specimens of red algae and globochaetids are also observed (Figs 8a, b). Large pellets and oncoids are relatively numerous. The coating of the oncoids is usually composed of sparry, irregular laminae, asymmetrically arranged around an ellipsoidal nucleus. The laminae are intercalated with twisted bundles of Girvanella filaments or encrusting bryozoans, which form the earliest stage of the coating process. Bioclastic limestones. This group is represented

by 14 clasts whose detrital character differs distinctly from the Visran carbonates described above. These rocks show a packstone to grainstone texture and are composed of very diversified bioclasts (Fig. 7e). Allochems include mainly crinoids and bryozoan hash, but fragments of coral colonies, heterocorals and brachiopods are also very common. Thick micritic envelopes are omnipresent on bioclasts. The limestones are normally non-graded and very poorly-sorted. The similarity of the assemblages of biotic elements present in the different Visran carbonates points to their close relationship. It is, therefore, not unlikely that the three recognized lithotypes occupied a relatively narrow facies belt, presumably along the margin of the carbonate platform. The varied biota suggests rather shallow-water deposition, open circulation and normal salinity. Fragments of biolithites document the presence of mudsupported buildups in this region. The clasts probably represent the 'core' facies of such buildups, which must have been similar in lithology, biota and depositional setting to the buildups described from different parts of the early Carboniferous shelf in Britain and Ireland (e.g. Adams 1984; Jameson 1987; Somerville et al. 1992; Pickard 1992). The peloidal fabric, probably due to microbial activity, is the most common feature in all these buildups. The role of auloporids in the formation of mud-supported buildups is still a matter of debate. Fagerstrom (1987) concluded that auloporid corals played only a subordinate role, acting as baffles or traps for carbonate mud. Fliagel & Krainer (1992), however, have shown that although these tabulates did not build an organic framework, they were of major importance in controlling mound development. The second recognized lithotype, peloidal limestones, may have originated in a relatively wide zone of

323

the back-reef environment. This corresponds to the SMF Type 22 (standard microfacies after FliJgel 1982). The micrite envelopes point to fairly slow sedimentation rates. In areas of more dynamic conditions, however, such as intermound channels or high-energy shoals, the third lithofacies was deposited. This is, first of all, indicated by the grain-supported fabric of the bioclastic limestones. Bowman (1979) has described a similar association of small mounds and surrounding detrital carbonates from the Upper Carboniferous of the Cantabrian Mountains. Besides the three facies identified within the clast material, large coral heads (over 1 m high) flourished locally on the Visran carbonate platform. They were formed by both rugose and tabulate coral colonies which can be found in the breccia of the Todowa Hill section (cf. Nowinski 1976).

Palaeogeography The Vis~an carbonate gravity-flow deposits exposed in the Ostrowka quarry are interpreted as representing part of a submarine, deep-water channelized slope fan (Belka & Skompski 1988). Based on the orientation of crinoid stems and coral fragments in some of the grain-flow beds, the transport direction appears to be towards the north (Belka & Skompski 1988). Thus it is evident that the source area from which the clast material of the investigated debrite was derived was located south of the Holy Cross Mountains. To the south, however, in the Nida Trough, Palaeozoic rocks lie at considerable depth covered by thick Permian, Mesozoic and Tertiary formations. Numerous deep boreholes have provided evidence of the overall structure of the sub-Permian, so that the distribution of the Carboniferous rocks is well known (Jurkiewicz & Zakowa 1972; Pozaryski 1977). The subsurface geological map of the Nida Trough (Fig. 9) shows that, in a southerly direction, there is no Lower Carboniferous close to the Galezice area. It has been removed due to post-Variscan erosion. The next Lower Carboniferous record is preserved southwest of Jedrzejow, about 35 km away from Galezice. The Visran succession penetrated by the Wegrzynow IG-1 borehole in this area is represented mainly by basinal deposits including carbonate gravity-flow deposits (Jurkiewicz 1973). It therefore seems reasonable to suppose that the carbonate platform, from which the submarine fan system of Galezice was supplied, existed in the area between Wegrzynow and Galezice. Its geographical extent cannot be

324

Z. B E L K A ET AL.

!

"~ ~:L"~r:~

.....

Fig. 8. (a) Aciniform association of Globochaete sp. in Vis6an peloidal wackestone. (b) Fragment of solenoporaceaean (red algae) colony. (c) Typical encrustation form of Pseudolituotuba gravata (Conil & Lys). (d) Polished slab of Vis6an auloporid biolithite. (e) Photomicrographs of auloporid bafflestone with Sinopora polonica Nowinski. (f)-(g) cryptalgal crusts (arrowed) on corals and spongiform matrix. Scale bars are 1 mm, except on (d) where ~,ar is 3 cm.

CARBONATE PLATFORM RECONSTRUCTION, POLAND precisely outlined, but most probably it corresponded broadly with the extent of the so-called Jedrzejow High, an elevated fragment of Precambrian basement (Fig. 9) from which the Palaeozoic rocks have been totally removed. Today, the Precambrian is directly overlain by Permian and/or Triassic rocks. According to Morawska & Stupnicka (1985), the Jedrzejow High played a significant role in controlling the tectonic evolution of the northern part of the Malopolska Massif during Variscan time. We propose to call the inferred carbonate platform the Nida Platform after the river that drains this area. The Nida Platform constituted an isolated element and was not connected with the vast Cracow Platform situated further to the south (cf. Belka 1987). The borehole data appear to indicate that there was another small carbonate platform between these two (Fig. 10), which has also been totally destroyed by erosion. It presumably covered the narrow, elevated basement block of Opatkowice (Paszkowski 1988). This is indicated by the record of Vis6an polymictic debrites found in the Wegrzynow IG-1 and Lobzow IG-1 boreholes (Jurkiewicz 1973; Migaszewski & Zakowa 1991), which were drilled on both sides of this basement high. The debrites contain a wide spectrum of clasts, including Precambrian, Lower Palaeozoic, Devonian and Carboniferous rocks. This material is assumed to have been derived from the

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325

Opatkowice Platform, which must have already been affected by strong erosion during the late Vis6an, since pre-Devonian rocks were exposed and eroded at that time. The reconstructed bathymetric relief of the shelf area south of the Holy Cross Mountains (Fig. 10) presented in this paper differs significantly from previously presented palaeogeographical interpretations, in which the Jedrzejow and Opatkowice highs were interpreted to represent land areas during the entire Early Carboniferous (Jurkiewicz & Zakowa 1972; Kicula & Zakowa 1972).

Depositional trends in the lost sedimentary sequence Carbonate clasts and fossils contained in the debrite layer provide basic information from which the lost sedimentary sequence of the Nida Platform can be reconstructed. Clearly, we realize that the inferred sequence includes only the sedimentary and stratigraphic record of a fragment of the platform margin that collapsed. Nevertheless, the sequence reveals depositional trends which certainly reflect the main stages in the sedimentary and tectonic evolution of the entire platform. In terms of lithology and stratigraphy, the reconstructed succession is in part surprisingly similar to that of the Galezice area. The latter represents a small, isolated carbonate platform which was submerged during Late Devonian and Early Carboniferous times (Szulczewski et al. 1995). Its drowning and transformation into a submerged seamount is expressed by a progressive, stepwise lowering of the sediment surface through time from a peritidal, lagoonal environment to a deep-water anoxic basin. The carbonate platform of Galezice is represented by shallow-water Frasnian carbonates which are separated from the post-platform deposits by an unconformity surface developed during subaerial exposure (Fig. 2). The rapid drowning is demonstrated by the Famennian pelagic condensed sequence, followed by increasingly deepwater Lower Carboniferous deposits. This process was complete in the early Vis6an with the deposition of the anoxic black shales of the Zareby Beds (Szulczewski et al. 1996). There is not only a similarity in lithology between the Galezice seamount and the Nida Platform, but also a concurrence in the general depositional trend, since we have not found any clasts from intervals that in the Galezice section are represented by stratigraphic gaps. In addition,

326

Z. BELKA E T AL.

Nida Platform

Cracow Ptatform Opatkowice Platform

Holy Cross l~nan

Fig. 10. Palaeogeographical reconstruction of the late Vis~an platform-to-basin topography of the Fennosarmatian shelf in southern Poland. Records of Visran debrites of the Ostrowka section (asterisked) and in the deep boreholes (black spots) are indicated. Depth and distances not to scale.

there is an obvious correlation between the thickness of particular units at Galezice and the number of clasts of the same age in the analysed material. This appears to indicate a comparable development in thickness. During Frasnian times the Nida Platform was characterized by shallow-water, lagoonal depositionaI conditions (Fig. 11). The lack of lower and middle Famennian rocks among the clasts suggests the presence of a stratigraphic gap in the inferred sequence, which could have resulted from platform emergence that began during the late Frasnian. The pelagic character of the Famennian and Tournaisian carbonates seems to indicate a similar drowning scenario to that of the Galezice section (cf. Szulczewski et al. 1996). The influence of the tectonic factor on drowning is documented by the presence of Tournaisian neptunian dykes cutting the Frasnian substrate. During Visran times, however, a totally reversed trend in the development of both platforms becomes pronounced. Although we do not have any clasts documenting the early Visran sedimentary development, it is evident that the Nida Platform recovered at this time. This is indicated by the deposition of shallow-water carbonates

on the Nida Platform in the middle Visran (Fig. 11), while in the Galezice section deepwater basinal deposits can be observed. A shallowing trend during the early Vis~an is found to occur in many sections worldwide (e.g. in the Dinant Basin) and is interpreted as a result of eustatic sea-level fall (Ross & Ross 1987). It was, however, not eustasy alone that induced the recovery of the shallow-water conditions on the Nida Platform. Rather, it seems that uplift of the platform contributed significantly to the lowering of the sediment surface. This can be inferred by comparison with the bathymetric evolution of the seamount of Galezice, which drowned more and more during the early Visran and does not show any impact of a eustatic fall on its depositional regime at that time (cf. Szulczewski et al. 1996).

Conclusions Stratigraphic and microfacies examination of debris material contained in the gravity-flow deposits of the Vis~an succession of the southwestern Holy Cross Mountains has enabled the

CARBONATE PLATFORM RECONSTRUCTION, POLAND

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reconstruction of the sedimentary sequence removed by erosion. It has thus supplied new data on Fennosarmatian shelf palaeogeography. The picture that emerges is one of a carbonate platform located south of the Holy Cross area that underwent sedimentation until the late Vis6an. The geographic extent of the inferred carbonate platform system, termed here the Nida Platform, presumably corresponded to the Jedrzejow High, an elevated fragment of the Precambrian basement. In addition, by comparison with the succession of the Galezice seamount, we are able to suggest that this platform drowned during Famennian and Tournaisian times, but later on, during the Vis+an, started to recover. This

reversed trend is interpreted to have resulted from the combined effect of eustatic sea-level fall and tectonic uplift. The new data significantly modify hitherto-proposed palaeogeographical reconstructions of the Fennosarmatian shelf in southern Poland, in which the existence of 'islands' on the shelf were suggested in places where the Devonian and Carboniferous rocks are now absent. We thank J. Dvorak (Brno), J. Nebelsick (Tiibingerl), P. Strogen (Dublin) and an anonymous reviewer for helpful comments and contributions to the text. A. Nowinski (Warsaw) is acknowledged for tabulate coral determinations and B. Waksmundzki (Warsaw) for drafting the figures.

Z. B E L K A E T AL.

328

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& ZAKOWA, H. 1972. Lithologic-palaeogeographic development of the Devonian and Lower Carboniferous in the Nida Trough. Kwartalnik Geologiczny, 16, 817-850. KALVODA, J. 1986. Upper Frasnian and Lower Tournaisian events and evolution of calcareous foraminifera close link to climatic changes. In: WALLISER, O. H. (ed.) Global Bio-Events. Lecture Notes in Earth Sciences, 8, 225-236. KICULA, J. & ZAKOWA, H. 1972. Devonian and Carboniferous in the basement of the southern part of the Miechow syncline. Annales de la Soci~t~ GOologique de Pologne, 42, 165-228. LARSEN, B. R., CHAN, M. A. & BERESKIN, S. R. 1988. Cyclic stratigraphy of the Upper Member of the Guilmette Formation (Uppermost Givetian, Frasnian), west-central Utah. In: MCMILLAN, N. J., EMBRY, A. F. & GLASS, D. J. (eds) Devonian of the World. 2. Sedimentation. Canadian Society of Petroleum Geologists Memoir, 14, 569-579. MALEC, J. & RACKI, G. 1993. Givetian and Frasnian ostracode associations from the Holy Cross Mountains. Acta Paleontologica Polonica, 37, 359-384. MIGASZEWSKI, Z. & ZAKOWA, H. 1991. Some remarks on the Permian basement in the vicinity of Szreniawa (Lobzow). Kwartalnik Geologiczny, 35, 163-188. MORAWSKA, A. & STUPNICKA, E. 1985. Polnocnozachodni zasieg masywu malopolskiego i pozycja tektoniczna wyniesienia Wloszczowej. Przeglad Geologiczny, 11,602-606. NOWINSK1, m. 1976. Tabulata and Chaetetida from the Devonian and Carboniferous of southern Poland. Palaeontologia Polonica, 35, 1-125. PASZKOWSKI, M. 1988. Dinantian basin in the Cracow area: an attempt of a synthesis. Przeglad Geologiczny, 4, 200-207. PICKARD, N. A. H. 1992. Depositional controls on Lower Carboniferous microbial buildups, eastern Midland Valley of Scotland. Sedimentology, 39, 1081-1100. POZARYSKI, W. 1977. The Variscan epoch in the epi-Gothian Platform and its border zone. In: KSIAZKIEWlCZ, M., OBERC, J. & POZARYSKI, W. (eds) Geology of Poland, I V - Tectonics. Wydawnictwa Geologiczne, Warszawa, 207-252. PRAJq, B. R. 1982. Stromatolitic framework of carbonate mud mounds. Journal of Sedimentary Petrology, 52, 1202 1227. PREAT, A. & MAMET, B. 1989. SGdimentation de la plate-forme givGtienne carbonatGe franco-belge.

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Geological Society of America Special Paper, 196, 179-194.

Contemporaneous erosion and reworking within the Dinantian of the South Munster Basin D. N A Y L O R l, G. D. S E V A S T O P U L O 2 & A. G. S L E E M A N 3

I E R A - M a p t e c Ltd, 5 South Leinster Street, Dublin 2, Ireland 2 Department o f Geology, Trinity College, Dublin 2, Ireland 3 Geological Survey of Ireland, Beggars Bush, Dublin 4, Ireland Abstract: Contemporaneous erosion and reworking of sediment is now known from a wide scattering of locations within the Dinantian rocks of the South Munster Basin. The phenomenon is most marked close to the northern margin of the basin, which lay within the Cloyne Syncline from Cork Harbour westwards to Ballygarvan. Exploration drilling further west in this syncline, near Inishannon, has found Dinantian/Namurian basinal sequences containing thicknesses of breccias with shelf carbonate clasts. The northern basin margin continues obliquely across strike westwards to cross the extension of the Cork Syncline near Crookstown, where recent drilling has located the basin margin. There is no evidence of reworking in the basin during the latest Devonian-early Tournaisian. A widespread sequence boundary is recognized in the middle Tournaisian (base of CourtmacsherryReenydonagan Formations). Within the Kinsale sub-basin, Member 2 Courtmacsherry Formation is a silt-sand unit within a mudstone-carbonate sequence; sourcing of the sand demands considerable intrabasinal or basin margin erosion. In the Bantry sub-basin, reworked conodonts occur at different levels within Member 3, Reenydonagan Formation (late Tournaisian to Arundian). There is indirect evidence for a carbonate shelf or intrabasinal high of Tournaisian to Asbian age in this western region.

Reworking of shelf material into basinal environments is commonplace within the Rhenohercynian zone of the Variscan orogen. Examples are found in Germany (Franke et al. 1975) and southwest England (Matthews & Thomas 1974). In some cases the provenance of the reworked material provides the only evidence of the location and nature of the shelf source area from which it was derived. Evidence for contemporaneous erosion and reworking of sediment is now known from a wide scattering of locations within the Dinantian rocks of the South Munster Basin. The distribution and nature of the reworked material may help in assessing the position of the shelf margin in those areas where the shelf deposits have since been removed by erosion or are now covered by the sea. The aim here is to record these data, to provide further evidence, to describe a critical piece of ground around Crookstown at the western limit of Dinantian limestones in the Cork Syncline, and to discuss mechanisms and provenance. The overall facies picture for the Upper Devonian and Dinantian rocks in southernmost Ireland has gradually emerged from more than a century of work in the region (Griffith 1839; Jukes 1865, 1866; Turner 1939, 1952; Naylor 1969). Figure 1 shows the general run of Old Red Sandstone and Lower Carboniferous outcrops in

southernmost Ireland. A line drawn from Cork Harbour to the Kenmare River separates two distinct geological provinces. To the north of the line, on the North Munster Shelf, the Old Red Sandstone is overlain conformably by a relatively thin marine clastic unit (the Lower Limestone Shales of the earlier literature), which is followed by a thick carbonate sequence of shelf and bank limestones. To the south, the red beds are succeeded by a thick marine sequence with relatively little carbonate, which ranges from latest Devonian to Namurian in age. This southern region was designated the South Munster Basin by George et al. (1976) to distinguish it from the earlier Old Red Sandstone Munster Basin, whose depocentre lay to the north (Fig. 2). In the last 25 years, the stratigraphy of the South Munster Basin and North Munster Shelf has been intensively investigated. The reviews of the stratigraphy by Naylor et al. (1983) and MacCarthy & Gardiner (1987) contain extensive references to earlier work. Biostratigraphical correlations in the region are based almost exclusively on the study of miospores and conodonts (Higgs et al. 1988). Within the basin, Devonian and Carboniferous lithostratigraphical units can be traced over wide areas. Naylor et al. (1974) demonstrated that the marine sequences of the South Munster Basin

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 331-343.

332

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Old Red Sandstone Dingle Group

0 t

,

45Km ,

,

Lower Palaeozoic

~ i

Fig. 1. Outline geological map of the south of Ireland, with the locations of places mentioned in the text.

N

~"~'g NORTIt MUNSTERSII ELF

~ . : : : :: !::::::::::::::::::::::: ::!:::i:::!::: : : : : :: : :::::::::::::::::::::::::::::::::::: KENMARE :::: , - G...~. :.:~;:, ;: .:::: ; .:e ~:: :: :": ~ :: ::: [:: .~ j: / .• ,• ...:.:.. ::::::: : : :::::::::::. ======================== •

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R ATIC

SItEEP'S HEAD IIlGlt MIZEN HEAD 0

,

Kilometres

a°

~

~

_l~_r~_.A ~ ~ "'-~NORT ',

....... , ,.IlERN MARGINOFTIIE . . . . j__~-" SOUTlt MUNSTERBASIN . . . . -.- - -. . . ~. - -.- . . . . . . . . . '" rA~ K L INS - % i ~' '~ ~,

'

., "

~ ....

CLANDORE ' ' It IGil

50

f MAIN DEPOCENTRES O.R.S. MUNSTER BASIN

Esso Marathon 48/30- l

Fig. 2. Sketch map showing the main features of the Munster and South Munster basins, the locations of places mentioned in the text, and the position of Esso-Marathon well 48/30-1. Cities Services well 63/4-1 lies approximately 170 km SSW of Mizen Head. ORS, Old Red Sandstone.

SOUTH

MUNSTER

were deposited in two sub-basins: an eastern Kinsale sub-basin (Naylor 1966), and a western Bantry sub-basin (Jones 1974; Naylor & Sevastopulo 1993), separated by the positive element of the Glandore High (Naylor et al. 1985, 1988; Sleeman 1987). The more important localities mentioned in the text are shown in Figs 1 and 2, and the relevant stratigraphical nomenclature in Fig. 3.

Kinsale sub-basin There is no evidence of downslope reworking within the sub-basin during the latest Devonian and early Tournaisian (Old Head Sandstone and Kinsale Formations), except for minor slumping. This is in line with other evidence that suggests that only gentle slopes were developed at this time, despite the contrast in thickness between shelf and basin (Naylor et al. 1989). Any deep fault accommodation to the thickness variations did not apparently result in emergent fault scarps. Throughout the basin a marked change occurs within the middle Tournaisian at the base of

SERIES/STAGE

MIOSPORE/ CONODONT BIOZONES

BRIGANTIAN

nodosus

BASIN

333

Courtmacsherry Member 1. Thick shallowmarine clastic successions are overlain abruptly by carbonate and mudrock sequences. In many sections there is evidence of slight erosion, where phosphate-bearing carbonate lenses cut down into Kinsale Formation. At North Ringabella, PC Biozone spores occur in the top few tens of centimetres of the Kinsale Formation, while elsewhere in the basin their first occurrence is in the base of Member 1 of the Courtmacsherry Formation (Higgs et al. 1988), perhaps indicating widespread erosion. Evidence from shelf areas throughout the south of Ireland (Phillips & Sevastopulo 1986) suggests that this is a regional phenomenon, i.e. a sequence boundary which can be recognized in both shelf and basin. However, in the northeast portion of the basin, the results of field mapping could be interpreted as indicating substantial pre-Courtmacsherry Formation erosion, which would have presumably been caused by active movement on basin margin faults. An outcrop section at Halfway, at the western end of the Cloyne Syncline (Fig. 2), exposes about 200 m through the Pig's Cove Member, the upper member of the Kinsale Formation.

KINSALE SUB-BASIN OLD HEAD RINGABELLA CLOYNE OF KINSALE SYNCLINE SYNCLINE

BANTRYSUB-BASI~

NORTH MUNSTER

SHELF

LITTLE

bilineatus ASBIAN

HOLKERIAN

LISPATRICK FM

REENYDONAGAN FM (MBR4)

LISPATRICK

FM

SHELF A N D

MUDBANK LST

ARUNDIAN

Z:

~[

L. VISEAN

REENYDONAGAN

FM (MBR3)

anchoralis REENYDONAGAN

carina

COURTMACSHERRY FM (MBR4) M I N A N E LOUGHBEGFM COURTMACSHERRY FM (MBR 3) COURTMACSHERRY

FM (MBR2)

FM (MBR2)

PC Si.phonodell~ R E E N Y D O N A G A N mornatus FM (MBR 1) TOURNAISIAN

BP HD

ARDNAMANAGH ~ FM u. REENAGOUGH ~ FM <

VI

o

LN

~ ARDATU RRIS H FM

CHEKr FM

WAULSORTIAN LST

RINGABEL.LA

LST FM

WAULSORT[ANL.ST BALLYSTEEN FM

BALLYSTEENFM BALLYMARTINFM

COURTMACSHERRY FM (MBR 1) PIGSCOVEMBR

RINGMOYLAN SHALEFM MELLONHOUSE FM

NARROWCOVEMBR CASTLE SLATE M B R O L D HEAD SANDSTONE FM

'OLD RED SANDSTONE'

Fig. 3. Latest D e v o n i a n and D i n a n t i a n stratigraphical units m e n t i o n e d in the text. Only those m i o s p o r e and c o n o d o n t biozones recognized in the S o u t h M u n s t e r Basin are shown.

334

D. NAYLOR ET AL.

This then appears to thin rapidly eastwards towards Ballea (Fig. 4), where only 75m is recorded (Sleeman 1987). At Raffeen, where the member is no more than 50 m thick, it is seen to pass up into the Courtmacsherry Formation. The member has not been mapped further north and east of Raffeen. While available evidence points to eastward thinning of the member, postKinsale Formation erosion, as suggested by the absence of the BP Miospore Biozone in the Ballea area (Sleeman 1987), and strike-parallel faulting may also have contributed to the apparent thinning and absence of the member through much of the Cloyne Syncline (Sleeman et al. 1986). No further evidence of reworked material or slope activity has been recorded within Member 1, Courtmacsherry Formation in the basin, despite thickness changes between basin and shelf. There are various manifestations of reworking higher in the Courtmacsherry Formation. In the main, these appear to be related to activity along the northern and southern margins of the basin.

~

Little Island Formation

~White

N o r t h e r n margin o f the Kinsale sub-basin

A number of papers in the last few decades have dealt with the nature of the facies transition from the basin to the northern shelf: in the Cork Harbour sector by Naylor (1969), Sleeman et al. (1986) and Sleeman (1987); and in the west by Naylor et al. (1981, 1983). The northern boundary of the South Munster Basin has generally been shown as a straight line from Cork Harbour, through the Crookstown area, to Kenmare in the west. However, a glance at Fig. 1 reveals that the level of present erosion across the major Variscan folds has removed all evidence at the critical levels of stratigraphy between Crookstown and Kenmare. Suggestions of a serrated northern margin controlled by offset along en echelon faults have been made by Naylor & Sevastopulo (1979, 1993) and Naylor et al. (1989), whilst Williams et al. (1989) have suggested a specific model for the position of the margin and its control by two major strike faults with a dextral offset along a NW-SE fault system.

Strand Formation

[-'---1 Pig's Cove Member

~ ' C a s t l e Slate & , I--F...iZ-I Cusklnny Members

~

Lou'ghbeg Formation

~

Lispatrick Formation

~

Waulsortian Limestone

~

Ringabella Limestone Member ~

~

Ballysteen Formation

~

Courtmacsherry Formation undivided m --- ~

Faults

~

Old Head Formation Old Red Sandstone

Approximate shelf margin

Fig. 4. Geological map of the Cloyne and Ringabella Synclines (after Sleeman 1987). The proposed position of the shelf margin is shown with a bold dashed line.

SOUTH MUNSTER BASIN In the Cork Harbour area three major synclines are cored by Dinantian and Namurian rocks, with Devonian Old Red Sandstone exposed in the intervening anticlines: these are, from south to north, the Ringabella Syncline, the Cloyne Syncline and the Cork Syncline (Figs 1 & 4). The Dinantian rocks of the Ringabella Syncline are in basin facies. Recent work, cited below, has shown that the northern margin of the South Munster Basin lay within the Cloyne Syncline as far west as Ballygarvan (Fig. 4). Microfaunal evidence of reworking has been recorded at a number of localities. The margin then crosses the Cork Syncline further west in the vicinity of Crookstown. Even in detail, the margin is believed to follow a zig-zag course, due to dextral offset along deep-seated N N W SSE faults (Sleeman 1987). Evidence of reworking in the South Munster Basin was first recorded by Matthews & Naylor (1973, Fig. 3) in the Ringabella Limestone Member, and this is also stratigraphically the earliest example of the phenomenon. The Ringabella Limestone is between 244m and 366m thick and contains substantial quantities of quartz sand and silt admixed with the carbonate. Here Polygnathus communis carina occurs in a few of the lowest beds of the unit (Matthews & Naylor 1973; subsequent bed-bybed collecting by G.D.S.). These conodont faunas do not contain any of the sub-biozonal indicators and this, in conjunction with the occurrence of a distinctive morphotype of Gnathodus euneiformis, strongly suggests that they should be assigned to the lowest part of the P. communis carina Biozone. Occurring with elements of P. communis carina Biozonal age are numerous reworked elements of Siphonodella Biozone age, which must have been derived from Courtmacsherry Member 1 or equivalent strata. Through the remainder of the Ringabella Limestone, as far as is known, the conodont fauna in the limestone beds, which contain significant amounts of phosphatic debris, remarkably consists entirely of reworked elements of Siphonodella Biozone age. Therefore the age of the highest exposed bed of the Ringabella Limestone can only be established as older than the overlying late Tournaisian/early Vis6an Minane Chert (Naylor et al. 1983). The Ringabella Limestone appears to be restricted to the present coast. An equivalent section along strike only 5 km to the west (Tracton: Fig. 4) contains no thick carbonate beds or significant sand content. Sleeman (1987) suggested that the carbonate-rich Ringabella unit was deposited by turbidity currents carrying

335

eroded material from the slope and shelf immediately to the north. An example of much less extreme reworking is seen in the Cloyne Syncline on the west side of Cork Harbour. In Ballygarvan quarry in the Cloyne Syncline (Figs 1 & 4; Sleeman et al. 1986) the lowest beds of limestone (Ballysteen Formation) overlying the Courtmacsherry Formation (Member 1) are of earliest P. c. carina Biozone age, but contain reworked conodonts of Siphonodella Biozone age. However, in contrast to the Ringabella Limestone, the higher parts of the Ballysteen limestones, also of P. c. carina age, do not contain reworked conodonts or quartz sand. A similar situation occurs at Raffeen, closer to the harbour (Fig. 4). A conspicuous feature of the Asbian succession at Ballygarvan is the presence of several horizons of limestone breccias. Some of the limestone clasts are of cobble grade. At Meadstown House, immediately west of Ballygarvan, rocks of late Tournaisian-Asbian age are absent. Late Asbian limestone breccias of the Little Island Formation rest directly on the Courtmacsherry Formation (P. inornatus Biozone) and interdigitate with black carbonaceous basinal mudstones (Lispatrick Formation). Sleeman (1987) envisaged that the Meadstown area suffered erosion at this time, associated with its palaeogeographical position on the basin slope. Turbidity currents and debris flows may have carried material down the slope in response to tectonic activity on deepseated structures along the Cork-Kenmare line. A suggested course for the northern margin of the basin between Cork Harbour and Meadstown is shown in Fig. 4. Exploration drilling a further 10km southwest in this syncline at Rag Bridge, approximately 4kin ENE of Inishannon (Fig. 1), has proved basinal sequences of Dinantian/Namurian age, which differ in several respects from the time-equivalent type section on the Old Head of Kinsale (Fig. 5). The boreholes penetrated black cherty shales resting on limestones and calcareous shales, the 'Inishannon Limestones', which contain substantial amounts of resedimented limestone clasts as well as quartz sand similar to that in the Ringabella Limestone. Only an outline of the stratigraphy (Fig. 5), with critical faunal evidence, is presented here. Boreholes BS10 (125 m E of Rag Bridge) and BS11 (400 m S of Rag Bridge) provide representative sections. The rocks penetrated in both boreholes exhibit variable, in places steep, dips. The successions shown in Fig. 5 give the best estimate of true stratigraphical thickness.

D. NAYLOR ET AL.

336 BS 11

BS 10

1~)0

4----'-- 80/1432 109.fun Briganfian or younger

( ~ / o s ~ Z)

I"

ql--'- 80/1440 2ff/.6m Late Asbiaa or younger

(bilintat~s Z)

'dl'~

~r/r | 4'4 3 242.65m

Asbian

80/1413 152.5m Late Asbian or younge~ (bilineams Z) 80/1411 165.6m

0 "I.D.292.45 m

8011409 169.8m

LIMESTONE CONGLOMERATE

Middle Toumaisian

(Siphonodella/ inornatus Z)

I~

LIMESTONE

1~

GREY/BLACKMUDROCK

"I.D. t95.40 m

Fig. 5. Carboniferous successionpenetrated in Riofinex boreholes BS10 and BS11 (Rag Bridge, approximately 4 km ENE of Inishannon: see Fig. 1), showing the horizons of critical micropalaeontologicalsamples. The sections have been corrected for tectonic dip.

Borehole BS10 shows the older part of the succession. The lowest horizons which have yielded conodonts are thin limestones in Samples 80/1409 & 80/1411. Both contained P. inornatus and the former, in addition, yielded Siphonodella sp. These faunas are clearly of Siphonodella Biozone age, which is consistent with the assignment of this lower part of the section on the basis of the lithology (predominantly calcareous and non-calcareous grey mudrock) to Member 1 of the Courtmacsherry Formation. The next highest horizon that has yielded stratigraphically significant conodonts is a limestone conglomerate, with pebble and granule grade clasts (Sample 80/1413). The fauna includes Gnathodus bilineatus, which indicates a level substantially above the base of the Asbian Stage. The stratigraphical section

between the middle Tournaisian and Asbian horizons, which as yet has not yielded any microfauna, consists of an estimated 10 m of dark mudrock, cherty and calcareous in part, with some beds of fine-grained carbonate that contain phosphate granules. It is not clear from the core whether the middle Tournaisian to Asbian succession is highly condensed, or is missing across a disconformity surface. The remaining upper part of the core contains dark mudrocks that are commonly pyritous and laminated, and horizons of limestone, some of them coarse and many of them silicified. Conodonts from the limestone horizons are at least as young as Asbian; no diagnostic Brigantian forms have been found. Borehole BS11, which was drilled to 292.45 m, did not reach Member 1 of the Courtmacsherry Formation. The succession consists of a lower carbonate-rich unit and a higher unit of dominantly dark grey to black shale, with subordinate chert and fine-grained carbonate containing small goniatites. The carbonate-rich section is varied, and includes calcarenites and calcirudites of granule grade, interbedded with dark shale. The limestone beds contain lime-mud clasts and shallow-water material such as ooids. The lowest fauna is from Sample 80/1443 and contains a poorly preserved foraminifer identified as either Howchinia or Vissariotaxis. These taxa indicate an Asbian or younger age. Sample 80/1440 yielded the conodont G. bilineatus, indicating a maximum age significantly above the base of the Asbian. Sample 80/1432 yielded the conodont Lochriea nodosus, indicating a Brigantian or younger age. The black shale at the top of the borehole is undated and may extend into the Namurian. The Courtmacsherry Formation has recently been mapped by AGS in the area immediately to the northwest of Inishannon. A particularly interesting section has been logged in the old Bandon and South Coast Railway cutting at Rockfort House, Brinny, 3kin northwest of Inishannon. The formation is here about 200 m thick and passes up into the Lispatrick Formation. The lower part of the section comprises silty and variably calcareous mudstones (containing PC biozone miospores) with thin crinoidal bioclastic limestones and crinoid-free dolomitized calcisiltites. The upper part of the sequence is dominated by blocky, nodular, cherty, dolomitized calcisiltites, and argillaceous, decalcified and cleaved mudstones. Abundant quartz silt is present in the ferroan dolomitized calcisiltites. Detailed micropalaeontological dating has not yet been completed, so

SOUTH MUNSTER BASIN

337 .~-:::i:::

.

- . . : : : i .~..... .. :. ::-: .... . ..... -.-.-~-,-..~, ,~,,,,-,.. :

I

.

.

......

::

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,- .

.

.

.

~

~

======================================== ,-,-,-,-,-,~.,.~ ~ ' - ~-, .

.

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.

....

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.::i::i!i " ~

. -- ..:-:... . --L~

.- a-~

0

Kilometres

" . . . . . ~ ~::::! :::! i :ii:i l ~ i........ l I S L A N D FORMATII "O1 FORMATION

z

.~

~

OLD

~

O L D RED S A N D S T O N E

HEAD

FORMATION

Fig. 6. Geological map of the area between Crookstown and Killumney in the Cork Syncline, showing the locations of boreholes mentioned in the text.

correlation is tentative. The topmost 15 m of the section, cherts and splintery mudstones, is provisionally assigned to the Lispatrick Formation (cf. the Minane Chert Member: Naylor 1969; Sleeman 1987). No dates have been obtained from this uppermost section; as yet all samples have proved barren. The geology of the Crookstown area at the western end of the Cork Syncline is not well known, due to poor exposure. The core of the Cork Syncline, containing sparse outcrops of Dinantian limestones, plunges gently westwards, with the limestone exposure closing at Crookstown to reveal scattered outcrops of the underlying Cork Group. Jukes & Du Noyer (1862) gave a brief description of poorly exposed black or grey shale, some 20-35 m thick, overlying the Old Red Sandstone and overlain in turn by Carboniferous Limestone. They contrasted this sequence with the 2000m or so of 'Carboniferous Slate' (Cork Group) exposed 28 km south of Crookstown on the Old Head of Kinsale. This is a critical area for the understanding of the northern margin of the basin, and the most westerly control available until the Kenmare syncline is reached. Recent work (Sleeman et al. 1994, Fig. 6) has given an improved understanding of the area. Details will be published elsewhere, but an outline of the results is presented here. The drilling of several boreholes for base metal exploration by Inco/Oliver at Stickstown and Aherla (Figs 6 & 7; IN7, IN8 • and IN 12; Sleeman 1986) has revealed a hitherto unexposed section through the Loughbeg Formation, a unit which at Ballygarvan and

Loughbeg occurs at the basin margin (Figs 1 & 4; Sleeman et al. 1986; Sleeman 1987). The stratigraphic units in the core of the Cork Syncline in the Crookstown area are described below. Waulsortian Limestones. In the area between Killumney, Aherla and Crookstown (Fig. 6), Waulsortian mudbank limestones are exposed in a number of old quarries, notably the nowflooded old Castlemore Quarry, and were encountered in Borehole IN8 at Aherla. They range in age from the P. communis carina Zone (as in borehole IN8) to the Sc. anchoralis Zone. Extensive outcrops of Waulsortian Limestones are also exposed further eastwards towards Cork City. Loughbeg Formation. This formation has been encountered in Borehole IN7 at Stickstown, where a true thickness of about 30m was penetrated, and in Borehole IN12 (about 35m seen). In neither case is the base of the formation seen. The succession comprises cherty mudstones and silicified limestones, non-cherty crinoidal packstones, and intervals of darkgrey to black or sooty carbonaceous mudstones. The lower half of the penetrated section contains conodonts belonging to the Se. anchoralis Zone, but the upper part lacks diagnostic faunas and could be Vis6an in age. It is clear that the Loughbeg Formation at Stickstown is the time-equivalent of some of the Waulsortian Limestones in the Castlemore quarries and also of outcrops of Waulsortian

338

D. NAYLOR E T AL.

Z

Little Island Formation. This formation is generally poorly exposed in the Crookstown district, but is seen in the new Castlemore Quarry and was penetrated in boreholes IN7 (about 47 m seen) and INI2 (about 66 m seen). Most of the limestones in the boreholes are pale to medium grey 'mudbank' calcilutite wackestones. Occasional calcarenitic or calcirudite wackestones and crinoidal packstones occur. There are mottled calcilutite breccias in borehole IN7, and similar breccias are present in the Castlemore and Ballygarvan quarries (Sleeman 1987).

IN7 OB 16.5m r~

.<

100

| 82.4m

IN 8 OB 2.6m

Southern margin o f the Kinsale sub-basin

On Galley Head (Fig. 1) the exposed marine stratigraphic units, from latest Devonian onwards, are considerably thinned relative to those of Dunmanus-Bantry to the west, and Old Head of Kinsale-Seven Heads to the east (Naylor et al. 1983). This was regarded by Naylor et al. (1974) as evidence for a basementcontrolled positive element - named the Glandore High - separating eastern and western subbasins. However, Matthews in Naylor et al. (1974) suggested that the element could form part of the southern platform of the South Munster Basin. This interpretation was pre"I.D. 128m T.D. 47m ferred by Williams et al. (1989), who envisaged (their fig. 7) the edge of the southern platform as ~ CALCARENITE ~ LSTBRECCIA extending east-west across the landward end of the Galley Head and Mizen peninsulas. How~ CHERTYCALCARENITE ~ MUDROCK ever, the marked difference of stratigraphy CALCISILTITE ~ WAULSORTIAN between the east and west coasts of the Seven Heads peninsula (Naylor et al. 1988) suggests CALCILUTITE that the margin of the positive element had a Fig. 7. Carboniferous succession penetrated in northerly trend, at least in this area, and did not Boreholes IN 7 and IN 8 (for locations see Fig. 6). The strike east-west. sections have not been corrected for tectonic dip. Solid The position of the southern margin of the lines indicate the confidently identified extent of South Munster Basin is c o n j e c t u r a l - as is the biozones and stages. The dotted lines indicate the possibility of deriving sediment from this source. tentatively identified extensions of the ranges. MacCarthy & Gardiner (1987) postulated a 'southern platform' to the South Munster Basin, controlled by faults, lying immediately south of the present coastline and broadly on the south limb of the Cork Syncline. This parallel to it. The only direct stratigraphic relationship is analogous to that found around evidence relating to the southern margin of the Ballygarvan, and indicates that Stickstown lay Kinsale Sub-basin comes from an offshore well, close to the shelf-basin margin. The margin is Esso-Marathon 48/30-1 (Fig. 2). This penetrated likely to have been controlled by a deep 154 m of strata beneath the sub-Mesozoic structure, a possible course for this is along the unconformity, which were dated by Higgs present scarp to the south of the boreholes. The (1983) as belonging to the VI Subzone of the edge of the Waulsortian mudbanks should lie NV Biozone (probably Tn2a). Following Higgs immediately to the east of Stickstown, as has (ibid. p. 111) we interpret this sequence as being been identified at Aherla. part of the Kinsale Formation. The thickness of s

s

s

~

s J s s ,

s s s r ~

t s s ¢ ~

s s s s ~

:ii!iiiii

SOUTH MUNSTER BASIN the Kinsale Formation intersected in 48/30-1 argues that the early Tournaisian shelf margin lay well south of the present coastline. Within the southern part of the Kinsale sub-basin, Member 2 of the Courtmacsherry Formation is a sand-silt dominant unit within an otherwise mudstone-carbonate sequence. Member 2 is 36 m thick at the type section on the Old Head of Kinsale (Naylor 1966) and 120 m thick on the east coast of the Seven Heads peninsula (Naylor et al. 1988). The member is not known further north, and is absent on the west side of Seven Heads and on Galley Head, where the Courtmacsherry Formation thins out against the margin of the Glandore High; nor is it present within the Bantry sub-basin. Sourcing of substantial thicknesses of sand into the southern part of the Kinsale sub-basin clearly demands considerable intrabasinal or basin margin erosion, to at least sub-Courtmacsherry Formation levels in the stratigraphy. The age of Member 2 is known only from Seven Heads (Naylor et al. 1988) where late Tournaisian-aged conodont faunas occur above it. Coeval sequences over a broad shelf area to the north are in carbonate facies. The sedimentary characteristics of the member suggest relatively shallow deposition (Naylor 1966). This may indicate a zone of shallowing in the southern part of the sub-basin, either due to proximity to the southern basin margin, or to an intra-basinal emergent feature. Despite the influence of the Glandore High on the thickness of the Kinsale and Courtmacsherry Formations, there is no evidence of significant erosion of Kinsale strata on the exposed onshore portion of the high. However, the thickening of Member 2 westwards between the Old Head and east Seven Heads could indicate sand input related to an eroded offshore portion of the Glandore structure. The Ringabella Limestone and Member 2, Courtmacsherry Formation may be in part coeval, and the sand in both units could have the same source. Winnowing of sufficient sand from a northerly derived carbonate-quartz admixture to produce Member 2 seems unlikely. Equally, special pleading is required to envisage sand from a southern source being swept into the carbonate sediments at Ringabella, whilst not being seen in intervening sections. It is probably simpler to envisage the Ringabella Limestone as the product of localized vigorous erosion of fault slivers on the northern basin margin east or north of Cork Harbour, in which Kinsale Formation or older rocks were exposed, and the sand of Member 2 as having a separate southern provenance.

339

Bantry sub-basin Naylor & Sevastopulo (1993) reviewed the stratigraphy of the Bantry sub-basin. Member 2 of the Reenydonagan Formation, which is probably of carina Zone age, is very condensed, but with no evidence of reworking. Member 3 consists of grey mudrock and interbedded, generally dolomitic, limestone. In its lower part the limestones are fine-grained and contain conodont faunas assigned to the anchoralis Biozone, apparently without reworked elements. The highest bed of the member is a crinoidal limestone which has yielded foraminifera and algae of Arundian age as well as conodonts, some of which (Bispathodus aculeatus aculeatus, P. communis carina, Eotaphrus cf. bultyncki and Dollymae bouckaerti) are clearly older. Jones (1974) has interpreted the sedimentary features of this member as indicating sedimentation by turbidites flowing into a relatively deep basin. The marine Old Head Sandstone and Kinsale Formations are known to thin from south to north across Dunmanus Bay onto the southern flank of the Sheeps Head Anticline (Naylor et al. 1977). There is some field evidence to suggest that the succession on the northern flank of the anticline is similarly thinned. Naylor et al. (1983) envisaged this as a further positive element, the Sheeps Head High, possibly controlled by strike-parallel faults north and south of the Head. On their isopachyte map for the Kinsale Formation (their fig. 2.6) they contoured the Sheeps Head High as a westward prolongation from the Glandore structure. Williams et al. (1989, fig. 7), on the other hand, envisaged Sheeps Head to be an area of relatively thin stratigraphy on the footwall of a major fault extending along Dunmanus Bay and inland. The identification of this fault inland appears to rest primarily on the recognition of a substantial thickening of the Old Head Sandstone Formation from the thin, poorly exposed sequence in the Cork Syncline southwards to a 1.1 km thick sequence in the Bandon Syncline (the depocentre shown on Williams et al. 1989, fig. 7). There are few good control sections inland, however, and work in progress in the Bandon Syncline shows that the scattered exposures of the Old Head Sandstone Formation can easily be confused with the sandier parts of the overlying Kinsale Formation, unless palynological control is available. Indeed, the outcrop width of the Cork Group in this district is dominated, and increased, by folding of the Kinsale Formation. While the thickening of the Old Head Sandstone Formation could well

340

D. NAYLOR ET AL.

extend significantly westwards from Ringabella, where it is well controlled by palynology, this thickening does not, in our opinion, justify the interpretation placed on it by Williams et al. (1989). We are not aware, either, of any data pointing to contemporaneous steep fault scarps during deposition of the Old Head Sandstone and Kinsale Formations. In these circumstances we prefer the interpretation showing two sub-basins separated by the positive Glandore element (Fig. 2), whilst recognizing that there is poor control inland on the precise shape of both the Glandore and Sheeps Head Highs.

Northern margin o f the Bantry sub-basin At Kenmare itself, shelf carbonate sequences are known in the syncline (Wingfield 1968; Naylor et al. 1983). However, thick Cork Group successions of the South Munster Basin crop out to the west along the southern shoreline of Kenmare River (Naylor et al. 1969), thus constraining the position of the basin margin on the west coast.

Southern margin o f the Bantry sub-basin The carbonate beds of Member 3, Reenydonagan Formation in the Bantry sub-basin, were interpreted by Jones (1974) as turbidites, probably derived from a shelf to the north or northeast. Naylor & Sevastopulo (1993) pointed out that present information was consistent with either of two interpretations. The first was that the Member 3 sediments were derived from the north, in which case the occurrence of reworked conodonts from different horizons of the P. c. carina Biozone might indicate that the up-slope equivalent of Member 2 was also condensed, so that a small depth of erosion would have yielded faunas spanning a range of ages. The second was that the turbidites were derived from the southwest. This would be consistent with the suggestion by Naylor & Sevastopulo (1993) that the varied limestone boulders described by Farrington (1936) and Sevastopulo & Naylor (1981), which occur in Quaternary deposits on the south side of Bantry Bay, were derived locally from the floor of the bay to the west. The boulders, which are very localized in distribution, were initially considered to have been derived from the northern shelf, in the Kenmare region, but improved biostratigraphic data have made this appear unlikely. The boulders, which range from Waulsortian Limestone of carina

Biozone (Eotaphrus bultyncki Sub-biozone) age to shelf limestone of Asbian age, may thus be representative of a former limestone shelf area to the west or southwest, which during the Dinantian was the source of the carbonate in the Reenydonagan Formation turbidites. This may indicate a western closure to the basin (as hinted by the northward swing of the regional gravity contours), or alternatively the existence of an intrabasinal high. The only direct offshore evidence is from the Cities Services well 63/4-1 drilled in the Fastnet Basin (Reeves et al. 1978; Robinson et al. 1981), which proved late Tournaisian shelf carbonates. The 63/4-1 results demonstrate the existence of a southern carbonate shelf, albeit 170 km SSW of Mizen Head (Fig. 2).

Discussion The trace of the northern margin of the South Munster Basin between Cork Harbour and Crookstown is now known in some detail. Its position was controlled by strike faults offset by NNW-SSE cross-faults, to produce a serrated margin. There is little evidence of basin margin fault activity, significant depositional slopes or intrabasinal erosion during deposition of the Old Head Sandstone and Kinsale Formations. A widespread hiatus took place prior to deposition of the Courtmacsherry Formation. In the Cloyne Syncline, near the northern margin of the basin, there is a zone of apparently significant erosion at this level, with uppermost Kinsale Formation (BP biozone) rocks being eroded. Evidence for reworking in the Courtmacsherry Formation begins in the early P. c. carina Biozone (Ringabella Limestone). Much of the other evidence in the Cloyne Syncline - absence of late Tournaisian to Asbian sequences, interbedding of coarse limestone conglomerates into the basinal Lispatrick Formation - appears to be related to activity along the northern margin of the basin. Further south in the Kinsale sub-basin, the presence of sand in Member 2, Courtmacsherry Formation and the general westward thinning of stratigraphic units against the Glandore High indicate more widepread activity of intrabasinal fault blocks or proximity to a southern basin margin. In the Bantry sub-basin there is similar evidence for erosion and reworking. The base of the Reenydonagan Formation again represents a marked and widespread facies change. Notably there are various levels of reworked

SOUTH MUNSTER BASIN conodonts within Member 3 (late Tournaisian to Arundian) of the Reenydonagan Formation. There is some evidence that this material may have been deposited by turbidite flows derived from the west or southwest. There is also indirect evidence from carbonate boulder erratics for a carbonate shelf, or intrabasinal high, of Tournaisian to Asbian age in this western region. The Sheeps Head positive element also points to the existence of deep fault-controlled blocks. One question to be asked is: what was the trigger for basinward transport of shelf material? Was it tectonic, linked to the evolution of the basin, perhaps footwall uplift? Or was it related to sea-level changes? In the latter case the episodes of sedimentation of exotic material into the basin might correspond to lowstands when the shelf was a terrigeneous shelf, or with highstands when the shelf was a carbonate shelf. Or was shedding of shelf sediment merely a consequence of the development of a slope at the faulted shelf margin, which periodically became unstable, or down which storm-generated sediment flows were transported? The evidence presented here suggests that the base Courtmacsherry-Reenydonagan event represents a sequence boundary which can be recognized throughout the basin and traced onto the northern shelf (base Ringmoylan Shale Formation), and which was followed by a flooding event. There is no evidence of major breaks in the succession on the North Munster Shelf that would equate with the reworking in the Ringabella Limestone, which therefore owes its origin to local tectonic movement at the basin margin. The episode of major tectonic activity in the early Vis6an and Arundian identified in the Irish Midlands (Nolan 1989; Philcox 1989) did not apparently affect the South Munster Basin, where this period was one of basin starvation. During the Asbian and Brigantian, the North Munster Shelf was affected by changes of sea level, which led to the development of cyclic sequences of limestones (Heselden 1991). However, further north in the Irish Midlands there is widespread evidence of tectonically controlled influxes of debris flows in the Dublin Basin (Nolan 1989). It is interesting to speculate that this tectonic activity may have been linked to that on the northern margin of the South Munster Basin. The other events in the South Munster basin appear to be more localized and do not have northern shelf correlatives. They appear to be related mainly to the jostling and differential erosion of fault slivers along the Cork-Kenmare

341

line, and possibly to similar movements on intrabasinal highs or the southern basin margin. The recognition of allochthonous material is easy where carbonate breccias occur within black mudrock sequences, or where there is fossil evidence of stratigraphical mixing; but it is less easy where fine-grained carbonates are interbedded with basinal shales, and w h e r e they show no, or rare, fossil evidence of stratigraphical mixing. Reworking and intrabasinal erosion may prove to be a more widespread phenomenon in the South Munster Basin than the present evidence suggests. AGS publishes with the permission of the Director, Geological Survey of Ireland

References FARRINGTON, A. 1936. The glaciation of the Bantry District. Scientific Proceedings of the Royal Dublin Society, 21,345-361. FRANKE, W., EDER, W. & ENGEL, W. 1975. Sedimentology of a Lower Carboniferous shelfmargin (Velbert anticline, Rheinisches Schiefergebirge, W. Germany). Neues Jahrbuch fur Geologie und Palaeontologia, 150, 314-353. GEORGE, T. N., JOHNSON, G. A. t., MITCHELL, M., PRENTICE, J. E., RAMSBOTTOM, W. H. C., SEVASTOPULO, G. D. & WILSON, R. B. 1976. A Correlation of Dinantian Rocks in the British Isles. Geological Society of London Special Report 7. GRIFFITH, R. 1839. On the principle of colouring adopted for the Geological Map of Ireland, and on the geological structure of the South of Ireland and accompanying map. Journal of the Geological Society of Dublin, 2, 78-93. HESELDEN, R. G. W. 1991. Sedimentology and Stratigraphy of the Courceyan-Asbian Limestones (Dinantian, Lower Carboniferous) of the Cork Harbour Area, Southern Ireland. PhD Thesis, National University of Ireland. HIGGS, K. 1983. Palynological evidence from two wells drilled in the Celtic Sea area. Geological Survey of Ireland Bulletin, 3, 107-112. --, CLAYTON, G. & KEEGAN, J. B. 1988. Stratigraphic and Systematic Palynology of the Tournaisian Rocks of Ireland. Geological Survey of Ireland Special Paper 7. JONES, P. C. 1974. Marine transgression and facies distribution in the Cork Beds (Devonian-Carboniferous) of west Cork and Kerry, Ireland. Proceedings of the Geologists' Association, 85, 59-188. JUKES, J. B. 1865. Notes for a comparison between the rocks of the South-West of Ireland, and those of North Devon, and of Rhenish Prussia (in the neighbourhood of Coblentz). Journal of the Geological Society of Ireland, 1, 103-143.

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- - 1 8 6 6 . On the Carboniferous Slate (or Devonian rocks) and the Old Red Sandstone of South Ireland and North Devon. Quarterly Journal of the Geological Society of London, 22, 320-371. & DU NOYER, G. V. 1862. Explanation to accompany Sheets 194, 201,202 of the Maps of the Geological Survey of Ireland Illustrating a Portion of the County of Cork. Memoirs of the Geological Survey of Ireland. MATTHEWS, S. C. & NAYLOR, D. 1973. Lower Carboniferous conodont faunas from south-west Ireland. Palaeontology, 16, 335-380. -& THOMAS, J. M. 1974. Lower Carboniferous conodont faunas from northeast Devonshire. Palaeontology, 17, 371-385. MACCARTHY, I. m. J. & GARDINER, P. R. R. 1987. Dinantian cyclicity: a case history from the Munster basin of southern Ireland. In: MILLER, J., ADAMS, A. E. & WRIGHT, V. P. (eds) European Dinantian Environments. John Wiley, Chichester, 199-237. NAYLOR, D. 1966. The Upper Devonian and Carboniferous geology of the Old Head of Kinsale, County Cork. Scientific Proceedings of the Royal Dublin Society, Series A, 2, 229-249. - - 1 9 6 9 . Facies change in Upper Devonian and Lower Carboniferous rocks of Southern Ireland. Geological Journal, 6, 307-328. -& SEVASTOPULO, G. D. 1979. The Hercynian 'Front' in Ireland. Krystalinikum, 14, 77-90. & 1993. The Reenydonagan Formation (Dinantian) of the Bantry and Dunmanus synclines, County Cork. Irish Journal of Earth Sciences, 12, 191-203. , HIGGS, K. & BOLAND, M. A. 1977. The stratigraphy on the north flank of the Dunmanus Syncline, West Cork. Geological Survey of lreland Bulletin, 2, 143-157. , REILLY, T. A. & SEVASTOPULO, G. D. ]988. Dinantian and Namurian stratigraphy, Seven Heads Peninsula, County Cork. Irish Journal of Earth Sciences, 9, 1-17. , JONES, P. C. & CLARKE, M. G. 1969. The stratigraphy of the Cork Beds (Upper Devonian and Carboniferous) in southwest Ireland. Scientific Proceedings of the Royal Dublin Society, Series A, 3, 171-191. , & MATTHEWS, S. C. 1974. Upper Devonian to Lower Carboniferous depositional environments in south-west Ireland and adjacent regions. Geological Journal, 9, 77-95. --, NEVILL, W. E., RAMSBOTTOM, W. H. C. & SEVASTOPULO, G. D. 1985. Upper Dinantian stratigraphy and faunas of the Old Head of Kinsale and Galley Head, south County Cork. Irish Journal of Earth Sciences, 7, 47-58. --, REILLY, T. A., SEVASTOPULO, G. D. & SLEEMAN, A. G. 1983. Stratigraphy and structure in the Irish Variscides. In: HANCOCK, P. L. (ed.) The Variscan FoM Belt in the British Isles. Adam Hilger Ltd, Bristol. 20-46. --, SEVASTOPULO, G. D. & SLEEMAN, A. G. 1989. Subsidence history of the South Munster Basin, -

-

Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation. Yorkshire Geological Society Occasional Publication, 6, 99-109. -& REILLY, T. A. 1981. The Variscan Fold Belt in Ireland (The Variscan Orogen in Europe: final report of the ICG Study Group). Geologie en Mijnbouw, 60, 49-66. NOLAN, S. C. 1989. The style and timing of Dinantian syn-sedimentary tectonics in the eastern part of the Dublin Basin, Ireland. In: ARTHURTON, R. S., GUTI'ERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation. Yorkshire Geological Society Occasional Publication, 6, 83-97. PHILCOX, M. E. 1989. The mid-Dinantian unconformity at Navan, Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation. Yorkshire Geological Society Occasional Publication, 6, 67-81. PHILLIPS, W. E. A. & SEVASTOPULO,G. D. 1986. The stratigraphic and structural setting of Irish mineral deposits. In: ANDREW C. J., CROWE, R. W. A., FINLAY, S. PENNELL, W. M. & PYNE, J. F. (eds) Geology and Genesis of Mineral Deposits in Ireland. Irish Association for Economic Geology, 1-30. REEVES, T. J., ROBINSON, K. W., & NAYLOR, D. 1978. Ireland's offshore geology. Irish Offshore Review, May 1978, 25-28. ROBINSON, K. W., SHANNON, P. M. & YOUNG, D. G. G. 1981. The Fastnet Basin: an integrated analysis. In: ILLING, G. D. & HOBSON, G. D. (eds) The Geology of the Continental Shelf of North-West Europe. Heyden & Son, London, 444-454. SEVASTOPULO, G. D. & NAYLOR, D. 1981. Erratic Carboniferous boulders in Bantry Bay, County Cork. Geological Survey of Ireland Bulletin, 3, 79-84. SLEEMAN, A. G. 1986. Logs of boreholes drilled by Inco/Oliver Prospecting Ltd., through the Dinantian Limestones of the Crookstown District, Co. Cork. Geological Survey of Ireland Unpublished Report. 1987. The Dinantian and Namurian rocks of the Ringabella and Cloyne Synclines, west of Cork Harbour, County Cork. Geological Survey of Ireland Bulletin, 4, 67-76. --, PRACHT, M., DALY, E., FLEGG, m. M., O'CONNOR, P. & WARREN, W. P. 1994. Geology of South Cork. A Geological Description of South Cork and Adjoining Parts of Waterford to accompany the Bedrock Geology 1:100000 Scale Map Series, Sheet 25, South Cork. Geological Survey of Ireland. , THORNBURY, B. & SEVASTOPULO, G. D. 1986. The stratigraphy of the Courceyan (Carboniferous: Dinantian) rocks of the Cloyne Syncline, west of Cork Harbour. Irish Journal of Earth Sciences, 8, 21-40.

SOUTH M U N S T E R BASIN TURNER, J. S. 1939. The Upper Devonian and Carboniferous rocks of the Cork District. Proceedings of the Geologists" Association, 50, 319-323. - - 1 9 5 2 . The Lower Carboniferous rocks of Ireland. Liverpool and Manchester Geological Journal, l, 113-147. WILLIAMS, E. A., BAMFORD, M. L. F., COOPER, M. m., EDWARDS, H. E., FORD, M., GRANT, G. G., MACCARTHY, I. A. J., MCAFEE, A. M. &

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O'SULLIVAN, M. J. 1989. Tectonic controls and sedimentary response in the Devonian-Carboniferous Munster and South Munster basins, southwest Ireland. In: ARTHURTON, R. S., GUTTERIDGE, P. & NOLAN, S. C. (eds) The Role of Tectonics in Devonian and Carboniferous Sedimentation. Yorkshire Geological Society Occasional Publication, 6, 123-141. WINGF1ELD, R. T. R. 1968. The Geology of Kenmare and Killarney. PhD Thesis, University of Dublin.

Siliceous rocks of the Culm Basin, Germany HANS-JI~IRGEN GURSKY

Geologisch-Pal~iontologisches Institut der Technischen Hochschule Darmstadt, Schnittspahnstrasse 9, D-64287 Darmstadt, Germany

Abstract: The classic Dinantian cherts of the Rhenish Massif and Harz Mountains, Germany, were analysed for their petrography and sedimentology. They occur in five lithostratigraphic units: the Lower Alum Shale Formation, the Black Chert Formation, the Pale Chert Formation, the Siliceous Limestone Formation, and the Siliceous Transitional Formation (upper Hastarian to Asbian, Cull to CulIIo~ of the German cephalopod zonation). Mafic volcanic rocks are locally associated with these strata. Four regional facies zones have been defined. Radiolarian chert, spiculitic chert, homogeneous chert, and silicified quartz-trachytic tephra and tuffite are the main chert types. These mostly form current-laminated grey, black, greenish or reddish beds rhythmically alternating with layers of siliceous mudstone and, in places, variably intercalated with grey and black mudstones and siltstones, phosphorites, metabentonites, limestone turbidites, greywackes and quartz arenites. The siliceous rocks were deposited in a moderately deep (upper bathyal) strait of an equatorial sea that was part of palaeo-Tethys. Westward-driven nutrient-rich currents from the palaeo-Pacific Ocean favoured high radiolarian fertility. Relatively pure biosiliceous ooze was deposited under temporarily anoxic conditions when, due to dry climate and high sea-level, detrital input decreased. Deposition of the siliceous facies ended with increased detrital input and breakdown of the ocean-current system due to the collision of Laurasia and Gondwana during the Variscan Orogeny.

Siliceous sedimentary rocks are widespread in the Palaeozoic rocks of central Europe. The most prominent occurrences are the bedded chert successions of Silurian, Devonian and Dinantian age in the Franconian Forest/ Thuringian Slate Mountains, Vosges Mountains, Rhenish Massif, and Harz Mountains. In contrast to the Silurian and Devonian occurrences, Dinantian chert exposures are relatively continuous because of the weak nature of the Variscan tectonic deformation; they have been lithostratigraphically and biostratigraphically well studied (e.g. Bischoff 1957; Nicolaus 1963; Braun 1990; Braun & Schmidt-Effing 1993). The Dinantian cherts ('Kulm-Kieselschiefer') thus represent an important stratigraphic marker horizon in the German Upper Palaeozoic which has been recognized since the beginning of the 20th century (Denckmann 1909). However, due to their monotonous appearance, few studies before the beginning of the last decade addressed in detail the sedimentology, petrology and origin of Dinantian siliceous rocks in central Europe (e.g. Schwarz 1928; Hoss 1957; E1 Tarabili 1962). Therefore the cherts were not well understood genetically and were considered problematic, anomalous rocks (Schwan 1952). Based on the enormous progress in research on

siliceous sedimentary rocks triggered by the Deep Sea Drilling Programme (e.g. Rad 1979; Jenkyns & Winterer 1982; Hein & Parrish 1987; Gursky 1988), sedimentological and petrological aspects of Dinantian cherts in Germany have been studied by several modern authors (e.g. Hausmann 1983; Jackson 1985; Noeltner 1986; Kubanek & Zimmerle 1986; Dehmer et al. 1989; overview in Braun & Gursky 1991). The present paper summarizes some results of an extensive comparative study on the sedimentology, facies, petrography and chemistry of Dinantian siliceous rocks in Germany (Gursky 1992).

Occurrence and geological setting Rock sequences of Dinantian age are present in all major units of the Variscan basement in Germany (Fig. 1). They include shallow-marine platform limestones in the western Rhenish Massif ('Kohlenkalk'='Carboniferous Limestone'), relatively deep-marine shales, terrigenous and calcareous turbidites, mafic volcanic units, as well as altered tufts in the eastern Rhenish Massif and the western Harz Mountains ('Kulm'), and large olistostromes in the eastern Harz Mountains.

From STROGEN, P., SOMERVILLE, I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 303-314.

304

H.-J. GURSKY 50

and very low grade regional metamorphism due to the Variscan Orogeny. Less deformed Dinantian siliceous rocks have also been found by drilling in areas covered by post-Variscan rocks (e.g. Wolburg 1963).

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In the German Culm basin, siliceous and associated non-siliceous sedimentary rocks were mostly deposited from Ivorian to Asbian time (approximately Cu lift to Cu IIIc~ of the German cephalopod zonation). The standard stratigraphic succession (Fig. 2) of the 'Kulm' in Germany was first described by Denckmann (1909) from the northern rim of the eastern Rhenish Massif. Later studies showed its validity for most areas of the eastern Rhenish Massif and western Harz Mountains. Different successions are observed in the transition zone to the Carboniferous Limestone shelf, close to palaeogeographic highs and in the H6rreGommern Zone (Fig. 3). In most areas, the siliceous rock-bearing sequences are underlain by Upper Devonian to lowermost Dinantian basinal mudstones and siltstones with intercalated sandstone turbidites,

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Siliceous sedimentary rocks are of major importance in the basinal sequences of the eastern Rhenish Massif and the western Harz Mountains (Fig. 3). In these areas, the rocks show different degrees of folding, cleavage, faulting,

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SILICEOUS ROCKS, C U L M BASIN

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Fig. 3. Zones of siliceous facies of the Culm Basin in the eastern Rhenish Massif and Harz Mountains, with selected representative lithostratigraphic sections (modified after Gursky 1992). Section localities: (1) Zippenhaus quarry (abandoned) near Neviges; (2) Muensterland I borehole near Muenster (modified after Wolburg 1963); (3) Rottenberg quarry (abandoned) near Adorf; (4) Galgenberg railway cutting near Herborn; (5) new road cutting Harzstrasse near Lerbach.

306

H.-J. GURSKY

or by nodular limestones (Hangenberg Formation and its equivalents). The first siliceous mudstone and chert beds occur in the upper part of the Lower Alum Shale Formation ('Liegende Alaunschiefer'). This formation, made up of black mudstones and siltstones, is up to 45 m thick, relatively monotonous and poorly stratified. The rocks are rich in organic carbon and pyrite, but poor in both macro- and microfauna. However, abundant early-diagenetic phosphorite concretions are rich in well-preserved radiolarians (Braun 1990), signifiying that the Lower Alum Shale Formation as a whole was originally rich in biogenic silica. The Lower Alum Shale Formation passes up into the Black Chert Formation ('Schwarze Kieselschiefer') with an increase in silica content. The Black Chert Formation is up to 20m thick and consists of a rhythmic alternation of cm to dm-thick black radiolarian chert beds and mm-thick black mudstones. Minor intercalations of metabentonites and calciturbidite beds are locally present. The content of organic carbon decreases upsection, and the black chert-mudstone alternation grades into the Pale Chert Formation ('Helle Kieselschiefer'). The Pale Chert Formation is up to 25 m thick and consists of alternating cm to dm-thick grey, greenish and reddish chert beds with thin mudstones. The chert beds consist of radiolarites, spiculitic cherts, homogeneous cherts and silicified tufts. Metabentonite interbeds are typical. In the northeastern Rhenish Massif, the Pale Chert Formation grades laterally into the Siliceous Limestone Formation ('Kulm-Kieselkalke') by intercalation of partly-silicified limestone turbidites. In the Siliceous Limestone Formation, which is generally up to 30 m thick, calcareous beds may be dominant. The siliceous facies passes gradually into the over-lying Siliceous Transitional Formation ('Kieselige Ubergangs-Schichten'), a thin-bedded alternation of grey to black mudstones, siltstones, minor cherts, limestones and metabentonites which is up to 15 m thick. The chert-bearing formations are overlain by the Culm Slate Formation ('Kulm-Tonschiefer'), generally more than 100 m thick, which is mostly composed of monotonous grey basinal mudstones and siltstones. It grades upwards into the diachronous Culm Greywacke Formation ('Kulm-Grauwacken'), several hundred metres thick, which represents the flysch facies of the Culm Basin. At the northern rim of the Rhenish Massif and north of it, the flysch facies is

followed by Upper Carboniferous molasse facies including coal measures, several thousand metres thick in all.

Stratigraphic and facies variations The chert-bearing formations show a number of regional variations which reflect the palaeogeography (Fig. 3). The thicknesses of the formations are highly variable. Individual ranges are: Lower Alum Shale Formation 0-45m, Black Chert Formation 0-20m, Pale Chert Formation 0-25m, Siliceous Limestone Formation 0 - > 150 m, Siliceous Transitional Formation 0-25m. The proportion of individual formations in the chert-bearing succession is also highly variable (Fig. 3). Furthermore, the proportion of limestone turbidites in the individual sections is variable because of differing distances from shelf areas (Carboniferous Limestone platform) and intrabasinal highs (Devonian reef ruins, submarine volcanoes, etc.) where carbonate was produced in Dinantian times. In the northwestern part of the eastern Rhenish Massif, the entire siliceous facies of the Dinantian is gradually replaced by thick-bedded, proximal limestone turbidites supplied from the carbonate shelf (section no. 1 of Fig. 3; cf. Franke et al. 1975). To the east, the uppermost part of the Siliceous Transitional Formation and the lower part of the Culm Slate Formation are locally replaced by a thickbedded sequence of limestone turbidites of the Culm Platy Limestone Formation ('KulmPlattenkalk'; cf. Eder et al. 1983). Along the eastern rim of the Rhenish Massif, the carbonate-free Pale Chert Formation grades laterally to the west into the Siliceous Limestone Formation by increasing intercalations of distal limestone turbidites shed from intrabasinal highs (Witten 1979). Carbonate debris flows are locally present close to such highs. In the southeastern Rhenish Massif and in the western Harz Mountains there occurs a lithologically variable series, up to 500 m thick, of extrusive and intrusive mafic volcanic rocks ('Deckdiabas' or Dinantian Diabase). The sills and dykes intrude all formations up to the Pale Chert Formation (section no. 4 of Fig. 3). In sections where diabase flows underlie the Lower Alum Shale Formation and the Black Chert Formation, the thicknesses of these formations may be reduced (sections no. 4 and 5 of Fig. 3). The amount of altered, partly silicified, finegrained tephra (now chert beds and metabentonite layers) is regionally variable. It increases

SILICEOUS ROCKS, CULM BASIN towards the east and southeast and reaches maxima in sections of the western Harz Mountains (section no. 5 of Fig. 3). The Elbingerode Complex, a Devonian-Dinantian volcanic centre in the north-central Harz Mountains, is suspected to have supplied most of the tephra beds in the German Culm Basin. The standard regional succession (Fig. 2) is not developed in parts of the southeasternmost Rhenish Massif and the central Harz Mountains. In the H6rre-Gommern Zone especially, and some adjacent areas, Upper Devonian to Dinantian mudstones and siltstones and siliceous rocks are associated with greywackes, mature quartz arenites and limestone turbidites (Bender 1989; Herbig & Bender 1992).

Regional facies zones Based on these regional variations, four major facies zones are distinguished below.

The Bergian Zone. The Bergian Zone includes rocks of the transition from the Carboniferous Limestone shelf into the German Culm Basin. The individual sections are composed of a variable proportion of proximal limestone turbidites and basinal pelagic rocks (section no. 1 of Fig. 3). Siliceous rocks are subordinate. The Westphalian Zonel In this zone the standard succession (Fig. 2) is best developed and includes siliceous rocks that range in age from Ivorian to Asbian (CuIIfl to CuIIIc~). In places, siliceous rocks and black shales of late Brigantian age (CuIII-,/) are present (Korn 1989). The Deckdiabas is absent. The Siliceous Limestone Formation is well developed and the Siliceous Transitional Formation is lithologically variable.

307

rocks. The sections only partly correspond to the succession in the Dill-Innerste Zone (cf. Bender et al. 1993). In the central and eastern Harz Mountains, the Dinantian is mostly represented by large olistostromes (e.g. Lutzens 1972) in which siliceous rocks of Dinantian age are absent.

Sedimentology Dinantian siliceous rocks in Germany include radiolarian cherts, spiculitic cherts, homogeneous cherts, and silicified tufts and tuffites. These rock types variably alternate with mudstones and siltstones, greywackes, quartz arenites, metabentonites, limestones, phosphorite nodules and layers, non-bedded quartz-hematite rocks, and stratiform manganese mineralizations (mostly Mn-carbonate and oxides). Rhythmic ('dyscyclic') bedding is the most conspicuous macroscopic feature; hard, splintery chert beds alternate with thinner and softer pelitic interlayers poorer in SiO2 (Fig. 4). In most sections, this basic rhythm is interrupted by intervals of rhythmic intercalations of metabentonite beds and limestone turbidites. In places, the limestone turbidites may be dominant. Several types of alternations occur (Gursky 1992). The most important macroscopic sedimentary structures in the siliceous beds are well-defined and even bedding surfaces with rare deep-water ichnofauna, graded top contacts of some tephra and limestone-turbidite beds with overlying cherts, mostly tabular bed

The Dill-Innerste Zone. Here the Deckdiabas is present. The Lower Alum Shale Formation and Black Chert Formation may be missing or substituted by diabase. The Pale Chert Formation is mostly free of carbonate turbidites and is locally variegated. The Siliceous Transitional Formation is poorly developed. The Lahn-Bode Zone. This zone, including the H6rre-Gommern Zone, has not been studied in detail and lies outside the scope of the present paper. In the southeasternmost Rhenish Massif, lithologically and lithostratigraphically variable sections are present and include units of siliceous

Fig. 4. Typical outcrop of the Dinantian siliceous rocks (Pale Chert Formation) in the Huettenteich roadcut at Lerbach, western Harz Mountains, with rhythmic bedding of fractured chert beds (mostly silicified tephra and radiolarites) and weathered interlayers (mostly siliceous mudstones and siltstones, some metabentonites).

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geometry, variable parallel lamination (Fig. 5), minor homogeneous ('structureless') beds, small-scale load casts and convolute bedding, and smooth pinch-and-swell structures. Small-scale lamination occurs in most chert beds and is mostly due to weak primary variations in biogenic and clastic sediment supply ('thinly interlayered bedding' of Reineck & Singh 1980). Thus, the fundamental sedimentary units of the siliceous rocks are the laminae. They are interpreted as the result of various deep-marine, low-energy, non-erosive bottom and suspension currents, and possible annual variations of sediment input (cf. Reineck & Singh 1980; Stow & Piper 1984). The existence of low-energy currents is indicated by weak normal grading and micro cross-lamination within laminae, concentrations of sponge spicules, and alternations of laminae rich in radiolarians and in quartz silt. Macroscopic rhythmic bedding is a classical feature in many Phanerozoic chert sequences. It reflects continuous oscillations of the silica-mud ratio which may be due to cyclical variations of bioproduction, dilution by fine-grained terrigenous detritus, and/or silica dissolution (Jenkyns & Winterer 1982; Decker 1991). Dilution cycles

result in thick mudstone interbeds (Einsele & Ricken 1991) and, in contrast to carbonate sediments, dissolution cycles are insignificant in siliceous sediments owing to the strong permanent undersaturation of sea water with respect to silica. Therefore, it is suggested that cyclic fluctuations of radiolarian production in the near-surface water layer were the main cause for the rhythmic bedding in these Dinantian siliceous rocks. These fluctuations may have resulted from orbitally-driven climatic oscillations. Approximately 25 m of net siliceous material is present in representative sections. This material was deposited during approximately 14Ma (main phase of siliceous sedimentation in the German Culm Basin). The average thickness of the chertmudstone couplets is 5 cm. Based on these data, a post-compaction sedimentation rate of 1.8 mm/a and a time span of 28 ka is indicated for each chert-mudstone couplet. Such a periodicity is of the same order as one component of the Milankovitch cycles of obliquity/tilt of the earth's axis (cf. Einsele & Ricken 1991).

Petrography Radiolarian cherts

Fig. 5. Typical lamination of a black radiolarite bed with alternating laminae of radiolarian-rich material and organic carbon-rich pelite; crossed polarizers. Radiolarians are filled by microquartz aggregates (white), cryptocrystalline quartz-pigment mixtures (dark), spherulitic chalcedony (bright and dark semicircular structures) and opal-CT lepispheres replaced by microquartz (tiny white roundish specks; e.g. within the lowermost radiolarian test).

The dominant radiolarian cherts are characterized by radiolarian contents of up to 60%, rarely more. They are composed of quartz/chalcedony (radiolarians, sponge spicules, bioclasts, silica cement), siliciclastic and authigenic minerals (quartz silt, feldspar, phyllosilicates), volcanielastic fragments (quartz, feldspar, mica, altered glass shards), organic carbon, metal oxides and sulphides (especially pyrite and hematite; Fig. 6) as well as heavy minerals and crystals of calcite, dolomite and apatite. Several microlithotypes are distinguished according to their contents of non-siliceous constituents (especially organic carbon, silicates, hematite, calcite) and diagenetic and metamorphic grain growth of quartz ('microquartzitic' structures; Fig. 7). The radiolaria have been studied by Braun (1990), Braun & Schmidt-Effing (1993) and others, mostly by scanning electron microscope (SEM) techniques. In thin section, specific determinations are generally impossible. However, mostly circular cross-sections up to 300 #m in diameter, internal skeletal elements, porous structures, spines and spine fragments up to 200 #m long and 60 #m thick are typical (Fig. 6). The degree and type of radiolarian preservation are both variable (Gursky 1988, 1990, 1992;

SILICEOUS ROCKS, CULM BASIN

309

Fig. 6. Extraordinary radiolarian preservation in reddish hematitic radiolarite bed; parallel polarizers. Radiolarians are mostly replaced by early diagenetic hematite (dark); microquartz is white. Gursky & Gursky 1988). Preservation depends mostly on diagenetic and metamorphic grain growth and selective dissolution, and ranges from excellent (Fig. 6) to poor (Fig. 7); extreme grain growth results in complete obliteration of the radiolarians and the formation ofmicroquartzitic textures (for details see Gursky 1992; Gursky & Gursky 1988). The type of radiolarian preservation is defined by diagenetic mineralogy and the crystal size of the replaced skeletons and their infillings. The skeletons are made up of predominantly quartz/chalcedony (<5-30#m, Fig. 5) with minor contributions of pyrite, hematite (Fig. 6), chlorite, calcite, dolomite and carbon. The test infillings consist of cryptocrystalline to microcrystalline quartz-pigment mixtures, quartz aggregates and mosaics, spherulitic chalcedony, chlorite, pyrite, hematite, calcite, carbon, and clay minerals. Some radiolarian infillings include

~

r Fig. 7. Diagenetically 'recrystallized' radiolarian chert with relatively coarse-grained microquartz (microquartzite texture) and few preserved radiolarians; crossed polarizers.

Fig. 8. Pressure solution in a laminated radiolarian chert bed; parallel polarizers. The middle lamina is especially affected by horizontal microstylolite bands due to slight enrichment in very fine-grained siliciclasticdetritus (mostly phyllosilicates). Microquartz-filled radiolarians are dissolved at upper and lower contacts with microstylolites.

early diagenetic cristobalite-tridymite (opal-CT) lepispheres replaced by microquartz (Fig. 5). Pressure solution resulted in characteristic microstylolitic banding subparallel to bedding (Fig. 8) and has accentuated the bedding surfaces and the typical macroscopic 'ribbon chert' fabric of the siliceous rocks. The present mineralogy of the radiolarian cherts, their texture and many of the sedimentary structures are diagenetic. However, microfauna and siliciclastic admixtures indicate the existence of an original radiolarian ooze 'contaminated' by fine-grained terrigenous detritus. During early diagenesis, the ooze containing radiolarian skeletons that had resisted dissolution during pelagic settling, was compacted. Phosphorite nodules grew during early diagenesis, in horizons characterized by anoxic conditions. The biogenic silica of the skeletons ('opal A') was partly dissolved and transformed to opal-CT (cf. Rad 1979). Silica cement (as opal-CT) formed by precipitation from pore waters, which resulted in lithification as porcellanite (opal-CT rock). Thermal metamorphism occurred at hot diabase contacts. During late diagenesis, opal-CT was transformed to quartz/ chalcedony and 'mature' quartz cherts were formed. Pressure solution and regional metamorphism followed.

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Spiculitic cherts Spiculitic cherts make up only a few per cent of the siliceous rocks and occur in the Pale Chert Formation and the Siliceous Limestone Formation. Individual spiculitic beds are rare and most spiculitic cherts form laminae within beds of radiolarian chert. Such spiculitic laminae may represent the tails of calciturbiditic currents (Herbig & Mamet 1994). Several microlithotypes are present, among which fine-grained, carbonate-free radiolarian-bearing spiculite is dominant. Diagenetic alteration of spiculitic and radiolarian cherts was similar.

Homogeneous cherts Homogeneous Cherts are pale-coloured and very fine-grained with less than 1% of microscopically identifiable non-siliceous components (mostly phyllosilicates). Compositionally they grade into siliceous mudstone. Typical microlithotypes include homogenized radiolarite (only a few radiolarian 'ghosts' visible), strongly silicified mudstones and siltstones, strongly silicified fine-grained tufts and tuffites, and extremely fine-grained cherts of unknown origin.

Silicified tephra Tephra layers are abundant in the Dinantian of the German Culm Basin. They are present as soft, weathered metabentonite beds reported by many earlier authors (e.g. Hoss 1957; Dehmer et al. 1989), and as massive chert beds. Many chert beds reveal their volcaniclastic origin only in thin section. They consist of silicified volcanic glass shards, fragments of alkali feldspar, plagioclase, quartz (with partly resorbed margins), biotite and heavy minerals; the groundmass is a cryptocrystalline to microcrystalline quartz cement. The microlithotypes include silicified vitric tufts (Fig. 9), silicified crystal tufts, silicified tuffites, and thermometamorphic microquartzitic tufts. Parallel lamination and graded bedding are typically present. Such silicified tufts are common in the Pale Chert Formation of the southeastern Rhenish Massif and the western Harz Mountains. These tufts are interpreted as water-lain fallout tephra resulting from subaerial eruptions of trachytic to quartz-trachytic composition (cf. Dehmer et al. 1989; Kubanek & Zimmerle 1986). The tufts were chemically unstable in the marine environment, and early diagenesis resulted in alteration and compaction of unstable

Fig. 9. Silicifiedquartz-trachytic tuff (massive chert bed from the Pale Chert Formation) mostly composed of microquartz-replaced glass shards cemented by microquartz; parallel polarizers.

silicate fragments. Glass shards were hydrated and oxidized, part of the glass was replaced by and cemented by silica, albitization of feldspars occurred, and some minerals were completely dissolved (e.g. augite, hornblende). The rest of the glass and unstable crystal fragments were possibly transformed to smectite and zeolites (cf. Iijima et al. 1980). As the result of subsequent regional metamorphism and weathering, the final mineralogy of the metabentonites is composed of quartz, illite, mixed-layer clay minerals, and chlorite (Hoss 1957). Many tephra beds, however, were further affected by late-diagenetic silicification due to silica supplied from adjacent biogenic siliceous rocks. Most of the silicate fragments were replaced and cemented by silica. This process

Fig. 10. Coarse-grained base of a limestone-turbidite bed from the Siliceous Limestone Formation. The echinoderm bioclasts are cemented and marginally replaced by microquartz imported from adjacent biosiliceous beds. Crossed polarizers.

SILICEOUS ROCKS, CULM BASIN

311

The late Hastarian (CuI/CulI boundary) was characterized by a basin-wide change from grey to black sediments. This onset of deposition of the Lower Alum Shales reflects the maximum of a rapid sea-level rise that started in the uppermost Devonian (Ross & Ross 1987; Fig. 2). This sea-level rise enlarged the biologically productive shelf areas. Increased production of organic carbon, reworking of nutrient-rich soils, and restriction of intrabasinal circulation resulted in increased deposition of organic carbon and, consequently, in anoxic conditions in the deeper basin. In the late Tournaisian and early Vis6an, the regional climate changed from semi-arid to arid (Zwan et al. 1985; Wright 1990; Peeters et al. 1992). Thus, the elastic supply from the hinterlands decreased, and owing to the wide shelf areas, only small amounts of terrigenous detritus reached the Culm Basin. Reduction of dilutant siliciclastic material and high radiolartan fertility in the near-surface sea water resulted in the transition into the siliceous oozes of the Black Chert Formation under anoxic, starvedbasin conditions. Anoxic conditions gradually ceased around the time of the Chadian/Arundian boundary (Cull"/) and the oozes of the Pale Chert

transformed many tephra layers to massive quartz chert (Fig. 9). Some bioclastic carbonate detritus in the limestone turbidites was likewise cemented and partly replaced by silica- the formation of 'siliciceous limestones' (Fig. 10). These observations, and the fact that the diagenesis of tephra releases only insignificant quantities of silica (e.g. Rad 1979), indicate that nearly all of the silica stored in the Dinantian rocks of Germany is biogenic in origin.

Palaeoceanographic interpretation The central European Culm Basin was an equatorial sea strait adjacent to a branch of the Dinantian palaeo-Tethys, situated off the passive southeastern continental margin of Laurussia (Fig. 11). The southeastern boundary of the basin was a subduction-zone plate margin related to the Variscan Orogeny, which resulted in the closure of the basin between Laurussia and Gondwana in the Late Carboniferous. Based on the circulation patterns of the modern oceans, Fig. 11 depicts probable major surface currents in Dinantian time. NE and SE trade winds probably caused nutrient-rich westbound surface currents between Laurussia and Gondwana.

I

=\

"-..... ,__-s, •

~__ --":'_7.~

/~ ::;:::::::::::

f;.":---:--7.

::::

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L__,

currents retotive H/L pressure zone

with surfnce winos

upwetUng

Dinontion

~ ~

chert occurrences cherts of the [uLm bosin

Fig. l l. Palaeocontinental world map of the Dinantian depicting major chert occurrences, ocean currents, upwelling areas and high (H) and low (L) pressure zones (north winter). The central European Culm Basin is represented by a star. Modified after Dietrich & Ulrich (1968), Parrish (1982), Hein & Parrish (1987) and Scotese & McKerrow (1990).

312

H.-J. GURSKY

Formation were deposited. The greater proportion of volcaniclastic layers (chert and metabentonite beds) indicates a maximum of subaerial tephra eruptions, probably in the Elbingerode area (Harz Mountains). In the Bergian and Westphalian Zones, carbonate turbidites were shed from the shelf and intrabasinal highs (Siliceous Limestones). These turbidites may reflect increased carbonate production due to sea-level highstands; however, sea-level variations during the time of deposition of the Black Chert Formation and Siliceous Limestone Formation are still difficult to interpret (Herbig & Bender 1992; Bender et al. 1993). During the early Asbian (CuIIIc0, sea level dropped temporarily and the climate became more humid; the siliciclastic input into the basin and the sedimentation rate increased (cf. Leeder 1987) and the Siliceous Transitional Formation was deposited. However, the main cause for the gradual cessation of siliceous sedimentation was tectonic. Hurley & Van Der Voo (1987) gave palaeomagnetic evidence that the onset of the collision between Laurussia and Gondwana occurred in mid-Vis~an time. The ocean current system in the European branch of palaeo-Tethys became gradually less vigorous, radiolarian biomass production decreased, and strongly increased clastic input from the nascent Variscan orogen in the southeast resulted in a high rate of flysch deposition in the narrowing trough (Culm Slate Formation and Culm Greywacke Formation).

Conclusions

The phase of biosiliceous sedimentation in the German Culm Basin was not extraordinary. It was the logical consequence of palaeogeographical and palaeoceanographical conditions which persisted from Ivorian to Asbian time (CuII-CuIIIo 0. Conditions were characterized by high radiolarian fertility combined with reduced siliciclastic input, and the absence of competitive high-production carbonate plankton. Thus, the Dinantian cherts are another example, modified by regional factors, of the typical bedded radiolarian chert sequences which characterize many deep to moderately deep-water continental margin successions deposited from Ordovician to Cretaceous times. Field and laboratory studies as well as evaluations have been supported by many people and institutions. I gratefully acknowledge a research grant from the Deutsche Forschungs-Gemeinschaft (Gu 289/2-1 ). Special thanks are given to M. Amler, P. Bender, A. Braun,

W. Franke, H.-G. Herbig, M. Horn, H. Knappe, E. Paproth, R. Schmidt-Effing, S. Schfiffler, M. Schwab, O. Tietz, E. Thomas, H. Weller, D. Weyer, A. Willner, H. Zankl and W. Zimmerle for information on outcrops, discussions, and field trips. Technical support was mostly provided by the Institut ffir Geologie und Pal/iontologie of the Philipps-Universitfit Marburg. Additional technical assistance came from M. Dukat, M. Gursky, and E. Wettengl. M. Gursky, S. Kempe, and D. Schumann made helpful comments on an early version of the manuscript. The paper was constructively reviewed by J. Hein and H.-G. Herbig.

References

ARBEITSGRUPPE FUER DINANTIUM-STRATIGRAPHIE. 1971. Unter-Karbon (Dinantium). Fortschritte der Geologie des Rheinlandes und Westfalens, 19, 5-18. BENDER, P. 1989. Die Hoerre und ihre Stellung im oestlichen Rheinischen Schiefergebirge. Jahresberichte und Mitteilungen des oberrheinischen geologischen Vereins, Neue Folge, 71,347 356. --, HERBIG, H.-G., GURSKY, H.-J. & AMEER, M. R. W. 1993. Beckensedimente im Oberdevon und Unterkarbon des oestlichen Rheinischen Schiefergebirges. Fazies, Pal/iogeographie und Meeresspiegelschwankungen. Geologica et Palaeontologica, 27, 332-355. BISCHOFF, G. 1957. Die Conodonten-Stratigraphie des rhenoherzynischen Unterkarbons mit Beruecksichtigung der Wocklumeria-Stufe und der Devon/Karbon-Grenze. Abhandlungen des hessischen Landesamtes ffir Bodenforschung, 19, 1-64. BRAUN, A. 1990. Radiolarien aus dem Unter-Karbon Deutschlands. Courier Forschungs-Institut Senckenberg, 133, 1 177. - & GURSKY, H.-J. 1990. Kieselige Sedimentgesteine des Unter-Karbons im Rhenoherzynikum eine Bestandsaufnahme. Geologica et Palaeontologica, 25, 57-77. & SCHMIDT-EFFING, R. 1993. Biozonation, diagenesis and evolution of radiolarians in the Lower Carboniferous of Germany. Marine Micropaleontology, 21, 369 383. DECKER, K. 1991. Rhythmic bedding in siliceous sediments - an overview. In: EINSELE, G., R[CKEN, W. & SEILACHER, A. (eds) Cycles and Events in Stratigraphy. Springer Verlag, Berlin, 464-479. DEHMER, J., HENTSCHEL,G., HORN, M., KUBANEK, F., NOELTNER, T., RIEKEN, R., WOLF, M. & ZIMMERLE, W. 1989. Die vulkanisch-kieselige Gesteinsassoziation am Beispiel der unterkarbonischen Kieselschiefer am Ostrand des Rheinischen Schiefergebirges. Geologie-PetrographieGeochemie. Geologisches Jahrbuch Hessen, 117, 79-138. DENCKMANN, A. 1909. Ueber eine Exkursion in das Devon- und Culmgebiet noerdlich von Letmathe. Jahrbuch der koeniglich-preussischen geologischen Landesanstalt, 27, 20-47.

S I L I C E O U S R O C K S , C U L M BASIN DIETRICH, G. & ULRICH, J. (eds) 1968. Atlas zur Ozeanographie. Bibliographisches Institut, Mannheim. EDER, F. W., ENGEL, W., FRANKE, W. & SADLER, P. M. 1983. Devonian and Carboniferous limestone turbidites of the Rheinisches Schiefergebirge and their tectonic significance. In: MARTIN, H. & EDER, F. W. (eds) lntracontinental Fold Belts. Springer Verlag, Berlin, 93-124. EINSELE, G. & RICKEN, W. 1991. Limestone-marl alternations - an overview. In: EINSELE, G., RICKEN, W. & SEILACHER, A. (eds) Cycles and Events in stratigraphy. Springer Verlag, Berlin, 23 47. EL TARABILI, E. S. 1962. Geologie des Devons und Kulms im Nordwest-Fluegel der Soese-Mulde (Oberharz) unter besonderer Beruecksichtigung der Petrographie der Kulmkieselschiefer. Roemeriana, 5, 1-115. FRANKE, W., EDER, W. & ENGEL, W. 1975. Sedimentology of a Lower Carboniferous shelf margin (Velbert anticline, Rheinisches Schiefergebirge, W Germany). Neues Jahrbuch fuer Geologie und Palaeontologie, Abhandlungen, 150-3, 314-353. GURSKY, H.-J. 1988. Gefuege, Zusammensetzung und Genese der Radiolarite im ophiolithischen NicoyaKomplex (Costa Rica). Muenstersche Forschungen zur Geologie und Palaeontologie, 68, 1-189. - - 1 9 9 0 . Radiolarian petrographic preservation types in Jurassic to Lower Tertiary cherts of Costa Rica. Marine Micropaleontology, 15, 249-263. - - 1 9 9 2 . Sedimentaere Dynamik und stoffliche Entwicklung kieseliger Sedimentgesteine im mitteleuropaeischen Unter-Karbon. Habilitation Thesis, University of Marburg. - - - & GURSKY, M. M. 1988. Thermal alteration of chert in the ophiolite basement of Southern Central America. In: HEIN, J. R. & OBRADOVIC, J. (eds) Siliceous Deposits of the Tethys and Pacific Regions. Springer-Verlag, New York, 217-233. HAUSMANN, R. 1983. Kieselsedimente unter besonderer Beruecksichtigung syndiagenetischer Gleitvorgaenge. PhD Thesis, University of Cologne. HEIN, J. R. & PARRISH, J. T. 1987. Distribution of siliceous rocks in space and time. In: HEIN, J. R. (ed.) Siliceous Sedimentary Rock-Hosted Ores and Petroleum. Van Nostrand Reinhold, New York, 10-57. HERBIG, H.-G. & BENDER, P. 1992. A eustatically driven calciturbidite sequence from the Dinantian II of the eastern Rheinisches Schiefergebirge. Facies, 27, 245-262. --& MAMET, B. 1994. Hydraulic sorting of microbiota in calciturbidites - a Dinantian case study from the Rheinische Schiefergebirge, Germany. Facies, 31, 93-104. HOSS, H. 1957. Untersuchungen ueber die Petrographie kulmischer Kieselschiefer. Beitraege zur Mineralogie und Petrographie, 6, 59-88. HURLEY, N. F. & VAN DER VOO, R. 1987. Paleomagnetism of Upper Devonian reefal limestones, Canning Basin, Western Australia. Geological Society of America Bulletin, 98, 138-146.

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The Apedale tufts, North Staffordshire: probable remnants of a late Asbian[Brigantian (Pla) volcanic centre J. G. REES, J. D. C O R N W E L L , Z. K. D A B E K & R. J. M E R R I M A N

British Geological Survey, Keyworth, Nottingham NG12 5GG, UK Astract: A borehole sunk in 1920-21 at Apedale, North Staffordshire, unexpectedly proved a tuff-dominated sequence, at least 840 m thick, below a thin cover of Silesian sedimentary rocks. The tufts were initially interpreted as Dinantian in age. However, there are similarities of form and lineament between an aeromagnetic anomaly centred upon Apedale and anomalies in the Welsh Marches associated with Neoproterozoic (Uriconian) rocks, suggesting that Neoproterozoic rocks occur at Apedale and that the tufts in Apedale Borehole might also be of this age. Recent modelling of geophysical data shows that Apedale is underlain by two (magnetic) volcanic bodies- a deeper ridge of probable Uriconian rock, and a shallower, broadly stratiform body which includes the tufts in the Apedale Borehole. Comparison of the tufts at Apedale with others in the region suggests that they are of Pla (late Asbian to early Brigantian) age. Although our knowledge of the extent of the Apedale tufts is poor, their thickness shows that they erupted from one of the largest centres of Dinantian volcanism in central England. The trace element geochemistry is very similar to that of Dinantian volcanic rocks in southwest England. During the First World War, imports of oil into the UK increased almost four-fold. Hence, in the following years the government funded the drilling of 11 boreholes in search of British Carboniferous reservoirs (Rowland & Cadman 1960; Torrens 1994). Two were drilled in North Staffordshire on anticlines bordering the Potteries Coalfield (Figs 1 & 2): one at Werrington, east of the coalfield, near which hydrocarbons

have been discovered (Besly 1993), and one west of the coalfield at Apedale, near Newcastleunder-Lyme (Giffard 1923). The Apedale Borehole (226m AOD; SJ 8074 4862) was positioned on Middle Coal Measures in the core of the Western Anticline of the Potteries Coalfield, a northwesterly verging, NNE-trending periclinal structure which may

'70

60 Cheshire

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~

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Fig. 2. Geological map of the area around Apedale.

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 345-357.

346 ~7o

J. G. REES E T AL. \ \ ",: a0~----"-'-~ 90 ~

00~

Fig. 3. Detailed aeromagnetic map of the area around Apedale (map covers the same area as Fig. 2). The lines show the positions of aeromagnetic profiles in Fig. 6. The aeromagnetic data were obtained by surveys flown in 1955 at 305 m mean terrain clearance, with E-W flight lines at a spacing of 1.61 km and N-S tie lines with a spacing of 9.66 km. Ticks towards magnetic lows.

be traced along the margin of the Cheshire Basin (Fig. 2). Following a first abortive attempt (Apedale No.1 Borehole), Apedale No.2 Borehole, hereafter referred to as the Apedale Borehole, was started in September 1920. It drilled through thin Westphalian and Namurian strata before entering a tuff sequence. By August 1921, when all hopes of finding oil had disappeared, the hole was abandoned and an unbottomed tuff-dominated sequence at least 847.3 m thick (down-hole) had been proved. At the time no study was made of the tufts (which probably represent a broad volcaniclastic spectrum from Surtseyan to Strombolian tufts, and reworked tufts), although a Dinantian age was suggested for them (Giffard 1923) because tufts of this age had been found at Astbury, 12 km N N E of Apedale (Gibson & Hind 1899; Hind 1904). This conclusion was adopted by Gibson (1925) who suggested that the tufts were underlain by Dinantian rocks; he interpreted the tufts to be only about 150m thick, but steeply disposed within the Western Anticline at Apedale. Neither Giffard nor Gibson gave any

Fig. 4. Shaded relief plot of the aeromagnetic data for the region between Apedale and the Welsh Marches. In this shaded relief plot the magnetic anomalies appear as a pseudotopography which have been illuminated from the southeast to enhance lineaments. The position of the Apedale Borehole is shown in Fig. 5. SOT, Stokeon-Trent; S, Stafford; W, Wolverhampton.

Fig. 5. Map of the same region as that of Fig. 4, showing the main geological features of the western microplate of the Midlands Microcraton. Location of the figure and the outline of the microcraton are shown in Fig. 1.

APEDALE TUFFS description of the composition or stratigraphy of the tufts. Within a few years the chippings from the borehole (one of the first to be completely rock-bitted) were lodged with the British Geological Survey (presently stored at Keyworth), and the tufts largely forgotten. The advent of regional aeromagnetic maps in the 1950s showed that a large magnetic anomaly is centred on Apedale (Fig. 3), indicating the probable presence of volcanic rocks at shallow depths. However, the aeromagnetic data initially suggested a pre-Dinantian age for the tufts. The maps show that the Apedale anomaly is not unique, but is the northernmost member of a set of anomalies trending NNE from the Welsh Marches (Fig. 4). The anomalies recognized in the Marches were shown by Brooks & Fenning (1968) and Wilson (1980) to be closely related to the distribution of near-surface Uriconian rocks. These consist of basaltic, basaltic-andesitic, dacitic and rhyolitic lavas and tufts (Pharaoh et al. 1987b) that were extruded in Neoproterozoic times, approximately between 680 and 560 Ma (Tucker & Pharaoh 1991). The Uriconian rocks occur within the western microplate of the Midlands Microcraton, a triangular area of ancient continental crust (Fig. 1; Pharaoh et al. 1987a). Where they are at, or near, surface it is thought that they may have been uplifted in flower structures generated along the margin of the microcraton during the Caledonian Orogeny (e.g. Lee et al. 1990). Outcrops are scattered over a wide area associated with the NNE-trending Church Stretton, Pontesford-Linley and Red Rock faults (Fig. 5). Susceptibility measurements on outcrops have subsequently confirmed the magnetic nature of several lithologies within the Uriconian suite, suggesting that they are the source of the high-amplitude aeromagnetic anomalies. Comparison between the aeromagnetic anomalies in the Welsh Marches and that centred upon Apedale suggested that shallow Uriconian rocks also lie below the latter, and that the tufts penetrated in the Apedale Borehole are not Dinantian but Neoproterozoic in age. This interpretation clearly contradicted that of Giffard (1923) and Gibson (1925). Consequently, during the recent revision of the 1 : 50 000 Sheet 123 and memoir for the Stoke-on-Trent district (Rees & Wilson in press) it was necessary for the British Geological Survey to clarify the stratigraphy, setting and origins of the Apedale tufts in view of their uncertain age. This was done using detailed aeromagnetic modelling, plus petrographical and trace element studies on chippings from the Apedale Borehole.

347

Aeromagnetic modelling of the Apedale anomaly The aeromagnetic data for the Newcastle-underLyme area were acquired in 1955 as part of the UK national coverage. The original analogue data were subsequently digitized, and this data set was used in the present study. The aeromagnetic map (Fig. 3) is based on data adjusted to a vertical magnetization direction ('reduced to pole') to centre the anomalies directly over the magnetic sources. The anomaly centred on Apedale (Figs 2, 3) closely follows the axis of the Western Anticline and parallels the Red Rock Fault. South of Apedale it is broader and, at its southwestern extremity, appears to follow the Red Rock Fault (Fig. 2). North of Apedale the anomaly is narrower and characterized by an abrupt change in gradient along its northwestern flank. To the northwest, the smooth, more widely spaced contours suggest deeply seated magnetic rocks, while the eastsoutheast elongated anomaly on the southeastern side indicates sources at an intermediate depth. The form of these anomalies suggests that magnetic source rocks occur at a wide range of depths. Depth estimates on the sharply defined central peak indicate that the shallowest magnetic rocks occur at about 0.4-1.0km. This evidence, and the confirmed presence of magnetic material in the chippings from the Apedale Borehole, indicates that the latter are responsible for at least the main anomaly maximum. Magnetic susceptibility measurements on the chippings show moderately high values of the order of 0.01 SI units, although the reliability of the results is reduced by the nature of the chippings. Measurements indicate that the intensity of the remanence is negligible compared with the induced component (P. B. Sadler pers. comm. 1997). On the basis of these susceptibility measurements, a thickness of volcanic rock measurable in hundreds of metres is likely to extend over an area of at least 100 km 2, the data being less sensitive to the effect of any additional deeper, or thinner, extensions. The extension of the anomaly to the ESE, for example, could reflect part of the volcanic complex. The volcanic rocks approach nearest the surface for 2 km, in a NE-SW direction, around the borehole site. Just northeast of the borehole the volcanic rocks appear to deepen (or thin) rapidly, but the anomaly peak remains more constant, suggesting that another culmination exists near Mow Cop (Fig. 2). North of this the anomaly disappears just to the southwest of the

348

J. G. REES E T AL.

Dinantian inlier at Astbury (Figs 2, 3), which contains lapilli tufts and tuff breccias (see below). The absence of an anomaly here, where preSilesian basement occurs at a shallower level than anywhere else in the area, suggests that at this locality magnetic rocks do not occur beneath the exposed tufts and tuff breccias. Quantitative modelling of the magnetic and gravity data along several profiles was carried out by Cornwell & Dabek (1993) using the GRAVMAG program (Busby 1987). This modelling attempted to determine whether the aeromagnetic anomaly at Apedale could be satisfactorily explained by: (a) deep magnetic rocks coming near to the surface locally, (b) by near-surface magnetic rocks, or (c) by a combination of the two. It is theoretically possible to model the anomaly solely using a deep magnetic source, which comes locally to intermediate depth, although it is not possible to make a realistic geological model in this manner. It is not possible to model the anomaly solely using shallow magnetic sources. The models that best fit the aeromagnetic and gravity data, and seem geologically reasonable, are those with magnetic bodies at two levels, one near the surface and the other at an intermediate depth (less than 10 km). The shallower body, comprising the Apedale Tufts, is broadly stratiform. Two profiles are illustrated in Fig. 6. Profile A-A' crosses the anomaly at the Apedale Borehole, where the interpreted maximum thickness of the Apedale Tufts is about 1 km. This estimate, whilst dependent upon the uncertain effect of the underlying magnetic basement, seems reasonable on geophysical grounds. However, the tufts appear to thin to the NW, towards the Were Fault, and seem to be truncated by the Apedale Fault (see also Fig. 2) to the SE. Profile B-B' lies about 6 km to the north and differs in the presence of the abrupt change in gradient at 8 km along profile, to the NW of the maximum anomaly, and a broad anomaly to the SE. The form of the northwestern flank could not be interpreted without extending the magnetic material at depth, and the preferred model, illustrated in Fig. 6, relates the faulted margin of the Apedale Tufts to the gradient change. The broader feature on the southeastern flank is believed to reflect an extension of the Apedale Tufts in that direction. The presence of magnetic basement at intermediate depth is still required to explain the part of the anomaly indicated by the smooth, more widely spaced contours (the long wavelength component).

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Fig. 6. Observed aeromagnetic profiles and models based on interpretations assuming finite strike extents. Locations of profiles A-A t and B-B' are shown in Fig. 3. Susceptibilities: Apedale tufts 0.02 SI, magnetic basement (Uriconian) 0.025 SI; remainder 0.00 SI. The thickness of Triassic and Carboniferous rocks is based on geological evidence. The Carboniferous rocks shown are mostly Silesian; the base of the Dinantian sequence is uncertain. BH, Apedale Borehole, WF, Wem Fault, RRF, Red Rock Fault, AF, Apedale Fault.

Given the similarity of the orientation and size of the aeromagnetic anomaly at Apedale with those associated with the Uriconian rocks in the Welsh Marches, there seems little doubt that the magnetic rocks at intermediate depth are Uriconian volcanic rocks. The origin of the shallower magnetic rocks, the Apedale Tufts, is discussed below.

S t r a t i g r a p h y and p e t r o g r a p h y o f the Apedale Tufts

The chippings from the Apedale Borehole were insufficient to add details about the stratigraphy

APEDALE TUFFS of the tuff sequence, but three mudstonedominated intervals (Fig. 7) may be recognized between c. 446-462m, 980-1000m and 10051200m down-hole. Sandstone fragments also occur (e.g. at 485m), although these may be cavings from the Silesian sequence in the borehole. The mudstones yielded no fossils; some material from 1008 m was prepared for palynology, but was barren. Petrographical examination of samples (BGS Nos E 35040-8, Fig. 8) shows that the rocks are basic tufts with subordinate lapilli tufts, comprising fragments of lava, feldspar crystals and pumice in a matrix of fine ash which now consists of chlorite and illite with minor quartz and haematite. Both vitric-lithic-crystal tufts and lithic-crystal tufts occur through much of the sequence. The pumice generally consists of highly vesicular glass containing undeformed

APEDALE BOREHOLE

J

Coal

Measures

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Duckmanlian

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G U N HILL BOREHOLE

ASTBURY SECTION

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349

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Fig. 7. Geological logs of the Apedale Borehole (by A. A. Wilson), and the Dinantian and early Namurian sequences exposed at Astbury and recorded in the Gun Hill Borehole (see Fig. 10 for locations).

Fig. 8. (a) Photomicrograph of Apedale Tufts. Highly vesicular basaltic pumice is cut by a thin calcite vein. Chlorite infills the vesicles and partially replaces the glassy groundmass. Width of field is 5.5 mm (E 35043; 492.6 m depth). (b) Photomicrograph of Apedale Tufts. Basaltic lava fragment containing chlorite pseudomorphs after olivine microphenocrysts in an altered hyalocrystalline groundmass. Width of field is 4.5 mm (E 35045; 506.3 m depth). (c) Photomicrograph of Gun Hill Tufts. Chloritized pumice and shards enclose a plagioclase phenocryst. Width of field is 5.5mm (E 46369; 731.52m depth).

350

J. G. REES ET AL.

gas bubbles up to 0.2mm across, with chlorite (and/or mixed-layer mafic phyllosilicates) and minor quartz entirely replacing and infilling the pyroclasts (Fig. 8a). Lava fragments are most commonly olivine basalt containing microphenocrysts of plagioclase and olivine, represented by chlorite pseudomorphs up to 0.5 mm across, in a holo- or hyalocrystalline groundmass (Fig. 8b). Although the latter is generally highly altered to chlorite, haematite and opaque Feoxide dust, groundmass textures of flow-aligned feldspar microlites up to 0.2mm long are preserved in a few lava lapiUi. Most pyroclasts are in the size range 0.64.0ram, but there is no evidence of systematic changes in grain size through the sequence. Pumice grains tend to be equant and angular in shape, whereas lava grains are commonly equant and rounded. The rarity of non-volcanic material - a few quartz grains and one grain of quartz-chlorite schist (E 35041) - suggests that explosive activity took place close to the palaeosurface.

Determination of the age of the Apedale Tufts using trace dement geochemistry X-ray fluorescence (XRF) analyses of the Apedale Tufts are shown in Table I. Although the major elements suggest a basaltic composition, the degree of alteration indicates that rock classification involving some of the more mobile elements, such as the alkalis and silica, would give suspect results. Rock classifications using ratios of less mobile trace elements give more reliable results, and in the case of the Zr/ TiO2 against Nb/Y plot of Winchester & Floyd (1977) also provide information on the magma source and degree of evolution (Fig. 9a). The Apedale volcanic rocks are basanites (olivinerich basalts), with some spread of data points into the alkaline basalt field (Fig. 9a). The composition of the Apedale Tufts differs markedly from that of the subalkaline Uriconian volcanic rocks, which have very low Nb contents (Fig. 9a), and have been derived from a mantle source much modified by subduction processes (Pharaoh et al. 1987b). The Apedale Tufts also contrast compositionally with largely tholeiitic to calcalkaline Ordovician volcanic rocks in the Welsh Marches area, described by Leat& Thorpe (1986). There is no possibility of confusion between the Apedale Tufts and the local Butterton-Swynnerton dykes of late CretaceousPalaeogene age (Thompson & Winchester 1995), as the former are clearly extrusive.

The distinctive trace element geochemistry of the Apedale Tufts supports the conclusion of the aeromagnetic modelling, that an upper magnetic body (comprising the tufts) may be distinguished from a deeper magnetic body represented by the Uriconian rocks. More significantly, however, the alkaline composition of the Apedale Tufts allies them closely with other volcanic suites of Carboniferous age in Britain (Francis 1978; Kirton 1984), suggesting that the Apedale Tufts are also of this age.

Comparison of the Apedale Tufts with possible Carboniferous equivalents in North Staffordshire Having identified a probable Carboniferous age for the Apedale Tufts it was desirable to determine their age more exactly. Palynomorphs from the basal I l m of Namurian rocks overlying the tufts in the Apedale Borehole include Bellisporus nitidus, Cirratriradites saturni(?) and Ibrahimispores brevispinosus, indicative of an early Chokierian age. The tufts are thus likely to be Arnsbergian in age or older. The aeromagnetic modelling suggests that the volume of extrusive material remaining in the Apedale volcanic centre can be measured in cubic kilometres. It is reasonable to expect that the products of such voluminous volcanism would be laterally extensive, so the occurrence of Dinantian and early Namurian volcanism within 25 km of Apedale was reviewed to try to find correlatives with the Apedale Tufts. During the Dinantian and early Namurian the site of Apedale lay on the Market Drayton Horst, a northern promontory of the WalesBrabant High, underlain by the western microplate of the Proterozoic Midlands Microcraton. Evidence for the pre-Chokierian stratigraphy of the horst is limited to outcrops at Astbury and the Bowsey Wood and Bitterns Wood boreholes (Fig. 10). The strata recorded are little over 300 m thick and are entirely of Asbian age or younger. A much thicker and more complete record is present in the North Staffordshire Basin to the east (Trewin & Holdsworth 1973). This basin, formed by Dinantian reactivation of the suture zone between the eastern and western microplates of the Midlands Microcraton (Lee 1988; Corfield 1991), was bounded on its eastern side by the Derbyshire Platform. The Courceyan to Chokierian sequence in the basin probably exceeds 1.5 km in thickness (Rees & Wilson in press). Together, the Gun Hill (SJ 9723 6182)

689 134 49 10 78 101 248 62 5 360 10 16 26 <1 215

748 160 46 4 80 109 218 65 6 611 8 14 26 2 236

49.47 14.66 2.49 8.71 10.27 7.78 1.13 2.23 0.18 0.65

640.1

541 115 40 7 77 82 190 91 7 337 6 13 25 3 214

51.23 14.73 2.37 8.12 7.15 8.83 1.93 2.62 0.14 0.53

701.0

711 80 44 6 55 70 144 23 9 650 9 17 25 4 178

47.61 13.00 2.61 9.65 9.18 9.91 2.60 0.70 0.21 0.63

760.5

703 114 46 4 75 90 233 33 15 467 <2 11 24 6 199

45.85 12.02 2.52 9.45 13.20 8.66 2.05 1.18 0.23 0.58

873.3

658 165 68 5 72 111 290 118 4 489 9 12 29 5 238

48.80 14.18 2.59 10.52 10.30 6.24 1.43 3.61 0.19 0.67

877.8

173 130 65 7 64 111 486 117 8 133 11 11 36 5 248

49.78 19.29 4.07 12.92 5.40 1.56 0.87 4.95 0.09 0.67

914.4

427 158 62 9 72 102 363 122 10 281 10 12 28 2 218

49.17 14.53 3.20 10.92 10.40 4.26 0.85 3.51 0.16 0.59

943.4

1066.8

348 177 436 8 82 107 256 120 <1 123 7 13 30 2 255

323 55 58 5 30 44 320 51 3 389 8 1 27 2 161

48.98 46.55 1 9 . 0 1 14.64 3.46 2.83 15.18 13.03 5.22 6.66 1.26 6.77 0.62 2 . 1 3 4.83 2.19 0.09 0.16 0.20 0.37

1004.3

697 96 44 7 53 82 249 75 9 438 3 8 22 <1 209

50.96 14.81 2.53 10.35 7.39 4.85 1.88 2.70 0.15 0.51

1126.2

1495 73 . . 41 66 159 30 <3 117 . 6 17 . 170

36.90 10.80 1.38 5.96 5.90 15.49 0.20 4.06 0.08 0.50

.

.

. .

183

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49 72 239 30 0 142

842 88

36.30 12.10 1.96 8.91 5.30 12.23 0.30 3.45 0.08 0.66

E70117" E70118"

.

.

. .

703 75 . . 51 71 200 33 2 126 . 6 21 . 198

39.80 11.00 1.68 6.50 5.10 13.14 0.10 3.59 0.09 0.66

E70119"

110

2 26

29 13 29 1 5 126

392 57

36.90 8.70 1.29 14.24 3.70 11.78 1.00 <0.05 0.12 0.31

E46369 °

466.3-1126.2 are sample depths (m) in the Apedale Borehole. E70117-9* are samples from Limekiln Quarry, Astbury. E46369 ° is the hyaloclastite sample from Gun Hill Borehole.

531 84 47 5 61 87 214 61 5 335 4 8 25 5 193

47.85 12.62 2.72 9.56 14.99 5.71 1.63 2.30 0.20 0.62

M ~ o r elements (wt%) 51.06 46.25 SiO2 13.85 12.49 A1203 2.51 2.55 TiO3 9.73 10.80 Fe303t 12.12 12.46 MgO 1.80 7.92 CaO 0.91 0.78 Na20 2.34 2.48 KzO 0.15 0.13 MnO 0.47 0.55 P~O5

Trace elements (ppm) Ba 1983 Ce 103 Co 47 Hf <3 La 52 Nb 82 Ni 199 Rb 76 Sm 10 Sr 263 Ta 5 Th 16 Y 28 Yb 4 Zr 198

580.6

+466.3

518.2

Sample

Table 1. M~or and trace element analyses of volcanic rocks from the Apedale Borehole, Gun Hill Borehole and Astbury

352

J. G. REES E T A L . ' 70

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APEDALE BORF_~HOLE GUN HILL BOREHOLE ASTBURY URICONIAN

1.0

10

Fig. 10. Map showing the Dinantian and early Namurian setting of the Apedale area (area is the same as that shown in Figs 2 and 3). BH, Borehole; RRF, Red Rock Fault. Brick ornament represents area of the structural block. Shaded area is the postulated extent of thick volcanic rocks. Tick marks indicate the downthrown side of faults.

• • • •

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Y

Fig. 9. Element discrimination diagrams. (a) Trace element data plotted on the Zr/TiO2 v. Nb/Y diagram of Winchester & Floyd (1977). Uriconian data from Pharaoh et al. (1987b). (b) Trace element data plotted on the Nb-Y-Zr diagram of Merchede (1986). WPA, within-plate alkaline basalt; WPT, within-plate tholeiite; P-MORB, Plume-modified mid-ocean ridge basalt; N-MORB, normal mid-ocean ridge basalt; VAB, volcanic arc basalt.

and Nooks Farm (SJ 9174 5803) boreholes (21km and 14km NE of Apedale), and the Caldon Low borehole (SK 0408 4822, 24 km E of Apedale; Fig. 10), penetrated the entire late Devonian to Chokierian sequence. Four horizons (a-d) of volcanic rocks or volcaniclastic debris are identified in these boreholes. (a) Thin pre-Holkerian lavas are interbedded with limestones between 1040 and 1080.5 m in

the Gun Hill borehole (Fig. 7; Hudson & Cotton 1945). No older volcanic rocks have been recorded in the Dinantian and late Devonian sequence preserved in this borehole, or in that at Caldon Low (Chisholm et aL 1988). The absence of tufts in association with these lavas suggests that the latter are unlikely to be correlatives of the Apedale Tufts and were probably sourced from a small volcanic centre near Gun Hill. (b) Holkerian tufts and lavas(?) over 60m thick occur in the Gun Hill borehole. Equivalent tufts have also been penetrated in the Nooks Farm borehole, and in the Alport borehole further to the NE (Hudson & Cotton 1945). (c) Late Asbian/early Brigantian tufts are exposed at Astbury and were cored in the Gun Hill and Nooks Farm boreholes. The faunas associated with the tufts are non-diagnostic, but are of P la age; it is unclear whether the latter biozone falls into the late Asbian or early Brigantian (Riley 1993). (d) Thin and impersistent bentonites have been found in the Pendleian and Arnsbergian sequence of North Staffordshire (Trewin 1968; Trewin & Holdsworth 1972; Chisholm et al. 1988). They have not been found locally, for instance at Astbury, and their relationship to the thick tuff sequence at Apedale is uncertain. This review of Carboniferous volcanic rocks in the vicinity of Apedale indicates that only the volcanic rocks of Holkerian and late Asbian/ early Brigantian age are likely equivalents of the

APEDALE TUFFS Apedale Tufts; consequently these were sampled for comparison with those at Apedale.

The Gun Hill (Holkerian) volcanic rocks The only remaining sample of the Gun Hill tufts (BGS No. E 46369; depth 731.52m) is greygreen and mostly consists of highly vesicular lava fragments, up to 2 cm across, and scattered plagioclase crystals replaced by albite and chlorite, in a sparry carbonate cement. Many of the chlorite (+quartz)-filled vesicles are spherical, up to 1 mm across, whereas others are contorted by flowage. Part of the sample consists of pseudomorphs after shards and pumice with plagioclase crystals (Fig. 8c), closely resembling basaltic textures described from the Haddon Fields borehole by Aitkenhead et al. (1985, plate 9). The latter interpreted these as hyaloclastic, or possibly autoclastic, in origin.

The Astbury (late Asbian/early Brigantian) volcanic rocks The Astbury volcanic rocks (Gibson & Hind 1899; Hind 1904; Evans et al. 1968) consist of tuft-breccias and tufts. The samples were collected from four 0.9 to 2.7m thick beds of tuff breccias, lapilli tufts and interbedded limestones from a 6.4m high scar 30m NE of Limekiln Lake (SJ 8626 5931). Three beds are crudely graded, with angular blocks of greenishbrown to dark reddish-brown basic lava up to 0.25 m across near their bases, and lapilli-size fragments in the top 0.1-0.2 m. Basaltic fragments for chemical analysis were hand-picked from tuff breccias at the bases of the lowermost and uppermost beds. Thin sections (BGS Nos E 70117-9) show that the lava fragments are hyalo- and holocrystalline olivine basalt containing microphenocrysts (<2mm) of olivine, plagioclase and augite, now represented by chlorite and calcite pseudomorphs. Some fragments show flow-aligned feldspar microlites (<0.1 mm), whereas others have numerous microphenocrysts in a groundmass of chlorite plus opaque Fe-oxide dust. Highly vesicular lava is most common in the upper bed, where numerous chloritized pumice lapilli are present. Vesicles, now chlorite or carbonate-filled, show little evidence of deformation or compaction. Whilst there is little evidence by way of pillow lavas or hyaloclastites that the Astbury volcanic

353

rocks originated as subaqueous eruptions, the presence of shell debris and interbedded limestones indicate that accumulation occurred in a marine environment. The size of the larger lava blocks indicates a proximal source that was possibly characterized by Strombolian eruptions, with each graded bed representing a renewed pulse of eruptive energy.

Trace element geochemistry of the Gun Hill, Astbury and Apedale volcanic rocks Chemical parameters Major element analysis identified that the Gun Hill and Astbury volcanic rocks have a broadly basaltic composition. However, the degree of alteration suggests that major element classification diagrams would not give useful results. Moreover, the samples have clearly suffered from carbonate dilution or enrichment of the major elements (Table 1). Hence, rock classifications using ratios of the less mobile elements were used to define composition. The Astbury volcanic rocks are compositionally similar to those of Apedale (Fig. 9a). In terms of Zr/TiOz ratios, two of the Astbury samples are more evolved than those from Apedale, but this may simply reflect variations in the proportion of glassy and crystalline lava fragments in the two deposits. Both groups of samples show high levels of Nb, indicating a mantle magma source similar to that of typical ocean island basalt (OIB; Weaver 1991). The hyaloclastite from the Gun Hill borehole has a significantly lower Nb concentration than the Apedale and Astbury basalts, and plots in the subalkaline basalt field. The tectonic setting of the basaltic volcanism was explored by plotting data on the Nb-Zr-Y diagram (Fig. 9b). Data from Apedale and Astbury mostly cluster in the top of the within-plate alkaline basalt field (WPA), whereas the Gun Hill tuff plots in the plumeinfluenced mid-ocean ridge basalt (P-type MORB) field, close to the within-plate tholeiite (WPT) field. However, Fig. 9b does not necessarily imply completely different tectonic settings for the Apedale and Astbury volcanism compared with that at Gun Hill. In both cases, within-plate rift volcanism was the most likely setting, erupting alkaline magmas at Apedale and Astbury with possibly more limited volumes of continental tholeiites in the Gun Hill area. Regardless, it would appear that the Apedale Tufts had the same magma source as the

J. G. REES ET AL.

354

Astbury volcanic rocks, and were erupted at broadly the same time; they too are likely to be of P 1a age. This conclusion is the same as that reached by Giffard (1923), although now supported by further evidence. The Apedale volcanic centre apears to be one of many late Asbian to early Brigantian volcanic centres in central England, all of which produced tufts and lavas of olivinebasaltic type. These are best known on the Dinantian Derbyshire Platform where four separate volcanic centres have been identified (Waiters & Ineson 1981) and the many individual lava and tuff horizons have been named (Aitkenhead et al. 1985). The volcanism may also be traced off the platform, for instance to the south (Falcon & Kent 1960). Even more extensive are the alteration products of the tufts, the potassium-bentonites which occur commonly in Asbian and Brigantian sequences, not only across Derbyshire, but throughout Britain and Ireland (Walkden 1972). The widespread extent of late Asbian/early Brigantian volcanism was probably caused by a phase of active rifting over much of northwestern Europe, including central England (Ebdon et al. 1990). Rifting at this time is incorporated into a modelled section across the Market Drayton Horst into the North Staffordshire Basin (Fig. 11). In this model both the Holkerian and late Asbian/early Brigantian tufts

SW

in the basin appear to split towards the footwall of the Lask Edge Fault, suggesting that normal movement on the fault was coincident with eruption of the tufts.

The m a g m a source o f the Apedale Tufts Analysis of immobile elements also provides evidence on the source of the magmas that produced the volcanic rocks at Apedale and Astbury. Enrichment in certain highly incompatible trace elements is characteristic of mantlederived OIB magmas (Weaver 1991; Hawkesworth & Gallagher 1993). These elements, including Ba, La, Nb, Rb, Ta and Th, have low bulk distribution coefficients in basic magmas (D < 0.01), and ratios of these elements are little affected by fractionation processes after the magma has separated from its mantle source. Hence variations in ratios of these elements largely reflect mantle heterogeneity. Ratios of La/Nb shown in Fig. 12 compare broadly with contemporaneous Dinantian volcanic rocks in England and Scotland. The Apedale and Astbury basalts show a narrow range of values indicating a common source. Dinantian basalts from SW England (Fig. 1; Rice-Birchall & Floyd 1988; Floyd et al. 1993) show a similar range of values, suggesting that the same asthenospheric mantle source possibly contributed to both

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volcanic rock, suggest that the borehole was sited close to a major subaerial volcano or a group of closely-related volcanoes. In contrast, the graded beds of tuff at Astbury appear to have accumulated in a marine environment which was proximal to a major vent or vents. On the basis of similarities in geochemistry and general proximity, the Apedale centre could have been the source of the Astbury tufts. However, it is perhaps more plausible that the volcanic rocks at Apedale and Astbury were derived from separate centres. Aeromagnetic evidence suggest that the vent or vents which produced the Apedale Tufts, and that which produced the Astbury volcanic rocks (which may be represented by the aeromagnetic culmination near Mow Cop, see above), and the contemporaneous volcanic rocks at Little Wenlock (Kirton 1984) coincide with the crop of the Wem-Red Rock Fault System, which is likely to have been active during the Dinantian (Rees & Wilson in press). The close association between Dinantian volcanism and the position of faults has been widely commented on previously (Stevenson & Gaunt 1971; Francis 1978).

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regions of rift volcanism, whereas Dinantian alkaline basalts from the Clyde Plateau show a wider range and higher ratios (Phillips 1994), indicating a different mantle source.

Location of late Asbian]early Brigantian volcanic centres Although coarse agglomeratic deposits are not recorded from the Apedale Borehole, the general absence of interbedded sediments or shell debris and limited occurrence of non-volcanic detritus, coupled with the considerable thickness of

Because of the uncertainties inherent in the aeromagnetic modelling, it is only possible to make very broad estimates of the size of the Apedale volcanic centre. However, on the basis of this modelling alone, it appears to be much larger than broadly contemporary centres elsewhere in Central England. No other Dinantian centre in England (Fig. 1) produced volcanic rocks of a thickness and volume comparable with those at Apedale, although tufts and lavas from the Derbyshire Platform centres are laterally extensive (Aitkenhead & Chisholm 1982, fig. 2). The likelihood that Apedale remained above sea level until the early Chokierian is perhaps some indication of the thickness of volcanic rocks that accumulated at this centre. The data do not allow any firm decisions about the shape and size of individual volcanic edifices. The thickness of the tufts and rarity of non-volcanic rocks, especially the lack of limestones, indicates that the volcanic complex was largely subaerial, but the very predominance of tuffaceous material strongly suggests that much of the volcanic activity was phreatomagmatic or hydromagmatic in origin. It appears that the Apedale Tufts may be only a small part of a much larger volcanic complex, as many of the volcanic rocks are

356

J. G. REES E T AL.

likely to have been eroded during the Brigantian (Ebdon et al. 1990) or Silesian to early Permian (Variscan) inversion. For instance, in both Bowsey Wood and Bitterns Wood Boreholes, south of Apedale (Figs 10 & 11), Namurian rocks unconformably overlie the Astbury Asbian limestones, the intervening sequence bearing the late Asbian/early Brigantian volcanic rocks possibly having been eroded. Furthermore, seismic, gravity and magnetic data (Cornwell & Dabek 1993) suggest that almost all of the Carboniferous rocks west of the W e m - R e d Rock Fault System were removed when the fault system was reactivated during Variscan inversion (Evans et al. 1993). Hence, it is likely that much of the Apedale volcanic centre and its products was uplifted and eroded. It may be the case that the volcanic rocks at Apedale, Astbury and Little Wenlock are merely the remnants of what was originally a large and distinct volcanic province developed along the W e m - R e d Rock Fault System in the late Asbian/early Brigantian.

Conclusions Apedale is underlain by two bodies of (magnetic) volcanic rocks, the lower of which appears to form a NNE-trending ridge with an upper surface at between 3 and 7 km depth; it probably consists of Uriconian rocks of Neoproterozoic age. The upper volcanic body consists of a broadly stratiform sequence of tufts of alkaline basaltic composition. The thickness of these tufts is likely to be about 1 km at Apedale and measurable in hundreds of metres over an area greater than 100 km 2. The tufts are likely to be of P la age, i.e. late Asbian/early Brigantian (late Visran), and are chemically very similar to contemporaneous volcanic rocks in SW England. Earlier work on the Apedale Borehole was done by previous colleagues at the BGS. A. A. Wilson logged the chippings, and R. K. Harrison carried out all of the petrographic work summarized here; their contribution is gratefully acknowledged, as is that of B. Owens who undertook the palynological work. J. I. Chisholm, M. K. Lee, E. H. Francis and an anonymous referee provided useful comments on the manuscript. We thank D. B. Thompson and B. M. Besly of Keele University for assistance with sampling and providing data on boreholes. This paper is published with the permission of the Director, British Geological Survey (NERC).

References AITKENHEAD, N. & CHISHOLM,J. I. 1982. A Standard Nomenclature for the Dinantian Formations of the Peak District of Derbyshire and Staffordshire. Report of the Institute of Geological Sciences No. 82/8. , - - & STEVENSON, I. P. 1985. Geology of the Country around Buxton, Leek and Bakewell. Memoir of the British Geological Survey, Sheet 111 (England and Wales). BESLY, B. M. 1993. Geology. In: PHILLIPS, A. D. M. (ed.) The Potteries: Continuity and Change in a Staffordshire Conurbation. Alan Sutton Publishing, Stroud, 17-36. BROOKS, M. & FENNING, P. J. 1968. Geophysical investigations. In: GRIEG, D. C., WRIGHT, J. E., HAINS, B. A. & MITCHELL, G. A. (eds) Geology of the Country around Church Stretton, Wenlock Edge and Brown Clee. Memoir of the Geological Survey of England and Wales, Sheet 166. BUSBY,J. P. 1987. An interactive Fortran 77 program using GKS graphics for 2.5D modelling of gravity and magnetic data. Computers and Geosciences, 13, 639-644. CHISHOLM, J. I., CHARSLEY, T. J. & AITKENHEAD, N. 1988. Geology of the country around Ashbourne and Cheadle. Memoir of the British Geological Survey, Sheet 124 (England and Wales). CORFIELD, S. M. 1991. The Upper Palaeozoic to Mesozoic Structural Evolution of the North Staffordshire Coalfield and Adjoining Areas. PhD Thesis, University of Keele. CORNWELL, J. D. & DABEK, Z. K. 1993. Geophysical Investigations in the Stoke-on-Trent District. British Geological Survey Technical Report, WK/94/04. EBDON, C. C., FRASER, A. J., HIGGINS, A. C., MITCHENER, B. C. & STRANK, A. R. E. 1990. The Dinantian stratigraphy of the East Midlands: a seismostratigraphic approach. Journal of the Geological Society, London, 147, 519-536. EVANS, D. J., REES, J. G. & HOLLOWAY,S. 1993. The Permian to Jurassic stratigraphy and structural evolution of the central Cheshire Basin. Journal of the Geological Society, London, 150, 857-870. EVANS, W. B., WILSON, A. m., TAYLOR, B. J. & PRICE, D. 1968. Geology of the country around Macclesfield, Congleton, Crewe and Middlewich. Memoir of the British Geological Survey, Sheet 110 (England and Wales). FALCON, N. L. & KENT, P. E. 1960. Geological results of petroleum exploration in Britain 1945-1957. Memoir of the Geological Society of London, No. 2. FLOYD, P. A., EXLEY, C. S. & STONE, M. 1993. Variscan magmatism in southwest England-Discussion and synthesis. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British lsles. Hilger, Bristol, 178-185. FRANCIS, E. H. 1978. Igneous activity in a fractured craton: Carboniferous volcanism in northern Britain. Geological Journal Special Issue, 10, 279-296.

APEDALE TUFFS GIBSON, W. 1925. Geology of the Country around Stoke-upon-Trent. Memoir of the Geological Survey of England and Wales. - & HIND, W. 1899. On the agglomerates and tufts in the Carboniferous Limestone Series of Congleton Edge. Quarterly Journal of the Geological Society of London, 55, 548-559. GIFFARD, H. P. W. 1923. The recent search for oil in Great Britain. Transactions of the Institute of Mining Engineers, 65, 221-250. HAWKESWORTH, C. J. & GALLAGHER, K. 1993. Mantle hotspots, plumes and regional tectonics as causes of intraplate magmatism. Terra Nova, 5, 552-559. HIND, W. 1904. In: Whitsuntide excursion to North Staffordshire. Proceedings of the Geologists' Association, 18, 173-184. HUDSON, R. G. S. & COTTON, G. 1945. The Lower Carboniferous in a boring at Alport, Derbyshire.

Proceedings of the Yorkshire Geological Society, 25, 254-330. KIRTON, S. R. 1984. Carboniferous vulcanicity in England with special reference to the Westphalian of the E and W Midlands. Journal of the Geological Society, London, 141, 161-170. LEAT, P. T. & THORPE, R. S. 1986. Ordovician volcanism in the Welsh Borderland. Geological Magazine, 123, 629-640. LEE, A. G. 1988. Carboniferous basin configuration of central and northern England modelled using gravity data. In: BESEY, B. M. & KEELING, G. (eds) Sedimentation in a Synorogenic Basin

Complex: the Upper Carboniferous of Northwest Europe. Blackie, London 69-84. LEE, M. K., PHARAOH, T. C. & SOPER, N. J. 1990. Structural trends in central Britain from images of gravity and aeromagnetic fields. Journal of the Geological Society, London, 147, 241-258. MERCHEDE, M. 1986. A method of discriminating between different types of mid-ocean ridge basalts and continental tholeiites with the N b - Z r - Y diagram. Chemical Geology, 56, 207 218. PHARAOH, T. C., MERRIMAN, R. J., WEBB, P. C. & BECKINSALE, R. D. 1987a. The concealed Caledonides of eastern England: preliminary results of a multidisciplinary study. Proceedings of the Yorkshire Geological Society, 46, 355-369. , WEBB, P. C., THORPE, R. S. & BECKINSALE, R. D. 1987b. Geochemical evidence for the tectonic setting of late Proterozoic volcanic suites in central England. In: PHARAOH, T. C., BECKINSALE, R. D. & RICKARD, D. (eds) Geochem-

istry and Mineralization of Proterozoic Volcanic Suites. Geological Society, London, Special Publication, 33, 541-552. PHILLIPS, E. R. 1994. Whole Rock Geochemistry of the

Alkaline Clyde Plateau lavas, Sheet 22 (Kilmarnock), Scotland. British Geological Survey Technical Report, WG/94/04. REES, J. G. & WILSON, A. A. Geology of the Country around Stoke-on-Trent. Memoir of the British Geological Survey, Sheet 123 (England and Wales), in press.

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RICE-BIRCHALL, B. & FLOYD, P. A. 1988. Geochemical and source characteristics of the Tintagel Volcanic Formation. Proceedings of the Ussher Society, 150, 427-446. RILEY, N. J. 1993. Dinantian (Lower Carboniferous) biostratigraphy and chronostratigraphy in the British Isles. Journal of the Geological Society, London, 150, 427-446. ROWLAND, J. & CADMAN, B. 1960. Ambassador for

Oil. The Life of John, First Baron Cadman. Herbert Jenkins, London. STEVENSON, I. P. & GAUNT, G. D. 1971. Geology of the Country around Chapel en le Frith. Memoir of the British Geological Survey, Sheet 112 (England and Wales). THOMPSON, D. B. & WINCHESTER, J. A. 1995. Chemical and field studies and the tectonic context of the largely Tertiary dyke suites in Staffordshire and Shropshire, Central England.

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TORRENS, H. S. 1994. 300 years of oil: mirrored by development in the West Midlands. In: The British Association Lectures (1993). Geological Society of London, 4-8. TREWIN, N. H. 1968. Potassium bentonites in the Namurian of Staffordshire and Derbyshire.

Proceedings of the Yorkshire Geological Society, 37, 73-91. & HOLDSWORTH, B. K. 1972. Further Kbentonites from the Namurian of Staffordshire.

Proceedings of the Yorkshire Geological Society, 39, 87-88. & 1973. Sedimentation in the lower Namurian rocks of the North Staffordshire Basin. Proceedings of the Yorkshire Geological Society, 39, 371-408. TUCKER, R. D. & PHARAOH, T. C. 1991. U-Pb zircon ages for Late Precambrian igneous rocks in southern Britain. Journal of the Geological Society, London, 148, 435-443. WALKDEN, G. M. 1972. The mineralogy and origin of interbedded clay wayboards in the Lower Carboniferous of the Derbyshire Dome. Geological Journal, 8, 143 160. WAETERS, S. G. & INESON, P. R. 1981. A review of the distribution and correlation of igneous rocks in Derbyshire, England. Mercian Geologist, 8, 81-132. WEAVER, B. L. 1991. Trace element evidence for the origin of ocean island basalts. Geology, 19, 123126. WILSON, C. D. V. 1980. An aeromagnetic survey of the Church Stretton area, Shropshire: a revised map. Proceedings of the Geologists' Association, 91,225-227. WINCHESTER, J. A. & FLOYD, P. A. 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements. Chemical Geology, 20, 325-343.

Cyclic stratigraphy, facies and fauna of the Lower Carboniferous (Dinantian) of the Moscow Syneclise and Voronezh Anteclise MARIYA

KH. MAKHLINA

Geocentre M o s c o w , Varshavskoye shosse, 39A, M o s c o w , Russia

Abstract: On the basis of sedimentoiogical investigation of rocks, together with bed-by-bed studies of foraminifers, conodonts and spores, rhythmic sedimentation has been demonstrated in the Lower Carboniferous of the Moscow Syneclise and Voronezh Anteclise. Cycles of different order have been established using the methods of rhythmic stratigraphy elaborated by Tikhomirov. Four transgressive-regressive cycles of the sixth order have been recognized in the Dinantian of the Russian Platform. The latter are time-stratigraphic equivalents of the newly proposed stages of the Khaninian and Shurinovkian in the Tournaisian, and Kozhimian and Okian Stages in the Vis6an. Each of these cycles includes two or three smaller cycles which represent the initial and maximum transgressive units and a regressive unit.

This paper briefly outlines the main results of an investigation of the Lower Carboniferous of the Russian Platform. For more detailed information the reader is referred to the monographic work of Makhlina et al. (1993) in which faunal characteristics of the Lower Carboniferous stratotype and reference sections for 15 horizons, together with the description of facies distribution are given. Horizons are understood as biostratigraphic units restricted to a palaeogeographical province or to some palaeobasin of sedimentation. Horizons by definition are close to regional stages. Stages are understood by the majority of Russian stratigraphers as units of international application. The monograph of Makhlina et al. (1993) also presents biozonal subdivision of the Russian Platform based on foraminifers, conodonts, spores (illustrated by numerous palaeontological plates) and, to a lesser degree, on brachiopods and corals. Stratigraphic ranges of important taxa are summarized in that work; only diagnostic zonal taxa are summarized here. Identifications of the fossils have been made by V. Zhulitova and M. Vdovenko (foraminifers), A. Alekseev and L. Kononova (conodonts), T. Byvsheva and N. Umnova (spores), A. Grigorieva and L. Donakova (brachiopods), and M. Hecker (corals). Petrographical and facies analyses have been made using macro- and microscopic rock descriptions, as well as analytical (geochemical, mineralogical, etc.) data.

Concept of cycles The Lower Carboniferous of the Moscow Syneclise and Voronezh Anteclise is represented mainly by marine deposits, i.e. by a

succession of transgressive deposits, in which sands and clays are gradually replaced by argillaceous limestones, dolomites and bioclastic (biomorphic detrital) limestones. Subhorizontal bedding and depositional cyclicity allows the opportunity to identify and follow horizons and relatively smaller units throughout the studied area, and to recognize cycles of different order. A detailed subdivision of Dinantian rocks in the Moscow Syneclise, based upon the methods of rhythmic stratigraphy first proposed by Shvetsov (1938) and developed by Tikhomirov (1988), was recently described by Makhlina et al. (1993). The Carboniferous System is considered to be a transgressive cycle of the third order; the Lower Carboniferous Subsystem represents a fourthorder cycle. The Tournaisian and Vis6an Series correspond to tectono-eustatic cycles of the fifth order, with Stages forming cycles of the sixth order (Fig. 1). On the Russian Platform these transgressional sixth-order cycles (Stages) consist of three units: an initial (a) transgressive unit, a maximum (/3) transgressive unit, and a regressive ('7) unit. In a shortened cycle, the initial (a) transgressive phase is missing (Tikhomirov 1988). These units can be traced in different facies throughout the region. Within units of the seventh order, smaller units can be distinguished which correlate with smaller subdivisions of the Dinantian. Each stratigraphic unit (cycle or rhythm) can be subdivided into two parts, the lower corresponding to a transgressive phase of marine sedimentation, and the upper one to a regressive phase. Contemporaneous stratigraphic units of different facies formed in the palaeobasin as a lateral sequence of rocks, reflecting nearshore, peripheral, shallow-water and relatively deep-water

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 359-364.

360

M. KH. MAKHLINA RUSSIAN PLATFORM (MOSCOW SYNECLISE , VORONEZH ANTECLISE Cyclostratigraphic scale 1994 Order of cycles IV

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Fig. 1. Correlation chart of the Lower Carboniferous for the Russian Platform compared with Western Europe. environments. A persistent feature of all simultaneously formed heterofacial units is a distributional pattern of faunal assemblages: they invariably show maximum diversity in a transgressive phase of a cycle, and low diversity in its regressive phase (Makhlina et al. 1993). Boundaries between stratigraphic units correspond to changes of sedimentation conditions, i.e. to changes in abiotic environment occurring

simultaneously in the whole region. The main criteria for the recognition and tracing of transgressive-regressive cycles are as follows: sedimentological and geochemical characteristics; rock textures; and faunal characteristics for age assessment of the stratigraphic units. In this respect, the importance of lithological and palaeoecological investigations should be emphasized (Osipova & Belskaya 1967).

LOWER CARBONIFEROUS OF RUSSIAN PLATFORM During the Early Carboniferous the studied area (Fig. 2) was situated in the western marginal part of the Russian Platform. During the Dinantian, four transgressions occurred here. In the post-Early Carboniferous (Mid-Carboniferous) period this tectonically and palaeogeographically uniform structure was divided into the Moscow Syneclise and the Voronezh Anteclise.

The first transgression

During the Early Tournaisian the first transgression (Fig. 3) is represented by argillaceous and carbonaceous rocks of lagoonal and shallowwater facies of the Khaninian Stage (Fig. 1). The initial phase (a) can be seen in the limestones with foraminifers of the Bisphaera malevkensis Zone and underlying black clays with spores (Vallatisporites pusillites) and conodont species Patrognathus crassus lying in the upper part of the Gumerovo Horizon (up to 5m thick).

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361

Overlying beds of the Malevka and Upa Horizons correspond to the maximum phase (/3) of the transgression and are represented by an alternation of bioclastic limestones and bluegreenish clays, which yielded foraminifers of the Bisphaera malevkensis and Prochernyshinella disputabilis-T, beata Zones, conodonts of the Patrognathus variabilis and P. andersoni Zones, and the Tumulispora malevkensis and Grandispora upensis Palynozones. The thickness of the Malevka Horizon is 6-14m, that of the Upa up to 35 m. The Karakuba Horizon corresponds to a regressive phase (7). The latter is developed in the Donets Basin and is absent both from the Moscow Syneclise and the Voronezh Anteclise (Fig. 3).

The second transgression

This trangression is coincident with the Shurinovkian Stage• The initial phase (o0 is absent. Maximum transgressive units (/3) include the Cherepet and the lower part of the Kizel Horizons (Fig. 1). The regressive unit (Kosva Horizon) is absent in the Moscow Syneclise and Voronezh Anteclise area (Fig. 3). The lower part of the Cherepet Horizon characterizes the brackish-lagoonal environment, and the upper part that of the shallow offshore marine environment. The black clays of the former (up to 20 m thick) yielded spores of the Apiculiretispora septalia Zone, and the latter (up to 22 m thick) consists of bioclastic limestones with foraminifers of the Chernyshinella glomiformis-Septabrunsiina

krainica-Palaeospiroplectammina tschernyshinensis Zone (see Rukina this volume)• The fauna

..... 520

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~\\\\\\,~ " ""

is diversified and abundant and contains many forms common with those widespread in the early to late Tournaisian of Western Europe. The Kizel Horizon, the lower part of which has been recognized only in the Voronezh Anteclise, is represented by bioclastic limestones (4.5-9.5m thick), with ostracodes, and foraminifers of the

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Belgorod----

Spinoendothyra costifera-Tuberendothyra tuberculata Zone. Spores of the Pustilatisporites uncatus Zone are also present. The upper part of the horizon is erosional (Fig. 3).

The third transgression Fig. 2. Tectonic scheme of the central Russian Platform. Moscow Syneclise: 1, western part; 2, southern part; 3, central part. Voronezh Anteclise" 4, Kursk Magnetic Anomaly (KMA); 5, ValuikiBoguchar Block; 6, boundary between tectonic elements.

This transgression occurred in the Early Vis6an and can be seen in beds of the Kozhimian Stage. It includes the Radaevka and Bobriki Horizons (cycles of the seventh order)• The former is interpreted as corresponding to the maximum

362

M. KH. MAKHLINA

Horizon

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I

Moscow Syneclise

Cyclostratigraphic Scale 1994 Orders of Cycle

Voronezh Anteclise

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8

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10

20

11

12

13

14

15

16

17

18

19

21

22

23

24

25

26

27

28

29

30

31

LOWER CARBONIFEROUS OF RUSSIAN PLATFORM (B) of the transgression, and the latter to the regressive ('7) phase (Fig. 1). The Radaevka Horizon is represented by continental strata in the Moscow Syneclise and by marine limestones in the Voronezh Anteclise (Fig. 3). The continental beds consist of clays with intercalated coals. They are characterized in the lower part by spores of the Cincturiasporites multiplicabilis Zone, and by those of the C. appendices Zone in the upper part (up to 20-30 m thick). The latter are also present in the marine deposits in the Voronezh Anteclise. The Bobriki Horizon in the Moscow Syneclise is composed of continental b e d s - alluvial and lacustrine deposits with coal layers of industrial importance (Fig. 3). In the east of the south flank of the Voronezh Anteclise, sands, clays and limestones are present, suggestive of lagoonal and shallow-marine environments. Alluvial deposits were laid down in palaeovalleys, cutting through the Famennian-Tournaisian basement to a depth of 100 m and extending for 50-120 km. The Bobriki Horizon is a rhythmic sequence with up to five smaller cycles. The sequence consists of sands and clays with coal beds, indicative of fluvial or lacustrine-boggy environments. In its lower part, spores of the Knoxisporites literatus Zone are present; the middle part yielded spores of the Densosporites intermedius Zone. The upper part contains spores of the Densosporites variabilis Zone. These three subdivisions are separated by hiatuses, as shown on Fig. 3 (see also Makhlina et al. 1993). The thicknesses are 30, 70 and 25 m respectively. In the Voronezh Anteclise only the upper part of this sequence is present (containing spores of the Densosporites variabilis Zone), being composed of sands with argillaceous-carbonaceous cement.

The fourth and largest transgression of the Dinantian This is represented by the deposits of the Okian Stage. It includes smaller units (cycles of the seventh order): the Tula Horizon (initial

363

transgressive unit a), the Aleksin and Mikhailov Horizons (maximum transgressive unit B), and the Venev Horizon (a regressive unit 7). In terrigenous-carbonaceous sections, rhythmic alternations of silts, clays of lagoonal and deltaic facies, and limestones, indicative of shallow-water conditions, are observed. Carbonaceous sections are characterized by various limestones: foraminiferal detrital types of shallow-water character, brecciated and algal limestones of lagoonal facies, and micrograined limestones with remains of Stigmaria indicative of uplift and shoaling. The top of this 'rhizoid limestone' is isochronous over the whole basin, and represents the so-called 'Hecker-Shvetsov' basal surface. The stratotype sections are located in the Oka River basin (Fig. 2). The Tula Horizon (cycle of the seventh order c~) is subdivided into three parts (Figs 1 and 3). The lower is alluvial and coal-bearing (30m thick) and was deposited in the palaeovalleys. The other two are composed of lagoonal and marine silts and lie transgressively on older rocks (up to 40 m thick). The Tula assemblage belongs to the Cingulazonates bialatus-Sirnozonotriletes brevispinosus Zone. The middle and upper parts are characterized by the Endothyr-

anopsis compressa-Archaediscus krestovnikovi foraminiferal Zone, and by the brachiopod genera Productus, Pugilis, Antiquatonia, Semiplanus and others. During the maximum of the transgression (Aleksin and Mikhailov Horizons) the basin widened considerably. In the west peripheral zone (deltaic facies), sands, silts and clays were deposited. On the remaining part, calcareous muds characteristic of an open shallow sea with diverse brachiopods, corals, and foraminifers predominated. During Aleksin and Mikhailov time, thick-walled gigantoproductids and large foraminifers were abundant. The Mikhailov Horizon contains foraminifers of the Eostaffella ikensis Zone, but the index species had appeared already in the uppermost part of the Aleksin Horizon in the Archaediscus gigas-Eostaffella proikensis Zone. Foraminiferal-rich limestones

Fig. 3. Stratigraphic cycles in the Lower Carboniferous of the Moscow Syneclise and Voronezh Anteclise. 1, limestone; 2, bioclastic (detrital or biomorphic) limestone; 3, argillaceous limestone; 4, marl; 5, dolomite; 6, argillaceous dolomite; 7, dolomite marl; 8, foraminiferal and biodetrital limestone; 9, biodetrital, micro-bedded foraminiferal-brachiopodal limestone; 10, argillaceous detrital limestone; 11, mottled limestone (syngenetic limestone breccia); 12, microgranular limestone penetrated by rhizoids of Stigmaria; 13, crinoidal limestone; 14, Calcifolium limestone; 15, Bisphaera limestone; 16, peloidal limestone; 17, ostracodal limestone; 18, polydetrital limestone; 19, microgranular polydetrital limestone with fragmented Bisphaera; 20, lime clays; 21, silt; 22, sand; 23, sandstone; 24, coals; 25, breccia; 26, brachiopods; 27, caverns; 28, hiatus; 29, initial phase of the transgression; 30, maximum phase of the transgression; 31, regressive phase of the transgression.

364

M. KH. MAKHLINA

with intercalating layers of Stigmar& limestones are predominant in the sequence. The rocks suggest periodic shoalings. The thickness of the Aleksin and Mikhailov Horizons is up to 40 m. The Venev Horizon (cycle of the seventh order) reflects a regressive phase (,~) of the transgression (Fig. 1). During Venev time the sea became more shallow. The siphonous alga Calcifolium was widely distributed. Among foraminifers, species surviving from older strata prevail, but a renewal of the fauna is also observed (the Eostaffella tenebrosa-Endothyranopsis sphaerica Zone). The sequence consists chiefly of bioclastic limestones with intercalated rhizoid limestones. Frequent shoalings and traces of pre-Serpukhovian karst are characteristic of the upper part of the section. The thickness is 6-16 m in the Moscow Syneclise, and up to 40 m in the Voronezh Anteclise.

Conclusions A brief review of the developmental history of the central part of the Russian Platform shows that in the Moscow Syneclise and Voronezh Anteclise the most important faunal, facies and structural changes are observed at the bases of the newly proposed stages: the Khaninian, Shurinovkian, Kozhimian and Okian Stages.

I should like to express my deep gratitude to the EDE 2 Organizing Committee and to the International Science Foundation in Washington (grant number 1922-I) for financial support, which gave me the opportunity to participate in the Symposium.

References MAKHLINA, M. KH., VDOVENKO, M. V., ALEKSEEV, A. S., BYVSHEVA, T. V., DONAKOVA, L.M., ZHULITOVA, V. E., KONONOVA, L. I., UMNOVA, N. I. & SHICK, E. M. 1993. The Lower Carboniferous of the Moscow Syneclise and Voronezh Anteclise. Nauka, Moscow [in Russian].

OSIPOVA, A. I. & BELSKAYA, T. N. 1967. An experience of lithologo-paleoecological investigations of Vis6an-Namurian deposits of the Moscow Syneclise. Litologia i polyeznye iscopaemye, fi, 118-142 [in Russian]. RUKINA, G. A. 1996. Sequence biostratigraphy of the Tournaisian-Lower Vis+an rocks of the Russian platform. This volume. SHVETZOV, M. S. 1938. History of the Moscow Carboniferous basin in the Dinantian. Trudy Moskovskogo Geologo-razvedochnogo, 12, 1-107 [in Russian]. TIKHOMIROV, S. V. 1988. The second edition of the manual 'Historical Geology. I. On the Methods of Historico-Geological Analysis'. Izvestiya Vuzov. Geologiyairazvedka, 10, 122-135 [in Russian].

Sequence biostratigraphy of the Tournaisian-Lower Vis6an rocks of the Russian Platform G. A. R U K I N A

Apt 16, Kirov Street, 41, Lyubertsy, Moscow Region, 140005, Russia

Abstract: A statistical analysis of foraminiferal communities was used to study Tournaisian to lowermost Vis6an rocks of the Moscow Syneclise, Voronezh Anteclise, Volga-Urals region and the northern part of the Pre-Caspian Depression. Eleven levels of foraminiferal community reorganizations were recognized and traced over the Russian Platform. These levels correspond to changes in foraminiferal dominance which often coincide with changes in lithology and thus reflect an environmental sequence. An understanding of regional events affecting the sedimentary basin, including foraminiferal reorganizations, leads to more precise correlation of depositional sequences from one area to another.

Eight foraminiferal zones corresponding to horizons were established in the Tournaisian to lowermost Vis6an rocks of the Russian Platform by previous studies (e.g. Vdovenko et aL 1990; Fig. 1). The lower boundaries of these zones or horizons coincided with the first appearance or extinction of species, or with increased abundances of particular species. However, the subdivision of Tournaisian and Vis6an rocks into horizons in the stratotype region (Moscow Syneclise) was based on the lithostratigraphy of the sedimentary basin (Schvetzov 1932). The boundaries between the horizons coincide with changes in lithologies that also correspond to changes in depositional environments (Makhlina this volume). These changes were accompanied by compositional reorganizations (i.e. changes in the proportions of different species) in the fossil communities, including foraminifers.

Method The compositional reorganizations of foraminiferal communities described in this paper were determined by the statistical methods outlined in Rukina (1992). These methods include calculation of foraminiferal population density per square centimetre of thin section, and calculation of the percentage of different taxa, with the results displayed diagrammatically. The information permits reconstruction of the composition of foraminiferal assemblages through time. When analysing the species percentage changes, it is important to pay attention to developmental trends within assemblages that are common to all studied foraminiferal sequences. These trends help to define the levels of compositional reorganizations of foraminiferal communities. The reorganizations often take

place simultaneously with changes in lithology. Comparison of the changes in foraminiferal communities and lithology in different sections, while recording features common to all sections, allows the recognition of palaeoenvironmental changes caused by various events in the sedimentary basin. An understanding of the nature of such events aids in the correlation of depositional sequences from one area to another. The approach mentioned above, which integrates biostratigraphy with sedimentology, was used for the study of borehole sections situated within the Moscow Syneclise, Voronezh Anteclise, Volga-Urals region, and the northern part of the Pre-Caspian Depression. It is based on the ideas of Krassilov (1970, 1977), Martinsson (1973) and Meien (1989), who suggested that the stratigraphic sequence reflects certain stages in the development of the preceding ecosystems. These stages are characterized by a unique palaeoenvironment and corresponding communities of organisms, and are separated from one another by levels of ecosystem reorganizations.

Foraminiferal reorganizations as the basis for detailed biostratigraphy From this study, 11 levels of foraminiferal reorganizations within the Tournaisian and lowermost Vis6an were recognized and traced over the Russian Platform (Fig. 1). Most of them reflect changes in the sedimentary basin. From numerous examples, four representative borehole sections are described here (Figs 2, 3). The uppermost Famennian deposits in the Volga-Urals region and the northern part of the Pre-Caspian Depression are represented by interbedded algal and micrograined limestones. The algal limestones include abundant tubular

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 365-369.

366

G. A. RUKINA

Russian Platform Vdovenkoet aL 1990 Sedes Horizon

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fragments of the genus Kamaena and other genera of green algae. Diverse Quasiendothyra species dominate the foraminiferal communities.

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The micrograined limestones consist of microgranular calcite and numerous primitive unilocular foraminifers with rare Quasiendothyra. These sediments were deposited in shallow-water, normal-salinity environments within the transitional zone from the open sea to littoral settings. The uppermost Famennian beds in the Moscow Syneclise differ from those of the Volga-Urals region in that the micrograined limestones contain fragments of small, thinwalled osrtacodes. Primitive unilocular foraminifers are rare. Makhlaev (1964) suggested that this type of carbonate accumulated in very shallowwater environments with increased salinity.

3o6

"k Ural (River

+ ,oreho.e.oca,,oo Fig. 2. Locations of investigated borehole sections in the Russian Platform.

The first level The first level of reorganization in the Tournaisian coincides with the base of the 'Bisphaera' beds of the Malevsky Horizon just above the Famennian-Tournaisian boundary (within T n l b in the France-Belgian B a s i n = l o w e r part of the H a s t a r i a n - see Conil et al. 1990, Jones & Somerville this volume). At this time similar conditions prevailed over the whole Russian Platform. Many groups of foraminifers

TOURNAISIAN STRATIGRAPHY OF THE RUSSIAN PLATFORM

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disappeared and Bisphaera became dominant. This genus was tolerant of brackish conditions. This level apparently corresponds to the culmination of an ecological crisis when numerous organisms disappeared. The crisis was probably caused by a decrease in salinity. During this time terrigenous beds were formed in several depressions of the VolgaUrals region. The terrigenous sedimentation was probably caused by a massive fresh-water influx which led to a decrease of salinity.

dominant in the foraminiferal communities (Fig. 3). This level corresponds to a major transgression which caused open marine conditions to extend across large areas (Fig. 1). This transgression has two oscillations. The first is recognized by the presence of fine polydetrital limestones with abundant foraminifers in the Volga-Urals region, and terrigenous beds in the Moscow Syneclise.

The second level

Following a brief regression, a second transgression defines the base of the upper Cherepetsky or the sixth level of compositional reorganization. It is characterized by an increase in Endothyra parakosvensis Lipina, and Tuberendothyra tuberculata (Lipina). Baituganella tchernyshinensis Lipina became abundant (Fig. 3).

This is situated at the top of the 'Bisphaera' beds (within Tnlb = Hastarian) and coincided with a fall in sea level (Fig. 1). The foraminiferal communities are characterized by a decreased population dominated by the unilocular foraminifer Vicinesphaera. Micrograined carbonate muds accumulated in the Volga-Urals region and terrigenous sediments accumulated in the Moscow Syneclise.

The third level This occurs at the Malevsky-Upinsky Horizon boundary (Tnlb/Tn2a=Hastarian). In most foraminiferal assemblages, the diversity and abundance increased. Eochernyshinella, Chernyshinella (Prochernyshinella) and rare Chernyshmella glomiformis (Lipina) forma minima can be seen in thin sections, along with many algae. These occurrences suggest a reversal to normal marine conditions following transgression (Fig. 1).

The fourth level This is marked by an increase in the abundance of Chernysh&ella glomiform& and in the diversity of Chernyshinella (Fig. 3). This level reflects shallowing at the Upinsky-Karakubsky Horizon boundary (Fig. 1; Tn2a/Tn2b = Hastarian) where algal sediments were replaced by polydetrital limestones.

The sixth level

The seventh level Regression at the end of the Cherepetsky Horizon and the following transgression at the beginning of the Kizelovsky Horizon (Tn3a & 3b = Ivorian of Conil et al. 1990 and within the Cf2 Foraminifera Zone) caused the seventh reorganization in foraminiferal communities. The Kizel transgression was not as extensive as the previous one. It was not accompanied by pronounced deepening. A shallow-water environment, in which algal sediments accumulated, remained up to the middle of the Kosvinsky Horizon (Fig. 1).

The eighth and ninth levels Changes in the composition of foraminiferal communities can be observed within the Kizelovsky Horizon (the eighth level) and at the base of the Kosvinsky Horizon (the ninth level). The former is marked by the dominance of various Spinoendothyra and Brunsia. The latter is characterized by an increase in Pseudoplanoendothyra and Dainella. Where algal limestones are replaced by fine detrital ones, Eoforschia increases (Rukina 1993).

The fifth level The fifth level of compositional reorganization was established at the base of the Cherepetsky Horizon (Tn2c=top of the Hastarian). Palaeospiroplectammina tchernyshinensis (Lipina), Chernyshinella glomiformis, Brunsiina spp. and Neoseptaglomospiranella spp. became

The tenth level By mid-Kosvinsky time, stability in the depositional environments was again interrupted by a transgression (Fig. 1). Here the tenth level can be recognized. The number of Eoforschia species decreases and Omphalotis ex gr. chariessa (Conil

TOURNAISIAN STRATIGRAPHY OF THE RUSSIAN PLATFORM & Lys) appears. Algal limestones were replaced by fine-grained and micritic-polydetrital limestones, often intercalated with terrigenous beds.

The eleventh level The eleventh level of foraminiferal reorganization at the base of the Radaevksy Horizon ( V I a = M o l i n i a c i a n of Conil et al. (1990) and within the Cf4c~2 Foraminifera Subzone) corresponds to extensive changes in foraminiferal communities. Eoparastaffella simplex Vdovenko, Plectogyranopsis, Pseudolituotubella, Pseudoglomospira and Pseudoammodiscus began to dominate. In the sections, fine-grained polydetrital limestones prevail. This level coincides with a transgression and pronounced deepening (Fig. 1).

Conclusions The study of foraminiferal community successions is a key to more detailed subdivision of many carbonate strata as well as for longdistance correlations. In this case correlation is based on an understanding of the sequence events which have led to compositional reorganizations in the foraminiferal communities. From this study, 11 levels of foraminiferal reorganizations within the Tournaisian and lowermost Vis~an were recognized and traced over the Russian Platform. These levels are characterized by the change of dominant groups in the foraminiferal communities and often coincide with changes in lithology and thus reflect the environmental sequence. I gratefully acknowledge the financial help of the Coordination Committee of the EDE '94 Symposium to attend the symposium in Dublin and present a paper. I am greatly indebted to referees P. Brenckle and M. Laloux for their useful comments.

369

References CONIL, R., GROESSENS, E., LALOUX, M., LEES, A. & TOURNEUR, F. 1991 (1990). Carboniferous guide foraminifera, corals and conodonts in the FrancoBelgian and Campine Basins: their potential for widespread correlation. Courier Forschungsinstitut Senckenberg, 130, 15-30. JONES, G. LL. & SOMERVILLE, I. D. 1996. Irish Dinantian biostratigraphy: practical applications. This volume. KRASSILOV, V. A. 1970. Paleoecosystems. Izvestia Academii Nauk SSSR. Seriya geologicheskaya, 4, 144-150 [in Russian]. 1977. Evolution and Biostratigraphy. Nauka, Moscow [in Russian]. MAKHLAEV,V. G. 1964. Depositional Environments in the Upper Famennian Basin of the Russian Platform. Nauka, Moscow [In Russian]. MAKHLINA, M. K. 1996. Cyclic stratigraphy, facies and fauna of the Lower Carboniferous (Dinantian) of the Moscow Syneclise and Voronezh Anteclise. This volume. MARTINSSON, A. 1973. Editor's column: Ecostratigraphy. Lethaia, 6, 441-443. MEIEN, S. V. 1989. Introduction to the Theory of Stratigraphy. Nauka, Moscow [in Russian]. RUKINA, G. m. 1992. The technique of correlation of deposits by foraminifers (using the Tournaisian Stage in the Volga-Urals region as example). Bulletin of Moscow Naturalists' Society Sect. Geology, 67, 99-107 [in Russian]. 1993. Development of foraminiferal assemblages at the Tournaisian-Vis6an Boundary. Stratigraphy and Geological Correlation, 1, 62-66. SCHVETZOV, M. S. 1932. General Geological Map of the European Part of USSR. Sheet 58 (Northwestern quarter of the sheet). Transactions of the United Geological and Prospecting Service of USSR, F83, Moscow. VDOVENKO, M. V., AISENVERG, D. YE., NEMIROVSKAYA, T. I. & POLETAEV, V. I. 1990. An overview of Lower Carboniferous biozones of the Russian platform. Journal of Foraminiferal Research, 20, 184-194.

Irish Dinantian biostratigraphy: practical applications G. LL. J O N E S

& I. D. S O M E R V I L L E

D e p a r t m e n t o f Geology, University College Dublin, Belfield, Dublin 4, Ireland

Abstract: For the last two decades there have been considerable advances in the dating and correlation of marine Dinantian carbonate sequences in Ireland, as elsewhere in Europe, primarily through the increased precision made possible by the use of microfossils, such as conodonts, foraminifers and miospores, together with macrofossils such as rugose corals. The advance was accelerated by the availability of boreholes up to 2 km deep drilled by mineral exploration companies in Ireland who routinely use biostratigraphy. Detailed biostratigraphic biozonations are now established and are used for geological mapping and borehole zonation. Data are assessed and attempts are made to highlight some of the practical problems encountered in locating Dinantian stage boundaries, recognizing biozones in Ireland, and suggesting correlations with other biozonation schemes in Europe. Although there are many difficulties in recognizing each of the stage boundaries in Ireland, two in particular present major problems-the Courceyan/Chadian and Holkerian/Asbian boundaries. From the work of Conil, Groessens and co-workers in Belgium, conodont and foraminiferal biozonation has been applied to the British and Irish Dinantian stages. Unfortunately, the bases of both the Chadian and Asbian stages rarely contain the zonal taxa. Also, at both stratigraphic levels, there are difficulties in comparing basinal and platform faunas. New biostratigraphic data in Ireland have permitted the recognition of two new intervals within the late Asbian Cf6-~ Subzone, referred to informally as Cf6~l and Cf6-y2. Until the Dinantian stages in Britain and Ireland are redefined biostratigraphically in the existing stratotype sections, or new stratotype sections are defined with faunal criteria, it is becoming more expedient and practical to recognize and define biozones in Ireland which can be identified and correlated with other Dinantian sections in continental Europe. This paper presents correlations with Belgian and Russian biostratigraphic schemes.

Since the 1970s there have been considerable advances in the dating and correlation of Dinantian carbonate sequences in Ireland (Fig. 1), primarily through the increased precision made possible by the use of microfossils, such as conodonts and foraminifers. These advances were enhanced by the release of confidential micropalaeontological data from deep boreholes drilled by base-metal exploration and mining companies operating in Ireland in the 1980s and 1990s. The principal areas of interest to these mining companies, from where much of these new microfossil data are derived, include the Dublin Basin, the Shannon Trough and the surrounding platform areas of the Irish Midlands (Fig. 1). Much of the information is now published and a considerable microfaunal database has been compiled (see Clayton et al. 1980; Strogen & Somerville 1984; Somerville & Jones 1985; Jones et al. 1988; Strogen et al. 1990, 1995; Pickard et al. 1992, 1994; Somerville & Strogen 1992; Somerville et al. 1992a, b, c). This paper assesses the data and attempts to highlight some of the practical problems encountered in: (i) using different macrofossil and microfossil groups in dating borehole and

outcrop successions; (ii) locating Dinantian stage boundaries and establishing faunal/floral criteria for recognizing biozones (Fig. 2); and (iii) correlation with other biozonation schemes in Europe, particularly in Belgium and Russia (Fig. 3), where much recent data are now available.

Historical review" of Irish Dinantian biostratigraphy Until the late 1960s most biostratigraphic work on Dinantian successions in Ireland was on well exposed, but not always continuous, coastal sections in eastern Ireland (Matley & Vaughan 1906, 1908; Smyth 1915, 1920, 1950; Hudson et al. 1966b), the south of Ireland (Smyth 1930; Hudson & Philcox 1965) and the west and northwest of Ireland (Douglas 1909; George & Oswald 1957; Caldwell 1959; Shephard-Thorn 1963; Hudson et al. 1966a). In virtually all cases this involved the use of macrofossils: corals and brachiopods in shallow-water platform facies, and goniatites in deeper water mud-mounds and basinal sequences.

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 371-385.

372

G. LL. JONES & I. D. SOMERVILLE

(starved) Basin Sheet mudbanksand inter-bankshales ~

andDeepershaleShelf limestone Basinal calciturbidites and shales Predominantly shale

1

n.v~e.~

sandstone and shale

~ Evaporitesand L,u,., peritidallimestone Shallowshelf limestoneoverlain J~ by deltaic sandstone and shale ~ Shallowshelf limestone ~ and sandstone ~ Shallow shelf bioclastic limestone

Fig. 1. Generalized Irish palaeogeographic provinces in the Dinantian. C, Cork; D, Dublin; K, Kingscourt; L, Limerick; N, Navan. (After Sevastopulo 1981; Strogen & Somerville 1984; Somerville & Jones 1985; Phillips & Sevastopulo 1986; Nolan 1986, 1989; Jones et al. 1988; Strogen 1988; Strogen et al. 1990, 1995; Somerville et al. 1992a,b; Somerville & Strogen 1992; Pickard et al. 1992.) More recently, studies in rugose corals have upheld their biostratigraphic value (Fig. 2) with the recognition of many short-range genera and species (see Poty 1981, 1985, 1994; Somerville & Strank 1984a, b; Somerville et al. 1986, 1989, 1992a, b; Nudds & Somerville 1987; Sevastopulo & Nudds 1987; Mitchell & Somerville 1988; Mitchell 1989; Somerville 1994; Strogen et al. 1995). Brachiopods, on the other hand, have a much more limited value with only occasional diagnostic taxa with short stratigraphic range (see Riley 1993 for a recent summary). Many species of Courceyan brachiopod are not known in Vis6an strata, and several Vis+an brachiopod species show strong facies control, being

confined to mud-bank facies (Brunton & Tilsley 1991). Goniatites have been studied in basinal sequences in Ireland, notably in the South Munster Basin (Naylor 1969) and Dublin Basin (Smyth 1951), but little biostratigraphic refinement has been achieved such as that in the Craven Basin (Riley 1990, 1993). From the late 1960s the emphasis on biostratigraphic studies in Ireland has switched to microfossils, particularly conodonts and foraminifers. Important published works on conodonts and/or foraminifers include: Aldridge et al. (1968); Austin (1968, 1974, 1976); Austin & Aldridge (1969); Mamet (1969); Austin et al. (1970); Sheridan (1972); Matthews & Naylor (1973); Austin & Husri (1974); Sleeman et al. (1974, 1983, 1986); Marchant (1974); Conil & Lees (1974); Austin & Mitchell (1975); Conil (1976); Conil & Lys (1977); Clayton et al. (1977, 1980); Johnston & Higgins (1981); Keeley (1983); Marchant et al. (1984); Strogen & Somerville (1984); Somerville & Jones (1985); Sevastopulo & Nudds (1987); Jones et al. (1988); Strogen et al. (1990, 1995); Jones (1991); Pickard et al. (1992); Kelly & Somerville (1992); Somerville et al. (1992a, b,c) and Rees (1992). Conodonts and foraminifers have allowed detailed biostratigraphic biozonations to be established which are of direct use both for geological mapping and in borehole zonation. In addition, a number of unpublished theses contain important biostratigraphical data on Irish Dinantian faunas and floras. These include: Johnston (1976); Whitaker (1976); Jones (1977); Marchant (1978); Keeley (1980); Polgar (1980); Browne (1981); ten Have (1982); Thornbury (1985); Lewis (1986); Nolan (1986); Rees (1987); Shearley (1988); Kelly (1989); Clipstone (1992) and Gallagher (1992).

Fossil groups used in Irish biostratigraphy The fossils groups used in biostratigraphy are dependant on the facies. In non-marine and shale-rich transgressive facies miospores are used. However, in the limestone facies that dominates the Dinantian succession, conodonts and foraminifers are preferred, supported by corals, ostracodes, algae and problematica.

Miospores

In siliciclastic facies, miospores are recovered from dark grey to black argillaceous rocks. Irish workers have established a good palynological

IRISH DINANTIAN BIOSTRATIGRAPHY

373

IRISH DINANTIAN B I O Z O N A T I O N UK&IRL BELG I-IAN STAGES is ~

BRIGAN -TIAN

CONODONTS SHELF outer I SHELF main Gnathodus girtyi collinsoni Lochreia nodosa tVlestognathus bipluti Gnalhodus girtyi girtyi Gnathodus bilineatus bilineatus

(late)

-ASBIAN' (early) E

D

.

HOLKE -RIAN

Taphrognathus

z ,~ oO ~>

ARUN -DIAN

Nodosar~iscus, Paraarcha~iscus @ 8 involutus, Palmospiroptectammina s~/zranica, Uralodiscus, Glomodiscus ~-'~ Viseidiscus all @ involutus Eostaffetla

Lochreia commutata

Gnathodus homopunclatus

(late) .,CHAD -IAN

Mestognathus beckmanni

(early)

• Meslognathus prmbeckmanni P. bischoffi

(late) "--" Z<: <: Z n-

Eotextularia diversa Tetrataxis

Cf2

P. inornatus P. sp~tus ......

Paraendothyra Toumayella discoidea Eoforschia, Eblanaia michoti

I. N . D .C . Siphonodeila

Pu

Mediocns, Omphalotis, Siphonophyllia 0~1 Pseudoammodiscus, E.laxa cylindrica Ps. gravaCa, Brunsia, Sychnoelasma Valvulinella, Dainella urbanowitschi

Dol. latus E D. bouckaerti o~ Polygna~us E. bul~ncki mehli lafus Eotaphrus Ps.oxypageus cf. bultyncki Doilymm Ps., P. c. carina hass~ multisb'iatus

D Siphonodendron sociale Siphonodendron C marlini Siphonophyllia garwoodi B Dorlodotia briarti

Eoendothyranopsis, GioboDorlodotia pseu 0{2 endothyra, Plectogyranopsis -dovermiculare Pseudolituotubella A Cart. compacta Eoparastaffella simplex Ax. simplex

Cf3

O

(early)

Cf4

P.bischoffi

A j ,o

COURC -EYAN

RUGOSE MIOS CORALS PORE NC Loeblichia paraamminoides J-K Ork~naslrea Warnanteila I Corwenia mgosa pars Palastrea regia Janischewsldna H A. floriformis VF Asteroarchaediscus Dibunophyllum 72 Brady/n.aro.tula, bipartitum ,Howbh/n/a zxaoyana G Haplolasma cf. NM Cf6 i Cribroslomum lecomptei densum ¥1 Pseudoendothyra Siphonodendron Neoarcha~iscus, pauciradiale Nodasperodiscus, F Siphonodendron TC 3~-~ Archa~iscus _. junceum all @ angulatus Diounopnyllum Vissariotaxis oourtonense Koskinotextularia Lilhostrotion Pojarkoveila nibelis E vomcale TS Paraar~iscus Cf5 Lithostro~on @ cof~cavus araneum Palmotextularia (monolam) FORAMINIFERS

Cfl

Caninophyllum modavense Sychnoelasma hawbankense rSychnoetasma konincki Sychnoelasma clevedonensis Cyathaxonia comu Caninia comucopim Zaphrentites delanouel Zaphrentites .v.augha.nj. ""

CM

PC

BP HD VI

Fig. 2. Irish Dinantian biozonation schemes. Conodonts modified from Johnston (1976), Somerville & Jones (1985), Varker & Sevastopulo (1985), Sleeman et al. (1986), Jones et al. (1988), Somerville et al. (1992b, c) and Riley (1993). Foraminifera from Conil et al. (1980,1989,1991), Strogen et al. (1990), Gallagher (1992), Pickard et al. (1992) and Somerville et al. (1992b). Miospores from Higgs (1984) and Higgs et al. (1988). Corals adapted from Somerville & Jones (1985), Jones et al. (1988), Mitchell & Somerville (1988), Mitchell (1989), Gallagher (1992) Kelly & Somerville (1992), Somerville et al. (1992a, b), Somerville (1994) and Strogen et al. (1995). Conodonts: D., Dollymae; Dol., Doliognathus; E., Eotaphrus; G., Gnathodus; P., Polygnathus; Ps., Pseudopolygnathus; c., cornmunis; m., rnehli; S., Scaliognathus; I.N.D.C., interval of no diagnostic conodonts. Foraminifers: E., Endothyra; Ps., Pseudolituotuba. Corals: A., Actinocyathus; Ax., Axophyllum; Carr., Carruthersella.

374

G. LL. JONES & I. D. SOMERVILLE RUSSIAN PLATFORM

BELGIUM

IRELAND FORAM STAGE INIFER

FORAMtNIFERS

BRIGAb -TIAN

~5

HORIZON

CONODONTS

Cf6 i5

FORAMINIFERS

Paragnathodus nodosus

Venevsky Endothyranopsis crassa Archagdiscus gigas

Gnathodus late

Neoarchaediscus

Cf 1' 2 6

7

bifineatus

Mikhailovsky

z re

y1 -ASBIAI~ early HOLK E -RIAN ARUN -DIAN (late) CHAD -IAN (early)

Cf5 Koskinotextulafia Pojarkovella nibelis Mestognathus 8 ~5 Cf4 beckmanni Y 13-7 [3 Eopara - praebeck Cf -staffella -manni 4 et2 ct2 Cf5

Tulsky

Endothyranopsis compressa Propermodiscus krestovnikovi

Bobrikovsky

Uralodiscus rotundus Planodiscus primaevus

z

< > J

~ < --, 0

Tetrataxis Eotextularia

Kosvinsky

z Polygnathus

~_

communis carina

--

Endothyra elegia Palaeospiroplect -ammina D , . E. diversa bouckaerti Tetrataxis -

Cf2

COURC -EYAN

Paraendothyra Cfl

Kiselovsky

Cfl

Chernyshinella

Spinoendothyra costifera Tuberendothyra tuberculata

z _< Cherepetsky co

<

(early)

Mestognathus beckmanni

Eoparastaffella simplex Eoendothyranopsis

Radaevsky

Sc. anchorafis

Cf3

Cf2

Gnathodus texanus

ctll

Cf3 (late)

Gnathodus b. bifineatus

Aleksinsky z< .u.l

et-[~ 13

, ctl

CONODONTS

Siphonodella

z

z re

•:

o

< i-

Karakubsky

I-

-i-

Chernyshinella disputabilis Chernyshinella glomiformis

Upinsky

Malevsky Gumerovsky

Bisphaera malevkensis Earlandia minima

Gnathodus typicus

P. c. carina .,

S, isosticha Siphonodella quadrupfic -ata 'Siphono Patro-della gnathus kononovae' andersoni S. duplicata Pa.variabifis PatroSiphono gnathus -della crassus sulcata

Fig. 3. Comparison of foraminiferal and conodont biozonation schemes for Ireland, Belgium and the Russian

Platform. Belgian scheme after Conil et al. (1991). Russian Platform after Vdovenko et al. (1990), Makhlina (this volume) and Rukina (this volume). D., Dollymae; P., Polygnathus; Pa., Patrognathus; S., Siphonodetla; Sc., Scaliognathus; c., communis; E, Eotextularia.

zonation of the Dinantian of Ireland and western Europe (Fig. 2; Clayton & Higgs 1979; Clayton et al. 1977, 1978, 1980; Keegan 1981; Higgs 1984; Marchant et al. 1984; Higgs et al. 1988a, b). However, for most of the Dinantian limestone succession, the miospores recovered tend to be too oxidized and poorly preserved to be useful. They are routinely used in the early to mid-Courceyan and in the Brigantian, but for the rest of the Dinantian they are rarely used,

except in northwest Ireland where there are numerous siliciclastic intervals interbedded with limestones (Higgs 1984).

Conodonts

Conodonts are very useful in dating and correlating Courceyan and Chadian successions, and since much of the Irish mining drillcore covers this

IRISH DINANTIAN BIOSTRATIGRAPHY part of the Dinantian succession, 1 m + channel samples (selected, cumulative, split-core) are routinely taken for dating. However, Arundian to mid-Asbian (middle Vis6an) conodonts are remarkably scarce and mostly comprise longranging taxa. It is apparent that after radiation and peak diversity, late Courceyan (Tournaisian) conodonts suffered a serious decline of the 12 genera and >26 species, with only a few surviving Courceyan genera (Varker & Sevastopulo 1985; Sweet 1988; Webster & Groessens 1990). A second, mid-Asbian to Brigantian (late Vis6an), radiation event occurred, and once again conodonts become useful in biostratigraphic correlations. In the late Courceyan (Tournaisian) rocks in Ireland, conodonts show a marked facies control, which has resulted in the recognition of two broadly coeval biozonal schemes for two contrasting provinces (Fig. 2; Somerville & Jones 1985; Varker & Sevastopulo 1985): (i) an outer shelf/deep-water ramp association centred on the South Munster Basin (Fig. 1) and continued across to Belgium, and characterized by Scaliognathus anchoralis, Dollymae bouckaerti, Doliognathus latus and gnathodids (see Sleeman et al. 1986); and (ii) an inner to mid-shelf association centred on the Dublin Basin and south central Midlands Platform, and characterized by polygnathids and pseudopolygnathids. It is thus apparent that there are two schemes, one applying to the shelf edge and the other to the main shallowwater shelf facies (Figs 1, 2).

Foraminifers

Foraminifers are encountered routinely from drillcore in spot samples for petrographical studies and can be locally abundant. Repeated sampling at closely spaced intervals is often undertaken in Ireland for the precise location of biozonal and stage boundaries. Apart from mudstone-dominant sequences, or intervals that have suffered dolomitization or enhanced diagenesis affecting the wall-structure, foraminifers are usually identifiable at generic level, and in correctly oriented sections, at specific level. Unlike conodonts, Tournaisian foraminifers are sparse and first become important in late Tournaisian rocks (Fig. 2). Like conodonts, foraminifers diminish rapidly in the latest Tournaisian times, but unlike conodonts, many new early Vis6an genera appear, especially the diagnostic Eoparastaffella. In the Arundian and Holkerian (mid-Vis~an) time interval, when conodonts have little biostratigraphic

375

significance, archaediscids become prolific in platform sequences, but then show a decline in diversity (but not abundance) in the late Holkerian to early Asbian interval. During this period the fine-grained basinal limestones contain archaediscids represented by dwarf forms. Like conodonts, foraminifers show a major radiation event in the late Asbian to early Brigantian interval, with the appearance of many new genera and species such as Howchinia bradyana. The two microfossil groups used in conjunction can achieve great precision in dating and correlating sequences (Fig. 2).

Other micro f o s s i l groups

In thin section other microfossils are used to supplement the foraminiferal record, particularly when diagnostic foraminifera are sparse or absent. Special mention should be made of the appearance of the calcareous algae Globochaetes (Fig. 4h) in the late Courceyan and of Koninckopora in the Chadian, and the demise of the latter at the end of the Asbian (Marchant et al. 1984; Strogen et al. 1990, 1995; Pickard et al. 1992; Somerville et al. 1992b). Also the calcisphere Mendipsia leesi (Fig. 4a) and the microproblematicum Sphaerinvia piai (Figs 4d, f) have a limited range in the late Courceyan and Chadian (Somerville et al. 1992c, Strogen et al. 1990). A closely related taxon is Sphaerinvia sp. A, characterized by a thick wall with strongly developed trefoil septa (Fig. 4i), which occurs in the early to mid-Vis6an. The algae Fasciella, Kulikia and Ungdarella (see Somerville et al. this volume) are important in Asbian rocks, while Calcifolium is restricted to Brigantian strata. Probable red algae (aoujgaliids) are first recorded in the late Tournaisian with forms such as Mamatella (Fig. 4e) and ?Aoujgalia (Fig. 4c). Other locally useful microproblematica include Salebra sibirica (Fig. 4b), Draffania (Somerville et al. 1992b; Strogen et al. 1995) and Luteotubulus licis (Fig. 4g). The latter, a tubular issinellid (Order Beresellida), has a distinctive thick wall of yellow hyaline calcite and is characterized by rare complete diaphragms. In Ireland this taxon is restricted to early Vis6an (late Chadian to Arundian) strata (Pickard et al. 1992; Strogen et al. 1995) and is typically found in high-energy grainstones. It is known from age-equivalent Moliniacian rocks in Belgium (Fig. 3; Vachard 1994). From heavy residues, silicified and phosphatized ostracode valves are occasionally recovered and can be useful adjuncts to

I I

- KZ-"

~

~

~

~,:~ i¸~i

!

m

.<

©

rn

©

IRISH DINANTIAN BIOSTRATIGRAPHY biostratigraphy (ten Have 1982). Also of note is the frequent recovery of ichthyoliths, which current research (see both Ivanov and Lebedev this volume) has shown will make a significant contribution to future biostratigraphic studies in the Dinantian.

Corals Courceyan rugose corals, as with foraminifers, are represented by only a few genera and many of these are long-ranging, extending into the Vis6an. However, in the late Courceyan several short-range coral genera and species appear: Cyathoclisia modavense, Cyathaxonia cornu, Sychnoelasma hawbankense and Caninophyllum patulum patulum which correspond to the Cf3 Foraminifera Zone (Fig. 2). They are succeeded by Siphonophyllia cylindrica and Sychnoelasma urbanowitschi in the early Chadian (latest Tournaisian; Somerville 1994). In the Vis6an, ten rugose coral assemblage biozones (A-K) have been defined in Britain by Mitchell (1989) which can be readily identified in Ireland. Many new rugose corals appeared associated with the widespread development of shallow-water platforms across western Europe. In Ireland these include solitary dissepimented forms such as Carruthersella compacta, Spirophyllum praecursor, Axophyllum simplex (Somerville et al. 1992a) and the first fasciculate colonial coral Dorlodotia pseudovermiculare. These late Chadian corals characterize the basal Vis~an rocks (Cf4a2 Foraminifera Zone). In Arundian time there is a diversification of fasciculate forms with Dorlodotia briarti and several species of Siphonodendron including the large forms: S. scaleberense, S. sociale, and S. martini (Nudds 1980; Kelly 1989). Large caniniids such as Siphonophyllia garwoodi are restricted to this stage (Kelly 1989; Kelly & Somerville 1992; Somerville et al. 1992b). This coral assemblage is associated with a rich

377

foraminiferal assemblage dominated by archaediscids at the involutus stage (Cf4~-/5 subzones). The Holkerian Stage is marked by the appearance of large cerioid corals such as Lithostrotion araneum and L. vorticale associated with palaeotextulariid foraminifers and archaediscids at the concavus stage (Cf5 Foraminifera Zone; Gallagher 1992; Somerville et al. 1992b). In the Asbian Stage the first appearance of Dibunophyllum and small forms of Siphonodendron such as S. junceum and S. pauciradiale coincides with the appearance of archaediscids at the angulatus stage (Cf6c~ foraminifera subzone). The late Asbian is marked by the appearance of diphyphyllids, Lithostrotion maccoyanum, Haplolasma cf. densum and the ubiquitous Dibunophyllum bipartitum (Gallagher 1992; Somerville et al. 1992b; Strogen et al. 1995). This is matched by a similar increase in the diversity of foraminifers (Cf6"/ foraminifera subzone). The Brigantian Stage is readily identified by rugose corals, with the appearance of Actinocyathus floriformis, Palastraea regia and Corwenia rugosa (Strogen et al. 1995), followed in younger Brigantian rocks by Orionastraea (Nudds 1979).

The recognition of Dinantian stage boundaries and biozones in Ireland, and their comparison with other European biozonations The Dinantian stages and the stratotype sections defined for the British Isles in George et al. (1976), and supplemented by additional faunal data (Ramsbottom 1981), were reviewed by Austin & Davies (1984) and Austin & Moore (1989), and have recently been critically assessed by Riley (1993). It is now quite evident that many stage boundaries lack diagnostic faunal evidence and constraints. Moreover, key biostratigraphic markers often first appear some distance above the defined base of a stage in

Fig. 4. Microproblematica and calcareous algae from the Dinantian of Ireland. (a) Mendipsia leesi Conil & Longerstaey 1980, Slane Castle Formation, late Tournaisian, Athboy Borehole (1349/2), Co. Meath, 1423.8m. Scale bar 100 #m. (b) Salebra sibirica Bogush 1976, Moathill Formation, late Tournaisian, Athboy Borehole (1349/2), Co. Meath, 1644.4m. Scale bar 200 #m. (c) Aoujgaliid (?Aoujgalia), Feltrim Limestone Formation (Waulsortian facies), late Tournaisian, Bray Hill Quarry, Co. Meath. Scale bar 200 #m. (d) Sphaerinviapiai Vachard 1980 (longitudinal section showing ?tabulae), Slane Castle Formation, late Tournaisian, Woodtown Borehole (91-3347/1), Co. Meath, 693.6m. Scale bar 200#m. (e) Aoujgaliid (?Mamatella), Feltrim Limestone Formation (Waulsortian facies), late Tournaisian, Bray Hill Quarry, Co. Meath. Scale bar 200 #m. (f) Sphaerinv& piai Vachard 1980, Feltrim Limestone Formation (Waulsortian facies), late Tournaisian, Athboy Borehole (1349/2), Co. Meath, 1176.4m. Scale bar 100#m. (g) Luteotubulus licis Malakhova 1975, Arundian, Urlingford Borehole (3246/6), Co. Tipperary, 119.5m. Scale bar 500/~m. (h) Globochaetes, Slane Castle Formation, late Tournaisian, Athboy Borehole (1349/2), Co. Meath, 1401.8m. Scale bar 200#m. (i) Sphaerinvia sp. A, early Vis6an, Bunbrosna Borehole, BU-I, Co. Westmeath, 876ft (267.9 m). Scale bar 100#m.

378

G. LL. JONES & I. D. SOMERVILLE

stratotype sections (e.g. Arundian in South W a l e s - s e e Simpson & Kalvoda 1987). Without faunal criteria it is quite difficult to recognize stage boundaries away from the stratotype sections. Thus, until such time as the Dinantian stages in Britain are redefined biostratigraphically in the existing stratotype sections, or new stratotype sections defined with faunal criteria, it is becoming more expedient and practical to recognize and define biozones in Ireland which will be identified and correlated with other Dinantian biozonal schemes in Europe, particularly in Belgium and Russia (Figs 2, 3). The following is a summary of some of the problems encountered in recognizing the Dinantian stage boundaries in Ireland, and of the faunal criteria used to recognize biozones and sub-biozones. Famennian/ Courceyan boundary

The base of the Courceyan Stage is defined at the stratotype section in the south of Ireland (South Munster Basin; Fig. 1) on microfloral criteria (George et al. 1976). The basal nonmarine and marine clastic facies are well zoned with microflora (Higgs 1984; Marchant et al. 1984; Higgs et al. 1988a, b). The Courceyan Stage

In Ireland foraminifers are sparse in marine Courceyan strata, and most of the stage is zoned on conodonts, which are used extensively in dating and correlating sequences in boreholes (Somerville & Jones 1985; Jones et al. 1988; Pickard et al. 1992; Somerville et al. 1992b) and outcrop (Austin et al. 1970; Sleeman et al. 1974; Clayton et al. 1977; Johnston & Higgins 1981). In these sections precision is achieved with the recognition of conodont biozones and subbiozones. Furthermore, these conodont zones can be used as a basis for the construction of isopach maps which can identify areas of rapid subsidence for given time slices (Jones et al. 1988; Pickard et al. 1994), or palaeogeographic reconstructions for intervals within the Courceyan Stage (Sevastopulo 1981, 1982; Nolan 1989; Strogen et al. 1990). Early~late Courceyan boundary

The base of the late Courceyan in Belgium is marked by the incoming of the conodont Polygnathus communis carina. However, in Ireland this biozone is only recognized in the shelf-edge faunas in the South Munster Basin (Figs 1, 2).

Over the remainder of the shelf, P.c. carina forms a sub-biozone of the Pseudopolygnathus multistriatus Biozone, whose incoming marks the base of the late Courceyan (Fig. 2). This boundary is also recognized by the incoming of the Cf2 Foraminifera Zone of Conil et al. (1989, 1991), based on a sparse assemblage including Tournayella discoidea, Eoforschia, Paraendothyra and Eblanaia michoti. This contrasts with the Russian Platform (Moscow Basin) where foraminifers are abundant and the Cf2 biozone can be correlated with the Kiselovsky Horizon (Fig. 3; Vdovenko et al. 1990; Makhlina this volume; Rukina this volume). In Belgium the Hastarian/Ivorian boundary (see Groessens 1975; Conil et al. 1977; Paproth et al. 1983) equates with the early/late Courceyan boundary (Fig. 3). The former boundary was claimed to be recognized in the UK (Ramsbottom & Mitchell 1980), but its usage is now not generally accepted (see Riley 1993). Over the main shelf the Ps. multistriatus Biozone is succeeded by the Polygnathus mehli Biozone, whilst somewhat later the P.c. carina Biozone is succeeded by the Scaliognathus anchoralis Biozone at the shelf edge. At the same level an assemblage including Eotextularia diversa, Palaeospiropectammina and Tetrataxis is diagnostic of the Cf3 Foraminifera Biozone of Conil et al. (1980, 1989, 1991). In the Russian Platform (Fig. 3) this zone is equivalent to the Kosvinsky Horizon (see Vdovenko et al. 1990; Rukina this volume). In the Dublin Basin (Fig. 1) this biozone coincides approximately with the appearance of the Caninophyllum patulum coral assemblage biozone (Jones et al. 1988; Strogen et al. 1990; Somerville 1994).

Courceyan/ Chadian boundary

The Courceyan/Chadian stage boundary is the least satisfactory and by current definition the most difficult boundary to identify away from the stratotype. As recently highlighted by Riley (1990, 1993), the base of the Chadian Stage at the stratotype is an arbitrary 'spike' without faunal constraint. It was allegedly defined at a horizon below the first appearance of Eoparastaffella in the stratotype section (George et al. 1976). This has not been verified (Fewtrell et al. 1981; Riley 1990), and it is quite clear that other important Chadian faunal markers appear some distance above the base. In the Dublin Basin (Fig. 1), late Courceyan (Cf3) foraminiferal faunas are often indistinguishable from those of the early part of the Chadian (Cf4o~l), and are now clearly of late

IRISH DINANTIAN BIOSTRATIGRAPHY Tournaisian age (Conil et al. 1989, 1991) as they occur stratigraphically below the basal Vis6an marker Eoparastaffella simplex (Strogen et al. 1990, 1995). In Ireland this late Courceyan to early Chadian interval coincides with a period of widespread faunal restriction affecting many faunal groups (conodonts, foraminifers, corals) and major changes in depositional environments. In the Dublin Basin (Fig. 1) the latter results primarily from increased tectonic influence and a change from predominantly ramp sedimentation, with Waulsortian facies, to platform and basin conditions (Jones et al. 1988; Strogen et al. 1990, 1995). In Belgium this same time interval is equivalent to the basal part of the Moliniacian (Conil et al. 1988, 1989, 1991; Hance et al. 1994) and recognized biostratigraphically by the Mestognathus praebeckmanni - beckmanni conodont lineage. We cannot accurately identify the Courceyan/Chadian boundary in Ireland.

379

Chadian/Arundian boundary. Close to this boundary, at the base of the Arundian Stage in Ireland and well documented in the Athboy Borehole and Navan Mine boreholes (Strogen et al. 1990), is a rich and diverse archaediscid foraminiferal assemblage dominated by Uralodiscus, Glomodiscus and Viseidiscus at involutus stage, together with Eostaffella (Cf4/3-6 subzones: Strogen et al. 1990; Pickard et al. 1992; Somerville et al. 1992b). Eostaffella cf. parastruvei characterizes the Arundian Stage here, and partially nodose forms of archaediscids along with Paraarchaediscus at involutus stage, appear in the late Arundian (Cf4~ subzone). In the North Caspian Syneclise and Dneiper-Donetz Depression a very similar suite of archaediscids characterizes the Bobrikovsky Horizon (Fig. 3; see Vdovenko et al. 1990). This diversity of foraminifers is matched in the late Arundian by the rugose corals with abundant fasciculate Siphonodendron species and Dorlodotia briarti (Rugose Coral Zones B-D).

Early Chadian/late Chadian boundary

This boundary is recognized in Ireland in Limerick (Somerville et al. 1992b) and in the Dublin Basin (Strogen et al. 1990; Pickard et al. 1992; Somerville et al. 1992a, c) by the first appearance of a rich and diverse Cf4o~2 foraminiferal suite including the diagnostic Eoparastaffella simplex together with Pseudolituotubella and Eoendothyranopsis and the conodont Gnathodus homopunctatus. These are accompanied by solitary dissepimented rugose corals (Coral Zone A) and rare calcareous algae Koninckopora tenuiramosa and Koninckopora minuta (bilaminar wall forms). In the Russian Platform the same boundary defines the base of the Radaevsky Horizon (Fig. 3), which contains a similar foraminiferal assemblage (see Vdovenko et al. 1990; Rukina this volume). L a t e Chadian to Brigant&n interval

The subdivision of the succeeding Vis6an rocks (late Chadian to Brigantian) and the recognition of Dinantian stages is, in practice, heavily dependent on the evolution of foraminifers of the family Archaediscidae (from the Arundian onwards), as described by Conil et al. (1980) and amended by Brenckle et al. (1987). The various types of structure and coiling in archaediscids (involutus, concavus, angulatus and tenuis stages) are biostratigraphically important, as well as the recognition of genera belonging to the two subfamilies Kasachstanodiscinae and Archaediscinae (Brenckle et al. 1987).

Arundian/Holkerian boundary. The base of the Holkerian Stage in Ireland is taken at the first appearance of monolaminar palaeotextulariids such as Koskinotextularia and the rare Pojarkovella nibelis, which is often missing in basinal sequences (Strogen et al. 1990; Pickard et al. 1992; Somerville et al. 1992b). They are diagnostic elements of the Cf5 Foraminifera Zone of Conil et al. (1980, 1989, 1991). In Ireland the Holkerian Stage also marks the appearance of cerioid colonial corals with large species of Lithostrotion (Rugose Coral Zone E). This boundary marks a subtle change in the archaediscid population here, with the demise of several of the Arundian genera (e.g. Uralodiscus and Glomodiscus). In practice it is the dominance of Paraarchaediscus at concavus stage and frequent nodose archaediscids at concavus stage which are diagnostic. Holkerian/Asbian boundary. This is a very difficult boundary to detect away from the stratotype, because of the scarcity of diagnostic foraminiferal taxa, particularly in shallow-water platform sequences. It is now recognized that in the Asbian boundary stratotype section the Cf6 zonal genera Archaediscus and Neoarchaediscus at the angulatus stage (see Conil et al. 1991) first appear some distance above the base, together with the rugose corals Dibunophyllum bourtonense, Siphonodendron junceum and S. pauciradiale (Rugose Coral Biozone F). In 'practice, particularly in borehole sections, the base of the

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Asbian Stage and the Cf6 zone in Ireland is placed below the incoming of Vissariotaxis (Strogen et al. 1990; Pickard et al. 1992; Somerville et al. 1992b). Unfortunately, although fairly common in deepwater basinal sequences, this taxon is very rare in shallow-water platform successions and hence has severe limitations as a zonal indicator. Nevertheless, in the Russian Platform this genus is one of several new constituents that characterize the Tulsky Horizon (Fig. 3; see Vdovenko et al. 1990). Moreover, gigantoproductid brachiopods also appear for the first time in the Russian Platform in the Tulsky Horizon. In platform sequences in Ireland, it is often virtually impossible to distinguish early Asbian from late Holkerian (Cf5) foraminiferal assemblages when the diagnostic Asbian taxa are absent. Early Asbian/Late Asbian boundary. In the late Asbian in the Limerick and Dublin regions, a rich and diverse foraminiferal suite is always developed (Strogen et al. 1990, 1995; Pickard et al. 1992; Somerville et al. 1992b; Gallagher this volume), characteristic of the Cf6"/subzone and very similar to that described in Belgium (Laloux 1988; Conil et al. 1991). In the lower part, within the newly recognized Cf6'71 subzone (Somerville et al. 1992b; Strogen et al. 1995), there are numerous Cribrostomum lecompteii and Pseudoendothyra. In the upper interval (Cf6-y2 subzone) are locally abundant Bradyina rotula, and occasional Howchinia bradyina, Saceamminopsis and Asteroarchaediscus of latest Asbian age (Somerville et al. 1992b; Strogen et al. 1995; Fig. 2). Also, the conodonts Gnathodus bilineatus bilineatus and Gnathodus girtyi girtyi and the rugose coral Dibunophyllum bipartitum (Rugose Coral Biozone G) occur in the same subzone (Strogen et al. 1995). The Cf67 subzone fauna is similar to that of the Aleksinsky Horizon (Fig. 3) in the Russian Platform (see Vdovenko et al. 1990), in which Howchinia, Bradyina and G.b. bilineatus are first recorded. Asbian/Brigantian boundary. In the early Brigantian in Ireland, diagnostic foraminifers of the Cf66 Subzone (see Conil et al. 1991) such as Loebliehia paraamminoides and Janischewskina are sparse, although Howchinia bradyina and Asteroarchaediscus can be locally very abundant. A notable extinction close to the boundary is that of the calcareous alga Koninckopora as seen in the Limerick region and the Dublin Basin (Fig. 1; Pickard et al. 1992; Somerville et al. 1992b; Strogen et al. 1995). However, in platform

sequences as at Kingscourt (Strogen et al. 1995), a diverse Brigantian coral assemblage is recorded including Actinocyathus floriformis, Palastraea regia and Corwenia rugosa (Rugose Coral Biozones H-I), together with the conodont Mestognathus bipluti (Fig. 2). The Brigantian calcareous alga Calcifolium is also recorded in the north Cork area (Gallagher 1992). Younger Brigantian rocks in Ireland are characterized by the conodont Lochreia nodosa (Somerville et al. 1992b) and the coral Orionastraea (Rugose Coral Biozones J-K; Nudds 1979). The Mikhailovsky Horizon on the Russian Platform (Fig. 3) sees the first appearance of Janischewskina, Climacammina, Asteroarchaediscus bashkirikus and the alga Calcifolium similar to that known from the type Brigantian section in northern England (George et al. 1976; Ramsbottom 1981). The conodont Gnathodus girtyi collinsoni characterizes the upper Brigantian (Varker & Sevastopulo 1985), but significantly has been recorded from late Asbian rocks in Co. Leitrim (Kelly 1989). The co-occurrence of the two Brigantian marker conodonts G. g. collinsoni and Lochreia nodosa in these late Asbian rocks (Kelly 1989) is unusual. The Asbian age has been confirmed by the goniatite assemblage (Brandon 1977) and the foraminiferal assemblage (Kelly 1989). It is perhaps significant that the fauna occurs in association with an intertidal evaporitic facies of the Meenymore Formation. The Venevsky Horizon on the Russian Platform is impoverished in foraminifers but also contains the important conodonts Paragnathodus nodosus (Lochreia nodosa) and P. mononodosus (Fig. 3).

Conclusions The recent advances in subdividing the late Courceyan with conodonts have been followed by advances in subdividing the late Asbian Cf67 Subzone with foraminifers. Further developments in the biostratigraphy of ostracodes, ichthyoliths, algae, problematica, and so on, are also proving useful. These, and other advances of the last two decades have produced detailed biostratigraphic biozonations for use in borehole zonation and geological mapping. The problems in recognizing the stage boundaries, especially the Courceyan/Chadian and Holkerian/Asbian, have made it more pertinent to use the conflated biozonation scheme shown here. This also allows correlation with other Dinantian sections in Europe, such as the Belgian and Russian biostratigraphic schemes.

IRISH DINANTIAN BIOSTRATIGRAPHY We wish to thank D. Clipstone, S. Gallagher, K. Higgs, J. Kelly, M. Makhlina, N. Riley and G. Rukina for discussion, information and access to unpublished data. We would also like to acknowledge our gratitude to all the mining companies operating in Ireland who granted us access to exploration drillcores and permission to publish palaeontological data from them. We are particularly indebted to our two reviewers E. Paproth and R. Austin, who made valuable comments which significantly improved the paper.

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posium on the Environmental Setting and Distribution of the Waulsortian Facies. E1 Paso Geological Society and University of Texas at E1 Paso, 17-33. NUDDS, J. R. 1987. Courceyan (early Dinantian) biostratigraphy of Britain and Ireland: Coral and conodont zones compared. Courier Forschungsinstitut Senckenberg, 98, 39-46. SHEARLEY, E. 1988. The Geology of the Mitchelstown Syncline, Counties Cork and Tipperary. PhD Thesis, University of Dublin, SHEPHARD-THORN, E. R. 1963. The Carboniferous Limestone succession in north-west County Limerick, Ireland. Proceedings of the Royal Irish Academy, 62B, 267-294. SHERIDAN, D. 1972. The Upper Old Red Sandstone and Lower Carboniferous of the Slieve Beagh Syncline and its setting in the Northwest Carboniferous Basin of Ireland. Geological Survey of Ireland Special Paper, 2, 1-129. SLEEMAN, A. G., HIGGS, K. & SEVASTOPULO,G. D. 1983. The stratigraphy of the Late DevonianEarly Carboniferous rocks of south County Wexford. Geological Survey of Ireland Bulletin, 3, 141-158. , JOHNSTON, I. S., NAYLOR, D. & SEVASTOPULO, G. D. 1974. The stratigraphy of the Carboniferous rocks of Hook Head, Co. Wexford. Proceedings of the Royal Irish Academy, 74B, 227-243. , THORNBURY, B. & SEVASTOPULO, G. D. 1986. The stratigraphy of the Courceyan (Carboniferous: Dinantian) rocks of the Cloyne Syncline, west of Cork Harbour. Irish Journal of Earth Sciences, 8, 21-40. SIMPSON, J. & KALVODA, J. 1987. Sedimentoiogy and foraminiferal biostratigraphy of the Arundian (Dinantian) stratotype. In: HART, M. B. (ed.)

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& JONES, G. LL. 1985. The Courceyan stratigraphy of the Pallaskenry Borehole, Co. Limerick, Ireland. Geological Journal, 20, 377-400. -8~ STRANK, A. R. E. 1984a. The discovery of Arundian and Holkerian faunas from a Dinantian platform succession in North Wales. Geological Journal, 19, 85-105. & 1984b. The recognition of the Asbian/ Brigantian boundary fauna and marker horizons in the Dinantian of North Wales. Geological Journal, 19, 227-237. & STROGEN, P. 1992. Ramp sedimentation in the Dinantian limestones of the Shannon Trough, Co. Limerick, Ireland. Sedimentary Geology, 79, 59 75. --, MITCHELL, M. & STRANK, A. R. E. 1986. An Arundian fauna from the Dyserth area, North Wales and its correlation within the British Isles. -

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Proceedings of the Yorkshire Geological Society, 46, 57-75. --, PICKARD, N. A. H., STROGEN, P. & JONES, G. LL. 1992a. Early to mid-Vis6an platform buildups, north Co. Dublin, Ireland. Geological Journal, 27, 151-172. --, STRANK, A. R. E. & WELSH, A. 1989. Chadian faunas and flora from Dyserth: Depositional environments and palaeogeographic setting of Visban strata in northeast Wales. Geological Journal, 24, 49 66. --, STROGEN, P. 8¢ JONES, G. LL. 1992b. Biostratigraphy of Dinantian limestones and associated volcanic rocks in the Limerick Syncline, Ireland. Geological Journal 27 201-220. Mld-Dlnantlan Waulsortian buildups in the Dublin Basin, Ireland. Sedimentary Geology, 79, 91-116. & SOMERVILLE, H. E. A. 1996. Late Vis6an buildups at Kingscourt, Ireland: possible precursors for Upper Carboniferous bioherms. ,

Micropalaeontology of Carbonate Environments. British Micropalaeontological Society Series. Ellis Horwood, Chichester, 226-237. SMYTH, L. B. 1915. On the faunal zones of the Rush-Skerries Carboniferous section, Co. Dublin.

Scientific Proceedings of the Royal Dublin Society, 14, 535-562. - - 1 9 2 0 . The Carboniferous coast section at Malahide, Co. Dublin. Scientific Proceedings of the Royal Dublin Society, 16, 9-24. - - 1 9 3 0 . The Carboniferous rocks of Hook Head, County Wexford. Proceedings of the Royal Irish Academy, 29B, 523-566. - - 1 9 5 0 . The Carboniferous system in north County Dublin. Quarterly Journal of the Geological Society of London, 105, 295-324. - - 1 9 5 1 . A Vis6an cephalopod fauna in the Rush Slates of County Dublin. Proceedings of the Royal Irish Academy, 45, 25-32. SOMERVILLE, I. O. 1994. Early Carboniferous rugose coral assemblages from the Dublin Basin, Ireland: possible bathymetric and palaeoecological indicators. Courier Forschungsinstitut Senckenberg, 172, 223-229.

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This volume. STROGEN, P. 1988. The Carboniferous lithostratigraphy of southeast County Limerick, Ireland, and the origin of the Shannon Trough. Geological Journal, 23, 121-137 & SOMERVILLE, I. D. 1984. The stratigraphy of the Upper Palaeozoic rocks of the Lyons Hill area, Co. Kildare. Irish Journal of Earth Sciences, 6, 155-173. , JONES, G. LL. & SOMERVILLE, I. D. 1990. Stratigraphy and sedimentology of Lower Carboniferous (Dinantian) boreholes from west Co. Meath, Ireland. Geological Journal, 25, 103-137. - - - , SOMERVILLE, I. D., JONES, G. LL. & PICKARD, N. A. H. 1995. The Lower Carboniferous (Dinantian) stratigraphy and structure of the Kingscourt Outlier, Ireland. Geological Journal, 30, 1-23. SWEET, W. C. 1988. The Conodonta." Morphology, -

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Taxonomy, Palaeoeeology and Evolutionary History of a Long Extinct Animal Phylum. Oxford Monographs on Geology and Geophysics No. 10, Clarendon Press, Oxford.

IRISH DINANTIAN BIOSTRATIGRAPHY TEN HAVE, M. R. 1982. Studies of Irish Lower Carboniferous Ostracods. PhD Thesis, University of Dublin. THORNBURY, B. M. 1985. Conodont Biostratigraphy of Dinantian Rocks from the Cloyne Syncline, Co. Cork. MSc Thesis, University of Dublin. VACHARD, D. 1994. R6vision du genre Luteotubulus Vachard, 1977 (Pseudo-algue Issinellidae), Grace au mat6riel de R. Conil. M~moires de l'lnstitut G~ologique de l'UniversitOde Louvain, 35, 213-219. VARKER, W. J. & SEVASTOPULO, G. D. 1985. The Carboniferous System: Part 1 - Conodonts of the Dinantian Subsystem from Great Britain and Ireland. In: HIGGINS, A. C. & AUSTIN, R. L.

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(eds) A Stratigraphical Index of Conodonts. Ellis Horwood, Chichester, 167-209. VDOVENKO, M. V., AISENVERG, D. Y., NEMIROVSKAYA, T. I. & POLETAEV, V. I. 1990. An overview of Lower Carboniferous biozones of the Russian Platform. Journal of ForaminiferaI Research, 20, 184-194. WEBSTER, G. D. & GROESSENS, E. 1990. Conodont subdivision of the Lower Carboniferous. Courier Forschungsinstitut Senckenberg, 130, 31-40. WHITAKER, M. F. 1976. The Palynology of the Carboniferous Sediments in Ireland, with Special Reference to the Ballycastle and Leitrim Areas. PhD Thesis, University of Aston.

Fish assemblages in the Tournaisian-Vis~an environments of the East European Platform O. A. L E B E D E V

Palaeontological Institute of the Russian Academy of Sciences, 123 Profsoyuznaya Street, 117647, Moscow, Russia

Abstract: New fish material, including microremains from the Tournaisian-Vis6an interval of the Moscow Syneclise, Northern Urals and Voronezh Anteclise of the East European Platform, comprises new species of Chondrichthyes: Lissodus pectinatus sp. nov., (?)Diplodoselache antiqua sp. nov., Eunemacanthus krapivnensis sp. nov. Other fish taxa include Stethacanthus obtusus (Trautschold 1874), "Orodus'tumidus (Trautschold 1874) and Taeniolepis trautscholdi (Chabakov 1927) are revised. Tournaisian-Vis6an deposits on the East European Platform are mostly represented by shallow marine or near-shore hyposaline facies with alternating salinity levels. Four major types of aquatic vertebrate environments are suggested: (1) continental fresh- and brackishwater; (2) lagoonal and estuarine with unstable salinity levels; (3) near-shore intertidal; and (4) off-shore neritic. Osteolepidids and actinopterygians are found in two or even three environmental zones, which suggests that they were eurybiontic and tolerant to changes of salinity, temperatures, hydrodynamics and other parameters of the basin. The most stenobiont forms are: Diplodoselache, Ageleodus, Pycnoctenion and dipnoans (1); Deltodus, Streblodus, Copodus and petalodontids (2); Lissodus and actinistians (4). The environmental distribution of the other groups is not as well expressed.

Current research on Early Carboniferous fish material, including microremains, incorporates Tournaisian-Vis6an material from the southern (Tula, Kaluga and Ryasan regions) and northwestern (Novgorod region) parts of the Moscow Syneclise, and the Northern Urals (Komi Republic and Arkhangelsk region), as well as from Belgorod and Volgograd regions (southern and eastern slopes of the Voronezh Anteclise) (Fig. 1). These regions are unevenly investigated with respect to vertebrates. Tournaisian fishes from the central part of Russia were first described in the second half of the last century (Pander 1858; Eichwald 1860; Semenov & M611er 1864a, b; Trautschold 1874; Shchurovsky 1878; Rohon 1893). These studies were mostly dedicated to the fauna of the Malevka Horizon coalbearing strata of Tula and Ryasan, but most of the Upper Tournaisian and Vis6an fauna remained undescribed until now. The finds of the Early Carboniferous fishes from the northeastern part of the platform were only mentioned once (Kalashnikov 1962). This paper presents preliminary systematic descriptions of previously existing data, and follows the dependence of certain fish groups on facial conditions. Collections used in this paper are stored in the Palaeontological Institute of the Russian Academy of Sciences, Moscow, Russia (PIN).

Systematic description Phylum Vertebrata Class A c a n t h o d e i 'Acanthodes' indet.

Description.

Isolated scales with smooth unornamented rhomboid crown, slightly depressed in the posterior half and hardly elevated in the anterior, well-formed neck and moderately high base. Remarks. Scales of such appearance are widely spread in the Lower Carboniferous strata of the Russian Platform and apparently belong to various taxa; this name may only be regardedas a form taxon. The same type is almost globally distributed in the Carboniferous-Permian deposits. This material is currently under description by P.-Y. Gagnier (Parc de Miguasha, Nouvelle, Quebec).

Material, age and geographical distribution. About 40 isolated scales in various states of preservation from bioclastic limestones of the Malevka Horizon, Tula Region; four scales from black sandy limestones of Novy Oskol borehole 86, 196.16-196.33 m: Malevka or Upa Horizons, Lower Tournaisian, Belgorod Region; three isolated scales from biogenic limestones of the Mikhailov Horizon, Upper Vis6an, Polotnyany Zavod quarry, Kaluga Region, Central Russia.

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 387-415.

388

O. A. LEBEDEV

f/ ~

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:MOSCOW

VORONEZH BELGOROD

10'0 Kl~i Fig. 1. Locality map. (a) The north-central part of the Russian platform. Rectangles show the positions of maps (b) and (c). Dotted filling indicates Tournaisian subcrop, hatched filling Visban subcrop; white circles represent Tournaisian, black circles Vis6an localities. (b) Central part of the Russian platform. Tournaisian: 1, SnetkiPavlovskoye; 2, Zolotoy Verkh ravine; 3, Suvorov borehole 814; 4, Novy Oskol borehole 86; 5, Andreyevka-1; 6, Krasnoye (Mokraya Tabola River); 7, Prisady; 8, Chernyshino; 9, Znamenskoye. Vis6an: 1, Polotnyany Zavod; 2, Vitsa; 3, Priksha River; 4, Kamenka; 5, Gryzlovo quarry; 6, Azarovo; 7, Erino; 8, Gorenki; 9, Ust-Buzuluk borehole 41. (c) Timan- Pechora Province. Vis6an: 1, Ovin Kamen, Shchugor River; 2, Kozhim River; 3, KirpichKyrty, Podcheryom River; 4, Klina-Shor Creek, Kozhva River; 5, cement factory quarry, Vorkuta River; 6, Kamenka River; 7, Talata River; 8, Vangyr River. After Makhlina et al. (1993) and Chernov (1972).

DINANTIAN

FISHES OF RUSSIA

389

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Fig. 2. Lissodus pectinatus sp. nov., teeth. (a)-(c) PIN 2878/429;.(d)-(f) Holotype PIN 2878/432; (g)-(h) PIN 2878/431; (i)-(j) PIN 2878/430; (k)-(m) PIN 2878/433; (n) PIN 2878/434. (a), (f), (g), (j) and (1) lingual; (e), (d), (k) labial; (b), (e), (h), (i), (m), (n) occlusal view. Scale bar is 300 #m.

390

O. A. LEBEDEV

Class C h o n d r i c h t h y e s Subclass E l a s m o b r a n c h i i O r d e r Euselachii H a y , 1902 S u p e r f a m i l y H y b o d o n t o i d e a Zangerl, 1981 Lissodus pectinatus sp. nov. Fig. 2 a - n

Etymology. The species name originates from the comb-like appearance of the crown. Holotype. PIN 2878/432, tooth. Diagnosis, Small Lissodus teeth characterized by strongly developed accessory tubercles at the lingual edge of the crown, moderately developed median crest, lateral cusps and massive vertical labial ridges. Median cusp well developed. The labial side is deeply depressed forming a pocketlike hollow, so that the labial edge of the crown overhangs it. Description. The teeth length ranges from 2 to 4mm. The crown is elongated, with a well developed principal tubercle in most cases supported with a massive labial buttress; that is not supported with a labial peg, as in Lissodus wirksworthensis (Duffin 1985). The lingual edge of the crown bears a row of well developed tubercles, whose position does not necessarily correspond to that of the lateral cones or the ridges at the labial side (Fig. 2b, e,h,n). The crown edge strongly overhangs the root from all sides. The median crest is moderately developed. The most outstanding features of the labial side

of the crown are massive vertical ridges that tend to extend towards the cusps; they may be sharpedged or rounded (Fig. 2d, 1). The aboral side of the base is smooth and slightly concave. The lingual surface bears large, irregularly disposed vascular foramina (about ten) in the upper part, delimited with massive swollen crosspieces. The lingual surface is strongly concave, and there is a row of large foramina underneath the crown-base contact. Some specimens were very much worn in life on the crown surface, but the major characteristic features, such as lingual tubercles, median crest and the principal cusp supported by the labial buttress, still persist (PIN 2878/429, Fig. 2a-c). Material, age and geographical distribution. Eight isolated teeth from Novy Oskol borehole 86, 196.16-196.33 m: black sandy limestones of Malevka or Upa Horizons, Lower Tournaisian, Belgorod Region, Central Russia.

Euselachii indet. Fig. 3b Description. A single scale with a wide, flat and thin, slightly concave base. The crown consists of about six thorn-like, ridged and almost vertical irregular projections, fused to different extents with their bases. Remarks. The systematic position of this kind of scale, very common in the Carboniferous

A

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Fig. 3. (a) "Orodus" tumidus (Trautschold 1874). Scale, crown view, PIN 4359/65. (b) Euselachii indet., scale, crown view, PIN 4472/59; (c) (d) ? Desmiodus sp., tooth, PIN 2876/14, coronal view from the (c) lingual side and (d) from the labial side. Scale bar is 300 #m.

DINANTIAN FISHES OF RUSSIA deposits, cannot be clarified with certainty in the absence of articulated material, or at least large quantities of specimens allowing the study of histological construction and variability. This specimen superficially resembles the so-called 'Ctenacanthus' or hybodontoid-type scales, for example described by Turner (1993) from the Narrien range, central Queensland, Australia. Material, age and geographical distribution. An isolated scale, PIN 4472/59, from detrital bituminous limestones of the Tula Horizon, Upper Vis6an, Ovin Kamen, Shchugor River, Komi Republic, Northern Russia.

? Order D e s m i o d o n t i d a Zangerl, 1981 ? Desmiodus sp. Fig. 3c-d Description. The only specimen, PIN 2876/14, is part of a tooth whorl composed of two elements fused at their bases. The crown of both is triangular in outline, bordered by an elevated ridge that forms the crown margin. There is only one cone in both elements, flanked by a well pronounced median crest on both sides, the lateral cusps being abraded or missing. The crown is separated from the base by a distinctive groove. The base is deeper than the crown, its lateral edges projecting beyond the crown

391

margins (Fig. 3d). The lingual surface of the base is slightly convex. Remarks. The attribution of this specimen to Desmiodus is tentative. The specimen does not show any special resemblance to any of the species described by St. John & Worthen (1875), because the lateral cusps are missing from the crown surface. Nevertheless, the general construction of the whorl does not differ much from the specimen illustrated by Zangerl (1981, fig. 66D) from the Salem Formation, Missouri, USA. Material, age and geographical distribution. One isolated tooth whorl from argillaceous limestones of the Aleksin Horizon, Upper Vis6an, right bank of Msta River, downstream from Vitsa Village, Novgorod Region, centralwest Russia. Order S y m m o r i i d a Zangerl, 1981 F a m i l y Stethacanthidae Lund, 1974 Stethacanthus obtusus (Trautschold, 1874) Fig. 4 a - e 1864, Cladodus simplex: Semenov & M611er, p. 235, pl. I, fig. 12. 1874, Cladodus simplex: Trautschold, p. 265, pl. XXVI, fig. 1. 1889, Cladodus simplex: Woodward, p. 25.

Fig. 4. Stethacanthus obtusus (Trautschold 1874), teeth. (a) PIN 4359/74, occlusal. (b) PIN 4359/77, lingual. (c) PIN 4359/78, lingual. (d) PIN 2921/3138, lingual. (e) PIN 2921/3139, labiobasal. Scale bar is l mm.

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O. A. LEBEDEV

1874, Cladodus obtusus: Trautschold, p. 266, pl. XXVI, fig. 2. 1889, Cladodus obtusus: Woodward, p. 24. 1874, Cladodus primigenius: Trautschold, p. 266, pl. XXVI, fig. 3. 1889, Cladodus primigenius: Woodward, p. 25. 1982, Cladodus thomasi: Turner, p. 126, Figs 6c &8j. 1984, 'Cladodus' (? = Stethacanthus) thomasi: Long & Turner, p. 237. 1990, Stethacanthus thomasi: Turner, p. 70.

Holotype. Unnumbered, isolated tooth. Diagnosis. The base is an elongated lozenge in the plan view, its lateral terminations being shifted somewhat lingually so that the lateral cusp tips project beyond the base. The contact swelling at the crown-base edge of the labial surface extends laterally up to the median pair of side cusplets base. The lateral cutting edges are well developed at the principal cone. Up to four pairs of lateral cusps may be present, but two pairs are common. Description. The button is pierced lingually with 2-6 openings of large nutrient foramina that are connected by short canals with the openings on the labial side of the button. The contact swelling at the crown-base edge of the labial surface is wide, extending laterally opposite the base of the lateral side of the median pair of side cusplets. The labial half of the aboral side of the base is deeply depressed and bears two or more large vascular foramina. The principal and lateral cusps are slightly bent lingually and ornamented with coarse longitudinal ridges, sometimes bifurcating towards the base and becoming intercalated with short crests. There may be up to four pairs of side cusps (PIN 4359/77; Fig. 4b). There are several specimens of teeth with a principal and a lateral cone (PIN 4359/78; Fig. 4c), these are sometimes fused with their bases. In this case the button is irregularly shaped.

Remarks. Trautschold (1874) based his descriptions on isolated teeth collected from the bank of a creek close to Malyovka village, Tula region. At that time these deposits were regarded as Devonian in age. This author did not indicate the storage place, but figured specimens for the publications in the Transactions of the Moscow Naturalists' Society were usually deposited in the Museum of the Society itself. At the turn of the century, these collections were transferred to the Moscow University Museum (later Moscow Geological Prospecting Institute, now Vernadsky State Geological Museum). Unfortunately the location of these specimens is not known, but there is hope that they may be found soon due to the Museum reconstruction and registration of collections, which is why the neotype is not indicated at this stage. Turner (1982) described teeth from the Broken River embayment, Northern Queensland, Australia (Lower Carboniferous, (?) Tournaisian, original description of similar material by Thomas (1957, 1959) from Western Australia) which look practically the same as the specimens described here from the locotypic area, and thus C. thomasi is considered here as a junior synonym to S. obtusus (Trautschold 1874). Material, age and geographical distribution. About 40 isolated teeth in various states of preservation from numerous localities of bioclastic and argillaceous limestones of the Malevka-Upa Horizons, Lower Tournaisian of Tula Region, Central Russia. Stethacanthus altonensis (St J o h n & W o r t h e n , 1875) Fig. 5a c Description. The tooth base is oval or lozengeshaped in plan view. The button is pierced at the base lingually by a single or a pair of large

Fig. 5. Stethacanthus altonensis (St John & Worthen 1875), teeth. (a) Median (?) symmertic element, PIN 4359/76, labiolateral. (b) Tooth, PIN 4359/73, linguolateral. (c) Lingual view of the same tooth. Scale bar is 1mm.

DINANTIAN FISHES OF RUSSIA vascular canals. The contact swelling at the crown-base edge of the labial surface is short, extending laterally opposite the base of the principal cusp. In the plan view the tips of the lateral cusps do not project beyond the base. The cusps are rounded in cross-section and the lateral cut~ng edges are only slightly developed in the largest specimens. The surface is evenly striated with fine subparallel ridges sometimes intercalated with ridges extending from the base up to approximately the middle of the cusp. This

393

is the only character that disagrees with Williams' (1985) description of teeth in this species. This author stated, that intercalation of accessory ribs are characteristic to Symmorium, but material represented here shows, apparently, wider variability range. Specimen PIN 4359/76 (Fig. 5a) shows a shortened base and single principal cusp, whilst the base is subquadrate and massive. Teeth of this type might be situated in the jaw between the families of regular multicuspid teeth.

Fig. 6. 'Stemmatiid' denticles and 'Ctenacanthus'-type scales. (a)-(b) 'Periplectrodus warreni' type, (a) PIN 4359/ 89, (b) PIN 4359/87, (e) 'Stemmatias bicristatus' type, PIN 4359/90. (d) (h) scales, (d)(e) from the trunk region, (f) (h) from the head: (d) PIN 4359/86, (e) PIN 4359/85, (f) PIN 4359/92, (g) PIN 4359/93, (h) PIN 4359/88, (i) 'Listracanthus'-type scale, PIN 4359/91. Scale bar is 100 itm.

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O. A. LEBEDEV

Material, age and geographical distribution. About 150 isolated teeth of various states of preservation from bioclastic limestones of the Malevka Horizon, Lower Tournaisian of Tula region, Central Russia. ' S t e m m a t i i d ' denticles and ' Ctenacanthus'-type scales There are about 50 examples in the collection representing 'Stemmatias'-type and 'Periplectrodus'-type elements of the buccopharyngeal or branchial cavity, described, for example, by St John & Worthen (1875), Zangerl (1981) and Zidek (1993), which cannot be even tentatively assigned to either of the Stethacanthus species described above. The spiral-shaped elements of the 'Periplectrodus warreni' type may be symmetric or asymmetric. Both types may bear a single row of cusps or possess accessory lateral cusps (PIN 4359/87, PIN 4359/89; Fig. 6a-b). Numerous specimens of 'Stemmatias simplex' and 'Stemmatias bicristatus' type are present in the collection (PIN 4359/90; Fig. 6c). There are many transitional types, such as pairs of partly fused 'simplex' elements, as well as 'bicristatus' with intermediate denticles turning to 'Stemmatias cheiriformes'. 'Ctenacanthus'-type scales (as designated by Karatajute-Talimaa, 1992) are known from the Malevka Horizon, including the trunk and head types (PIN 4359/85-86, 88, 92-93; Fig. 6d-h).

sigmoidally curved apex, 2-3 pairs of lateral cusps and a pair of articulation buttons on both oral and aboral sides of the tooth base. The largest nutrient foramina are situated at the lingual edge of the base, between articulation knobs and also in a hollow lingually from the labial edge of the aboral surface of the base (PIN 2878/424). Remarks. These minute teeth do not differ morphologically in any respect from the specimens described under the names 'Cladodus' occidentalis (for example, Leidy 1873; Zidek 1973), 'Ctenacanthus' occidentalis (Glikman 1964) or Symmorium reniforme (Cope 1893; Williams 1985; Zidek 1993). Thus, as the generic name 'Cladodus' was demonstrated to be nornen dubium (Chorn & Whetstone 1978; Zangerl 1981) it should be replaced with a valid name Symmorium with a retention of the original species name occidentalis, as reniforme is a junior synonym. This suggestion is indirectly supported by Zidek's practice, who originally named a tooth from the Francis Formation, Missourian, Oklahoma (OMNH 00259) Cladodus occidentalis (Zidek 1973) and then replaced the original definition with Symmorium reniforme (Zidek 1993). Material, age and geographical distribution. Two isolated teeth, one crushed, from biogenic limestones of the Mikhailov Horizon, Upper Vis6an, Polotnyany Zavod quarry, Kaluga Region, Central Russia.

F a m i l y S y m m o r i i d a e Dean, 1909 Symmorium occidentalis (Leidy, 1873) Fig. 7 a - b

Order O r o d o n t i d a Zangerl, 1981 F a m i l y O r o d o n t i d a e De Koninck, 1878 'Orodus' tumidus (Trautschold, 1874) Fig. 8 a - g

Description. Small 'cladodont'-type teeth with a trapezoid base, large coarsely striated principal cusp with well-defined cutting edges and slightly

1874, Orodus tumidus: Trautschold, p. 267, pl. XXVI, fig. 5a-c. 1889, Orodus tumidus: Woodward, p. 238.

A

B

...... ,

Fig. 7. Symmorium occidentalis (Leidy 1873), tooth, PIN 2878/424. (a) Occlusal view. (b) Linguolateral view. Scale bar is 300 #m.

DINANTIAN FISHES OF RUSSIA

395

Fig. 8. "Orodus' tumidus (Trautschold 1874), teeth and scales. (a)-(f) teeth. (a) PIN 4359/83, lingual; (b)-(c) PIN 4359/2, (b) labial, (c) lingual. (d) PIN 4359/82, lingual. (e) Two elements fused with their bases, occlusal view, PIN 2878/426. (f) Occlusal view, PIN 2878/427. (g) Median (?) symmetric tooth PIN 4359/81, linguolateral. (h) (i) Scales, (h) PIN 4359/84, (i) PIN 2878/425: note the disposition of odontods in the crown. Scale bar is 300 #m, if not indicated otherwise. 1874, Orodus excentricus: Trautschold, p. 268, pl. XXVI, fig. 6a-d. 1889, Orodus excentricus: Woodward, p. 237. 1874, Orodus sublaevis: Trautschold, p. 268, pl. XXVI, fig. 7a-b. 1889, Orodus sublaevis: Woodward, p. 238. 1874, Helodus aversus: Trautschold, p. 268, pl. XXVI, fig. 8a-e. 1889, Helodus aversus: Woodward, p. 226. 1874, Helodus gibberulus: Trautschold, p. 269, pl. XXVI, fig. 9a-c.

1874, Helodus contractus: Trautschold, p. 270, pl. XXVI, fig. 10a-h; 1889, Helodus contractus: Woodward, p. 227. 1874, Helodus angustus: Trautschold, p. 276 (err. typ.). 1874, Psammodus porosus: Trautschold, p. 270, pl. XXVII, fig. 11 a-c. 1864a, Psammodus inflexus: Semenov & M611er, p. 235, pl. I, fig. 9. 1874, Psammodus inflexus: Trautschold, p. 271, pl. XXVII, fig. 12a-e.

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O. A. LEBEDEV

1889, Psammodus inflexus: Woodward, p. 107. 1874, Psammodus linearis: Trautschold, p. 271, pl. XXVII, fig. 13a-f.

Description. Teeth elongated, symmetric or often asymmetric. Coronal depth is approximately equal to that of the base. The crown is separated from the base by a well-marked groove. Basal side of the base smooth and slightly concave transversely, the labial side is perforated by large vascular foramina; similar foramina over the lingual side are separated by prominent ridges. The crown may be designed in 'Orodus' or 'Helodus' type; various transitional constructions may be observed. The principal cone is situated in the centre or displaced towards the lateral end; there may be up to two pairs of variously developed flanking cones laterally. The sagittal crest, if present, may be continuous or interrupted in the middle if the principal cone is well developed. The crest is often but not necessarily accompanied by vertical ridges branching occasionally towards the base of the crown; intercalation ridges may be rarely observed (Fig. 8d). There are some examples (PIN 4359/81, Fig. 8g) of short,

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symmetric and high teeth with a triangleshaped crown in the plan view. Remarks. Trautschold (1874) described several species of Orodus, Helodus and Psammodus from the same strata of the Malevka Horizon. Variation of sculpturing have been chosen as diagnostic features. The material described here from the same horizon consists of hundreds of specimens, making it possible to trace transitions within variation series and assign all teeth to a single species 'Orodus' tumidus (Trautschold 1874). The generic name is in inverted commas, since the generic attribution needs revision. Material, age and geographical distribution. About 300 teeth and more than 100 scales from Snetki-Pavlovskoye locality, Odoev District, 18 teeth and two scales from Zolotoy Verkh Ravine, Shchekino District; about 100 teeth and more than 30 scales from the Suvorov borehole, 68.0 m depth; all from bioclastic limestones of the Malevka Horizon, Lower Tournaisian of Tula Region. Also, seven teeth in various states of preservation from Novy Oskol borehole 86, 196.16-196.33m: black sandy limestones of Malevka or Upa Horizons, Lower Tournaisian; Belgorod Region, Central Russia.

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Fig. 9. ? Diplodoselache antiqua sp. nov., teeth, branchial denticles (?) and fin spine. (a)-(d) Teeth. (a) Holotype, PIN 2921/3161. (b) PIN 2921/3162. (c) PIN 2921/3160. (d) PIN 2921/3159. (a) (c) Lingual, (d) basal. (e)-(f) Branchial (?) denticles. (e) PIN 2921/3164. (f) PIN 2921/3163. (g)-(h) Median fin spine, PIN 2921/3137, (g) lateral, (h) posterior. Scale bar is 300 mm, if not indicated otherwise.

DINANTIAN FISHES OF RUSSIA Scales, t e n t a t i v e l y assigned to

'Orodus' tumidus Figs 3a, 8 h - i The isolated scales of 'Orodus'-type are tentatively assigned to the same species as teeth. The situation will be completely resolved only when an articulated skeleton is found. Scales are high, with a cushion-shaped base and fiat crown. The horizontal cross-section

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100p~

397

varies from quadrate to rectangular and diamond-shaped depending on the horizontal disposition of the composing odontods (PIN 4359/84; PIN 2878/425); subparallel orientation results in the rectangle shape and curved around the primordial odontod in the guttiform outline (Fig. 3a). Scales assigned to 'Orodus' tumidus are similar to those of Protacrodus wellsi and Holmsella sp. (Orvig 1966; Gross, 1973).

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Fig. 10. ?Diplodoselacheantiqua sp. nov., scales. (a)-(d) Scales with a funnel-shaped base, (a) PIN 2921/3148, (b) PIN 2921/3157, (c) PIN 2921/3151, (d) PIN 2921/3153. (a)-(c) Crown view, (d) basal view. (e)-(f) PIN 2921/3144, (f) the magnified area of the crown. (g)-(i) Scales with a massive base and fused odontods of the crown, all crown view (g) PIN 2921/3146. (h) PIN 2921/3149. (i) PIN 2921/3154. Scale bar is 300 #m, if not indicated otherwise.

398

O. A. LEBEDEV

Order X e n a c a n t h i f o r m e s Berg, 1955 F a m i l y Diplodoselachiidae Dick, 1981 ? Diplodoselache antiqua sp. nov. Figs 9 a - h , 10a-i

Etymology. The species name originates from its earliest chronological position. Holotype. PIN 2921/3161, tooth. Diagnosis. Crown cusps are strongly inclined labially. Median cusp is well developed, possesses sharp cutting edges; its length is about half of that of the lateral ones. Cusps with smooth cutting edges, separate vertical crests at the lingual side of the cusps are not uncommon. The base oval-shaped, strongly extended lingually, very low. The button of the oral surface is strongly developed and may occupy about half of its area. The basal tubercle is in most cases very strongly developed as well. Description. The teeth of ?D. antiqua differ significantly from that of the nominal species D. woodi (Dick 1981). In the former the cutting edges are smooth and not in the least carinate. Some specimens bear 1-2 well defined separate sharp-edged ridges at the lingual surface, which do not reach the base of the cusp. The crown cusps are inclined labially at an angle of about 120 °. The cusps are rounded in cross-section at the base, slightly compressed apically. The apical button at the oral surface of the base may be very large, its position varies, but usually it is situated at the base of the cusps. Its lingual edge is in most cases pierced by 3-5 large nutrient foramina. The aboral surface of the base is concave in the middle, the basal button occupies its labial part. Lingually the concavity is bordered with a horseshoe-shaped flat ridge. Scales are extremely variable in the shape and construction of both the crown and base. The base, as in Diplodoselache woodi, may be funnelshaped or cushion-shaped. In the former variety the crown consists of numerous odontods fused together with their walls (Fig. 10a-d). The direction of growth of the odontods varies from subparallel to fan-shaped. Separate odontods are superficially ornamented with rough longitudinally-directed ridges and furrows; their tips sometimes become loose at the end and look like projecting spikes (Fig. 10e-f). In the type transitional to that with a cushion-shaped base, the crown odontods become much shorter and wider at the base (Fig. 10i), sometimes irregularly fusing (Fig. 10g-h) and significantly decreasing in number. Some elements (Fig. 9e-f) may be tentatively regarded as branchial denticles because of their

poorly developed flat basal plate. The crown in these elements is composed of finger-like smooth odontods rounded in cross-section. The general crown shape varies from longitudinally to transversely elongated. The only known complete dorsal spine (PIN 2921/3137, Fig. 9g-h) is of the xenacanth type, about 3cm long, straight and slender, slightly compressed laterally in the apical part and oval in cross-section. The superficial ornamentation is composed of about ten longitudinal, compactly set smooth ridges and grooves. The surface lacks a shiny coating. The posterior surface is strongly convex and bordered with paired proximally hooked cusps, which extend for about a third of the total spine length. The opening of the internal cavity occupies less than half of that. Remarks. Morphologically and histologically the scales are very similar to those described in Diplodoselache woodi. Their variety matches that observed in this species, as described by Dick (1981). At the same time, the teeth are strikingly different, in possession of a large median cuspule, strong labial inclination of the crown labially in respect to the base, absence of carination at the lateral cutting edges, rounded cross-section of the cusps, low base and a prominent apical button. These features, except for an obtuse angle between the plane of the crown and the base, are characteristic of Xenacanthus, but in this genus the body scales are not known (Zangerl 1981). It might be possible that the taxon described here belongs to a new genus of Diplodoselachiidae, but this needs confirmation by a complete articulated skeleton. The spines in Diplodoselache woodi are poorly known because of unprepared material; nevertheless, it agrees with the new material in the presence of rows of paired denticles at the posterior side and lack of orthodentine cover. On the other hand, in ? D. antiqua sp. nov. the striations of the external ornamentation are regular, the spine is laterally compressed in the apical area and the basal opening for the insertion of a basal cartilage is present. Very similar teeth and scales were described by Turner (1993) from the Early Carboniferous of the Central Queensland; the author named the former Xenacanthoid cf. 'Diploselache' (print error) woodi and the latter 'Hybodontoid gen. et sp. indet'. It is likely that the Australian material is conspecific, or at least congeneric with the Russian.

Material, age and geographical distribution. About 100 isolated teeth in various states of

DINANTIAN FISHES OF RUSSIA

399

because of the generally poor knowledge of the material, especially Tournaisian. The major difficulty with the specimens under description are the large dimensions of the central cusp, which is equal or even larger than the lateral ones, making its attribution either to Phoebodontidae or Xenacanthidae tentative. In all other respects the construction of the base and the crown is similar to the Upper Devonian form Phoebodus australiensis Long 1990, Penn?nebraskensis, and sylvanian Xenacanthus ' Bransonella' (Dit todus, ?Thrinacodus) triden tata Harlton (for discussion of their systematic position see Zidek 1973; Johnson 1984), differing by such features of uncertain rank as welldefined porosity of the apical button. Until more material becomes known, these teeth from the Mikhailov Horizon of the Kaluga Region are temporarily assigned to this species. Material, age and geographical distribution. Six teeth in various states of preservation from biogenic limestones of the Mikhailov Horizon, Upper Vis6an; Polotnyany Zavod quarry, Kaluga Region, Central Russia.

preservation from argillaceous limestones of Malevka or Upa Horizons (Richterina latior Pseudoleperditia venulosa - Shivaella microphtalma ostracode Zone), Lower Tournaisian: left bank of the Tresna River opposite to Andreyevka Village (Andreyevka-1 locality), Suvorov District, Tula Region, Central Russia.

O r d e r undefined F a m i l y X e n a c a n t h i d a e Fritsch, 1889 X e n a c a n t h u s ?nebraskensis J o h n s o n , 1984 Fig. l l a - b Description. The base is thick, and rounded in plan view. There is a prominent, extremely porous apical button, bordering the base of the central cusp; a large foramen is disposed in its lower part lingually. The aboral surface in PIN 2878/421 demonstrates a pit in the central part bordered with a large transversely disposed articulation button labially (Fig. l la). The median cusp is as high, or even slightly higher, than the lateral ones, which diverge laterally. Their labial surfaces bear characteristic qanceolate' ornamentation, found in Phoebodus australiensis Long morphotype 1 (Ginter 1990) and Xenacanthus ?nebraskensis as described by Johnson (1984) from the Late Pennsylvanian of Nebraska. The cusps are rounded in crosssection and side cutting edges are absent. The lingual side of the cusps is smooth. Remarks. The exact position of the specimens described is hard to establish with certainty

? F a m i l y P h o e b o d o n t i d a e Williams, 1985 ? Thrinacodus sp.

Fig. 1 lc Description. The tooth possesses a three-cusped crown, the central cusp being slightly higher, than those of the lateral pair. The cusps are straight and parallel to each other, their tips are

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Fig. 11. (a)-(b) Xenacanthus ?nebraskensis Johnson, 1984: teeth. (a) PIN 2878/421, basal. (b) PIN 2878/422, linguotateral. (c) ?Thrinacodus sp.; PIN 2878/428, lingually. (d) ?Denaea sp., tooth, PIN 2878/423, lingual. Scale bar is 100#m.

400

O. A. LEBEDEV

slightly inclined lingually. In the basal part the cross-section is rounded, the tips are slightly compressed. Cutting edges are present; the lingual side bears 2-3 strong vertical parallel ridges. The labial side of the cusps is mostly smooth, in the basal part the ridges are barely visible. The base is roughly pentagonal in plan view, forming a well defined projection lingually. There is a large single foramen of the nutrient canal in the central position. There are no well defined articulation knobs; instead the oral side of the base is swollen, at the basal surface the articulation pit is shifted in the labial direction, and the basal opening is situated at its lingual margin. Remarks. This specimen is retained under this name provisionally until more material becomes known and the study of variability range becomes possible. In some of the features (absence of apical and basal buttons, dominance of the central cusp above lateral ones, well defined ribs at the lingual surface of the cusp and lingual projection of the base) this tooth strongly resembles those of Thrinacodus, at the same time differing from it by comparatively short base and lateral cusps parallel to the central one. This difference may be due to topographic variability in the tooth families, or indicate its separate taxonomic status. Material, age and geographical distribution. A single tooth PIN 2878/428 from Novy Oskol borehole 86, 196.16-196.33 m; black sandy limestones of Malevka or Upa Horizons, Lower Tournaisian, Belgorod Region.

? Denaea sp. Fig. l l d Description. The base is low, rounded pentagonal in plan view, the obtuse angle apex facing lingually. The oral side of the lingual projection bears a symmetric pair of large vascular openings; there are one or two vascular foramina at the aboral side. Two more large foramina are situated at the bases of the intermediate cusps lingually. A labio-lingual ridge at the aboral surface of the base separates, as in Denaea meccaensis two depressed areas (Williams 1985). The crown consists of a larger median cusp and two pairs of lateral ones, the median pair the smallest. The lateral cusps are approximately one third shorter than the principal cusp. The surface of the cusps is smooth or bears one or two longitudinal ridges near the base; lateral cutting edges are present. All cusps are almost parallel to each other.

Remarks. This specimen differs from Denaea meccaensis by the presence of paired vascular foramina at the lingual side of the base. At the same time, Williams (1985) figured a tooth clearly showing two symmetric openings at the aboral side, suggesting inconsistency of this feature within the genus. In all other respects (construction of the crown, shape of the base, absence of articulation knobs and lateral depressions at the aboral side of the base) this tooth is similar to those present in Denaea. Material, age and geographical distribution. Three isolated teeth, all damaged to some extent, from biogenic limestones of the Mikhailov Horizon, Upper Vis6an; Polotnyany Zavod quarry, Kaluga Region, Central Russia. Undefined o r d e r and family Eunemacanthus krapivnensis sp. nov. Fig. 12a-e

Etymology. The species is named after Krapivna town in the Tula Region, Central Russia. Holotype. PIN 4359/1, the apical portion of the spine. Diagnosis. Slender, small fin spines, slightly curving posteriorly with a triangular crosssection, being twice as long as wide. The thin anterior rib is practically as wide as the other ribs of the lateral sides. Side ribs are widely spaced, narrow and deep; the denticles of the neighbouring ribs do not meet. From 3-5 side ribs on both sides. Description. The sides are slightly convex, the posterior margin almost flat or concave and ornamented with thin ridges and furrows. Primary bifurcation occurs anteriorly. Lateral ribs are ornamented with comb-like denticles, as in Eunemacanthus costatus (St John & Worthen 1883; Maisey 1982). There are up to four marginal denticles disposed within 1 mm length of the posterior edge. Material, age and geographical distribution. Two fragments from Zolotoy Verkh Ravine, Shchekino District; six fragments from SnetkiPavlovskoye locality, Odoev District, one fragment from Suvorov borehole, 68.0m, from organoclastic limestones of Malevka Horizon, Lower Tournaisian; several fragments of the finspines from the argillaceous limestones at the left bank of the Tresna River opposite to Andreyevka Village (Andreyevka-1 locality), Suvorov District, Malevka-Upa Horizon (Richterina latior-Pseudoleperditia venulosa-Shivaella microphtalma ostracode Zone), all from the Tula Region, Central Russia.

DINANTIAN FISHES OF RUSSIA

-

401

300Pro

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300.~ Fig. 12. Eunemacanthuskrapivnensis sp. nov., fin spines. (a)-(d) Holotype, PIN 4359/1, (a) lateral, (b) slightly turned to show the posterior edge, (e)-(d) magnified, to show: (e) the character of the side ribs ornamentation; (d) the posterior edge and marginal denticles. (e) A fragment of an apical part, PIN 2921/3143. Scale bar is 1 mm, if not indicated otherwise.

Remarks. No Eunemacanthus species are known from associated remains, thus the attribution of these elements to any of the known natural groups is possible. Maisey (1982) suggests that this genus is closely allied to neoselachians. 'Ageleodus'-type denticles Fig. 1 3 a - c

Description.

The denticles are about 1-2mm high and wide. The base is narrower and consists of porous bone which is often broken off because of its loose structure. The neck is only slightly expressed. The upper margin of the base forms a well marked transverse ridge lingually, which separates the former from the crown. A rather deep furrow corresponds to it at the labial surface. The cusps are slightly turned lingually and flattened in the labio-lingual direction. The

individual cusps bear sharp cutting edges at their sides, joined to the neighbouring ones at the base. The quantity of cusps in the crown ranges from 5 to 11, the central ones being the largest (PIN 2921/3166, 2921/3167 and 2921/3168), although sometimes the central cusp and a pair of those closer to the margins are much smaller than a pair of the. biggest, which are neighbouring the central cusp. The cusp row may be straight or arched. Remarks. It is suggested that these elements represent specialized branchial denticles of ?Diplodoselache antiqua. This idea will be supported or refuted if Ageleodus teeth are found associated with more informative skeletal parts. It should be noted that in both localities (Narrien Range, Queensland and Andreyevka-1, Russia) the associations of Ageleodus elements, Diplodoselache teeth and typical diverse scales of Diplodoselache type are the same.

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O. A. LEBEDEV

Fig. 13. 'Ageleodus'-type denticles. (a) PIN 2921/3166. (b) PIN 2921/3168. (c) PIN 2921/3167. Scale bar is 300 #m.

Whether the described specimens belong to an enigmatic genus Ageleodus Owen 1867 is currently a matter of debate (Turner 1993). The major difference between the material presented here and classical teeth Ageleodus pectinatus Agassiz 1838 are the minute dimensions of the former and the absence of massive dentine tissue. Loose structure of the base is also uncharacteristic for 'true' Ageleodus teeth.

Material, age and geographical distribution. About 100 isolated denticles from argillaceous limestones of Malevka or Upa Horizons (Rich-

The smaller specimen is a quadrangular 'mandibular' toothplate (PIN 1488/32; Fig. 14d-e), also arched and strongly twisted. It bears crenulations along its labial and lingual margins.

Material, age and geographical distribution. Two isolated toothplates: PIN 1488/12 from Krasnoye village, Mokraya Tabola River, Epifan District and PIN 1488/32 from Prisady village, Upa River; both from argillaceous limestones of Upa Horizon, Lower Tournaisian; Tula Region, Central Russia.

terina latior-Pseudoleperditia venulosa-Shivaella microphtalma ostracode Zone), Lower Tournaisian; left bank of the Tresna River opposite to Andreyevka Village (Andreyevka-1 locality), Suvorov District, Tula Region, Central Russia. Subclass Holocephali Order Cochliodontiformes Obruchev, 1953 F a m i l y Psephodontidae Zangerl, 1981 Psephodus dentatus ( R o m a n o w s k y , 1864) Fig. 14a-e

Description. Two specimens are attributed to this species. The larger is a posterior 'upper' toothplate (PIN 1488/12; Fig. 14a-c), strongly arched and pentagonal in the plan view. Anteriorly and posteriorly the upper 'tubular dentine' layer slightly overhangs the basal layer, but no margins in this specimen bear the short vertical crests or tubercles that are characteristic of Psephodus magnus (Davis 1883). The occlusal surface bears a well-marked wear facet anteriorly and well marked incremental line, marking a gap in growth.

?Psephodus sp. and ?Helodus sp. Fig. 14f-1 Several isolated and, as a rule, poorly preserved specimens resemble Psephodus in appearance, but cannot be identified with certainty. 'Helodus'-type toothplates may also belong to Psephodus anterior dentition (Traquair 1888) and thus these elements are regarded together.

Material, age and geographical distribution. 'Helodus'-type toothplate PIN 4472/6 from the detrital bituminous limestones of the Tula Horizon, Upper Vis6an, right bank of Kozhim River; an abraded tooth plate PIN 4472/29 from bituminous argillaceous limestones of the Mikhailov Horizon, Upper Vis6an, KirpichKyrty, Podcheryom River; complete toothplate PIN 4472/37, Vis6an, Klina-Shor Creek, Kozhva River; very large damaged toothplate of 'Helodus' type PIN 4472/7, Vis6an, right bank of Vorkuta River, cement factory quarry; an abraded tooth plate PIN 4472/38, Lower Carboniferous, right bank of Kamenka River;

DINANTIAN FISHES OF RUSSIA

403

Fig. 14. Psephodus, 'Helodus' and Copodus toothplates. (a)-(e) Psephodus dentatus (Romanowsky 1864). (a)-(c) PIN 1488/12. (d)-(e) PIN 1488/32. (a), (d) occlusal view, (b)-(e) basal, (e) lingual. (f) (h) ?Psephodussp., all in occlusal view. (f) PIN 1488/21. (g) PIN 1488/15. (h) PIN 1488/36. (i) 'Helodus aversus' Trautschold 1874; PIN 1488/11, occlusal view. (j)-(l) 'Helodus'-type toothplates: (j) PIN 4472/6, (k) (!) 1488/29; (j)-(k) occlusal, (!) basal view. (m)-(n) Copodus auriculatus (Davis 1883), imperfect toothplate PIN 4472/24, clearly showing a concave contact margin at the left posteriorly; (m) occlusal view, (n) basal. Scale bar is 10ram.

all from the Komi Republic; an almost complete toothplate PIN 4472/12 from the upper part of the Vis+an, Talata River estuary, Sinkin Nos Peninsula, Arkhangelsk Region, Northern Russia; a small 'Helodus aversus'-type toothplate PIN 1488/11 from Krasnoye village, Mokraya Tabola River, Epifan District, argillaceous limestones of Upa Horizon, Lower Tournaisian; 'Helodus'-type element PIN 1488/29 from Gryzlovo quarry, Venev District, dark limestones of the lower unit of the Tula Horizon, Upper Vis6an, Tula Region, Central Russia; incomplete tooth plates PIN 1488/13-14 from Chernyshino quarry and PIN 1488/15, PIN 1488/21-23 from a quarry near Znamenskoye village, PIN 1488/16 from Cherepet River, Suvorov District, Tula Region; all from detrital limestones of Cherepet Horizon, Lower Tour-

naisian; PIN 1488/36, Gorenky quarry close to Mikhailov Town, Ryasan Region, argillaceous limestones of the Aleksin Horizon, Upper Vis6an, Central Russia. F a m i l y Cochliodontidae Owen, 1867 Cochliodus contortus Agassiz, 1838 Fig. 15a-b Description. A large well preserved strongly arched lower jaw toothplate lacking convoluted anterior part. The crown surface consists of three well marked ridges of unequal width. The posterior margin demonstrates a well preserved unfinished growth margin (Fig. 15b); there is a lifetime wear facet in the central part (see Fig. 15a-b).

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O. A. LEBEDEV

Fig. 15. Cochliodontidae and Psammodontidae toothplates. (a)-(b) Cochliodus contortus Agassiz 1838, (a) PIN 4472/5, (b) PIN 1488/30. (c)-(d) Deltodus sp., (c) PIN 2878/308. (e) Streblodus cf. oblongus (Portlock 1843), PIN 2876/8. (f) Poecilodusjonesii McCoy 1855, PIN 1988/7. (g) Psammodontidae indet., PIN 4472/21. (a), (c)-(g) in occlusal view, (b) posterior. Scale bar is 10 mm.

Material, age and geographical distribution. An isolated incomplete toothplate PIN 4472/5 from the Upper Vis6an (?); the left bank of Vangyr River, Komi Republic, Northern Russia.

Deltodus sp. Figs 15c-d

The materials in the collection do not permit specific identification, as they are represented by incomplete isolated teeth. All specimens demonstrate characteristic features of the genus: arched plates, subtriangular in the plan view with elevated rounded median area followed laterally by a shallow wide furrow and then elevation of the posterolateral corner; the medial margin is smooth and rather deep. Material, age and geographical distribution. An isolated fragmented tooth plate PIN 4472/36

from detrital bituminous limestones of the Tula Horizon, Upper Visdan, Ovin Kamen, Shchugor River; abraded tooth plate PIN 4472/34 from bituminous argillaceous limestones of the Mikhailov Horizon, Upper Vis6an, Podcheryom River, Komi Republic, Northern Russia; incomplete tooth plates PIN 2878/308 (Fig. 15c) from Azarovo quarry near Kaluga, Kaluga Region and PIN 1488/30 from Erino village, close to Mikhailov town, Ryasan Region; both from biogenic limestones of the Mikhailov Horizon, Upper Vis6an, Central Russia. Poecilodus jonesii M ' C o y , 1855 Fig. 15f Description. Toothplate much elongated antero-posteriorly. The coronal surface bears a prominent groove and several transverse ridges.

DINANTIAN FISHES OF RUSSIA Lingual edge rounded and sinuous in correspondence to the outline of the crown surface. The only available specimen is a fairly complete tooth in all respects similar to those described by Davis (1883) and elsewhere. Material, age and geographical distribution. An isolated tooth plate PIN 1988/7 from black clays of the Tula Horizon, Upper Vis6an, UstBuzuluk area, borehole 41, 81.5-85.5 m, Volgograd Region, Southern Russia.

Poecilodus aft. cestriensis St John & Worthen, 1883 The specimen is a partly broken tooth still permitting identification and very close to a specimen figured by St John & Worthen (1883, P1. VIII, Fig. 15). Material, age and geographical distribution. An isolated fragmented tooth plate PIN 4472/40 from the Vis6an of Bolshoy Patok River, Komi Republic, Northern Russia.

Poecilodus sp. An imperfect tooth with only partially preserved crown still makes possible its attribution to this genus. Material, age and geographical distribution. An isolated badly damaged tooth PIN 2876/6 from argillaceous limestones of the Aleksin Horizon, Upper Vis~an, Priksha River, Borovichi District, Novgorod Region, central-west Russia.

Streblodus cf. oblongus (Portlock, 1843) Fig. 15e Description. Much inrolled and antero-posteriorly elongated, incomplete tooth plate lacking the convoluted part. Two very broad and low ridges are separated by a shallow groove, transverse to the long axis of the tooth. In general features it is similar to the principal upper toothplates of the Lower Carboniferous Limestone of Armagh, Ireland, figured, for example, by Patterson (1992). Material, age and geographical distribution. An isolated incomplete toothplate PIN 2876/8 from argillaceous limestones of the Aleksin Horizon, Upper Vis6an, Kamenka River, Borovichi District, Novgorod Region, central-west Russia.

405

Order Psammodontiformes Obruchev, 1953 Family Psammodontidae De Koninck, 1878 Psammodontidae indet. Fig. 15g Very large asymmetric tooth plate with incompletely preserved margins and lacking the basal part. Orientation of these tooth plates in the mouth cavity is uncertain, although there have been attempts at reconstructions, e.g. MoyThomas (1939), and the tooth orientation in this paper is based on his illustration. The oral surface is almost flat, the lateral (medial?) part showing a shallow transverse depression. The posterior edge is deeply embayed laterally and tapers, the anterior one is slightly rugose and ends abruptly; the edge here forms marked vertical crenulations. Material, age and geographical distribution. An isolated fragmented tooth plate PIN 4472/21 from bituminous argillaceous limestones of the Mikhailov Horizon(?), Upper Visdan; KirpichKyrty, Podcheryom River, Komi Republic, Northern Russia. Order Copodontiformes Obruchev, 1953 Family Copodontidae Davis, 1883 Copodus auriculatus (Davis, 1883) Figs 14m-n The only incomplete tooth presents features typical for this species: two rounded symmetric notches in the posterior part of the crown. One at the left side is well preserved, the right side of the specimen is crushed and only a short portion of the notch is distinguishable. Material, age and geographical distribution. An isolated incomplete toothplate PIN 4472/24 from bituminous argillaceous limestones of the Mikhailov Horizon, Upper Vis6an of Podcheryom River, at the Ist Wall; Komi Republic, Northern Russia. Order Petalodontida Zangerl, 1981 Petalodontidae indet. Only a few specimens belonging to Petalodontidae are known from the Tournaisian-Vis6an interval of the East European Platform, all of them poorly preserved, except PIN 4472/25 which resembles Tanaodus, but more detailed identifications are not possible.

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O. A. LEBEDEV

Material, age and geographical distribution. Three incomplete teeth PIN 4472/25-26 and PIN 4472/29 from Kirpich-Kyrty, Podcheryom River; all from bituminous argillaceous limestones of the Mikhailov Horizon, Upper Vis6an, Komi Republic, Northern Russia.

Order M e n a s p i f o r m e s Obruchev, 1953 F a m i l y Deltoptychiidae Patterson, 1965 Oracanthus vetustus Leidy, 1856 (non Newberry, 1897) Fig. 16a-f Description. Large triangular-shaped, asymmetric plate slightly curved medially and

covered with tubercles of stellate shape, which tend to form transverse rows dorsally and longitudinal ones ventrally towards the apical part (PIN 2878/312, Fig. 16a). The ventrolateral side of the apical part was abraded during life. The medial edge of the specimen is almost completely preserved; its posterior half is ornamented with denticles and the anterior is smooth, separated from the dorsal surface by an abrupt ridge, suggesting a contact area for a neighbouring plate. Most of the ventral side of the plate is broken off, with the exception of the posterior quarter. The anterolateral edge corner is intact and smooth, suggesting the margin of a notch, probably for a fin. Numerous scales were associated with the plate. Their shape generally corresponds to that

Fig. 16. Oracanthus vetustus Leidy 1856. (a) Posterolateral plate of the trunk armour, PIN 2878/312, dorsal view (?). (b)-(f) scales, crown view. (b) PIN 2878/419. (e) PIN 2878,/418. (d)-(e) PIN 2878/417. (f) PIN 2878/420. Scale bar is 300 #m, if not indicated otherwise.

DINANTIAN FISHES OF RUSSIA of Deltoptychius armigerus (Traquair) figured by Patterson (1965, fig. 32), but is much more variable. Most of the scales are of the 'Petrodus' type; they have a fairly thick slightly concave elongated base and a high crown, which is also laterally compressed and ornamented with coarse vertical ridges extending from the base towards the apical part, that may form a crest. The crest is sinusoidally curved, as the vertical ridges do not oppose each other, but are disposed alternately (PIN 2878/418-419, Figs 16b-c). In the others the base is rounded, slightly concave and the crown is formed of vertical blades irregularly fused in the central part and forming thorn-like projections above (PIN 2878/417, Figs 16d-e). Another type is represented by scales with almost flat, irregular, star-shaped base and conical crown ornamented from the sides with irregularly disposed, rough vertical ridges, as in the previous types (PIN 2878/420, Fig. 16f). Remarks. The taxonomy of the numerous Oracanthus species was last discussed by Patterson (1965). His particular concern was the presence of osteocyte spaces in the spines of Oracanthus milleri, a character that might indicate relationships with the acanthodians. Up to now no associated remains were known

407

with the Oracanthus spines. The find of scales similar to those of Deltoptychius armigerus (Traquair) shows close affinity of at least Oracanthus vetustus Leidy, 1856 (non Newberry, 1897) specimens to Deltoptychiidae Patterson, 1965. This serves as grounds for the attribution of a group of species of this genus sensu stricto (sensu Patterson 1965) to this ;amily. Material, age and geographical distribution. Imperfect posterolateral plate PIN 2878/312 and about 50 isolated scales from biogenic limestones of the Mikhailov Horizon, Upper Vis6an, Polotnyany Zavod quarry, Kaluga Region, Central Russia. Class Osteichthyes Order Osteolepiformes F a m i l y Osteolepididae Greiserolepis tulensis Vorobyeva & Lebedev, 1986 Fig. 17a-d

Description. This imperfectly known small osteolepidid fish is characterized by small pineal foramen, elongated nostrils shifted backwards in

Fig. 17. Greiserolepis tulensis Vorobyeva & Lebedev 1986. (a)-(b) Ethmosphenoid blocks. (a) Holotype, PIN 2921/1, dorsal. (b) PIN 2921/2, ventral. (c)-(d) Scales. (c) PIN 2921/6a. (d) PIN 2921/6b. (a)-(c) From Vorobyeva & Lebedev (1986). Scale bar is 10mm.

408

O. A. LEBEDEV

respect to the long lacrymo-maxillary notches and a very wide lateral gular.

Material. age and geographical distribution. Three isolated ethmosphenoid blocks and isolated skull bones, numerous scales in various states of preservation (coll. PIN 2921) (Fig. 17) from argillaceous limestones of Malevka or Upa Horizon ( Richterina latior-Pseudoleperditia venulosa-Shivaella microphtalma ostracode Zone); left bank of the Tresna River opposite to Andreyevka Village (Andreyevka-1 locality), Suvorov District, Tula Region, Central Russia. F a m i l y Rhizodopsidae Berg, 1940 Taeniolepis trautscholdi ( C h a b a k o v , 1927) Fig. 18a-b 1874, Taeniolepis s. sp.: Trautschold, p. 272, Taf. XXVII, fig. 17. 1874, Glyptolepis leptopterus: Trautschold, p. 272, Taf. XXVII, fig. 14. 1927, Rhizodopsis trautseholdi: Chabakov, p. 302.

Description.

Scales rounded or elongated, approximately 8-10mm long. Ornamentation of the non-overlapped surface consists of concentric ridges composed of dentine, the space in between being occupied with radial crests. The overlapped surface is sculptured with longitudinally disposed, unevenly anastomosing and interwoven bony ridges being transformed into small tubercles towards the ossification centre. The visceral side of the scale is smooth and bears numerous, unevenly disposed vascular foramina. Teeth conical and slightly sigmoidally curved apically, rounded in cross-section. Tooth ornamentation is represented by extremely fine longitudinal striation, cutting edge is absent. The base is slightly folded.

Material, age and geographical distribution. About 30 isolated fragments of scales and teeth (coll. PIN 2921 and PIN 4359) from bioclastic and argillaceous limestones of Malyovka-Upa Horizons, Lower Tournaisian of Tula Region, Central Russia. F a m i l y R h i z o d o n t i d a e Traquair, 1881 Pycnoctenion aft. siberiacus (Chabakov, 1927) Fig. 18c-d

Description.

Scales mostly rounded or sometimes elongated, thin, approximately 4-10mm long. Ornamentation over the non-overlapped

surface consists of thin bony longitudinally directed parallel ridges; numerous scales bear deep furrows of sensory canal branches that are superimposed over the ornamentation. The overlapped surface is sculptured with high bony peg-like tubercles, disposed in concentric rows; posteriorly these rows tend to be gradually replaced by the anterior ends of the longitudinal ornamentation crests of non-overlapped surface. The visceral surface of the scale bears concentric increment lines and a knob in the centre of ossification, which is slightly shifted forwards. Teeth conical, high and slightly sigmoidally curved at the apex, rounded in cross-section. Tooth ornamentation is represented by fine striation directed towards the lateral sides, but cutting edges are absent. The base is slightly folded.

Material, age and geographical distribution. About ten fragmented scales and teeth and two complete scales (coll. PIN 2921) from argillaceous limestones of Malevka or Upa Horizons (Rich-

terina latior-Pseudoleperditia venulosa-Shivaella microphtalma ostracode Zone), Lower Tournaisian; left bank of the Tresna River opposite to Andreyevka Village (Andreyevka-1 locality), Suvorov District, Tula Region, Central Russia.

Discussion

Fish assemblages and facies Attempts to establish correlations between finds of organic remains, especially vertebrates, with facies is a risky procedure. We must remember that fossil lists never correspond precisely to the real faunal assemblages (biocenoses) that existed during a given time in certain conditions. The real community may either be inadequately reflected in the fossil record due to the incompleteness of studies and fossil record, as well as biased sampling, or may include alien elements introduced to the orictocenosis by incoming debris flows. Nevertheless, these two information blocks should correspond to a certain degree to the sedimentological conditions. Biased sampling might also affect the results if bulk sampling was not used to obtain vertebrate microremains. From the whole Tournaisian-Vis~an interval of the East European Platform the Gumerovo Horizon (Kupavna Formation of the Moscow Syneclise), the Kizele Horizon (upper part of the Cherepet Horizon of the Moscow Syneclise), the Kosva Horizon, the Radayevka Horizon (Glubokovo Formation of the Moscow Synclise),

DINANTIAN FISHES OF RUSSIA

409

Fig. 18. Osteolepiform remains. (a)-(b) Taeniotepis trautscholdi (Chabakov 1927), scale fragments, showing the characteristic ornament of the non-overlapped area. (a) PIN 2921/3140. (b) PIN 4359/79. (e)-(fl) Pycnoctenion aft. siberiacus (Chabakov 1927). (e) PIN 2921/3142, incomplete scale. (fl) Tooth, PIN 2921/3141. Scale bar is 1 mm.

and the Bobriki and Venev Horizons are not characterized with fish remains or there are no specimens in the collection, probably owing to very rare occurrence or poor coverage. For this reason these intervals are not discussed here (Fig. 19). The Lower Carboniferous deposits in the central parts of the Russian Platform were formed in semi-closed basins, only rarely connecting with the ocean, which is why the conodont assemblages are impoverished and represented by shallow-water taxa. The standard conodont zonation can only be tentatively applied to these deposits (Makhlina et at. 1993; Makhlina this volume).

Tournaisian Malevka Hor&on. The Malevka deposits are represented by interbedding of clays with thin layers of the bioclastic limestones, containing the remains of a rich assemblage of invertebrates and fishes. These sediments are interpreted as belonging to intertidal, lagoonal and shallow-

water facies with significant input of flesh water (Makhlina et al. 1993). Fish remains are represented by extremely small isolated remains: teeth, scales and spine fragments. Almost all elements are eroded to some extent. The absence of articulated skeletons, abrasion and apparent sorting by size suggests active hydrodynamic conditions in the basin. Patrognathus crassus conodont Zone includes the upper part of the Kupavna (Gumerovo) Horizon and the lower part of the Malevka Horizon. The upper part of the latter and the lower part of the Upa Horizon belong to the P. variabilis conodont Zone (for discussion of zonation, see Makhlina et al. 1993; Alekseev et al. 1994). The following species are known from numerous localities, both outcrops and subsurface of the southwest of the Tula and west of the Ryasan Region: Acanthodei: "Acanthodes' indet.; Chondrichthyes: 'Orodus' turn# dus Trautschold, Stethacanthus altonensis (St John & Worthen), S. obtusus (Trautschold), Eunernacanthus krapivnensis sp. nov.; Osteolepiformes: Taeniolepis trautscholdi (Chabakov),

410

O. A. LEBEDEV S

SE

S

R

[ M o s c o w Syneclise &

Horizons] Voronezh Anteclise [ . ~ a l conodont zone~

Fish Assemblages Moscow Syneclise

VENEV

Paragnathodus multinodosus

MIKHAILOV

Paragnathodu~ "Acanthodes" sp., nodosus Xenacanthus? nebraskensis? Denaea sp., Symmorium occidentalis, Ctenacanthidae ind., Deltodus sp., Oracanthus vetustus, Psephodus sp.

V C I A S R E

ALEKSIN

Gnathodus bilineatus

TULA

Deltodus sp., "Helodus" sp., Copodus auriculatus,

Petalodontidae ind., Psammodontidae ind AcfinopteTygii ind.

?Desmiodus sp., Psephodus sp., Poecilodus sp., Streblodus cf. oblongus Poecilodus jonesii

O N N I T F

BOBRIKI RADAYEVKA KOSVA KIZEL CHEREPET

E O UPA U

Patrognathus andersoni

R R O N U A MALYEVKA S

Patrognathus variabilis

I S

I

. . . . . . . Patrognathus ¢ra$$us

A N

Northern Urals

?b"-se~oa,,ssv.,

"Hdodus" sp.

B A

Voronezh Anteclise

t Psep~u~us sp. Psephodus dentatus, Helodus" aversus, Strepheoschema fouldensis, Aetheretraon valen t iacu m , Holurus parki "Acanthodes" sp., "Orodus "tumidus, Eunemacanthus krapivnensis, S tethacanthus altonensis, S. obtusus, Taeniolepis trautscholdi,

Stethacanthus obtusus, Diplodoselache antiqua, Ageleodus sp., Eu nemaca nt h us krapivnensis, Greiserolepis tulensis, Pycnoctenion afL siberiacus, Taeniolepis trautscholdi, Dipnoi indet., ?S t rcpheos chema Osteolepididae ind., sp., Str~u~osch~,ma sp., Aetheretmon sp. Rhadinichthys sp.

Euselachii indef., Deltodus sp., Ps~h0dus sp., "Hdodus" sp.

"A canthodes "

sp., Lissodus pectinatus, "Orodus " tumidus, ?Thrinacodus sp.,

Osteolepiclidac ind., Actinistia ind.

KUPAVNA

Fig. 19. Position of the fish assemblages in the stratigraphic chart of the Moscow Syneclise, Voronezh Anteclise and northern Urals. Local conodont zonation of the Moscow Syneclise and Voronezh Anteclise after Makhlina et al. (1993).

Osteolepididae gen. and sp. indet; Actinopterygii: Strepheoschema sp., Rhadinichthys sp. The Andreyevka-1 locality is very poorly characterized by invertebrates and its exact age is uncertain; it may be dated within the Upper Malevkian-Lower Upa interval by abundant ostracodes Glyptolichwinella cf. limbata Posner, characteristic of the ostracode Zone Richterina latior-Pseudoleperditia venulosa-Shivaella microphtalma of both the Malevka and the Upa Horizons. Yellow-grey laminated loose marls contain lycopod moulds, abundant ostracodes, probable spirorbiid worm shells and fishes. These marls are interbedded with several limestone layers, which also contain marine invertebrates: Camarotoechia (?)upensis Sokolskaja, Tulathyris vogdti (Peetz) and conodonts Bispathodus aculeatus aculeatus, B. aculeatus plumulus, Patrognathus crassus, Pa. variabilis, Polygnathus parapetus, and Pandorinellina nota.

Stagnant water conditions appeared in the early Upa Basin, on occasion resulting in the formation of thin-layered argillaceous limestones rich in organic carbon; the presence of lycopods indicates hyposalinity and a close land mass. These sediments are interpreted as shallow fresh- or brackish-water deposited in near-shore estuarine conditions. The following list of vertebrates originates from this locality: Chondrichthyes: Stethacanthus obtusus (Trautschold), Eunemacanthus krapivnensis sp. nov., ?Diplodoselache antiqua sp. nov., Ageleodus sp.; Sarcopterygii: Dipnoi indet., Greiserolepis tulensis Vorobyeva & Lebedev, Pycnoctenion aft. siberiacus Chabakov, Taeniolepis trautscholdi (Chabakov); Actinopterygii: ?Strepheoschema sp., ? Aetheretmon sp. The original description and dating by conodonts of this locality and the previously described Malevka Horzon ichthyoassemblage was carried out by Alekseev et al. (1994).

DINANTIAN FISHES OF RUSSIA Turner (1993) discussed numerous vertebrate Coal Measures and Pennsylvanian assemblages in the Northern Hemisphere that yield Ageleodus/ xenacanth/hybodont remains, stressing that the Narrien Range assemblage is the first of this kind found in the Southern Hemisphere; thus this fauna should be globally distributed. Nevertheless, as suggested above, microscopic 'Ageleodus' teeth and various scales, attributed by Turner (1993) to hybodontoids and ctenacanths, might well belong to the same Diplodoselache, thus making the Australian assemblage closer to that from Andreyevka-1 (Tula Region, Russia). Another assemblage originates from the Novy Oskol borehole 86 core 196.16-196.33m of the Belgorod Region which also cannot be dated with certainty. Conodonts found here identified as Polygnathus parapetus Druce are characteristic for both Malevka and Upa deposits (A. S. Alekseev pers. comm. 1994). The sample matrix consists of dark limestone rich in quartz sand particles and markasite, characteristic of offshore anaerobic conditions rich in organic matter. This list comprises: Acanthodei: 'Acanthodes' indet., Chondrichthyes: Lissodus pectinatus sp. nov., 'Orodus' tumidus Trautschold, ?Thrinacodus sp., Osteolepididae indet., Ctenacanthidae indet., Actinistia indet. Upa Horizon. The Upa limestones were deposited in a very shallow-water basin characterized by salinity close to normal marine or slightly brackish water. Shoal calcarenites are abundant. Stagnant water conditions appeared in the early Upa Basin on occasions, resulting in the formation of thin-layered argillaceous limestones rich in organic carbon. The upper part of the Upa Horizon probably indicates an increase in salinity. The major part of the Upa Horizon corresponds to the Patrognathus andersoni conodont Zone. It also does not contain a single species, making possible its correlation to the Standard Zonation. The index species is widely distributed, but is always found in the shallow-water facies (Makhlina et al. 1993). The fish assemblage is impoverished: Chondrichthyes: Psephodus dentatus (Romanovsky), 'Helodus aversus' ; Actinopterygii: Strepheoschema fouldensis White, Aetheretmon valentiacum White, Holurus parki Traquair. Cherepet Horizon. The Cherepet age on the Russian Platform was generally characterized by a marine transgression. Deposits of this age are mostly represented by biogenic open-sea detrital

411

limestones. Most fish remains are poorly preserved and only the chondrichthyan form Psephodus sp. may be determined with certainty.

Vis~an Tula Horizon. The Tula Horizon in the Moscow Synclise is represented by continental lacustrine and alluvial valley deposits. Only occasionally did the sea briefly invade this territory, thus almost no fish remains are yet known except for a chondrichthyan 'Helodus' sp. In the east of the Voronezh Anteclise (Volgograd Region), Poecilodusjonesii (M'Coy) is known from near-shore clays. In the western slope of the Cispolar Urals, near-shore and estuarine facies are represented by detrital bituminous limestones (Chermnykh 1976) which yield the remains of chondrichthyans Deltodus sp., Psephodus sp., 'Helodus' sp. and indeterminable isolated euselachian scales (Kalashnikov 1962). The upper part of the Tula and all of the Aleksin Horizons belong to Gnathodus bilineatus conodont Zone. In western Europe this Zone occupies the lower part of the Upper Vis6an (Makhlina et al. 1993). Aleksin Hor&on. Shallow open-sea conditions with currents predominated in the central and southern parts of the Moscow Syneclise. In the western part facies variability increases, with lagoonal, estuarine and near-shore facies appearing (Makhlina et al. 1993). Fish remains are comparatively rare and include chondrichthyans (?)Desmiodus sp., Psephodus sp., Poecilodus sp. and Streblodus cf. oblongus (Portlock). Mikhailov Horizon. The environmental conditions in the Moscow Syneclise during this age varied from lagoonal to open-sea. The conodont Zone index species, Paragnathodus nodosus, first appears in the middle of the Mikhailov Horizon. The zone extends into the Venev Horizon and is then replaced by the Paragnathodus multinodosus Zone, which extends into the Serpukhovian. The bulk of the information on the species list known from this region came from a sample from Polotnyany Zavod quarry, comprising a plate of Oracanthus vetustus Leidy prepared with acetic acid. This sample yielded microremains of: Acanthodei: 'Acanthodes' indet.; Chondrichthyes: Xenacanthus ?nebraskensis Johnson, Symmorium occidental& (Leidy), ?Denaea sp., Ctenacanthidae indet., Oracanthus vetustus Leidy scales. The deposits are interpreted as

412

O. A. LEBEDEV

being laid down in a near-shore environment with comparatively calm water. Deltodus sp. and Psephodus sp. are known from contemporaneous deposits of the Kaluga and Ryasan regions. In the Northern Urals the Mikhailov deposits are mostly shallow-water. They are composed of bituminous argillaceous limestones with banks of the brachiopod Gigantoproductus; sometimes these are replaced with variegated, subaerially deposited carbonaceous mudstones, intercalated with oolitic iron ores, with desiccation mudcracks over the bedding surfaces (Chermnykh 1976), suggesting nearshore intertidal, lagoonal and estuarine conditions. A comparatively rich fish assemblage is known from various localities of this time interval: Psephodus sp., Deltodus sp., 'Helodus' sp., Copodus auriculatus (Davis), Petalodontidae gen. indet., Actinopterygii indet. The Psammodontidae indet, toothplate, which originated from Kirpich-Kyrty, Podcheryom River locality might be of the same age.

A s s e m b l a g e analysis

Lithological and invertebrate data are usually used to reconstruct the environments and habitat of vertebrates. Nevertheless, in some cases such groups as, for example, xenacanthids, become the palaeoenvironmental indicators (e.g. Zangerl 1981; Dick 1981; Masson & Rust 1984). This point of view is refuted here as it is extremely difficult to prove the true fresh-water origin of sediments in the Palaeozoic. The absence of marine invertebrates can only indirectly suggest non-marine environments, including brackishwater conditions. On the other hand, the finds of fish remains in well documented marine sequences demonstrate that these animals could live in these environments, but cannot exclude the possibility of post mortem transportation by streams and currents into different habitats (Efremov 1950). The first attempts to trace the dependence of fish remains burial conditions from the facies in the Carboniferous deposits of the Russian Platform were made by Ivanova & Obruchev

Table 1. The distribution offish genera in the environments of the East European Platform Continental fresh and brackish-water (1)

Lagoonal and estuarine with unstable salinity levels (2)

Near-shore littoral

Off-shore neritic

(3)

(4)

Diplodoselache Ageleodus Pycnoctenion

Dipnoi Taeniolep&

Taeniolepis

Aetgerethmon Greiserolepis Strepheoschema Eunemacanthus Stethacanthus

Aetherethrnon

Rhadinichthys

Osteolepididae

Osteolepididae

Strepheoschema

Strepheoschema Eunemacanthus Stethacanthus ?Denaea Xenacanthus?

Osteolepididae Actinopterygii

?Thrinacodus Lissodus Poecilodus ' Orodus' Psephodus

Poecilodus 'Orodus" Psephodus Oracanthus

' Orodus' Psephodus

Deltodus Streblodus Copodus

Petalodontidae 'A canthodes'

'A canthodes"

Actinistia Disposition in the same line reflects eurybiont taxa.

DINANTIAN FISHES OF RUSSIA (t958) and Obruchev (1977), who suggested two major facially dependent groups: (1) neritic deep-water, including teeth of grinding and piercing types and fin spines; and (2) shallowwater lagoonal and fresh-water sediments yielding separate or sometimes articulated remains. The analysis of the composition of Tournaisian-Visran fish assemblages of the East European Platform makes it possible to suggest four major types of aquatic vertebrate environments: (1) continental fresh- and brackish-water; (2) lagoonal and estuarine with unstable salinity levels; (3) near-shore littoral; and (4) off-shore neritic (Table 1). It can be seen that some fishes are found in two or even three environmental zones, which suggests that they were eurybiontic and tolerant to changes of salinity, temperatures, hydrodynamics and other parameters of the basin. These are osteolepidids and actinopterygians. The most stenobiont forms are: Diplodoselache, Ageleodus, Pycnoctenion and dipnoans (1); Deltodus, Streblodus, Copodus and petalodontids (2); Lissodus and actinistians (4). The environmental distribution of the other groups is not as well expressed.

Conclusions Most of the fish material from the TournaisianVisran interval of the Moscow Syneclise, Northern Urals and Voronezh Anteclise of the East European Platform is being described for the first time. A new species, ?Diplodoselache antiqua sp. nov. is probably the oldest known member of this genus. Associations of fishes with various detailed facial conditions are given. These fish assemblages were highly dependant on the environmental conditions. The most diverse were the assemblages in the lagoonal and estuarine environments, probably due to active hydrodynamics and effective energy flow from the continent (Malevka and Mikhailov Horizons). Four major types of aquatic vertebrate environments are being established. Eurybiont and stenobiont taxa are suggested based on their facial distribution. The author is grateful to A. S. Alekseev and M. Kh. Makhlina for consultations and advice during the preparation of the manuscript. S. Turner made helpful critical remarks and useful suggestions, as well as correcting language problems. D. N. Esin identified all of the actinopterygian material. G. Zakharenko performed skillful preparation of the material and sorting of residues, as well as performing much of the illustration work. Photographic work was carried out by A. Kuzmin.

413

The report at the European Dinantian Environments II conference was possible with the support of the Russian National Committee of Geologists and UNESCO lUGS IGCP project 328: 'Palaeozoic microvertebrate biochronology and global marinenon-marine correlation'. This is a contribution to that project.

References ALEKSEEV, A. S., LEBEDEV, O. A., BARSKOV, I. S., BARSKOVA,M. I., KONONOVA,L. I. & CHIZHOVA, V. A. 1994. On the stratigraphic position of the Famennian and Tournaisian fossil vertebrate beds in Andreyevka, Tula Region, Central Russia. Proceedings of the Geologist's Association, 105, 41-52. CHERMNYKH, V. A. 1976. [Carboniferous Stratigraphy of Northern Urals.] Nauka, Leningrad [in Russian]. CHERNOV, G. A. 1972. [Palaeozoic of the Bolshaya

Zemlya Tundra and an Outlook of its Oil- and GasBearing.] Nauka, Moscow [in Russian]. CHORN, J. & WHETSTONE, K. N. 1978. On the use of the term nomen vanum in taxonomy. Journal of Paleontology, 52, 494. COPE, E. D. 1893. On Symmorium and the position of the cladodont sharks. American Naturalist, 27, 999-1001. DAVIS, J. W. 1883. On the fossil fishes of the Carboniferous Limestone series of Great Britain.

Scientific Transactions of the Royal Dublin Society, 1, 327-548. DICK, J. R. F. 1981. Diplodoselache woodi, gen. et sp. nov., an early Carboniferous shark from the Midland Valley of Scotland. Transactions of the Royal Society of Edinburgh, Earth Sciences, 72, 99-113. DUFFIN, C. 1985. Revision of the hybodont selachian genus Lissodus Brough (1935). Palaeontographica, 188, 105-152. EFREMOV, I. A. 1950. [Taphonomy and the Geological Record. Book 1. Burial of Terrestrial Faunas in the Palaeozoic]. Trudy Paleontologicheskogo Instituta AN SSSR, 24 [in Russian]. EICHWALD, E. 1860. Lethaea Rossica ou Palrontologie de la Russie decrite et figurre. Ancienne P(riode, Stuttgart, 1, 1493-1607. GINTER, M. 1990. Late Famennian shark teeth from the Holy Cross Mts, Central Poland. Acta Geologica Polonica, 40, 69-81. GLIKMAN, L. S. 1964. [Subclass Elasmobranchii. In: OBRUCHEV, D. V. (ed.) Osnovy Paleontologii. Beschetyustnye, Ryby.] Nauka, Moscow, 196-237 [in Russian]. GROSS, W. 1973. Kleinschuppen, Flossenstacheln und Z~ihne yon Fischen aus Europ~iischen und nordamerikanischen Bonebeds des Devons. Palaeontographica, 142, 51-155. IVANOVA, E. A. & OBRUCHEV, D. V. 1958. [Fishes.] In: IVANOVA,E. A. (ed.) [The Development of the Fauna in Connection with the Conditions of Living.] Trudy Paleontologicheskogo Instituta AN SSSR, 69, 144-146 [in Russian].

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JOHNSON, G. D. 1984. A new species of Xenacanthodii (Chondrichthyes, Elasmobranchii) from the Late Pennsylvanian of Nebraska. In: MENGEL, R. M. (ed.) Papers in Vertebrate Paleontology Honoring Robert Warren Wilson. Special Publication, Carnegie Museum of Natural History, Pittsburgh, 9, 178-186. KALASHNIKOV, N. V. 1962. [On the finds and stratigraphic distribution of vertebrates in the Palaeozoic of Northern Urals and Timan.] Izvestiya Komi filiala Vsesoyuznogo Geograficheskogo Obshchestva, 7, 37-44 [in Russian]. KARATAJUTE-TALIMAA, V. N. 1992. The early stages of the dermal skeleton formation in chondrichthyans. In: MARK-KUR1K, E. (ed.) Fossil Fishes as Living Animals. Proceedings of the II International Colloquium on the study of the Palaeozoic fishes, Tallinn, 1989. Academia, 1, 223-231. LEIDY, J. 1873. Contributions to the Extinct Vertebrate Fauna of the Western Territories. Reports of the US Geological Survey of Territories, 1. LONG, J. A. & TURNER, S. 1984. A checklist and bibliography of Australian fossil fish. In: ARCHER, M. & CLAYTON, G. (eds) Vertebrate Zoogeography and Evolution in Australasia. Hesperian Press, Western Australia, 235-254. MAISEY, J. G. 1982. Studies on the Paleozoic selachian genus Ctenacanthus Agassiz: No. 2. Bythiacanthus St. John and Worthen, Amelacanthus, new genus, Eunemacanthus St. John and Worthen, Sphenacanthus Agassiz, and Wodnika Munster. American Museum Novitates, 2722, 1-24. MAKHLINA, M. KH. 1996. Cyclic stratigraphy, facies and fauna of the Lower Carboniferous (Dinantian) of the Moscow Syneclise and Voronezh Anteclise. This volume. --, VDOVENKO, M. V., ALEKSEEV, A. S., BYVSHEVA, T. V., DONAKOVA, L. M., ZHULITOVA, V. E., KONONOVA, L. I., UMNOVA, N. I. & SHICK, E. M. 1993. [The Lower Carboniferous of the Moscow Syneclise and Voronezh Anteclise.] Nauka, Moscow [in Russian]. MASSON, A. G. & RUST, B. R. 1984. Freshwater shark teeth as paleoenvironmental indicators in the Upper Pennsylvanian Morien Group of the Sydney Basin, Nova Scotia. Canadian Journal of Earth Sciences, 21, 1151-1155. MoY-THOMAS, J. A. 1939. The early evolution and relationships of the elasmobranchs. Biological Reviews, 14, 1-26. OBRUCHEV, D. V. 1977. [On the fishes of the Carboniferous.] In: OBRUCHEV, D. V. (ed.) [Essays on the Phylogeny and Systematics of Fossil Fishes and Agnathans.] Nauka, Moscow, 6-13 [in Russian]. ~RVIG, T. 1966. Histologic studies of ostracoderms, placoderms and fossil elasmobranchs. 2. On the dermal skeleton of two late Palaeozoic elasmobranchs. Arkivfr"r Zoologie, 19, 1-39. PANDER, C. H. 1858. Ober die Ctenodipterinen des devonischen Systems. Kaiserliche Akademie der Wissenschaften, Sankt-Petersburg.

PATTERSON, C. 1965. The phylogeny of the chimaeroids. Philosophical Transactions of the Royal Society of London, Set. B, Biological Sciences, 249, 101-219. 1992. Interpretation of the toothplates of chimaeroid fishes. Zoological Journal of the Linnean Society, 106, 33-61. ROHON, V. 1893. [Report on the MalevkaMurayevnya stage fishes.] Trudy SanktPeterburgskogo Obshchestva Estestvoispytateley, 22, 2, Protocols, Sect. of Geology and Mineralogy, VI [in Russian]. SEMENOV, P. & MOLLER, V. 1864a. Ober die oberen devonischen Schichten des mittleren Russlands. Bulletin d'Acad~mie Imp~riale des Sciences, SanktPetersburg, 8, 227-264. - & 1864b.[On the upper Devonian layers of the Central Russia.] Gorny Zhurnal, 1, 187-233 [in Russian]. SHCHUROVSKY, G. E. 1878. [Excursions in the Ryasan Gouvernement.] Izvestiya Imperatorskogo Obshchestva Lyubiteley Estestvoznaniya, Antropologii i Etnografii, 23, 462-473 [in Russian]. ST JOHN, O. & WORTHEN, A. H. 1875. Descriptions of fossil fishes. Geological Survey of Illinois, Geology and Palaeontology, 6, 2:1,245-488. & 1883. Descriptions of fossil fishes. A partial revision of the cochliodonts and psammodonts; including notices of miscellaneous material acquired from the Carboniferous formations of the United States. In: WORTHEN, A. H., ST JOHN, O. & MILLER, S. A. (eds) Descriptions of Fossil Vertebrates. Geological Survey of Illinois, 7, 2:1, 55-264. THOMAS, G. A. 1957. Lower Carboniferous deposits in the Fitzroy Basin, Western Australia. The Australian Journal of Science, 19, 160-161. 1959. The Lower Carboniferous Laurel Formation of the Fitzroy Basin. Bureau of Mineral Resources Report, 38, 31-36. TRAQUAIR, R. H. 1888. Notes on Carboniferous Selachii. Geological Magazine, 5, 81-86. TRAUTSCHOLD, H. 1874. Fischreste aus dem Devonischen des Gouvernements Tula. Nouv~lle Memoires de la Societ~ ImpOriale Naturalistes de Moscou, 13, 263-275. TURNER, S. 1982. Middle Palaeozoic elasmobranch remains from Australia. Journal of Vertebrate Paleontology, 2, 117-131. 1990. Early Carboniferous shark remains from the Rockhampton, Queensland. Memoirs of the Queensland Museum, 28, 65-73. 1993. Early Carboniferous microvertebrates from the Narrien Range, central Queensland. Memoirs of the Association of Australasian Palaeontologists, 15, 289-304. VOROBYEVA, E. I. & LEBEDEV, O. A. 1986. [New osteolepidid crossopterygian fishes from the Devonian and Carboniferous of the East European Platform.] Palaeontologicheskiy Zhurnal, 1, 70-77 [in Russian].

D I N A N T I A N FISHES OF RUSSIA WILLIAMS, M. E. 1985. The 'cladodont level' sharks of the Pennsylvanian black shales of Central North America. Palaeontographica, A190, 83-158. ZANGERL, R. 1981. Chondrichthyes I. Paleozoic

Elasmobranchii. Handbook of Palaeoichthyology. G. Fischer Verlag, Stuttgart-New York.

415

ZIDEK, J. 1973. Oklahoma Paleoichthyology. Pt. II. Elasmobranchii (Cladodus, minute elements of cladoselachian derivation, Dittodus, Petrodus).

Oklahoma Geology Notes, Oklahoma Geological Survey, 33, 87-103. - - 1 9 9 3 . A large stethacanthid shark (Elasmobranchii: Symmoriida) from the Mississippian of Oklahoma. Oklahoma Geology Notes, 53, 4-15.

The Early Carboniferous chondrichthyans of the South Urals, Russia A. I V A N O V

Institute o f the Earth's Crust, St. Petersburg University, 16 Lin~]a 29, St. Petersburg 199178, Russia

Abstract: Vertebrate remains including chondrichthyans were found in the Devonian/Early Carboniferous carbonate beds on the western and eastern slopes of the South Urals. Some of the Late Devonian forms (species of Phoebodus, acanthodians (Devononchus), placoderms and onychodontid sarcopterygians) disappear by the middle Siphonodella praesulcata conodont Zone. The ichthyoassemblages are considerably replenished by the new chondrichthyan taxa in the Polygnathus communis carina and Gnathodus texanusMestognathus beckmanni conodont Zones. The chondrichthyans were abundant among the vertebrate faunas of deep-water facies. A new species of chondrichthyan Protacrodus aequalis from the Late Famennian-early Tournaisian of the south Urals is described.

This report presents the first detailed data on the Early Carboniferous chondrichthyans of the Urals. The vertebrates were collected from the carbonate sections on the western and eastern slopes of the South Urals. Localities are Burlya, Kushelga, Popovskiy, Ryauzyak, Sholokh-Sai, Sikaza, Termyantash, Usuili and others (Fig. 1). The investigated interval ranges from the Tournaisian Siphonodella sulcata conodont Zone to the Serpukhovian Gnathodus bilineatus bollandensis-Adetognathus unicornis local conodont Zone (Barskov et al. 1984). The Early Carboniferous deposits are dominated by organic and siliceous limestones with cherts, yielding conodonts, foraminifers, ostracodes, ammonoids, brachiopods, trilobites, etc., interbedded with crinoidal and argillaceous limestones (Kochetkova et al. 1981). The vertebrate assemblages include acanthodians and palaeoniscids, but are strongly dominated by chondrichthyans. These are a taxonomically diverse group represented by various isolated teeth, tooth plates, spines, scales and denticles. The vertebrate remains from the South Urals, usually well dated by conodonts, have been collected by V. Pazukhin (Ufa) and the author. They are extracted with acetic acid from carbonate rocks. Biostratigraphical data are used here based on the local conodont zonation for the Early Carboniferous of Barskov et al. (1984). Lithostratigraphic nomenclature (formations, horizons and beds) is not used, since there are many local subdivisions in the different tectonic- facies zones of the South Urals. The described specimens from the South Urals are housed in the Laboratory of Palaeontology, St. Petersburg University: abbreviated as LP, collections LP8 and L P l l , specimens LP8-2 and LP11-1 to

LPll-246. These investigations are a contribution to IGCP Project N 328 'Palaeozoic Microvertebrates'.

Vertebrates at the Devonian/Carboniferous boundary Most of the Early Carboniferous chondrichthyan taxa in the South Urals appear in the late Famennian and survive across the Devonian/ Carboniferous boundary (Fig. 2). The species of Phoebodus are very common and variable among chondrichthyans in the late Famennian (in the Palmatolepis postera-Palmatolepis expansa conodont Zones) and include Phoebodus australiensis Long, Ph. gothicus Ginter, Ph. limpidus Ginter, Ph. turnerae Ginter & Ivanov, and Phoebodus sp. B (Ginter & Ivanov 1992). They are occasionally present in the late Pa. expansaearly Siphonodella praesulcata conodont Zones and absent from the middle S. praesulcata Zone (Fig. 2). Symmoriids and protacrodontids appear in the Late Devonian and cross the Devonian/Carboniferous boundary. The latest placoderms-ptyctodonts were recorded in the middle Pa. expansa Zone. The Early Carboniferous of the South Urals yields only 'Acanthodes'-type scales from the Late Devonian acanthodians. The latest Devononchus was reported from the late Pa. expansa Zone (Fig. 2). Onychodontid sarcopterygians are very common in the Famennian and have not been found in the Early Carboniferous, the Siphonodella sulcata Zone and above. The vertebrate assemblage of the S. praesulcata-S, sulcata Zones is taxonomically less diverse than that from the zones below and above, but it is rich in chondrichthyan remains

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 417-425.

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Chondrichthyan assemblages of the Early Carboniferous The symmetrical and asymmetrical teeth of Thrinacodus were found starting from the early Pa. expansa Zone and become abundant from the S. sulcata Zone. Thrinacodus is distributed

worldwide in the Early Carboniferous of Australia, China, Poland, France, Spain, England, North America, and also Russia: Novaya Zemlya, the North Caucasus, Moscow Region and from the Polar to the South Urals. Three major types of Thrinacodus teeth may be distinguished: firstly, teeth with a symmetrical crown and an almost symmetrical base as with Thrinacodus nanus St John & Worthen (1875), 'Diplodus' incurvus Newberry & Worthen (1866), and Thrinacodus ferox (Turner) from the Holy

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A rich assemblage was recorded in the Polygnathus communis carina Zone (Fig. 3). The hybodont teeth of Mesodmodus (Figs 7e, f) and tooth plates of bradyodonts (e.g. Fig. 7g) resembling Psephodus and Psammodus appear first in this zone. Mesodmodus occurs in the Kinderhookian (=early Tournaisian) of Iowa (St John & Worthen 1875). Tooth plates of Psephodus and Psammodus are very common in the Early-Middle Carboniferous of the East European Platform, the Kuznetsk Basin and the British Isles (Obruchev 1964). A more extensive radiation of the Early Carboniferous chondrichthyans took place in the Vis6an G. texanus-M, beckmanni Zone. The ichthyoassemblages are considerably increased by new chondrichthyan groups with various types of dentition: Xenacanthus ?nebraskensis Johnson (Figs 4f-h), Lissodus (Fig. 7a-d), Orodus, and petalodontids were also recorded there (Fig. 3). Teeth of Xenacanthus ?nebraskensis occur in the late Pennsylvanian of North America (Johnson 1984), the late Vis6an of the Polar Urals (Kozhim River) and in the Serpukhovian (early Namurian) of the Moscow region. The remains of Lissodus resemble some varieties of L. wirksworthensis Duffin from the late Vis+an of England (Duffin 1985). Teeth of Denaea similar to those of D. meccaensis Williams from the late Pennsylva-

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Gnathodus bilineatus bollandensis-Adetognathus unicornis conodont Zones of the South Urals (Figs 3, 6h, i). These teeth are very similar to the ones of the recent frilled shark Chlamydoselachus anguineus Garman. Apart from the chondrichthyans, the vertebrate assemblages include remains of palaeoniscids and acanthodians.

Systematic

description

Protacrodus aequalis, sp. nov. Fig. 6a-g.

Protacrodus sp. 'C': Turner, 1982. 125-126, Fig. 7.

Etymology. Latin aequalis, equal. Holotype. LP 11-18, Laboratory of Palaeontology, St. Petersburg University. South Urals, Sikaza; Early Tournaisian, Siphonodella sulcata conodont Zone. Other material. 12 complete teeth: LP 11-19LP 11-23 and LP 11-33-LP 11-38. Occurrence. South Urals, Sikaza, Ryauzyak; Late Famennian-early Tournaisian, early S. praesulcata-S, sulcata conodont Zones. Diagnosis. The teeth have three separate cusps in the crown. These cusps are almost equal in size, short and wide, straight or laterally curved, labio-lingually flattened. The cusp striation

consists of clear ridges diverging from the cusp top (3-6 ridges on the labial or lingual surfaces of the cusps). The ridges are sigmoidal on the lateral cusps. The tooth base is directed lingually at nearly a right angle to the crown. There is a lamina on the labial border of the base. The base is triangular or round in form. One or two canals open on the apical border. The basal surface of the tooth is weakly concave. Remarks. This species was found in the Famennian of the Kuznetsk Basin, Tom' and Yaya Rivers and Australia, Queensland (Turner 1982). The striation of the new species is similar to that of some teeth of Xenacanthus ?nebraskensis Johnson. The teeth of Protacrodus vetustus Jaekel differ from the described species by the pyramidal crown with fused basal parts of the cusps and a short and wide base. Also the cusps of P. vetustus are rounded in cross-section.

Symmorium sp. Fig. 5a-1.

Material.

218 teeth: L P l l - 8 - L P l l - 1 7 and LP 11-39-LP 11-246. Occurrence. South Urals, Sikaza, Ryauzyak, Popovskiy; early Tournaisian, S. sulcata-S. duplicata conodont Zones. Description. The specimens vary from teeth with five cusps in the crown (Figs 5c, d) to teeth with 15 cusps (Fig. 51).

SOUTH U R A L S C A R B O N I F E R O U S C H O N D R I C H T H Y A N S

421

Fig. 4. (a-e) Thrinacodus sp., teeth, South Urals, Sikaza; (a, b, d, e) early Tournaisian, (e) late Famennian. (a) LP11-1, occlusal view, S. duplicata Zone. (b) LP 8-2, lateral view, S. sulcata Zone. (c) LP 11-2, occlusal view, Late Pa. expansa Zone. (d) LP 11-3, occlusal view, S. duplicata Zone. (e) LP 11-4, occlusal view, P. communis carina Zone. ([)-(h) Xenacanthus ?nebraskensis Johnson, teeth South Urals, Sikaza, Vis6an, G. texanus-M. beckmanni Zone: (f) LP 11-5, (g) LP 11-6, (h) LP 11-7, occlusal views. Scale bars are 0.25 mm. All the teeth have a strong concavity in the central part of the labial surface (Fig. 5c, f,j,1). There are two processes on each side of the concavity (Fig. 5a, c,e,g). All the cusps are striated with 3-5 straight ridges on the labial and lingual surfaces. The largest lateral cusps are approximately half or a third the height of the main cusp. The teeth with a crown of five cusps have a semicircular base with two indistinct buttons

on the apical surface (Fig. 5c, d). The base width is about l mm. There is no clear b o u n d a r y between the base and the crown on the labial and lingual surfaces. The main cusp bears a lateral carina separating the labial and lingual surfaces (Fig. 5c, d). The outermost lateral cusp pair is higher than the inner ones. The teeth with numerous lateral cusps have a wide crescent base (Fig. 5i,j, 1). The base width is up to 5 mm. There are no indistinct buttons on

422

A. I V A N O V

Fig. 5. Symmorium sp., teeth, South Urals, Sikaza, early Tournaisian. (a), (b) S. sulcata Zone; (c)-(l) S. duplicata Zone. (a) LP 11-8, labial view. (b) LP 11-9, basal view. (c) (d) LP 11-10: (c) occlusal and (d) lingual views. (e) LP 11-11, labial view. (f) LP 11-12, occlusal view. (g) LP 11-13, basal view. (h) LP 11-14, labial view; (i) LP 11-15, basal view; (j), (k) LPll-16. (!) LPII-17, occlusal views. Scale bars are 0.25 ram.

SOUTH URALS CARBONIFEROUS CHONDRICHTHYANS

423

Fig. 6. (a)-(g) Protacrodus aequalis sp. nov., teeth, South Urals, Sikaza, early Tournaisian, S. sulcata Zone. (a), (b) LP 11-18, holotype: (a) labial and (b) occlusal views. (c) LP 11-19, lingual view. (d) LP 11-20, labial view. (e) LP 11-21, occlusal view. (f) LP 11-22, labial view. (g) LP 11-23, labial view. (h)-(i) Denaea sp., teeth, South Urals, Sholokh-Sai, Serpukhovian, G. bilineatus bollandensis-A, unicornis Zone, occlusal views: (h) LP 11-24, (i) LP 11-25. Scale bars are 0.25 mm.

the apical surface. The rows of canal openings are located on the apical and basal surfaces of the base (Fig. 5b, g,i,j,1). The crown is clearly separated from the base, especially in the teeth bearing 15 cusps. Remarks. Wang (1989) described various teeth designated as 'Cladodus' spp. from the Devonian/Carboniferous boundary beds (S. praesulcata-S, duplicata Zones) of the Dapoushang Section, China. One specimen (Wang 1989, pl. 28, Fig. 3) is very similar to the species described here. Teeth of Symmorium occidentalis (Leidy), S. reniforme Cope and other Carboniferous species differ from specimens described here in having a higher and thicker central cusp, and clear and large paired buttons on the lingual

and basal surfaces of the base (Zidek 1973; Williams 1985). Teeth of specimens from the South Urals differ from the Famennian teeth of Symmorium (Long 1990) in possessing fused bases of cusps in the crown and a less concave labial surface.

Discussion and conclusions Chondrichthyan remains have great significance as indicators of facies, changes in environment, sea depth and relationship between basins. In the localities of the South Urals described above, chondrichthyans are recorded in increasing numbers and diversity from the Early

424

A. IVANOV

Fig. 7. Lissodus sp., teeth, South Urals, Sikaza, Vis+an, G. texanus-G, beckmanni Zone: (a)-(b) LP 11-26: (a) lingual and (b) occlusal views. (e) LP 11-27, labial view; (d) LP 11-28, occlusal views; (e)-(f) Mesodmodus sp., tooth, South Urals, Sikaza, Late Tournaisian, Pol. communis carina Zone, LP 11-29: (e) occlusal and (f) labial views. (g) Bradyodont tooth plate, South Urals, Sikaza, Vis+an, G. texanus-M, beckmanni Zone, LP 11-30, occlusal views. (h) Ctenacanthid scale, Sikaza, Early Tournaisian, S. sulcata Zone, LP 11-31. (i) Orodontid scale, Sikaza, Vis+an, G. texanus-M, beckmanni Zone, LP 11-32. Scale bars are 0.25 mm.

Tournaisian to the Vis6an. The environment of the deep part of the shelf with abundant benthic and planktonic invertebrates is believed to be typical for the Vis6an in the South Urals, especially in the G. texanus-M, beckmanni Zone where chondrichthyan remains are the most abundant (Smirnov & Plyusnin 1975; Kochetkova et al. 1981). Xenacanthid taxa such as Xenacanthus ? nebraskensis and X. luedersensis Berman were found in the marine sequences together with hybodontids and petalodontids

(Schultze 1985). However, other xenacanthid species occur in freshwater assemblages. Thrinacodus as well as other phoebodonts are common in the deposits with abundant conodont and ammonoid assemblages. Neither of these groups occur in the Tournaisian and Vis6an shallow water vertebrate assemblages of the Moscow Synclise, but there are numerous sclerophagous chondrichthyans such as Helodus, 'Orodus', Psephodus, Deltodus and others, as well as sarcopterygians. In the South Urals assemblages

SOUTH U R A L S C A R B O N I F E R O U S C H O N D R I C H T H Y A N S

predator sharks with high migration potentials predominated, e.g. Symmorium, Denaea, stethacanthids and ctenacanthids. The sclerophagous chondrichthyans are rare in the Early Carboniferous of the South Urals. I gratefully acknowledge the financial help of the project IGCP 328 and the coordination committee of EDE '94 Symposium to attend the symposium in Dublin and present a paper. I am greatly indebted to V. Pazukhin (Ufa) for the vertebrate materials and stratigraphical information.

References BARSKOV, I. S., ALEKSEEV, A. S., GOREVA, N. V., KONONOVA, L. I. & MIGDISOVA,A. V. 1984. [The Carboniferous conodont zonation of the East European Platform]. In: MENNER, V. V. (ed.) [Palaeontological Characteristics of the Carboniferous Stratotype and Supporting Sections of the Moscow Syneclise (Conodonts, Cephalopods)]. Moscow University Press, 143-150 [in Russian]. DUFFIN, C. J. 1985. Revision of the hybodont selachian genus Lissodus Brough (1935). Palaeontographica Abteilung, 188A, 105-152. GINTER, M. 1990. Late Famennian shark teeth from the Holy Cross Mts, Central Poland. Acta Geologica Polonica, 40, 69-81. - - & IVANOV, A. 1992. Devonian phoebodont shark teeth. Acta Palaeontologica Polonica, 37, 55-75. JOHNSON, G. D. 1984. A new species of Xenacanthodii (Chondriehthyes, Elasmobranehii) from the Late Pennsylvanian of Nebraska. In: MENGEL, R. M. (ed.) Papers in Vertebrate Paleontology" Honoring Robert Warren Wilson. Carnegie Museum of Natural History, Special Publication 9, 178-186. KOCHETKOVA, N. M., LUTFULLIN, YA. L. & PAZUKHIN, V. N. 1981. Scheme of stratigraphy and correlation of the Early Carboniferous on the South Urals. Ufa [in P.ussian].

425

KULAGINA, E. I., RUMYANTSEVA,Z. S., PAZUKHIN, V. N. & KOTCHETOVA, N. N. 1992. [Lower/ Middle Carboniferous Boundary in the South Urals and Central Tien Shan]. Science, Moscow [in Russian]. LONG, J. A. 1990. Late Devonian chondrichthyans and other microvertebrate remains from Northern Thailand. Journal of Vertebrate Paleontology, 10, 59-71. NEWBERRY, J. m. & WORTHEN, A. H. 1866. Descriptions of New Species of Vertebrates, Mainly from the Sub-Carboniferous Limestone and Coal Measures of Illinois. Geological Survey of Illinois Report, 2, 9 141. OBRUCHEV, D. V. 1964. Subclass Holocephali. In: OBRUCHEV, D. V. (ed.) Agnathans, fishes. Fundamentals of Paleontology, 11. Science, Moscow, 238-266 [in Russian]. ST JOHN, O. H. 8~: WORTHEN, A. H. 1875. Descriptions of fossil fishes. Geological Survey of Illinois, Paleontology, 6, 245-488. SCHULTZE, H.-P. 1985. Marine to onshore vertebrates in the Lower Permian of Kansas and their paleoenvironmental implications. University of Kansas Paleontological Contributions, 113, 1-17. SMIRNOV, G. A. & PLYUSNIN, K. P. 1975. History of the geological development of the Urals during the Carboniferous time. In: Carboniferous of the Urals. Sverdlovsk, 3 14 [in Russian]. TURNER, S. 1982. Middle Palaeozoic elasmobranch remains from Australia. Journal of Vertebrate Paleontology, 2, 117-131. WANG, S.-T. 1989. Vertebrate microfossils. In: JI QIANG (ed.) The Dapoushang Section. Science Press, Beijing, 103-108. WILLIAMS, M. E. 1985. The 'cladodont level' sharks of the Pennsylvanian black shales of Central North America. Palaeontographica Abteilung, 190A, 83-158. ZIDEK, J. 1973. Oklahoma paleoichthyology, Part 2: Elasmobranchii (Cladodus, minute elements of cladoselachian derivation, Dittodus, and Petrodus). Oklahoma Geology Notes, 33, 87 103.

Mid-Dinantian brachiopod biofacies from western Ireland D A V I D A. T. H A R P E R

& ANNA

L. J E F F R E Y

Department o f Geology, University College, Galway, Ireland

Abstract: New silicified brachiopod faunas from lower and middle Vis~an horizons have been retrieved from the vicinity of Loughs Carra, Corrib and Mask in west Connaught. In addition, carbonate shell beds dominated by Linoprotonia are recorded from mid County Galway. At present over 40 species of brachiopod are known, from current sampling programmes and the literature, variably associated with bryozoan, coral, echinoderm, mollusc and trilobite faunas; microvertebrates such as sharks teeth are locally common. Four main assemblages are recognized; in ascending order, Arundian strata at Ardnasillagh, Lough Carra, Kilbeg Wood and Ballintober in a northern belt are variably dominated by Rhipidomelta, Schizophoria, Leptagonia, Rugosochonetes, Krotovia, Echinoconchus, Tylothyris, Punctospirifer and Spiriferellina. In addition to genera recorded from Arundian strata elsewhere, Streptorhynchus, Plicochonetes, Dictyoclostus, Pleuropugnoides and 'Spirifer' also occur in the diverse Kiltiernan fauna assembled last century. In a southern belt, possible Holkerian horizons at Dunsandle are less diverse, with faunas dominated by Composita, whereas limestones of Holkerian-early Asbian age at Kiltullagh Bridge have quite different faunas with Schizophoria, Brochocarina, Minythyra and Cleiothyridina; these faunas are more similar to the diverse Asbian assemblages described from County Fermanagh in the north of Ireland. In a central belt, in the Bunoghanaun area, Holkerian to early Asbian horizons are dominated by relatively thick shell beds in pure limestones with mainly opportunist large linoproductids; this fauna is not silicified.

Mid-Dinantian carbonate facies developed adjacent to Loughs Carra, Corrib and Mask in counties Galway and Mayo have yielded locally abundant brachiopod faunas. Deposition occurred on the western part of an exterisive carbonate platform adjacent to an older Precambrian-Lower Palaeozoic massif to the west (Cope et al. 1992). Three main belts are defined in terms of both bio- and lithofacies (Fig. 1). A northern belt of impure, commonly cherty limestones of mainly Arundian age overlies basal Carboniferous sandstones around parts of Loughs Carra, Mask and Corrib. A central belt is dominated by relatively pure limestones, probably ranging in age from Holkerian to Asbian, whereas a southern belt developed east of Oranmore is characterized by a younger, Holkerian-Asbian development of impure muddy and commonly cherty limestones. The brachiopod fauna from the central belt is of low diversity, with shell beds dominated by large linoproductids, whereas the impure limestone facies contains diverse silicified assemblages with a range of bionomic shell types reflecting a spectrum of habitats in shallowwater, nearshore environments.

Stratigraphical setting Despite adequate exposure and many fossiliferous localities, the Lower Carboniferous rocks immediately east of Loughs Corrib and Mask

have been largely neglected in comparison with Dinantian successions elsewhere in Ireland (Sevastopulo 1981). In broad terms, basal clastic facies of Chadian age pass upwards through a variety of nearshore impure carbonate facies, probably ranging in age from late Chadian to Holkerian and succeeded by the purer carbonate facies of the Burren-type limestones. As a whole the succession youngs to the south although there is some local tectonism manifest in a number of open folds with shallowdipping limbs and some faulting. The initial mapping of the region by the Geological Survey (Kinahan 1865, 1869; Kinahan et al. 1867; Kinahan & Nolan 1870; Kinahan & Symes 1871) last century covered the areas around the three loughs in terms of four main units: in ascending order, the Carboniferous Sandstone and the Lower, Middle and Upper limestones. The upper boundary of the Lower Limestone was drawn north of Oranmore. The Middle or Upper Limestone developed in a 'black earthy facies with shales', exposed northwest of Athenry, was differentiated from the Upper Limestones in a pale grey, crinoidal facies. It is probably the 'black earthy facies with shales' that corresponds to the intercalation of muddy limestone noted here at Kiltullagh Bridge and Dunsandle Station. MacDermot & Sevastopulo (1972) presented a more detailed up-to-date analysis of the shelf

From STROGEN,P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 427-436.

428

D. A. T. HARPER & A. L. JEFFREY

IRELAND

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Fig. 1. Generalized map of the study area with localities of brachiopod faunas. facies of the area in a series of palaeogeographical maps and sections• Mid-Vis~an facies developed along a transect from Oughterard to Loughrea are dominated, in the west, by basal sandstones and conglomerates, tidal flat and pellet limestones, together with cherts within a carbonate-shelf setting; and farther east towards Loughrea, argillaceous limestones are succeeded

by deeper-water basinal facies (MacDermot & Sevastopulo 1972, fig. 5). In his lithological review of the Lower Carboniferous stratigraphy of the Irish Midlands, Philcox (1984) included the areas east of the loughs on the fringe of his Dunmore Province• This province is characterized by a thick basal sandstone overlain by condensed argillaceous,

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Fig. 2. Schematic stratigraphical section with locations of main brachiopod faunas. bioclastic limestones; mud-mounds are rare and isolated. More recently, Drew & Daly (1993), in the course of investigations of karstification and groundwater in the region, published a generalized bedrock map based on a variety of unpublished sources. Much of the region is mapped as 'Pure Limestone', although three main tracts are dominated by 'Muddy Limestone'; the latter occur west of Ballinrobe, County Mayo, and in two larger areas south of Kilmaine and east of Oranmore, County Galway. O'Raghallaigh et al. (1995) have described, in general terms, the Carboniferous succession in north Clare and east Galway within the context of the mineral deposits of the region. Throughout most of the area studied, Dinantian limestones appear faulted against the Dalradian and Lower Palaeozoic basement. Around Loughs Corrib and Mask the Dinantian is developed as massive to thick-bedded limestones generally lacking shale, but with minor chert horizons. The limestone succession ranges in age from Chadian to Asbian. Both east and south of Galway, particularly on the Burren and the Aran islands, similar lithological developments have been termed, informally, the Burren Limestone (Gallagher 1994). However, the overlying Brigantian Slievenaglasha Formation includes

over 100 m of coarse bioclastic limestones, dominated by crinoid debris and extensive chert horizons (Drew 1989; Gallagher 1994). Data for the younger Carboniferous strata in northwest Clare are noted in Fitzgerald et al. (1994). With the exception of the northern development of the Holkerian-Asbian pure limestone facies of the Burren Formation, no formal stratigraphical units have been defined or recog-nized in the area studied. The stratigraphical terminology defined here (Fig. 2) is preliminary, and provides the necessary framework to insert the new brachiopod data. Few conodonts have been extracted from this succession; age constraints have been provided by foraminiferans and algae (G. D. Sevastopulo & I. D. Somerville pers. comm.) together with some data from coral assemblages (I. D. Somerville pers. comm.). Detailed locality information together with taxonomic descriptions of the silicified brachiopod faunas are documented in Jeffrey (1992).

Brachiopod assemblages With the notable exception of Brunton's studies (1966, 1968, 1984, 1987) of the Asbian faunas of

430

D. A. T. HARPER & A. L. JEFFREY

the Glencar Limestone, County Fermanagh, little research has been focused on Irish Carboniferous brachiopods since M'Coy's (1844) extensive monograph on the Carboniferous fossils of Ireland. Silicified faunas are documented and illustrated for the first time from this part of Ireland. In broad terms, two main types of brachiopod assemblage are apparent: firstly, relatively diverse assemblages with a wide range of taxa and shell types, usually silicified and often associated with impure, cherty limestones; and secondly, lower diversity unsilicified faunas in purer limestones lacking chert, dominated by large linoproductids and, higher in the succession, gigantoproductids. These faunas, with the exception of the younger gigantoproductids, are discussed below in terms of bionomic shell types (see also Brunton 1987).

The Ardnasillagh, Ballintober, Kilbeg Wood and Lough Carra faunas-northern belt (Arundian) These silicified faunas (Figs 3 & 4) are associated with a variety of nearshore, impure, cherty limestone facies. Many of the limestones are typically skeletal grainstones, less commonly packstones, with low proportions of quartz and feldspar minerals. In general terms the faunas are dominated by the following genera (Table 1): Rhipidomella, Schizophoria, Leptago-

nia, Rugosochonetes, Krotovia, Echinoconchus, Tylothyris, Punctospirifer and Spiriferellina. A range of shell types is represented. Fixed pedunculate orthides such as Rhipidomella and Schizophoria together with the younger shells of the spiriferides Tylothyris, Punctospirifer and Spiriferellina required patches of hard substrate such as other shells or rocks during most of their life cycles. Both Echinoconchus and Krotovia had pedicle valves coated with spines providing some anchorage and stabilization within the sediment. The Rugosochonetes lay recumbent within the soft substrate, together with the adult shells of the spiriferides. For example, adult Tylothyris possessed imbricate stegidial plates with an atrophied pedicle; the shells rested with their posterior surfaces partly within the substrate (Brunton 1984, p. 80). This ambitopic strategy was probably followed by a variety of spiriferides (Legrand Blain 1986). The strophomenide Leptagonia, however, was ambitopic, initially fixed by a pedicle but in later life the stalk atrophied and the brachiopod pursued a recumbent life strategy, quasi-infaunally within a soft substrate (Brunton 1987).

The Kiltiernan fauna, Lough Mask-northern belt (Arundian) A large fauna assembled last century by William King from the area of Kiltiernan adjacent to Lough Mask (Northern Belt) has yet to be precisely localized (Harper 1988). This fauna is silicified and contains most of the brachiopods found in adjacent Arundian horizons noted above, but in addition Streptorhynchus, Plicochonetes, Dictyoclostus, Pleuropugnoides and 'Spirifer' occur (Table 1). The assemblage is diverse, with bryozoans, echinoderms and molluscs together with corals including Siphonophyllia garwoodi Ramsbottom & Mitchell. In addition to the ambitopic, pedunculate, recumbent and spinose forms noted in Arundian strata elsewhere, this Kiltiernan fauna contained further ambitopic taxa such as the rhynchonellide Pleuropugnoides, the recumbent Plicochonetes, and the spinose Dictyoclostus. The orthotetidine Streptorhynchus may have embedded the apex of its conical pedicle valve in the sediment for support.

The Dunsandle Station fauna-southern belt (probably Holkerian) A low diversity brachiopod fauna dominated by Composita? was etched from fine-grained limestones in the railway cutting adjacent to Dunsandle Station. The facies is mainly lime mudstone with occasional patches of skeletal wackestone.

The Kiltullagh Bridge fauna-southern belt ( Holkerian-early Asbian ) This silicified fauna was retrieved from impure, cherty limestones adjacent to Kiltullagh Bridge in the southern belt. The limestones are finegrained skeletal wackestones. The fauna is similar in diversity to the older Arundian assemblages of the northern belt; it is dominated by the genera Schizophoria, Brochocarina, Rugosochonetes, Minythyra and Cleiothyridina together with some indeterminate productids (Table 1). Pedunculate forms such as Schizophoria together with Cleiothyridina and Minythyra required a firm substrate for attachment, although the small valves of Minythyra may have been fixed to other shells (Brunton 1987). Rugosochonetes was a recumbent form, and the orthotetidine Brochocarina may have been cemented to rocks or other shells during at least some of its ontogeny.

MID-DINANTIAN BRACHIOPODS

431

Table 1. Overview of the main stratigraphically-located brachiopodfaunas discussed in this study. Age

Location

Holkerian-early Dunsandle Station Asbian Kiltullagh Bridge

Brachiopod species

Abundance Size range (mm)

Composita sp. 1 Schizophoria resupinata (Martin) cf. lata Demanet Brochocarina sp. cf. B. wexfordensis

Rare (5) Rare (3)

2-16 20-52

Rare (5)

5-9.5

(Smyth) Productide superfam, gen. et. sp. indet. Minythyra sp. nov.

Rare (1) Common

40 1.5-2.5

(>30) Bunoghanaun

Cleiothyridina sp. cf. C. fimbriata (Phillips) Rugosochonetes sp. Linoprotonia ashfellensis Ferguson

Common

,-,80

(>30) Arundian

Lough Carra

Craniscus? sp. Rhipidomella michelini (L6veill+) Leptagon& sp. cf. L. analoga (Phillips)

Rare (1) Common (~60) common

~12 3-15 1.5-11

(>40) Rugosochonetes sp. cf. R. celticus Muir-Wood Krotovia sp. (?spinulosa)

Rare (--4) Common (>10) Rare (~5) Eomarginifera sp. Echinoconchus sp. cf. E. punctatus (J. Sowerby) Common (~15) Common Spiriferellina insculpta subsp, nov.

2-7.5 1-6 12.5 2.5-6.5 2-7.5

(>20) Beecheria sp. 2 Kiltiernan Lough Mask

Ballynalacka Kilbeg Wood

1-13

Common (>15) Common (>2O) Common (>25) Common (>11) Common (-,~40) Common

30-50

Rhipidomella michelini (L6veill+) Plicochonetes buchiana (de Koninck) Echinoconchus punctatus (J. Sowerby) Dictyoclostus semireticulatus (Martin) Schellwienella crenistria (Phillips) Pleuropugnoides pleurodon (Phillips) Actinoconchus planosulcata (Phillips) Composita ambigua (J. Sowerby) Syringothyris cuspidatus (J. Sowerby) Spirifer bisulcatus (J. Sowerby) Tylothyris laminosa (M'Coy) Balanoconcha saccula (J. de C. Sowerby) No brachiopoda data Schizophoria resupinata (Martin) gigantea (Demanet) Cleiothyridina sp. cf. C. fimbriata (Phillips)

Composita sp. 2 Tylothyris laminosa (M'Coy) subsp, nov. Ardnasillagh

Common (13)

Cleiothyridina sp. Tylothyris sp.

17-24

7-18 3-13.5 2-15

(>40) Punctospirifer sp. Chadian

Lemonfield

Common (~30)

9-20

No brachiopoda data

Data for the Kiltiernan locality are based on collections and notes by William King in the James Mitchell Museum. University College Galwav (see Harrier 19gg~.

d

i

t

i

I

i i

r

MID-DINANTIAN BRACHIOPODS

433

Table 2. Data from the Mid-Dinantian locality in the Kiltiernan Townland

Age

Location

Brachiopod species

Abundance

Size range (mm)

Mid-Dinantian

Kiltiernan Townland

Orbiculoidea sp. Rhipidomella michelini (L6veill~) Globosochonetes sp.

Rare (2)

3

Common (>20) Rare (1) Rare (5)

2-4

Common (>30) Rare (~10) Common (~20) Common (~15) Rare (~7)

1.5-5.5

Pleuropugnoides pleurodon (Phillips) Hustedia radialis (Phillips) Cleiothyridinafimbriata (Phillips) Crurithyris urei (Fleming) Crurithyris sp. cf. C. urei (Fleming) Spiriferid gen. et. sp. indet. Unispirifer? sp. Beecheria sp. 1

The Bunoghanaun fauna-central belt ( Holkerian-early Asbian ) This brachiopod fauna was collected from a series of exposed limestone pavements in the townland of Bunoghanaun and adjacent areas, Corrandulla, County Galway. The limestones are relatively pure, typical of the Burren Formation, lacking terrigenous material, cherts and apparently silicified fossils. The limestones are mainly pure skeletal packstones. The fauna is of relatively low diversity, dominated by shell beds of the large linoproductid brachiopod Linoprotonia ashfellensis Ferguson, together with clumps of the massive cerioid coral Lithostrotion vorticale (Parkinson). These large productoids were clearly recumbent, but like many other Palaeozoic brachiopod shell beds such as those formed by Ordovician and Silurian trimerellides and Silurian pentamerides, the large shells probably formed cosupportive shell banks.

The Kiltiernan Townland fauna-southern belt ( Mid-Dinantian) A second Kiltiernan locality, in south County Galway (southern belt), yielded a remarkable

15 1.5-4

1.5-5 1.2-5 <30 6.5-8

micromorphic fauna dominated by minute specimens of Globosochonetes together with small disarticulate specimens of the trilobite Weberiphillipsia (Table 2). The limestones are fine-grained skeletal wackestones and lime mudstones. Although probably younger than most of the other faunas in the region, the precise age and geological setting of this quite different silicified fauna has yet to be established.

Discussion

M'Coy (1844) was the first to describe and illustrate many Irish Carboniferous brachiopods; unfortunately, most of his material is now poorly localized by modern standards and the work as a whole is in need of synoptic revision. More recently, the benchmark studies by Brunton (1966, 1968, 1984) on the exquisitely preserved late Asbian fauna from the upper part of the Glencar Limestone in County Fermanagh have formed the basis for many taxonomic reappraisals of European Dinantian brachiopod faunas; Brunton's studies have provided much of the foundation of the present study. The palaeoenvironments of the Fermanagh succession have been described in detail by Brunton &

Fig. 3. Arundian brachiopod fauna from counties Galway and Mayo (Orthida and Strophomenida). (a) Rhipidomella michelini (Lbveill6), dorsal interior, Lough Carra, JMM.Br506.1, ×4. (b) Schizophoria resupinata (Martin) gigantea (Demanet), dorsal interior, Kilbeg Wood, JMM.Br524.1, x2. (c)-(d) Leptagonia sp. cf. L. analoga (Phillips): (c) dorsal interior, (d) ventral interior, Lough Carra, JMM.Br504.1-2, both ×2. (e)-(f) Rugosochonetes sp. cf. R. celticus (Muir-Wood) ventral exterior and interior, Lough Carra, JMM.Br538.1-2, both ×6. (g)-(i) Echinoconchus sp. cf. E. punctatus (J. Sowerby) ventral interior, dorsal exterior and ventral exterior of immature shells, Lough Carra, JMM.Br522.1-3, all ×6. All specimens are reposited in the James Mitchell Museum (JMM), University College Galway.

434

D. A. T. HARPER & A. L. JEFFREY

Fig. 4. Arundian brachiopod fauna from counties Galway and Mayo (Spiriferida and Athyrida). (a) Tylothyris laminosa (M'Coy), posterior view of conjoined pair, Kilbeg Wood, JMM.Br527.1, × 6. (b)-(c) Punctospirifer sp., external and internal views of dorsal valve, Ardnasillagh, JMM.Br536.1, both × 3. (d)-(e) Spiriferellina insculpta (Phillips), dorsal exterior, JMM.Br507.1 and ventral interior, JMM.Br507.2, both Lough Carra, × 3. (f)-(g) Cleiothyridina sp. cf. C.fimbriata Phillips, dorsal and ventral views of conjoined pair, Kilbeg Wood, JMM.Br525.1, both × 1.5. All specimens are reposited in the James Mitchell Museum (JMM), University College Galway.

Mason (1979), and more recently Brunton (1987) has provided an .ecological analysis of the Glencar fauna. The abundant and diverse Glencar fauna occupied warm, shallow Asbian seas with varying current velocities. The Connaught platform, however, contained similar faunas to the Fermanagh assemblages but developed earlier. Diverse silicified brachiopod faunas with a variety of shell types characterize impure, cherty limestones at Arundian horizons in a northern belt adjacent to Loughs Carra, Corrib and Mask; ecologically similar but younger silicified faunas are associated with muddy limestones with cherts of

Holkerian-early Asbian age, in a southern belt developed east of Kiltullagh Bridge. Intervening pure limestone facies of Holkerian-early Asbian age is characterized by some horizons with shell beds of large linoproductid brachiopods apparently associated with lower diversity faunas. The brachiopod faunas clearly shadow changes in facies on the Connaught part of the platform. Both types of assemblage were probably deposited within depths between Benthic Assemblage 2-3 (Brett et al. 1993), with the more diverse communities, dominated by filter-feeding benthos, apparently developed on more muddy and sandy substrates.

MID-DINANTIAN BRACHIOPODS In a recent detailed study of macrofaunas from the Dinantian of central Scotland, Wilson (1989, fig. 9) charted the distribution of common fossils across a lithological gradient from mudstones to limestones. Gigantoproductus together with other productoids such as Antiquatonia, Latiproductus and Krotovia preferred clear-water environments in either near-shore or off-shore zones. Lithostrotion corals together with Aulophyllum, Caninia and Dibunophyllum had a similar ecological distribution. The majority of pedunculate brachiopods including Schizophoria, 'Spirifer', Cleiothyridina and Beecheria, together with echinoderms, trilobites and zaphrentoid corals, ranged into intermediate zones with increasing amounts of siliciclastic sediment. It is unlikely that water depth varied greatly on the Connaught platform during the midDinantian. Rather, water turbulence, substrate type and quality, together with the proportion of siliciclastic material in the system, may have controlled the distribution of brachiopod faunas through the sequence. Brachiopod faunas from the younger, Asbian-Brigantian parts of the succession in north Clare and on the Aran islands have yet to be documented in detail; reconnaissance studies in both areas have identified several horizons dominated by gigantoproductids s.s. in a variety of pure limestone facies with correlatives in the north of England (Pattison 1981). We thank H. Brunton and M. Legrand Blain for advice during the course of this study. G. Sevastopulo and I. Somerville both provided biostratigraphical data for parts of the sequence, as well as useful discussion. ALJ acknowledges a Postgraduate Fellowship from University College Galway during part of this work. The paper has benefited considerably from reviews by M. Legrand Blain and an anonymous referee.

References BRETT, C. E. BOUCOT, A. J. & JONES, B. 1993. Absolute depths of Silurian benthic assemblages. Lethaia, 26, 25-40. BRUNTON, C. H. C. 1966. Silicified productoids from the Vis6an of County Fermanagh. Bulletin of the British Museum of Natural History (Geology), 12, 175-243. - - 1 9 6 8 . Silicified brachiopods from the Vis6an of County Fermanagh (II). Bulletin of the British Museum of Natural History (Geology), 16, 1-70. - - 1 9 8 4 . Silicified brachiopods from the Vis6an of County Fermanagh (1II). Rhynchonellids, spiriferids and terebratulids. Bulletin of the British Museum of Natural History (Geology), 38, 27-130.

435

1987. The palaeoecology of brachiopods and other faunas of Lower Carboniferous (Asbian) limestones in west Fermanagh. Irish Journal of Earth Sciences, 8, 97-112. - - & MASON, T. R. 1979. Palaeoenvironments and correlations of the Carboniferous rocks in west Fermanagh, Ireland. Bulletin of the British Museum of Natural History (Geology), 32, 91-108. COPE, J. C. W., GUION, P. D., SEVASTOPULO,G. D. & SWAN, A. R. H. 1992. Carboniferous. In: COPE, J. C. W., INGHAM, J. K. & RAWSON, P. F. (eds) A tlas of Palaeogeography and Lithofacies. Memoir of the Geological Society, 13, 67-86. DREW, D. 1989. New caves in the Burren? Irish Speleology, 13, 16-19. - - & DALY, D. 1993. Groundwater and karstification in mid Galway, South Mayo and North Clare. Geological Survey of Ireland Report RS 93/3. FITZGERALD, E., FELLY, M., JOHNSTON, J. D., CLAYTON, G, FITZGERALD, L. J. & SEVASTOPULO, G. D. 1994. The Variscan thermal history of west Clare, Ireland. Geological Magazine, 131, 545-558. GALLAGHER, S. J. 1994. The stratigraphy and cyclicity of the upper Dinantian platform carbonates in parts of southern and western Ireland. European Dinantian Environments II, University College Dublin, 6th-8th September 1994, Abstracts, 9-10. HARPER, D. A. T. 1988. 'The King of Queen's College', William King DSc, first professor of geology at Galway. In: HARPER, D. A. T. (ed.) William King DSc, a Palaeontological Tribute. Galway University Press. JEFFREY, A. L. 1992. Silicified Dinantian Faunasfrom Western Ireland with Emphasis on the Brachiopoda. PhD Thesis, National University of Ireland. KINAHAN, G. H. 1865. Explanatory Memoir to Accompany Sheets 115 and 116. Geological Survey of Ireland. 1869. Explanatory Memoir to Accompany Sheet 105 and Part of l14. Geological Survey of Ireland. - - , FOOT, F. J. & SYMES, R. G. 1867. Explanatory memoir to Accompany Sheets 96, 97 and 106. Geological Survey of Ireland. & NOLAN, J. 1870. Explanatory Memoir to Accompany Sheet 95. Geological Survey of Ireland. -& SYMES, R. G. 1871. Explanatory Memoir to Accompany Sheets 86, 87, 88 and Eastern Part of 85. Geological Survey of Ireland. LEGRAND BLAIN, M. 1986. Morphologie et functiones de la coverture st6gidiale chez quelques Spiriferidina du Pal6ozoique sup6rieur. In RACHBOEUF, P. R. & EMIG, C. C. (eds) Les Brachiopodes Fossiles et Actuels. Biostratigraphie du Pal6ozoique, 4, 331-338. M'CoY, F. 1844. A Synopsis of the Characters of the Carboniferous Limestone Fossils of Ireland. University Press, Dublin. MACDERMOT, C. V. & SEVASTOPULO, G. D. 1972. Upper Devonian and Lower Carboniferous stratigraphical setting of Irish mineralization. Bulletin of the Geological Survey of Ireland, 1, 267-280.

436

D. A. T. H A R P E R & A. L. J E F F R E Y

O'RAGHALLAIGH, C., FEELY, M., MCARDLE, P., MACDERMOT, C., GEOGHEGAN, M. & KEARY, R. 1995. Mineral Localities in the Galway Bay area. Geological Survey of Ireland Report Series RS 90, in press. PATTISON, J. 1981. The Stratigraphical Distribution of Gigantoproductid Brachiopods in VisOan and Namurian Rocks of Some Areas in Northern England. Report of the Institute of Geological Sciences 81/9.

PHILCOX, M. E. 1984. Lower Carboniferous stratigraphy of the Irish Midlands. Irish Association for Economic Geology. SEVASTOPULO, G. O. 1981. Lower Carboniferous. In: HOLLAND, C. H. (ed.) A Geology of Ireland. Scottish Academic Press, Edinburgh, 147 171. WILSON, R. B. 1989. A study of the Dinantian marine macrofossils of central Scotland. Transactions of the Royal Society of Edinburgh." Earth Sciences, 80, 91 126.

A palynofacies analysis of the Dinantian (Asbian) Glenade Sandstone Formation of the Leitrim Group, northwest Ireland JIM SMITH

Department of Geology, University College Cork, Cork, Ireland Abstract: A quantitative palynofacies analysis of the Dinantian (Asbian) Glenade

Sandstone Formation in two boreholes from the Glangevlin area, northwest Ireland, was carried out. It was found that the sands of the Glenade Sandstone Formation prograded rapidly over an arid lagoonal environment in the area of the more proximal northern borehole, but were deposited directly on top of the laminated dolomites of the underlying Meenymore Formation in the area of the southern borehole. Fluvial conditions predominated in the northern area, while lagoonal conditions were established in the southern area, before a major anoxic marine transgression, 12.7m above the base of the Glenade Sandstone Formation in both areas, submerged the flat landscape. In both areas the dominantly fluvial environment above the marine transgression was periodically replaced by more lagoonal environments. The flora at the time of deposition of the Glenade Sandstone Formation may have been dominated by ferns, tree-ferns and small seed-ferns, while herbaceous and arborescent lycopods, typically present in Carboniferous coal swamps , may have been absent.

The accumulation of particulate organic matter (POM) in sediments results from a complex combination of production, transport, reworking and preservation. The study of the distribution of POM in sediments led to the development of palynofacies analysis, details of which can be found in Pocock (1982), Tyson (1993) and references therein. The purpose of this paper is to demonstrate the use ofpalynofacies analysis as carried out on the Dinantian Glenade Sandstone Formation recovered from two borehole cores drilled in 1987 by North West Exploration Plc. near the village of Glangevlin, northwest Ireland (Fig. 1). This study forms part of a larger palynofacies investigation by the author of the entire Leitrim Group, the stratigraphy and sedimentology of which has been described in detail by Brandon & Hodson (1984). Previous palynological studies (Whitaker 1976; Whitaker & Butterworth 1978; Higgs 1984) carried out on the Leitrim Group were restricted to stratigraphic investigations. The Glenade Sandstone Formation occurs in the lower part of the Leitrim Group between the basal Meenymore Formation and the overlying Bellavally Formation (Fig. 2). The regional setting of the Leitrim Group was discussed by Philcox et al. (1992) and Mitchell (1992). The Glenade Sandstone Formation varies in thickness from 120-180m in the Fermanagh Highlands in the north, to less than 4 m in the Arigna area in the south, over a distance of 43km (Brandon & Hodson 1984).

Sli

Spotsamples

/

Borehole

N

Bore

1 Fig. 1. Map showing the location of boreholes and spot samples referred to in the text,

From STROGEN, P., SOMERVILLE,I. D. & JONES, G. LL. (eds), 1996, Recent Advances in Lower Carboniferous Geology, Geological Society Special Publication No. 107, pp. 437-448.

438

J. SMITH

Goniatite Zones. E2b

Bencroy Shale Fm.

E2a

Lackagh Sandstone Fm.

Elc-E2a Sic

Gowlaun Shale Fm. Briscloonagh Sandstone Fm.

Ela_c

Dergvone Shale Fm.

Plb/c'P2c B2-Plb

Carraun Shale Fm. Bellavally Fro.

B2

Glenade Sandstone Fm.

B2

Meenymore Fm.

Mostly shale. ~-~ Mostly sandstone. Evaporite bearing. Fig. 2. Stratigraphic log of the Leitrim Group. Data from Brandon & Hodson (1984).

Materials and methods Two North West Exploration Plc. borehole cores, 87-GL-9 and 87-GL-26, were logged and sampled, the lithologies sampled being restricted to siltstones and mudstones. A total of eight samples over a stratigraphic interval of 30.3 m. from the core of borehole 87-GL-9, and 21 samples over a stratigraphic interval of 57.45 m from the core of borehole 87-GL-26 were collected (Fig. 3). The Glenade Sandstone is composed mainly of orthoquartzitic, typically fine to medium-grained, white sandstone. The sandstones are typically massive, but occasional cross-bedding and slump structures can be seen. Interbedded micaceous, dark grey-green and black, occasionally pyritic siltstones and mudstones occur throughout the formation. Pebble conglomerates with clasts of quartz and angular silt and mudstone fragments occur at some localities. The base of the Glenade Sandstone

Formation is defined as the first sandstone unit above the striped dolomites of the Meenymore Formation. The petrology and regional variation of the Glenade Sandstone Formation was described in detail by Brandon & Hodson (1984) who, based on the absence of marine macrofossils and carbonates, consider it to be of deltaic or brackish water origin, removed from the open sea. In addition, the overlying Namurian Lackagh Sandstone Formation (Fig. 2) from the Geological Survey of Ireland (GSI) borehole core AR90-5 (Fig. 1) was logged and sampled for comparative purposes. A total of 22 samples over a stratigraphic interval of 67.75m were collected (Fig. 4). The lithology of the Lackagh Sandstone Formation is similar to that of the Glenade Sandstone Formation except that it contains coal horizons (AR-5-15, AR-5-10 and AR-5-8) and seat-earths (AR-5-14 and AR-5-7). The petrology and regional variation of the Lackagh Sandstone Formation was described in detail by Brandon & Hodson (1984) who considered it to be deltaic in origin. The Lackagh Sandstone Formation in the GSI borehole is not complete but begins just below the main coal. The deltaic sub-environments present in the borehole include: channels, channel-top environments, coals, seat-earths, backswamps and upper delta slope environments (Fig. 4). Spot samples of two goniatite-bearing marine bands (the Eumorphoceras bisulcatum and Cravenoceras cowlingense marine bands) of the Namurian Gowlaun Shale Formation (Fig. 2) were also collected from stream section 15 of Brandon (1968) to the north of Lough Allen (Fig. I). These samples were taken to compare oxygen levels at the sediment-water interface at the time of deposition of the mudrocks associated with the Glenade Sandstone Formation, with those of a more basinal marine setting.

Sample processing techniques Ten grams of each sample was crushed to fragments 5 mm in diameter. The sample was then processed using standard palynological preparation techniques (Traverse 1988; Van Bergen et al. 1990) to remove all mineral components. After sieving through a 10#m mesh, the > 10 #m residue was spiked by adding eight Lycopodium spore tablets (see Stockmarr 1971 for details of this technique) and boiled for one minute in 40% HCI. The samples were not bleached by any oxidizing agent. A uniform concentration of residue was then cold-mounted

439

PALYNOFACIES OF THE GLENADE SANDSTONE FM.

Key -']

Grain size:CI Clay. S~ Silt. Sd Sand. Gr Gravel.

Mudrock Sandstone.

r-~

Drillcore 87-GL-26 CI S~ Sd Gr

II

I I

Conglomerate.

~

Dolomite. Wavey lamination. Parallel lamination.

/,f j~

Planar cross bedding. Trough cross bedding. Sample number.

GL-26-1

Base of the Glenade Sandstone Formation.

/// ,~jj

Drillcore 87-GL-9. CI SI Sd Gr

II

GL-9-1 GL-9-2

I I

~ . . . . . . .

GL-,-3 - - - ' I - - 7 GL-9-4 ~ '

"

/// scale.

GL-9-5

.,.,.,

GL-9-~ GL-9-7

GL-9-8

Fig. 3. Sedimentary logs of the Glenade Sandstone Formation from the North West Exploration Plc drillcores, Glangevlin, Co. Cavan.

440

J. SMITH Depositional Environment

CI Sl Sd Gr AR.5-22 AR.5-21

I I

I

I

Key D

Channel and channel top

AR.5-20 AR-S-19 AR-5-18

Sandstone.

I Backswamp

AR-,¢,-17 AR.5-16

AR.5-1S AR.5-14 AR.5-13

.

.

.

.

.

.

.

.

.

.

.

.

.

.

Channel top ;-_-_--------_---

Coal and seatearth

Channel top

AR.5-12 AR.5-11 AR.5-10

Channel ::::::::-:--::

Mudrock.

Coal

Coal.

Sty~olite. ...,,.,,~ Wavey lamination. Parallel lamination. / / / Planar cross bedding. .,.,., Troughcross bedding. Erosive boundary. (Marked on lithology symbols). AR.5-22 Sample number. Grain size :cI Clay. Sl Silt. Sd Sand. Gr GraveL

Channel top

AR.5-9 AR-~-8 AR.5-7 AR.5-8

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

Coal and seatearth

Channel top

AR.5.5 Channel top / channel /upper delta slope .J.~.t

AR-5-4

Delta slope AR.54 AR.5-2 AR.5-1

Fig. 4. Sedimentary log and environmental interpretation of the Lackagh Sandstone Formation from the GSI drillcore AR-90-5, Arigna, Co. Roscommon.

in water, without a dispersal agent, on 22 x 22 mm coverslips and allowed to dry. Counts of 500 particles, not including the spike, were made from three slides per sample using a point counter. The distance between each horizontal count line was such that the whole of each coverslip was covered by the count.

PO M Classification The POM classification system used (Table 1) is based on the Amsterdam Palynological Organic Matter Classification (2nd edn) as defined at Aix-en-Provence, France (September 1992). No attempt was made to classify quantitatively the preservation style and degree of degradation

P A L Y N O F A C I E S OF T H E G L E N A D E S A N D S T O N E FM.

441

Table 1. Palynofacies classification scheme Particle

Description

Black woody material (BlWdy)

Thick-walled, sharp-edged wood-like material with an elongate and/or angular shape. Particles are black over their entire surface. Similar to the above apart from the edges of the particle which are brown in colour. Wood-like material, brown in colour, thick-walled, sharp-edged material with one or more of the following characteristics: pits, tracheids, cell structure, elongate and/or angular shape. Terrestrial plant material, excluding spores, showing cellular structure which cannot be classified as woody, epidermal or cuticular material. It is typically thin, sharp-edged, sheet-like material but also includes tube-like structures. Spores of terrestrial plants lacking saccae. Clumps of organic matter that are bright orange-brown in colour, with a well-defined outline, which show no evidence of structure or shape reflecting cellular organization. Similar to the above apart from being a dull grey-brown in colour.

Black-brown woody material (BIBrWdy) Brown woody material (BrWdy) Other plant tissue (OPT)

Non-saccate miospores (NSMio) Orange-brown structureless organic (OrBrSOM) Grey-brown structureless organic matter (GrBrSOM) Fluffy structureless organic matter (FlySOM) Black particular organic matter (BIPOM)

Organic material that shows no evidence of structure or shape reflecting cellular organization and lacks a well defined outline. Thick-walled, sharp-edged material with a roughly equidimensional, angular to sub-angular shape. Particles are black over their entire surface.

700000

600000 E

O) 500000

o u

0 ~

400000

¢,j

300000

0

m

200000

w,,,-

o

6

Z

~00000

.. nn lnL nnl !

o

~,

,o

~

,.o

~

~

,~ ,

~

~

¢~

~

,

..

cz

cz

~

~,

~,

~,

"7,1

~

,...,

~

'I"

,,?, o~, ,,"

t9

,,, ',

,0

~,

,,,,, ~ -, ,

!"

~

'.9

*:'

,

,

o

,,,~ ,,,'~ ~

~

~

'~-

~

~o

~

co

~

o

,

Sample Fig. 5. Number of black woody particles per gram of sample from the Lackagh Sandstone and Glenade Sandstone Formations in drillcores 87-GL-9, 87-GL-26 and AR-90-5. Note that the coal samples (AR-5-15, AR-5-10 and AR-5-8) have been omitted.

442

J.

SMITH

160000

14OOOO

E •-

120000

O 100OOO Q. 0

E

80ooo

m o u m

60000

?

C 0 c ~,. 0

40000

6

Z

20000

,,-

<

<

<

<

~

r~.

tO

1 oon_n ¢~

,-4

~

If"4

i..J

<

iO

lnnnn, n ?

,,

,

(.9

0

(.9

(.9

Sample Fig. 6. Number of non-saccate miospores per gram of sample from the Lackagh Sandstone and Glenade Sandstone Formations in drillcores 87-GL-9, 87-GL-26 and AR-90-5. Note that the coal samples (AR-5-15, AR-5-10 and AR-5-8) have been omitted.

of the organic matter. Although quantitative classification systems have been proposed (e.g. Hart 1986) they may prove ambiguous. Instead, the preservation and degree of degradation of the organic matter was assessed qualitatively for each sample.

Statistical analysis Statistical analysis of the data was carried out using MVSP-A Multivariate Statistical Package (Kovach 1990). The data was not standardized, but was square-root transformed to normalize the data set and a centred principal components analysis (PCA) was performed (Gauch 1982; Pielou 1984; Kovach 1993).

Results

Black woody mater&l and non-saccate miospores The number of black woody particles and nonsaccate miospores per gram of sample from the

Glenade Sandstone Formation, along with data from the Lackagh Sandstone Formation, are shown in Fig. 5 and Fig. 6 respectively. Samples from the coals within the Lackagh Sandstone Formation (samples AR-5-15, AR-510 and AR-5-8) were not included since they would have required significantly different processing techniques. The number of black woody particles in the Glenade Sandstone Formation samples is, with the exception of samples GL-9-3, GL-9-6, GL-9-7 and GL-2618, comparable with samples from the channel top environments of the Lackagh Sandstone Formation (samples AR-5-22, AR-5-20, AR-519, AR-5-18, AR-5-16 and AR-5-13) and lower than delta slope environments (samples AR-53, AR-5-2 and AR-5-1). The number of non-saccate miospores in the Glenade Sandstone Formation samples is comparable with the widely varying numbers found in the Namurian samples, with the exception of sample AR-5-17 from a backswamp deposit, and samples AR-5-14 and AR-5-7, both from seatearths, which give much higher values.

PALYNOFACIES OF THE GLENADE SANDSTONE FM. The miospore genera recovered from the Glenade Sandstone Formation and their palaeobotanical affinity (where known) are documented in Table 2. Notably absent are the genera Densosporites and Lycospora, both of which are abundant in samples from the Lackagh Sandstone Formation.

Preservation and grey-brown structureless matter The preservation and degree of degradation of the organic matter is controlled by its rate of burial and the amount of oxygen in the environment in which it is deposited (Killops & Killops 1993). If the overlying water and uppermost layers of sediment are oxic then only woody material, miospores and other more refractory material tends to be preserved. Disoxic and anoxic conditions result in the preservation of structureless material, often in very large amounts in the case of anoxic environments, and other less refractory material as well as the more refractory material associated with oxic conditions. Structureless organic matter is associated with marine anoxic environments such as goniatite-bearing marine bands (Van De Laar & Fermont 1990). In disoxic and anoxic environments, miospore walls often show pitting associated with pyrite formation. The amount of grey-brown structureless organic matter (GrBrSOM) present in the

Table 2. Miospore genera recoveredfrom the Glenade

443

Glenade Sandstone Formation is shown in Fig. 7. It varies widely, but sharp peaks occur at samples GL-9-4 and GL-26-9 where the values are comparable with samples from the E. bisulcatum and C. cowlingense marine bands of the Gowlaun Shale Formation (Table 3). The degree of pyrite formation and pitting of miospore walls is high in samples GL-9-1, GL-92, GL-9-4, GL-9-8, GL-26-1, GL-26-2, GL-26-9, GL-26-12 and GL-26-17, and shows a positive correlation with the amount of GrBrSOM.

Principal components analysis The results of the PCA indicate that over 93% of the variation in the data set can be accounted for by the first three principal component axes (Table 4). The eigenvectors (Table 5) for each axis can be considered as new independent variables composed of linear combinations of the original variables (Kovach 1993). The principal component score for the first three axes is shown in Fig. 8.

Axis 1. Axis 1 scores show a positive correlation with terrestrial plant material (black, blackbrown and brown woody material, other plant tissue and non-saccate miospores) and a negative correlation with grey-brown structureless organic matter (GrBrSOM). Large amounts of woody material are associated with proximity to fluvial systems (Tyson 1993) and GrBrSOM is associated with anoxic conditions. Axis 1 can, therefore, be considered a measure of the proximity to a fluvial system or the amount of fluvial input into the area of deposition.

Sandstone Formation and their palaeobotanical affinity Spore genera

Palaebotanical affinity

? Apriculatisporis Unknown Apiculiretusispora Unknown Calamospora Various sphenopsid Convolutispora Filicalean Cyclogranisporites Various filicalean, marattialean and lyginopterid Deltoidospora Various filicalean Dictyotriletes Unknown Orbisporis Unknown Perotrilites Unknown Proprisporites Unknown Punctatisporites Various filicalean and marattialean Verrucosisporites Various filicalean, marattialean and lyginopterid Waltzispora Unknown Palaeobotanical data from Smith & Butterworth (1967) and Ravn (1986).

Axis 2. Axis 2 scores show a strong positive correlation with grey-brown structureless organic matter and, as such, can be considered a measure of the influence of marine anoxic conditions in the environment of deposition. Axis 3. Axis 3 scores show a strong positive correlation with non-saccate miospores and a strong negative correlation with black woody material. High axis 3 scores indicate terrestrial input into an area of deposition that is removed from fluvial systems. Discussion

The number of black woody particles per gram of sediment in the Glenade Sandstone Formation samples, relative to samples from the

444

J. SMITH Drillcore 87-GL-26 No. of GrBrSOM particles per gram. ~

? GL-2$-1 GL-25-2 GL-2$-3

No. of GrBrSOM particles per gram. --= O

...i, O O O

O

O

8 O

8 O

I

l

I

l

Drillcore 87-GL-9.

8

8

L

!

° z

GL-25-4

GL-25-6 GL-2rP$ GL-26-7 GL-2¢~.8 GL-25-9 GL-2$-10

8k:t1:1t GL-~-13

GL-26-14 GL-2$-IS GL-2$-IS GL-2(;-17

GL-25-18 GL-26-19 GL-2(;-20 GL-25-21

Fig. 7. The amount of grey-brown structureless organic matter (GrBrSOM) in the Glenade Sandstone Formation samples from the North West Exploration Plc drillcores, Glangevlin, Co. Cavan. Note the logarithmic scale. Ornament and scale as in Fig. 3.

Lackagh Sandstone Formation, indicates that the Glenade Sandstone Formation sandstones were deposited mainly as channel top deposits without the development of a delta slope environment towards the base of the formation.

As black woody material is highly refractory and thus easily reworked, the reduced amount of woody material in the Glenade Sandstone Formation samples relative to the Lackagh Sandstone Formation samples may be due to a

PALYNOFACIES OF THE GLENADE SANDSTONE FM.

445

Table 3. No. of GrBrSOM particles in the Glenade

Table 5. Eigenvectors for a PCA of samples of the

Sandstone Fm andmarine bands of the GowlaunShale Fm

Glenade Sandstone Fm

Sample

GrBrSOM particles per gram

Particle type

GL-9-4 GL-26-9 E. bisulcatum marine band C. cowlingense marine band

193 805 +18931 182413 -t-15 845 143931 +14789

BlWdy B1BrWdy BrWdy OPT NSMio OrBrSOM GrBrSOM FIySOM BIPOM

263 790 d:32 053

The number of grey-brown structureless organic matter (GrBrSOM) particles per gram in two marine band spot samples and two samples from boreholes 87-GL-9 and 87-GL-26.

lower amount of reworking of material caused by a lower flow rate in the fluvial system associated with the Glenade Sandstone Formation compared with that associated with the Lackagh Sandstone Formation. Another possibility, based on differing floral inputs, is discussed below. Although the knowledge of the natural affinities of dispersed miospore genera in the Carboniferous is extremely limited, the possible palaeobotanical affinities of some of the miospores recovered from the Glenade Sandstone Formation is given in Table 2. The genera Densosporites and Lycospora, absent from the Glenade Sandstone Formation but common in the Lackagh Sandstone Formation, have palaeobotanical affinities with small, herbaceous lycopods and with arborescent lycopods respectively (data from Ravn 1986). The known palaeobotanical affinities of the miospore genera recovered from the Glenade Sandstone Formation suggests that the flora contributing the miospores found in the Glenade Sandstone Formation may have been mainly ferns, treeferns and small seed-ferns, while herbaceous and

Table 4. Eigenvalues for a PCA of samples of the

Glenale Sandstone Fm

Axis

Eigenvalue

% of total variation

1 2 3

19809.1 12 326.6 3 308.5

52.1 32.4 8.7

Eigenvalues and variation associated with the first three axes of a PCA carried out on samples from North West Exploration Plc boreholes 87-GL-9 and 87-GL-26.

Axis 1 0.66 0.20 0.30 0.37 0.38 0 -0.37 -0.03 0.12

Axis 2 0.16 0.04 0.14 0.22 0.26 0.05 0.91 0 -0.03

Axis 3 -0.60 0.02 0.11 0.25 0.61 0.30 -0.16 0.26 -0.10

Component loadings of the first three axes of a PCA carried out on samples from North West Exploration Plc boreholes 87-GL-9 and 87-GL-26.

arborescent lycopods, typically present in Carboniferous 'coal swamps', may have been absent. The arborescent lycopods were typically very large, up to 50m in height, (Thomas & Spicer 1987) and their absence may be partly responsible for the reduced amount of woody material in the Glenade Sandstone Formation. The distribution of grey-brown structureless organic matter (GrBrSOM) in the Glenade Sandstone Formation (Fig. 7) indicates a major change in environment at samples GL-94 and GL-26-9 with the incoming of marine anoxic conditions similar to that found in the marine bands. Since both samples are 12.7m above the base of the Glenade Sandstone Formation in their respective cores, this represents a single marine incursion in the area. Other anoxic horizons occur in both cores above and below the main horizon, notably at samples GL9-1, GL-9-2, GL-9-8 and GL-26-17, but these represent disoxic/anoxic lagoonal environments rather than major marine incursions (see below). The principal component axes plots (Fig. 8) give a graphic indication of the changing depositional environments in the Glenade Sandstone Formation. Axis 1 (indicative of fluvial input) shows a marked increase at samples GL-9-3, GL-9-6, GL-9-7, GL-26-18, GL-26-19, GL-26-20 and GL-26-21. The increase at sample GL-9-3 represents a progradation after the marine incursion at GL-9-4 and GL-26-9. The increase is much less in the more distal core 87-GL-26. Axis 1 values are high at the base of core 87-GL-26, but do not increase until some 4 m above the base of core 87-GL-9, indicating a non-fluvial facies in core 87-GL-9 below sample GL-9-7 that is not present in core 87-GL-26.

446

J. SMITH Drillcore 87-GL-26.

Principal Component

Scores

PCA Axis 1 ............

PCA Axis 2

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

PCA Axis 3

8L: t GL-26-3

GL-26-4

Driilcore 87-GL-9.

Principal Component Scores -40-20 0 20 40 60 80 i

,

,

,

,

,

,

, /

I

,'1

GL-26-5 GL-26-6

I

1

/

/

/ /

GL-26-7 GL-26-8 GL-26-9

o°-"

GL-25-I0 GL-26-11 GL-25-12 GL-26-13 GL-26-14 GL-26-15 GL-26-16

/ !

GL-26-17

GL-26-18 GL-26-19 GL-26-20 GL-26-21

Fig. 8. The principal component scores for the first three axes in the Glenade Sandstone Formation samples from the North West Exploration Plc drillcores, Glangevlin, Co. Cavan. Note the logarithmic scale. Ornament and scale as in Fig. 3.

An axis 3 value greater than an axis 1 value represents an environment where the supply of woody material has been cut off while the input of non-saccate miospore has either increased or remained constant, such as a lagoonal environment or a lacustrine environment removed from

any fluvial input. Samples where axis 3 values are greater than axis 1 values also tend to have increased amounts of grey-brown structureless organic matter (but not as high as in the major marine anoxic horizon at samples GL-9-4 and GL-26-9) indicating an disoxic/anoxic lagoonal

PALYNOFACIES OF THE GLENADE SANDSTONE FM. Drillcore 87-GL-9.

447

Drillcore 87-GL-26. B

Mudrocks of the main marine anoxic horizon. Undifferentiated Sandstones, siltstones and mudrocks of the

Glenade Sandstone Formation, Fluvial mudrocks. D

Lagoonal mudrocks.

L ~ J Striped dolomites of the Meenymore Formation. Base of the Glenade Sandstone Fonmation.

;cale.

5m.

GL-9-8----~

I

N

I

29 km.

S

Fig. 9. The location of samples mentioned in the text and the environmental interpretation of the lower part of the Glenade Sandstone Formation in the North West Exploration Pie drillcores, Glangevlin, Co. Cavan. environment of deposition for samples GL-9-1, GL-9-2, GL-9-8, GL-26-1, GL-26-2, GL-26-7 and GL-26-17.

Conclusions The lagoonal environment of deposition at the base of borehole 87-GL-9, and the uniform

thickness of sediment beneath the major marine anoxic horizon at samples GL-9-4 and GL-26-9, suggests that the Glenade Sandstone Formation prograded rapidly over an arid (indicated by the presence of laminated dolomites in the underlying Meenymore Formation), relatively flat landscape. A lagoonal environment of deposition was re-established in the more distal borehole 87-GL-26 before a rapid anoxic

448

J. S M I T H

marine transgression 12.7m above the base of the Glenade Sandstone Formation in both boreholes (see Fig. 9). Above the marine transgression a mainly fluvial environment of deposition was re-established, indicated by the cross-bedding, conglomerates and PCA axis 1 values, being replaced towards the top of both boreholes by a more lagoonal environment of deposition. Throughout the deposition of the Glenade Sandstone Formation the flora may have been represented mainly by ferns, treeferns and small seed-ferns, while herbaceous and arborescent lycopods, typically present in Carboniferous coal swamps, may have been absent. I wish to thank North West Exploration Plc., in particular E. Grennan, for access to the core material. Thanks are also due to K. Higgs for his constructive comments on the manuscript, D. Madden for help in drafting the diagrams, S. McCall, I. O'Connell and C. Roche for technical assistance. G. Clayton and an anonymous referee are also thanked for the constructive comments which greatly aided the clarity of this paper. Finally, special thanks must go to Hugh & Mary Sheridan of Carrick West, Glangevlin for their warm hospitality during sample collection.

References BRANDON, A. 1968. The Geology of the Carboniferous Strata (Vis~an-Namurian) & Parts of Counties Leitrim and Cavan. PhD Thesis, University of Southampton. & HODSON, F. 1984. The Stratigraphy and Palaeontology of the late Vis~an and Early Namurian Rocks of North-east Connaught. Geological Survey of Ireland Special Paper No. 6. GAUCH, H. G. JR. 1982. Multivariate Analysis in Community Ecology. Cambridge University Press, New York. HART, G. F. 1986. Origin and classification of organic matter in clastic systems. Palynology, 10, 1-23. HIGGS, K. 1984. Stratigraphic palynology of the Carboniferous rocks in northwest Ireland. Geological Survey of Ireland Bulletin, 3, 171-202. KILLOPS, S. D. & KILLOPS, V. J. 1993. An Introduction to Organic Geochemistry. Longman Scientific & Technical, Harlow. KOVACH, W. L. 1990. M V S P - A MultiVariate Statistical Package, version 2. INQUA Working Group on data-handling Methods Newsletter, 4, 1-3. - - 1 9 9 3 . Multivariate techniques for biostratigraphical correlation. Journal of the Geological Society, London, 150, 697-705.

MITCHELL, W. I. 1992. The origin of Upper Palaeozoic sedimentary basins in Northern Ireland and relationships with the Canadian Maritime Provinces. In: PARNELL, J. (ed.) Basins on the Atlantic Seaboard." Petroleum Geology, Sedimentology and Basin Evolution. Geological Society Special Publication, 62, 191-202. PHILCOX, M. E., BAILY, H., CLAYTON, G. & SEVASTOPULO, G. D. 1992. Evolution of the Carboniferous Lough Allen Basin, Northwest Ireland. In: PARNELL, J. (ed.) Basins on the Atlantic Seaboard." Petroleum Geology, Sedimentology and Basin Evolution. Geological Society Special Publication, 62, 203-215. PIELOU, E. C. 1984. The Interpretation of Ecological Data. Wiley-Interscience, New York. POCOCK, S. A. J. 1982. Identification and recording of particulate sedimentary organic matter. In: STAPLIN, D. J., DOW, W. G., MILNER, C. W. D., O'CONNOR, D. I., POCOCK, S. A. J., VAN GIJZEL, P., WELTE, D. H. & YI3KLER, M. A. (eds) How to assess Maturation and Paleotemperatures. Society of Economic Paleontologists and Mineralogists, Short Course, 7, 13-131. RAVN, R. L. 1986. Palynostratigraphy of the Lower and Middle Pennsylvanian Coals of Iowa. Iowa Geological Survey Technical Paper No. 7. SMITH, m. V. H. & BUTTERWORTH, M. m. 1967. Miospores in the coal seams of the Carboniferous of Great Britain. Special Papers in Palaeontology, 1. STOCKMARR, J. 1971. Tablets with spore used in absolute pollen analysis. Pollen et Spores, 13, 615-621. THOMAS, B. A. & SPICER, R. A. 1987. The Evolution and Palaeobotany of Land Plants. Croom Helm, London. TRAVERSE, A. 1988. Paleopalynology. Unwin Hyman, Boston. TYSON, R. V. 1993. Palynofacies analysis. In: JENKINS, D. G (ed.) Applied Micropaleontology. Kluwer Academic Publishers, Dordrecht, 153-191. VAN BERGEN, P. F., JANSSEN, N. M. M, ALFERINK, M. & KERP, J. H. F. 1990. Recognition of organic matter types in standard palynological slides. Mededelingen Rijks Geologische Dienst, 45, 9-22. VAN DE LAAR, J. G. M. & EERMONT, W. J. J. 1990. The impact of marine transgressions on palynofacies: the Carboniferous Aegir Marine Band in borehole Kemperkoul- 1. Mededelingen Rijks Geologische Dienst, 45, 75-83. WHITAKER, M. F. 1976. The Palynology of the Carboniferous Sediments in Ireland, with Special Reference to the Ballycastle and Leitrim Areas. PhD Thesis, University of Aston. -& BUTTERWORTH, M. A. 1978. Palynology of Arnsbergian (Upper Carboniferous) strata in Co. Leitrim. Journal of Earth Sciences, Royal Dublin Society, 1, 163-171.

Index Acanthodei, 387, 409, 411,412 'Aeanthodes' indet., 387, 409, 41 l, 412, 417 acanthodians, 417, 420 Actinistia indet, 411 Actinoeyathus, 373 Actinoeyathusfloriformis, 241,377, 380 Actinopterygii, 410, 411,412, 413 aeromagnetic anomalies, Apedale, 346, 347-8 Aetheretmon sp., 411 Aetheretmon valentiacum White, 411 'Ageleodus' sp., 401-402, 411,413 agglutinated stromatolites, 71 Aherlow, metal deposits, 1, 3 Alamo Canyon, 100, 119 Alamogordo Member, 106, 112-14 biotic gradients on homoclinal ramp, 111-12, 114-25

level-bottom beds, 83, 84, 85-6, 89-91, 93, 125 mounds, 84, 85, 86-9, 90-3, 97 Albuera River, coral reef-facies, 146, 147, 151 Aleksin Horizon, fish assemblages, 391,402, 403,405, 409 Aleksin(-sky) Horizon, 362, 363, 364, 380 algae Alamogordo Member, 115, 117, 119, 120, 122, 123, 124 England, 173, 225 as guide fossils, 229, 232 Holy Cross Mountains, 324 Ireland, 241,246, 249, 372, 380 Kingscourt, 133, 135, 141, 142 South Munster Basin, 339 Russian Platform, 364 southwest Spain, 145, 146, 147, 151 algal biostratigraphic ranges, Ireland, 242 algal buildups, 72, 75, 76, 77 algal depth limits, 120, 123 algal filaments, 122 algosponges, 258, 259 Allenwood, metal deposits, 14 alluvial fan lithofacies, Solway Basin, 169-70, 176, 180 alluvial plain lithofacies, Solway Basin, 166-8, 176-80 ammonoids, 423 Anatolipora, 227, 228-9, 232 anehoralis Biozone, South Munster Basin, 339 anchoralis-latus Zones, Lake Valley Formation, 84 anchoralis Zone, Holy Cross Mountains, 322 Andrecito Member, 84, 89, 113, 114 Anglesey, exposure surfaces, 282, 291,292, 294, 298 Antiquatonia, 433 Aoujgalia, 135, 259, 375, 376, 377 aoujgaliids Alamogordo Member, 121-2, 123 Ireland, 130, 133, 135, 139, 375, 376, 377 Apedale magnetic anomalies, 346, 347-8 tufts, 345-7, 348-56 Apedale Borehole, 345, 346, 347-8, 351,352, 355 Aphralysia, 135, 136, 137, 139, 140 Apiculiretispora septalia Zone, Russian Platform, 361 Appalachian Basin, 298 aragonite, 70, 72

Arbigland Group, 170, 171, 172, 174, 175-6, 180 Arcente Formation, 106, 107 Archaediscidae, 225, 226-7, 230-1,232, 379 archaediscids, 241,249, 255, 257, 375, 377, 379 Archaedis¢inae, 379 Archaediscus, 322, 379 Archaediscus gigas-Eostaffella proikensis Zone, Russian Platform, 363 Archaediscus karreri, 256, 257 Archaediscus krestovnikovi, 256, 257 Archaediscus reditus, 257 Archaediscus stilus, 256, 257 Archaediscus varsanofievae, 256, 257 Archaeolithophyllum, 134-5, 136-7, 139, 140, 141, 142 Ardagh Platform, 270, 271 Ardagh Quarry, 130, 131, 132, 134-5 Ardagh, Vis6an buildups, 128-33 biota, 133-42 Ardnasillagh, biota, 430, 431,434 Arenicolites, 168 Argillaceous Bioclastic Limestone Group (ABL), 208, 209, 217 arsenopyrite, Ireland, 32 Arundian buildups, 68, 69, 75, 260-1 Arundian/Holkerian boundary, Ireland, 379 Asbian boundary, early/late, Ireland, 380 buildups, 67, 68, 75, 128, 140, 260-1 carbonate ramp, Ireland, 253-61 carbonates, Urswick Limestone Formation, 221-36 exposure surfaces, 283, 291,293-4 palaeogeography, Ireland, 239, 249 platform carbonates, Ireland, 245, 246, 248, 249 volcanic centres, 355, 356 Asbian-Brigantian palaeogeography, England and Wales, 282 Asbian-Brigantian succession, Ireland, 240 biostratigraphy, 241-3, 249 Asbian/Brigantian boundary, Ireland, 380 cyclicity and sedimentation, 247-8 Asperodiscus, 241 Asperodiscus (=Neoarehaediseus) Zone, Ireland, 241 Asphaltinella, 115, 116, 117, 118, 121 Astbury Limestone Shales, 354 Astbury volcanic rocks, 353-4, 355, 356 Asteroarehaediscus, 380 Asteroarchaediscus bashkirikus, 380 A tractyliopsis, 140 Aulophyllum, 433 auloporid tabulates, 322, 323 Axophyllum, 133, 373 Axophyllum simplex, 377 Axophyllum sp., 257, 259 Axophytlum vaughani, 241 bacterial buildups, 65-77 Bairdiocypris samsonowiczi, 320 Baituganella tchernyshinensis Lipina, 367, 368 Baixo Alentejo Flysch Group (Culm Group), 154 Balbriggan Block, 266, 267, 269, 277 Ballina Limestone Formation, 185, 201,203 Ballinalack, metal deposits, 1, 3, 4

450

INDEX

Ballinleeny volcanic centre, 276 Ballintober, biota, 428 Ballyclogh Limestone Formation, 245 Ballydonnell Member, 246 Ballygarvan Quarry, 338, 339 Ballymartin Formation, 26, 273 Ballysteen Limestone Formation, 24, 26, 31,273, 335 Bandon Syncline, 339 Bantry sub-basin, 333, 339-41 barite (baryte), 32, 36, 39, 40, 41, 47, 48, 264 barium, Derbyshire Platform, 39, 42, 43, 45, 46, 47, 48 Barland Quarry, 285, 287 Barlocco Heugh Formation, 166, 167, 168, 180 Barnalisheen Fault, 28 Basal Cementstones, 54, 170 base metal deposits see metal deposits and mineralization basement structural controls, metal mineral deposits, Ireland, 1-21 basinal sedimentation controls, Ireland, 263-77, Poland 315-27 bathymetry Alamogordo Member assemblages related to, 120-1, 124-5 and exposure surface development, 291-5, 298 Beaconites, 186, 192 Bee Low Limestone, 287, 292 Beecheria, 431,433 Bellisporus nitidus, 350 Benbulben Shale Formation, 185, 204, 254-5, 256 Beresellidae, 122 Bergian Zone, Culm Basin, 307, 312 Bewcastle, 179, 181 Beyrichoceras micronotum, 256 Bibradya grandis, 243 bilaminar palaeotextulariids, 135 bilineatus Zone, Holy Cross Mountains, 322 bioclasts, 'diagnostic', 235 see also guide-fossils bioherms Ireland, 127, 139, 140-1 south west Spain, 150-1 biolithites, 322 biostratigraphy Germany, 304 Ireland, 241-3, 244, 245, 246-7, 249, 253-60, 261, 262, 371-80 Russian Platform, 365-9 biostromal beds, southwest Spain, 146-51 biotic gradients, Alamogordo Member, 111-12, 114-25 Birrenswark volcanic horizon, 170 Bispathodus aculeatus aculeatus, 339, 410 Bispathodus aculeatus plumulus, 410 Bisphaera, 363, 366, 367, 368 Bisphaera malevkensis, 361 bitumen nodules see thoriferous bitumen nodules, Solway Basin bivalves England, 168, 227 Ireland, 131, 133, 135, 137 southwest Spain, 146, 151 Urswick Limestone Formation, 225, 226-7, 228-9, 233

Black Chert Formation (Scwharze Kieselschiefer), 306, 307, 311,312 black matrix breccia (BMB), 24, 26, 30-1, 32 Black Neuk Formation, 53, 166, 167-8, 169-70, 180 block-and-basin sedimentation, Ireland, 276 blue-green algae, 146 Bobriki (Bobrikovsky) Horizon, 361-2, 363, 379 bored grains, Urswick Limestone Formation, 225, 229, 232 Boulder Conglomerate, Navan, 8, 10, 208, 210, 215 BP Miospore Biozone, South Munster Basin, 334 brachiopod shells, 87, 88 brachiopods Alamogordo Member, 115, t19, 123 Holy Cross Mountains, 322, 323 Ireland, 241,246, 258, 259, 260, 273, 371,372, 427-35 Ardagh buildup, 131, 133, 134-5, 137 northern England, 172, 229 Russian Platform, 363 Scotland, 433 southwest Spain, 145, 146, 151 Bradyina rotula, 241,243, 249, 380 bradyodonts, tooth plate, 419, 424 braided fluvial lithofacies, Solway Basin, 177-80 Bray Hill Quarry, 376, 377 Bretonic deformation, 153, 159 Brigantian buildups, 74, 75, 128, 140, 141,260-1,291 exposure surfaces, 283, 284, 291,293, 294, 295 palaeogeography, Ireland, 239 platform carbonates, Ireland, 243, 245, 246, 247, 248, 249 volcanic centres, Apedale, 355, 356 Brochocarina, 430, 431 Brunsia, 368 Brunsiina spp., 368 bryozoan buildups, 77 bryozoans, 67, 141,224, 226-7, 229, 232, 235 Holy Cross Mountains, 322, 323 Ireland, 247, 248, 258, 259, 260, 430 Ardagh buildup, 133, 134-5, 137-8, 139, 140, 141, 142 Sacramento Mountains, 87, 88, 89, 90, 91, 93, 104, 105 Alamogordo Member, 115, 116, 117, 119, 123 southwest Spain, 146, 147, 149, 151 buildups (mounds) bacterial, 65-77 Derbyshire Platform, 291 Ireland, 127-42, 260-1,266 New Mexico, 83-95, 99-108 Bundoran Shale Formation, 197 Bunoghanaun, fauna, 429, 431 burial calcite cements see calcite cements Burren, platform carbonates, 239-40, 244, 245, 246, 247, 249 Burren Formation, 244, 427, 431 Buttevant, platform carbonates, 239-40

INDEX calcareous algae, 118, 120, 141,227, 235 as environmental indicators, 235 Holy Cross Mountains, 320 Ireland, 241,243, 246, 249, 258, 375, 376, 377, 379, 380 Ardagh buildup, 134-5, 137, 139, 140 southwest Spain, 145, 146, 151 Calcifolium, 364, 375, 380 calcimicrobes, 67, 71, 72 calcispheres (-sphaerids), 87, 88, 91,227, 258,259, 320, 321,375 calcite, 30, 31, 32, 70, 71, 72, 86 calcite cements, 66, 68, 71-2, 73 Derbyshire Platform, 35-6, 38-44, 46-7 mineralizing fluids, 37, 44-5, 47-8 trace element release, 45-7 see also cements calcium, Solway Basin, 55 calcium carbonate, precipitation, 70, 71 calcrete mottle profiles, 284, 285, 286, 288, 289, 290, 291,292, 293, 294, 295, 298 calcrete/palaeokarst relationship, 284-8 calcretes, 284, 285, 290, 293, 294, 298 calcretization, 282, 288-90, 291,294, 295, 298 Callan, platform carbonates, 239-40, 244, 246 Camarotoechia (?) upensis Sokolskaja, 410 Canada, Arctic buildups, 74 Caninia, 433 Caninia cf. cornucopiae, 241 Caninophyllum patulurn Biozone, Ireland, 378 Caninophyllum patulum patulum, 377 carbonate buildups (mounds) see buildups (mounds) carbonate mud, formation, 77 carbonate mudbanks, 260-1 carbonate platforms, 35-49, 74, 75, 232, 239-251, 263-279 reconstruction, Fennosarmatia, 315, 318-23, 325-7 see also exposure surfaces, Dinantian carbonate precipitation, bacterially mediated, 72 carbonate ramps Cuilcagh Mountains, 253-61 see also buildups (mounds); carbonate platforms; ramps carbonate-producing organisms, 246 Carboniferous Limestone, South Munster Basin, 337 Carboniferous Limestone shelf, Germany, 304, 307 Carboniferous Slate, South Munster Basin, 337 Carn Limestone Member, 256, 258, 261 Carrickittle, metal deposits, 1, 3 Carrowmoran inlier, 199, 201 Carruthersella, 373 Carruthersella compacta, 377 Castle Archdale-Belhavel Fault Zone, 260 Castle Point Outlier, bitumen nodules, 52, 53-4, 56, 58, 59, 60, 61 Castle Point-Gutcher's Isle Outlier, 164, 165 Castlemore Quarry, 337, 338 catastrophism, Iberian Pyrite Belt, 153-60 'cave popcorn' speleothem deposits, 290-1 'cavity stromatolites', in buildups, 72 Cefn Mawr Limestone, 287, 292

451

cements microbial mediated, 71-2 Navan limestones, 214 vadose, 284, 290-1 see also calcite cements cerioid corals, 246, 377, 431 cerium, Derbyshire Platform, 39, 42, 43, 45, 48 Chadian, base, Solway Basin, 165, 181 Chadian boundary, early/late, Ireland, 379 Chadian succession, Ireland, dating, 374-5 Chadian to Brigantian interval, Ireland, 379 Chadian/Arundian boundary, Ireland, 379 chaetetids, 141 chalcopyrite, 14, 32 channels, Navan limestones, 208, 209, 215, 216 Cherepet(-sky) Horizon, 361,362, 368, 411 Chernyshinella glomiformis, 367, 368 Chernyshinella glomiformis-Septabrunsiina krainicaPalaeospiroptectamrnina tschernyshinensis Zone, Russian Platform, 361 Chernyshinella spp., 367, 368 cherts, Dinantian, Germany, 303-12 chlorophyte thalli, 116 chlorophytes, 133, 134-5, 139, 140, 246 chondrichthyans, 409, 415-23 Chondrichthyes, 390-409, 41 l, 412 Chondrites, 107, 168, 172, 175 Church Stretton Fault, 347 Cincturiasporites appendices Zone, Russian Platform, 363 Cincturiasporites multiplicabilis Zone, Russian Platform, 363 Cingulazonates bialatus-Simozonotriletes brevispinosus Zone, Russian Platform, 363 Cirratriradites saturni, 350 clastic marine facies Solway Basin, 170-6 Spain, 156 clay diagenesis, Derbyshire Platform, 45, 46 clay palaeosols, 283-4 Cleiothyridinafimbriata, 428, 429, 431,432, 433 Climacammina, 380 climatic changes, 304 effect on exposure surfaces, 288, 290, 299 sedimentation and, Culm Basin, 311, 312 Clintz Quarry, 285, 286 Clisiophyllum delicatum nanum, 254 Clisiophyllum garwoodi, 241 Clisiophyllum keyserlingi, 257, 259 Cloghan Hill Limestone Member, 256-7, 258, 259, 260, 261 Cloghany Limestone Member, 256, 257, 258, 259, 261 Clogherboy, metal deposits, 3, l0 clotted microfabrics (micrites), 66, 72, 74, 75 Cloyne Syncline, 333, 334, 340, 345 coal, 173, 174, 290 coastal sedimentation, northwest Ireland, 183-205 Cochliodontidae Owen, 403-4 Cochliodontiformes Obruchev, 402-5 Cochliodus contortus Agassiz, 403-4 Coelosporella, 224, 225, 226-7, 229, 232 Colleen Formation, 271 Conglomerate Group Ore, 10

452

INDEX

Connaught Platform, fauna, 434, 435 conodonts Ireland, 272, 371,373, 374-5, 379, 380 South Munster Basin, 335, 336, 339 Poland, 319, 320, 322 Russia, 410, 411,425 Copodontidae Davis, 405 Copodontiformes Obruchev, 405 Copodus auriculatus, 403, 405, 412, 413 Copodus tooth plates, 403 copper Derbyshire Platform, 39, 42, 43, 45, 47, 48 Ireland, 1, 3, 6, 17 coral reef-facies southwest Spain, 145-51 see also reefs corallites, 230-1 corals, 141 Holy Cross Mountains, 319, 322, 323, 324 Ireland, 137, 241,246, 256, 258, 259, 260, 261,272, 273, 319, 371,372, 373, 377, 379, 428 Ardagh buildup, 133, 135, 139, 140, 141, 142 New Mexico, 87, 88, 90, 115, 117, 119 northern England, 168, 172, 229 Russian Platform, 363 southwest Spain, 145-51 Urswick Limestone Formation, 227 Cork, 244, 245, 246, 247, 249 Cork Group, 339, 340 Cork Syncline, 331,335, 337, 338, 339 Corwenia rugosa, 377, 380 Courceyan boundary, early/late, Ireland, 378 buildups, 67 Courceyan succession, Ireland, 208, 210, 374-5 Courceyan/Chadian boundary, Ireland, 378, 380 Courtbrown, metal deposits, 1, 3 Courtmacsherry Formation, 333, 334, 335, 336, 339, 340 Cravenoceras cowlingense marine band, Ireland, 438, 443, 445 Cregg, Vis6an buildups, 128-9, 131, 140, 141-2 crenulata Zone, Holy Cross Mountains, 322 Cribrospira, 243 Cribrospira panderi, 243 Cribrostornum, 241 Cribrostomum lecomptei, 241,243, 249, 380 Criffel-Dalbeattie Granite, 56, 58, 60 crinoids, 227 Holy Cross Mountains, 320, 321,322, 323 Ireland, 244, 246, 247, 248, 249, 258, 259, 260 Ardagh buildup, 134-5, 140 Sacramento Mountains, 87, 88, 89, 90, 91 Alamogordo Member, 115, 116, 117, 119, 123, 124 Muleshoe Mound, 104, 106 Solway Basin, 172 Croghan Hill, 272 Crosspatrick Formation, 26, 28, 29, 30, 31 cryptalgal fabrics, 87, 88, 91, 130, 135, 324 wave-resistant, 140 cryptic encrustations, 71, 72, 74 "cryptic microbial carbonates', 70

Ctenacanthidae indet, 412 ctenacanthids, 419, 423, 424, 425 'Ctenacanthus'-type scales, 393, 394 Cuilcagh Mountains, carbonate ramp, 253-261 Culm Basin, siliceous rocks, 303-12 Culm Greywacke Formation (Kulm-Grauwacken), 306, 312 Culm Group, 153, 154, 155 Culm Platy Limestone Formation (Kulm-Plattenkalk), 306 Culm Slate Formation (Kulm-Tonschiefer), 306, 312 Cumbria, exposure surfaces, 283, 284, 285, 287, 291, 292, 293, 295 Cumbrian Platform, 295 cuneiformis Zone, Holy Cross Mountains, 322 current-influenced mound growth, 97, 99, 107-8 cutans, 214 cyanophytes, 133, 135, 137, 139, 140, 141, 142 Cyathaxonia cornu, 377 Cyathoclisia modavense, 377 cyclicity palaeoclimatic, 295-8 sedimentary, 282, 295-8 Urswick Limestone Formation, 221,222-31,233, 235 Ireland, 239, 244-49 Dainella, 368 Dartry Limestone Formation, 252, 256, 257, 258, 259, 260, 261 dasycladacean algae, 133, 229, 244 Davidsonina, 232 Davidsonina septosa, 241 Deadman Waulsortian buildup, 100 debrites, 318-19 deep-water buildups, 67, 77 Deer Park Formation, 141 delicatus Zone, Holy Cross Mountains, 322 delta environments and facies, Spain, 156-9 Deltodus sp., 404, 411,412, 413,425 Deltoptychiidae Patterson, 406-7 ?Denaea sp., 400, 412, 420, 425 Densosporites, 445 Densosporites intermedius Zone, Russian Platform, 363 Densosporites variabilis Zone, Russian Platform, 363 depth-related assemblages, 119, 120, 121, 123-4, 125, 140, 146 Derbyshire, exposure surfaces, 284, 285, 287, 291,292, 295 Derbyshire Platform hydrocarbon emplacement, 37, 39, 40, 41, 45, 46-8 mineralization, 35-49 Derryville Fault, 24, 26, 28-9, 30, 31, 32 ?Desmiodontida Zangerl, 391 ?Desmiodus sp., 391, 411 Devonian-Carboniferous boundary Iberian Pyrite Belt, 159, 160 South Urals, 417-18 Devononchus, 415 diagenesis, 26-33, 35-50, 51-63 Dibunophyllurn, 435 Dibunophyllum bipartitum, 133, 241,377, 380 Dibunophyllum bourtonense, 379

INDEX

Dibunophyllum sp., 256, 257, 377 Dictyoclostus, 428, 429 Dictyoclostus teres, 165 Dill-Innerste Zone, 307 Dinantian biostratigraphy, Ireland, 371-9 buildups, 65-77 carbonate ramp, Ireland, 253-61 causes of cyclicity, 248-9 cyclic stratigraphy, Russian Platform, 359-64 emersion surfaces and channels, Navan, 211-17 exposure surfaces controls, 282-3, 291-8 development, 288-91,298-9 lithostratigraphy, Ireland, 243-7 metal-organic interactions, Solway Basin, 51, 54-63 palaeogeography, Ireland, 239-40 platform carbonates, Ireland, 239-49 ramp, southwest England, 124 sedimentation, Ireland, 183-205, 263-77, 331, 333-41 siliceous rocks, Germany, 303-12 diphyphyllids, 377 Diphyphyllum lateseptatum, 241 Diplocraterion, 168, 175 ?Diplodoselache antiqua sp.nov., 394, 395, 396-7, 409, 411 Diplodoselachiidae Dick, 398-9 Diplosphaerina, 320 dissolution cavities, Navan limestones, 21 t-12 Dneiper-Donetz Depression, 379 Doliognathus, 373 Doliognathus latus, 375 Dollymae, 373, 374 Dollymae bouckaerti, 339, 375 dolomite, 10, 11, 30-1, 32, 33, 40 dolomitization, Ireland, 26, 28, 29, 33 Donegal Bay, 197 Donegal succession, 183-92 Donezella, 231 Doorin Shales, 195 Dorlodotia briarti, 377, 379 Dorlodotia pseudoverrniculare, 377 Downpatrick Formation, 185, 194-5, 196 Draffania, 375 Dromdowney Member, 244 Dromkeen Formation, 241,243 Drumanagh Member, 271-2 Drumman More Sandstone Formation, 204 Dublin Basin, 74, 92, 129, 140 biostratigraphy, 371,372, 378, 379, 380 sedimentation controls, 263-72, 276-7 Dunsandle Station, fauna, 427, 430, 431 Durnish Limestones, 273

Earlandia, 226-7, 230-1 Sacramento Mountains, 87, 88, 115, 117, 119, 123 Early Mississippian buildups, New Mexico, 83-95 level-bottom beds, New Mexico, 85-6, 88, 89-91, 93, 125

453

East European Platform, fish assemblages, 385-411

Eblanaia michoti, 378 Echinoconchus, 430, 431,432, 433 Echinoconchus punctatus, 432, 433 echinoderms, 226-7, 228-9, 232 Alamogordo Member, 115, 116, 117, 119, 124 Ireland, 430 Scotland, 435 southwest Spain, 146, 151 echinoids Ireland, 258, 259 Sacramento Mountains, 87, 88, 89, 90, 116, 117, 123 El Almendro, coral reef-facies, 145, 146, 147, 151 El Portezuelo, coral reef-facies, 145, 146, 147 El Torre6n, coral reef-facies, 145, 146, 147, 149, 150, 151 Elasmobranchii, 390-402 Elbingerode Complex, 307, 312 emersion surfaces, Navan, 211-17 encrusting bryozoans, 116, 119, 135, 141, 142, 258, 259, 323 Endostaffella fucoides, 257 Endothrya parakosvensis Lipina, 367, 368

Endothryanopsis compressa-Archaediscus krestovnikovi Zone, Russian Platform, 363

Endothyra/Endothyracea, 322, 373 Endothyridae, 227, 231,232 endothyrids, 226-7, 228-9 environmental indicators, 235 Eochernyshinella spp., 367, 368 Eoendothyranopsis, 379 Eoforschia, 367, 368, 378 Eogeinitzina devonica rara, 320, 321 Eonodosaria rausereae, 320 Eoparastaffella, 375, 378, 379 Eoparastaffella simplex, 369, 379 Eostaffella, 135, 137, 379 Eostaffella ikensis Zone, Russian Platform, 363 Eostaffella parastruvei, 257, 379

Eostaffella tenebrosa-Endothryanopsis sphaerica Zone, Russian Platform, 364

Eotaphrus, 373 Eotaphrus cf. bultyncki, 339, 340 Eotextularia diversa, 374, 378 Epistacheoides, 136, 137, 225, 226-7 erosion and reworking, South Munster Basin, 331, 333-41 Eskett Quarry, 284 Etruria Formation, 298 Eumorphoceras bisulcatum marine band, Ireland, 438, 443, 445 Eunemacanthus krapivnensis sp. nov, 400-401,409, 411 Euselachii Hay, 390-91 Euselachii indet., 390-91 eustacy Holy Cross Mountains, 326 and sedimentary cyclicity, 207, 216, 221,233, 235, 248 evaporite surface, Navan limestones, 212, 215 'event beds', 153 expansa Zone, Holy Cross Mountains, 320

454

INDEX

exposure surfaces controls, 282-3, 291-8 development, 288-91,298-9 extinction episodes, reef development arrested by, 73 Famennian clasts, Holy Cross Mountains, 320-1 Famennian/Courceyan boundary, Ireland, 378 Famennian/Dinantian boundary, Iberian Pyrite Belt, 153-60 Farleton Fell, 227, 230, 231,232 fasciculate corals, southwest Spain, 148 Fasciella, 136-7, 247, 274, 375 Fasciella kizilia, 244 fault-controlled fluid flow, 2, 33, 47, 48 faulting and mineralization, 2, 8, 10-11, 28, 32, 33, 48 Feltrim Limestone Formation, 4, 266, 268, 274, 376, 377 fenestellid (fenestrate) bryozoans, 115, 116, 117, 119, 124, 134-5, 258, 259 Holy Cross Mountains, 322 Ireland, 244, 245, 246, 247, 249, 258 Kingscourt, 130, 133, 137-8, 140 Sacramento Mountains, 86, 87, 88, 89, 90, 91, 93 Muleshoe Mound, 104, 105 Fennosarmatia, 315, 318-23, 325-7 Fermanagh, fauna, 433, 434 ferroan carbonate, 32 ferroan dolomite, 31 fish assemblages, Eastern European Platform, 387-413 Fistulipora, 135, 141 flora, Glenade Sandstone Formation, 448 fluid inclusions, calcite cements, Derbyshire Platform, 44-5 fluorine, Derbyshire Platform, 39, 42, 43, 45, 47 fluorite, Derbyshire Platform, 36, 39, 40, 41, 47, 48 fluvial channel lithofacies, Solway Basin, 177-80 Ireland, 198-9 'Footwall Green Shale', Navan, 214 foraminifera Holy Cross Mountains, 320, 321,322, 323 Ireland, 244, 246, 256, 258, 259, 261,272, 371,372, 373, 374, 375, 377 Ardagh buildup, 133, 135, 137, 139, 141, 142 South Munster Basin, 336, 339 Russian Platform, 361,363 Sacramento Mountains, 87, 88, 115, 117, 119-20, 123, 124 southwest Spain, 146, 151 Urswick Limestone Formation, 225, 226-7, 228-9, 230-1,235 foraminiferal zones Ireland, 242, 378 Russian Platform, 365-9 framebuilders, 140 Frasnian clasts, Holy Cross Mountains, 320, 321 Furfooz, Waulsortian mounds, 111, 112, 114 fusain, northwest Ireland, 190, 191, 192, 195 galena, 32, 36, 41, 47, 48 Galezice, 316, 323, 325-7 Galmoy, metal deposits, 3, 14 Galway, fauna, 431,432

gastropods Holy Cross Mountains, 322 Ireland, 189, 194, 258, 259 Ardagh buildup, 131, 133, 137 Sacramento Mountains, 87, 88 Alamogordo Member, 115, 117, 119, 123, 124 Solway Basin, 168 southwest Spain, 146, 149 Urswick Limestone Formation, 225, 227, 232, 233 Germany, siliceous rocks, 303-12 geochemistry, 41-44, 345, 350-54 gigantoproductids, 146, 147, 149, 151 Gigantoproductus, 410, 433 Gigantoproductus aft. semiglobosus, 151 Gigantoproductus cf. edelbergensis, 241 Gigantoproductus cf. giganteus, 241 Gigantoproductus maximus, 133 Gigasbia gigas, 225, 256, 257 Gillfoot Beds, 54, 170, 174, 175, 176, 180 Girvanella Ardagh buildup, 130, 133, 135, 136-7, 140 Holy Cross Mountains, 323 Sacramento Mountains, 87, 88, 89, 91 Urswick Limestone Formation, 225, 227, 229, 232, 235 girvanellids, 323 glacio-eustacy Ireland, 249 sedimentary cyclicity, Urswick Limestone Formation, 221 sedimentation and, 295, 296 Glandore High, 333, 338, 339, 340 Glenade Sandstone Formation, 437--48 Glencar Limestone Formation, 255-6, 258, 260, 261, 296, 433,434 Globochaetes, 87, 88, 91,324, 375, 376, 377, 423 globochaetids, 323 Globosochonetes, 433 Glomodiscus, 379 Glyptolichwinella cf. limbata Posner, 410 gnathodids, 375 Gnathodus, 373 Gnathodus bilineatus, 336 Gnathodus bilineatus bilineatus, 380 Gnathodus bilineatus, conodont Zone, East European Platform, 411 Gnathodus bilineatus bollandensis-Adetognathus unicornis conodont Zone, Urals, 417, 423 Gnathodus cuneiformis, 335 Gnathodus girtyi collinsoni, 380 Gnathodus girtyi girtyi, 380 Gnathodus homopunctatus, 379 Gnathodus texanus-Mestognathus beckmanni Zone, Urals, 419, 420, 421,424, 425 Gnathodus typicus Zone, Urals, 421 Goat Springs, 119 goniatites, 273, 320, 372, 438 Gortalughany Townland, 257-8 Gortdrum metal deposits, 3, 6-7, 8 Gower Peninsula, 282, 283, 293, 295 biota, 222, 233 Gowlaun Shale Formation, 438, 443 Graig Quarry, 285, 287

INDEX

Grandispora upensis Palynozone, Russian Platform, 361 Grange Mill Quarry, 285, 287 gravity anomalies, Ireland, 5, 10, 13 gravity-flow deposits, Spain, 159, 323 green algae, 115, 117, 119, 120, 124, 133, 246, 366 green mudstones, Navan, 213-14, 216 green-blue algae, 146 Greiserolepis tulensis Vorobyeva & Lebedev, 407-8, 411 Guadajira River, coral reef-facies, 145, 146, 147 guide-fossils, 229, 232, 235, 243 Gumerovo Horizon, 361,362 Gun Hill volcanic rocks, 349, 350, 351,352, 353-4 Haplolasma cf. densum, 377 Haplolasma densum, 256 Haplophragmella, 243 Harberton Bridge, metal deposits, 14 Harz Mountains, 303-4, 305, 306, 307, 310, 312 'Hecker-Shvetsov', basal surface, 363 Helodus, 403, 419, 425 'Helodus aversus', 403, 411 ?Helodus sp., 402-3, 411 Hendre Quarry, 287 Hercynian orogeny, Iberian Pyrite Belt, 153, 154, 159 heterocorals, 133, 259, 322, 323 hexactinellid sponges, 115, 116, 120, 134-5 Hexaphyllia, 133, 259 high-magnesium calcite, microbial formation, 70 Holkerian, base, Benbulben Shale Formation, 204 Holkerian/Asbian boundary, Ireland, 379, 380 Holme Park Quarry, 232, 287 Holmpatrick Formation, 271 Holocephali, 402-7 Holurus parki Traquair, 411 Holy Cross Mountains, 315-18 carbonate platform reconstruction, 315-329 palaeogeography, 323-5, 327 homoclinal ramp biotic gradients, New Mexico, 111-12, 114-25 Ireland, 269 Hrrre-Gommern Zone, 304, 307 Howchinia, 225, 235, 322, 336 Howchinia bradyana, 241,243, 249, 257, 375, 380 hyalosteliids, 90, 91, 123 hybodontids, 417 Hybodontoidea Zangerl, 390-91 hydrocarbons emplacement, Derbyshire Platform, 37, 39, 40, 41, 45, 46-8 groundwater interactions, 51 migration, Solway Basin, 60-1, 62 thorium interactions, 56, 59, 60 hydrothermal alteration, Ireland, 26, 30-1, 33 Iapetus Suture, Ireland, 276, 277 Iberian Pyrite Belt, catastrophism, 153-60 Ibrahimispores brevispinosus, 350 ichthyoliths, 380 illite, 32, 45, 48 inclined heterolithic stratification (IHS), 189, 191-2, 199-200

455

Indian Wells Canyon, 119, 120 'Inishannon Limestones', 335 intraclasts, 227, 232, 233, 247, 260, 322 Ireland Asbian carbonate ramp, 253-61 biostratigraphy, 333, 335, 336, 339, 371-80 brachiopod facies, 427-35 carbonates and shales, 195-6 Dinantian platform carbonates, 239-249 exposure surfaces, 295, 296 lithostratigraphy, 243-7 mineral deposits, 1-21, 23-34, 264 palynofacies analysis, 437-48 fiver systems and coastal sedimentation, 183-205 sedimentation controls, 263-77 see also Kingscourt Outlier; Navan; South Munster Basin Irish Midlands, 371 metal deposits, 3-4, 4-5, 6-10, 18, 23-4, 26 iron Derbyshire Platform, 42, 45, 48 Ireland, 10 iron sulphides, Ireland, 14, 32 issinellid, 375

Janischewskina, 380 Jazwiny Hill, 318, 319 Jedrzejow High, 325 Junceum Limestone, 292 Kamaena, 130, 135, 139, 140, 225, 228-9, 232, 366 Kamaenella Ireland, 244, 246, 247, 248 Kingscourt, 130, 135, 136-7, 139, 142 Urswick Limestone Formation, 225, 226-7, 228-9, 230-1,232, 233, 235 kamaenids, 133, 135, 138-9 Karakuba Horizon, 361,362 karstification, 282, 289, 290-1,294, 295, 298 Kasachstanodiscinae, 379 Keel, ore deposits, 3, 10, 11, 15 Keel Fault, 10 Keele Formation, 298 Kentstown Block, 269, 271 Khaninian Stage, Russian Platform, 361,364 Kilbeg Wood, fauna, 430, 431,433, 434 Kilbride Limestone Formation, 266 Kilcummin Head, 185, 200, 202 Killala Oolite Member, 201 Killoran Fault, 14, 24, 26, 28-9, 30, 31, 32 Kiltiernan Townland, fauna, 430, 431,433 Kiltullagh Bridge, fauna, 427, 430, 431 Kingscourt, biostratigraphy, 380 Kingscourt Block, 267, 270 Kingscourt Outlier, 266, 269 Visban buildups, 127-42 Kinsale Formation, 333, 334, 338-9, 340 Kinsale sub-basin, 333-9 Kirkbean Outlier, 164, 165, 170-81 bitumen nodules, 52, 54, 55, 56, 58, 59, 60, 62 Kirkbya, 91 Kizel (-ovsky) Horizon, 361,362 (368, 378) Knipe Scar Limestone, 287, 292

456

INDEX

Knockfeerina volcanic centre, 276 Knockmore Limestone Member, 256, 259, 260, 261 Knockroe Formation, 274, 275 Knockseefin Formation, 274 Knoxisporites literatus Zone, Russian Platform, 363 Koninckophyllum vaughani, 254, 255, 257 Koninckopora Ireland, 375, 380 Ardagh buildup, 130, 133, 134-5, 136-7, 139, 140 Urswick Limestone Formation, 227, 229, 232, 233 Koninckopora minuta, 379 Koninckopora sp.B, 243, 244, 246, 248, 249 Koninckopora tenuirarnosa, 379 Koskinobigenerina, 241,257 Koskinobigenerina breviseptata, 243 Koskinotextularia, 379 Kosva (Kosvinsky) Horizon, 361,362 (368, 378) Kozhimian Stage, Russian Platform, 361,364 Krotovia, 430, 431,435 Kulikia, 224-5, 375 Lackagh Sandstone Formation, 438,440, 441,443,445 Lahn-Bode Zone, Germany, 307 Lake District see Cumbria; Urswick Limestone Formation Lake Valley Formation, New Mexico mounds and level-bottom beds, 84, 85-91, 92, 93 see also Alamogordo Member; Sacramento Mountains Lake Valley ramp, 99 laminar calcretes, 284, 285, 290, 293, 294, 298 Laminated Beds, Navan, 208, 211, 212, 213, 215, 216, 217 Langholm, 179, 181 lanthanum, Derbyshire Platform, 39, 42, 43, 45, 48 Largysillagh Sandstone Formation, 185, 186-8, 189, 192 Lasiodiscidae, 225, 231,232 Lask Edge Fault, 354 Latiproductus, Scotland, 433 Latiproductus latissimus, 241 lavas, 275-76 lead Derbyshire Platform, 39, 42, 43, 45, 47, 48 Ireland, 32 lead-zinc deposits, Ireland, 10, 11,207, 217 lead-zinc-silver deposits, Ireland, 23-33, 264 Leaper's Wood Quarry, 233 Lechowek Beds, 316, 317 Legacurragh Gap, 258 Leinster Massif, 269, 277 Leitrim Group, palynofacies, 435-46 Leptagonia, 430, 431,433 Leptagonia analoga, 432, 433 level-bottom beds, New Mexico, 85-6, 88, 89-91, 93, 125 lime mud, 86, 139 Limekiln Lake, 353 Limekiln Quarry, 351 Limerick Limestone Formation, 26, 266, 273, 274, 276 Limerick Province, metal deposits, 3 Limerick Syncline, 243, 266, 273, 274-6, 277 limestone breccias, Holy Cross Mountains, 322

linoproductid brachiopods, 432 Lotoprotonia ashfellensis, Ferguson, 431 Liscarrol Limestone Formation, 245 Liscartan Formation, 266, 269 Lisduff Oolite Member, 24, 26-7, 31, 32 Lisheen metal deposits, 1, 3, 14, 15, 18, 23-34, 26 volcanic rocks, 272 Lismalin Member, 246 Lispatrick Formation, 335, 336, 337, 340 Lissodus, 418, 422 Lissodus expansa Zone, Urals, 419 Lissodus pectinatus sp.nov, 389, 390, 41 I, 413 'Listracanthus'-type, scales, 393 Lithostrotion, 133, 172, 175, 246, 319, 379, 435 Lithostrotion araneum, 232, 377 Lithostrotion maccoyanum, 241,377 Lithostrotion vorticale, 377, 433 Lithostrotionidae, 322 Little Island Formation, 335, 338 Little Wenlock volcanic rocks, 356 Lochenling Formation, bitumen nodules, 53, 56, 58, 60, 61 Lochreia nodosus/nodosa, 336, 380 Loeblichia, 373 Loeblichia paraamminoides, 380 Loeblichiidae, 227 Loggerheads Limestone, 292 Lonsdaleia duplicata, 241 Lophoctenium, 155 Los Santos de Maimona, coral reef-facies, 145-51 Lough Allen Basin, 261 Lough Carra, fauna, 427, 430, 431,432, 434 Lough Corrib, fauna, 427, 429, 434 Lough Mask, fauna, 427, 429, 430, 431,434 Loughbeg Formation, 337-8 Loughshinny Formation, 138, 140, 271,272 Lower Alum Shale (Liegende Alaunschiefer), 306, 307, 311 Lower Border Group, 181 Lower Calcarenite Member, 26 Lower Crinoidal Limestones, 245 Lower Limestone Shales, 331 Lucan Formation, 4, 269, 270, 271 Luteotubulus licis, 375, 376, 377 Lycospora, 445 McCoss constructions, 17 Magcobar, barytes deposit, 3, 10-12 magnetic anomalies Apedale, 346, 347-8 Ireland, 5-6, 10, 12 Malahide Limestone Formation, 265, 266, 269 Malevka(-evsky) Horizon, 361,362, 366 Malevsky-Upinsky Horizon, 368 Malyovka Formation, fish assemblages, 387, 390, 392, 394, 396, 399, 400, 402, 408, 409-11,413 Mametella, 87, 88, 91, 115, 118-19, 123, 375 manganese, Derbyshire Platform, 42, 45, 48 marcasite, Ireland, 32 marginifera Zone, Holy Cross Mountains, 322 Market Drayton Horst, 350, 354 Maryport-Stublick Fault System, 165

INDEX matrix peloids, in buildups, 68, 69, 70, 72 Mayo, 183, 185, 197, 198, 431,432 Meath Formation, 266, 269 Mediocris mediocris, 257 Meenymore Formation, 257, 258, 295, 380, 447 megabreccias, Muleshoe Mound, 105-6 Mellon House Formation, 26, 272-3 Menaspiformes Obruchev, 406-7 Mendipsia leesi, 375 Mesodmodus teeth, 419, 424 Mestognathus bipluti, 380 Mestognathus praebeckmanni-beckmanni, 379 metal deposits and mineralization Derbyshire Platform, 35-49 Ireland, 1-21, 23-34, 207, 264 metal-organic interactions, Solway Basin, 51, 54-63 Micrite Unit, Navan, 209, 210, 211,212, 213, 215, 216 'microatoll' structures, southwest Spain, 148, 150 microbial origin, mud-dominated buildups, 67-77 microbial peloids, Sacramento Mountains, 87, 88, 89 Microconglomerate, Navan, 209, 215 microfacies associations Urswick Limestone Formation, 222-4, 233-6 allochem distribution, 222-5 stratigraphic controls, 233 microproblematica, 225, 232, 375, 376, 377 Middle Coal Measures, North Staffordshire, 345 Midlands Microcraton, 345, 346, 347, 350 Midlands Province (Ireland), metal deposits, 3-4, 4-5, 6-18 Mikhailov Horizon, fish assemblages, 385, 392, 399, 400, 402, 404, 405, 406, 407, 408, 411-12, 413 Mikhailov(-sky) Horizon, 362, 363, 364, 380 Milverton Group, 269, 270, 271 mineral deposits see metal deposits and mineralization mineralizing fluids conduits for, Ireland, 2, 33 Derbyshire Platform, 36-7, 46, 47, 48 Minnaun Formation, 185, 193-4 Minythyra, 430, 431 miospores, 372, 373, 374, 441,442-3, 445, 446 Mississippi Valley-type mineralization, 36, 39, 40, 41 Mississippian see Early Mississippian Moathill Formation, 4, 266, 376, 377 Moelfre Limestone, 292 Mold, exposure surfaces, 284 molluscs, 115, 117, 119, 123, 172, 430 shells, 87, 88 moravamminids, 87, 88, 112, 122 Moscow Syneclise biostratigraphy, 368 cyclic stratigraphy, 359-64 fish assemblages see East European Platform mound v. ramp-related biotic assemblages, 123, 124 mounds, mud see buildups (mounds); mud-mounds Moyny Limestone Formation, 194-5, 197, 198 Moyvoughly, 3, 15, 216 Muckros Head, 184, 191 mud generation, 73, 77 mud-mounds, 65, 67, 70, 74, 127, 266 formation, 323 see also buildups (mounds)

457

Muddy Limestone, Navan, 208-9, 210, 215, 216, 217, 429 Muleshoe Mound, 85, 86, 87, 89, 92, 97-108 Mullaghfin Formation, 128, 129, 141,270 Mullaghmore Sandstone Formation, 185, 196-204, 205, 255 Multithecopora, 141,322 Mungret Limestone Formation, 273 Munster Basin, 331 Namurian buildups, 140 Nanicella ex gr galloway, 320 Navafria, coral reef-facies, 146, 147, 150 Navan, 66, 67, 266 emersion surfaces and channels, 211-17 lithostratigraphy and sedimentology, 207-11 metal deposits, 1, 3, 7-10, 18, 207 Navan Group, 4, 208, 269 Neoarchaediscus, 241,379 Neoarchaediscus incertus, 243, 249 Neoarchaediscus stellatus, 243, 249 Neoarchaediscus subzone, 241 Neoseptaglornospiranella spp., 367, 368 Nereites, 155 Nevillea dytica, 243 New Mexico buildups, 83-95, 99-108 see also Alamogordo Member; Sacramento Mountains Newcastle Formation, 298 Newtown Cashel, metal deposits, 3, 10, 11 Nida Platform, 325, 326, 327 Nida Trough, 323 Nodasperodiscus sp., 257 Nodosarchadiscus sp., 257 Nodular Micrite Unit, 26 non-ferroan calcite cements, 214 non-saccate miospores, 441,442-3, 445, 446 non-skeletal buildups, 65-77, 133-7 non-skeletal microbialite structures, 141 Nooks Farm Borehole, 352 North Caspian Syneclise, 379 North Mayo succession, 192-5, 197 North Munster Shelf, 331,341 North Solway Fault, 181 North Solway-Water Beck--Gilnockie Fault System, 165 North Staffordshire see Apedale North Staffordshire Basin, 350, 354 North Wales, exposure surfaces, 284, 285, 287, 291, 292 North Wales Shelf, 282, 294, 295 North West Province, metal deposits, 4 Northern Urals see East European Platform fish assemblages Northwest Basin, 277 northwest Ireland see Ireland Nunn Member, 113, 121 mounds and level-bottom beds, 84, 86-91, 92, 93, 97 Oakport Limestone Formation, 27I oil migration studies, Solway Basin, 51-63 Okian Stage, Russian Platform, 363, 364

458

INDEX

Old Head of Kinsale, 335, 337, 339 Old Head Sandstone Formation, 333, 339-40 Old Red Sandstone Solway Basin, 170 South Munster Basin, 331,332, 337 Omphalotis ex gr. chariessa, 368 oncoids, 130, 133, 225, 227, 323 onychodontid sarcopterygians, 417 ooids, 146, 227, 228-9,~233, 235 Opatkowice Platform, 325 Oracanthus vetustus Leidy, 406-7, 412 orbital-forcing and climatic changes, 296-7, 308 organic compounds, metals and, 45-6, 48 organic matter, 437, 440-43, 445-47 Orionastraea, 141,377, 380 Orodontida Zangerl, 392-5 Orodontidae De Koninck, 392-5 orodontids, 420, 422, 424 Orodus, 420, 425 'Orodus' tumidus, 390, 394-7, 409, 411 Orroland Group, 52, 53, 165, 166, 167 Orroland Limestone Beds, 165 Orrolond Lodge Formation, 168, 180 orthoceratid, 173 orthotetidine, 430 Ortonella, 130, 133, 137, 140, 227, 232 Osteichthyes, 407-8 Osteolepididae, 407-8, 409-10, 411,413 Osteolepiformes, 407-8, 409 ostracodes Holy Cross Mountains, 320, 321 Ireland, 189, 194, 258, 372, 375, 380 Ardagh buildup, 134-5 Russian Platform, 365, 366, 410 Sacramento Mountains, 87, 88, 90, 91 Alamogordo Member, 115, 117, 119, 123, 124 Solway Basin, 173, 174 southwest Spain, 146, 149, 151 Urswick Limestone Formation, 225, 226-7, 228-9, 230-1 Ostrowka Quarry, 315, 316-18, 323 Oxwich Head Limestone, 222, 287, 292 Ozawainellidae, 227, 231

Pachysphaerina, 320 Palaeoaplysina buildups, 77 palaeobathymetric criteria, 124 palaeoberesellid algae, 225, 229, 246, 249 palaeoceanography, Culm Basin, 311-12 palaeoclimatic cyclicity, 295-8 palaeoenvironmental indicators, 229 palaeogeography Dinantian, Ireland, 239-40, 249, 264 Holy Cross Mountains, 323-6, 327 palaeokarst/calcrete relationship, 284-8 palaeoniscids, 417, 420 Palaeophycus, 156 Palaeospiroplectammina, 378 Palaeospiroplectammina tchernyshinensis, 367, 368 Palaeotextularia ex. gr. longiseptata, 241,243 Palaeothethys, 303, 311-12 palaeotextulariids, 241,243, 249, 377, 379

Palaeozoic, Upper, Germany, 303 Palastraea regia, 241,377, 380 Pale Beds, Ireland, 8, 208, 209, 211,215, 216, 217 metal deposits, 4, 7, 10, 15 Pale Chert Formation (Helle Kieselschiefer), 306, 307, 310, 311-12 Palmatolepis expansa conodont Zone, Urals, 417, 418, 421 Palmatolepis postera-Palmatolepsis expansa conodont Zone, Urals, 417 palynofacies, Glenade Sandstone Formation, 437, 438-48 Pandorinellina nota, 410 Paraarchaediscus, 379 Parachaetetes johnsoni, 147, 151 Paraendothyra, 378 Paragnathodus mononodosus, 380 Paragnathodus multinodosus Zone, East European Platform, 412 Paragnathodus nodosus, 380, 412 Paragnathodus nodosus-Gnathodus bilineatus bollandensis-Adetognathus unicornis conodont Zone, Urals, 420 Parsonage Beds, 273 Patrognathus, 374 Patrognathus andersoni Zone, Russian Platform, 361, 411 Patrognathus crassus, 361,410 Patrognathus crassus conodont Zone, East European Platform, 409 Patrognathus variabilis, 361,409, 410 pectinid bivalves, 233 pedunculate brachiopods, 435 pedunculate orthides, 430 peloidal micrites, in buildups, 66, 67, 68, 69, 70-1, 72, 74, 75, 76, 77 peloids Sacramento Mountains, 87, 88, 89 Urswick Limestone Formation, 225, 226-7, 228-9, 230-1,232 'Periplectrodus warreni'-type scales, 393, 394 Permian buildups, 77 Petalodontida Zangeri, 405-6, 412 Petalodontidae, 405-6, 412, 413 petalodontids, 420, 425 peyssonelid, 139 Phanerozoic chert sequences, Culm Basin, 310 ?Phoebodontidae Williams, 399-400 phoebodonts, 425 Phoebodus australiensis Long, 415 Phoebodus gothicus Ginter, 417 Phoebodus limpidus Ginter, 417 Phoebodus sp., 417 Phoebodus turnerae Ginter & Ivanov, 417 photic zone delineation, 123 Phyllite-Quartzite (PQ) Group (Iberia), 153, 155, 159 phylloid algae, 135 phylloid algal buildups, 75, 76, 77 Phylum Vertebrata, 385 Pig's Cove Member, 333 placoderms-ptyctodonts, 415 Planolites, 156

INDEX platforms, sedimentation controls, Ireland, 263, 269-72, 274, 276-7 Plectogyranopsis, 369 Pleuropugnoides pleurodon, 430, 431,433 Pleurosiphonella, 149, 151 Plicochonetes, 430 plurilocular foraminifers, 87, 88, 115, 117, 119-120, 123, 124, 259 Poecilodus aft. cestriensis, 405 Poecilodusjonesii (M'Coy), 404-5, 411 Poecilodus sp., 405, 411 Pojarkovella nibelis, 379 Poland, carbonate platform reconstruction, 315-27 polygnathids, 375 Polygnathus, 373, 374 Polygnathus communis carina, 339 Polygnathus communis carina Biozones, 335, 337, 340, 378, 419, 424 Polygnathus inornatus, 336 Polygnathus mehli Biozone, Ireland, 378 Polygnathus parapetus, 410, 411 Polymorphocodium, 227, 232 Portling Outlier, bitumen nodules, 52, 53-4, 55, 56, 57, 58, 59-60, 61, 62 Potholes Limestone, 292 Powillimount Beds, 170, 172-4, 180 bitumen nodules, 54, 55, 56, 59 praesulcata Zone, Holy Cross Mountains, 320, 322 Pre-Caspian Depression, 365-7 problematica, 225, 372, 380 Prochernyshinella disputabilis-T, beata Zone, Russian Platform, 361 productoids, 435 protacrodontids, 417, 419 Protacrodus aequalis sp. nov., 419, 420-21,422, 423 Psammodontidae De Koninck, 405 Psammodontidae indet, 405, 412 Psammodontiformes Obruchev, 405-6 Psammodus tooth plates, 419, 420 Psephodontidae Zangerl, 402 Psephodus, 403,419-20, 425 Psephodus dentatus, 402, 403, 411 ?Psephodus sp., 402-3, 411, 412 Pseudoammodiscus, 369 Pseudochaetetes, 147, 151 Pseudoendothyra, 380 Pseudoglomospira, 369 Pseudolituotuba, 373 Pseudolituotuba gravata, 323, 324 Pseudolituotubella, 369, 379 Pseudomonas atlantica, 72 Pseudoplanoendothyra, 368 pseudopolygnathids, 375 Pseudopolygnathus, 373 Pseudopolygnathus rnultistriatus Biozone, Ireland, 378 Pseudozaphrentoides juddi, 257 Punctospirifer, 430, 431,434 Pustilatisporites uncatus, Zone, Russian Platform, 361 Pycnoctenion aft, siberiacus, 408, 409, 411,412 pyrite Derbyshire Platform, 40 Ireland, 10, 32

459

'Quartz Marker', Navan limestones, 212 Quasiendothyra spp., 366, 367 Radaevka(-aevsky) Horizon, 361-2, 363, 369, 379 radiolarian cherts, Culm Basin, 308-9 radiolarians, 311, 312 Radiosphaera, 320 rainfall regimes and calcrete formation, 288, 290 ramose bryozoans, 87, 88, 89, 115, 116, 117, 119 ramp-related v. mound biotic assemblages, 123, 124 ramps biotic gradients, 111-12, 114-25 growth, Ireland, 253-61 sedimentation controls, Ireland, 263, 263-9, 272-4, 276-7 southwest Britain, 74 tectonic control, Ireland, 253-61,263-73 Rascarrel Group, 165 bitumen nodules, 53, 56, 60 Rathdowney Trend, 23, 24, 26, 28, 29, 32, 33 Rathkeale Beds (Formation), 273 red algae, 141 Alamogordo Member, 119, 120, 123, 124 Holy Cross Mountains, 323, 324 Ireland, 133, 135, 139, 142, 241,375 southwest Spain, 146, 147, 151 Urswick Limestone Formation, 232 Red Beds, Navan, 208, 217 Red Rock Fault, 347, 348 Red Wharf Cherty Limestone, 292 reef-building organisms, 145, 146, 151 reef-framework buildups, transition to, 142 reefs arrested development, Phanzeroic, 73 construction, hiatus in, 140 definitions of, 148, 150 southwest Spain, 145-51 true, 65, 145, 148, 150 Reenydonagan Formation, 333, 340-1 regional dolomitization, 29-30, 33 Renalcis, 67, 71, 72, 140 Rerrick Outlier, 164, 165, 166-70 bitumen nodules, 52-3, 56, 60, 61, 62 Rerrick succession, 180 Rhenish Massif, 303, 304, 305, 306, 307, 310 Rhipidomella michelini, 430, 431,432, 433 rhizocretions, 284, 286, 288, 289, 290, 291,292, 293, 294, 295, 298 Rhizodontidae Traquair, 408 Rhizodopsidae Berg, 408 rhodophytes, 135, 139 rhynchonellides, 430 rhythmic (°discyclic') bedding, Culm Basin, 307-8

Richterina latior-Pseudoleperditia venulosa-Shivaella microphtalma Zone, East European Platform, 408 Ringabella Limestone, 339, 340 Ringabella Syncline, 334, 335 Ringmoylan Shale Formation, 26, 272-3 Rinn Point, 185, 191 Rinn Point Limestone Formation, 185, 187, 190, 191 river systems northwest Ireland, 183-205 Roelough Conglomerate Formation, 183, 185-6, 187, 192

460

INDEX

Rough Limestone, 292 Rugose Coral Zones, Ireland, 379, 380 rugose corals, 141 Holy Cross Mountains, 319, 322 Ireland, 246, 372, 377, 379 Cuilcagh Mountains, 258, 259, 260 Kingscourt, 133, 137, 139, 141 southwest Spain, 145, 146, 149, 151 Rugosochonetes, 430, 431,432, 433 Rugosochonetes celticus, 432, 433 Rush Conglomerate, 269 Russian Platform biostratigraphy, 365-9 cyclic stratigraphy, 359-64 see also East European Platform Saccamina Limestone, 292

Saccamminopsis, 224, 225, 232, 235 Saccamminopsis sp., 243, 249, 380 Sacramento Mountains, 84-5 buildups (mounds), 83-4, 85, 88-9, 92, 93, 97-108 biota, 87-9, 90-1, 93 compared with European buildups, 91-3 level-bottom beds, 85-6, 89-90 biota, 88, 90-1, 93 see also Alamogordo Member Salebra sibirica, 375, 376, 377 salebrids, 115, 116, 117, 119, 122, 123, 124 sandbergi Zone, Holy Cross Mountains, 322 sarcopterygians, 425 Sarcopterygii: Dipnoi indet, 411 Scaliognathus, 373, 374 Scaliognathus anchoralis, 375 Scaliognathus anchoralis Biozone, Ireland, 337, 378 Scar Heugh Formation, 166, 167, 180 Schizophoria, 430, 431,435 Schizophoria resupinata, 432, 433 sclerophagous chondrichthyans, 425 sea-level changes, 304 and biodiversity, Ardagh buildup, 139 and cyclic emersion surfaces, Ireland, 215, 216, 217 and cyclicity of platform carbonates, Ireland, 248-9 and exposure surface development, 291,293, 295, 296, 298 and sedimentation Culm Basin, 311,312, 326 Ireland, 198, 202, 204, 269, 341 New Mexico, 124 northern England, 229, 233, 235, 236 sediment gravity-flow deposition, Iberian Pyrite Belt, 159 sediment stability, influenced by bacterial mats, 72 sedimentation Asbian/Brigantian boundary, Ireland, 247-8 climatic change and, Culm Basin, 311, 312 coastal, Ireland, 183-205 controls on, Ireland, 263-77 cyclicity, 282, 295-8 Urswick Limestone Formation, 221,222-31,233 Navan, 207-11 sea-level changes and, Culm Basin, 269, 311,312, 326 sediments, reworking, South Munster Basin, 331, 333-41

seismic pumping, 48 Semiplanus fragilis, 241 Septabrunsiina spp., 367

Septarinia leuchtenbergensis, 241 sequence biostratigraphy, Russian Platform, 365-9 serpulid worm tubes, 118, 121, 123 Seven Heads Peninsula, 338, 339 Shaley Pale Limestones, 208, 209, 217 shallow-water buildups, 67 Shalwy Formation, 184, 185, 186, 187, 188-92 Shalwy Point, 186, 191-2 Shanagolden Limestones, 273 Shannon Trough, 263, 265, 272-7, 274, 371 Shap Fell Quarry, 287 sharks, Urals, 423 Sheeps Head Anticline, 339, 340, 341 Shivaella microphtalma, 410 Shoreface deposits, 168-9 Shurinovkian Stage, Russian Platform, 361,364 siderite, Ireland, 32 silica, Derbyshire Platform, 40 Siliceous Limestone Formation (Kulm-Kieselkalke), 306, 307, 310, 312 siliceous rocks, Culm Basin, 303-12 Siliceous Transitional Formation (Kieselige 13bergangs-Schichten), 306, 307, 312 siliciclastic sediments, Navan, 213-16 silicification, Ireland, 31 silicified tephra, Culm Basin, 310-11,312 silver, Ireland, 3, 32 Silvermines, metal deposits, 1, 3, 10-12, 15, 17 Sinopora polonica Nowinski, 322, 324 Siphonodella, 374 Siphonodella Biozone, South Munster Basin, 335 Siphonodella duplicata Zone, Urals, 419, 421,422 Siphonodella praesulcata conodont Zone, Urals, 417 Siphonodella praesulcata-Siphonodella sulcata conodont Zones, Urals, 417, 419, 421

Siphonodella sp., 336 Siphonodella sulcata conodont Zone, Urals, 417, 418, 420, 421,422

Siphonodella sulcata-Siphondella duplicata conodont Zones, Urals, 419

Siphonodendron Ireland, 133, 141,246, 258, 259, 377, 379 southwest Spain, 145, 146, 149 Urswick Limestone Formation, 227, 230-1 Siphonodendron intermedium, 256, 257

Siphonodendron irregulare, 151 Siphonodendronjunceum, 241,377, 379 Siphonodendron Limestone, southwest Spain, 145, 146-8, 150, 151

Siphonodendron martini, 151,254, 256, 257, 377 Siphonodendron pauciradiale, 256, 257, 377, 379 Siphonodendron scaleberense, 254, 256, 377 Siphonodendron sociale, 241,254, 256, 377 Siphonophyllia benburbensis, 255, 256, 257 Siphonophyllia cylindrica, 377 Siphonophyllia garwoodi, 377, 430 Siphonophyllia siblyi, 133 skeletal stromatolites, in buildups, 71, 72, 73, 74 Skolithos, 155, 168, 169 Skreen Fault, 270

INDEX Slane Castle Formation, 4, 266, 272, 376, 377 Slievenaglasha Formation, 245, 427 Sligo, 197, 198, 199, 239 smectite, 45, 48, 283 smectitic soils, 290 Snetki-Pavlovskoye, fish assemblages, 396, 400 Solenopora, 135, 137 Solenoporaceae, 149, 324 solenoporids, 140, 147 Solway Basin, 51-4, 60, 163-5 depositional environments, 165-82 see also thoriferous bitumen nodules, Solway Basin south Cumbria/north Lancashire Shelf, 282 South Munster Basin, 18, 277, 331-2, 375 biostratigraphy, 333, 335, 336, 339, 372, 377 erosion and reworking, 331,333-41 metal deposits, 3, 4, 5, 17 South Portugese Zone, 153, 154 South Urals, chondrichthyans, 415-23 South Wales, exposure surfaces, 283, 285, 292, 293, 298 South Wales Shelf, 295 Southern Irish Platform, platform carbonates, 240 Southerness Beds, 54, 56, 165, 170-2, 174-5, 180 southwest Spain, coral reef-facies, 146-51 'sparmieritization', 72 Sphaerinvia, 87, 88, 115, 116, 117-18, 119, 122, 375, 376, 377 Sphaerinvia piai, 375, 376, 377 sphalerite Derbyshire Platform, 36, 39, 40, 41, 47, 48 Ireland, 10, 14, 32 spiculiferous micrite, Sacramento Mountains, 93 spiculitic cherts, Culm Basin, 310 Spinoendothrya, 368 Spinoendothrya costifera- Tuberendothrya tuberculata Zone, Russian Platform, 361 "Spirifer', 430, 431,433,435 Spiriferellina, 430, 431,434 Spiriferellina insculpta, 434 spiriferides, 430 Spirophyllum praecursor, 377 spirorbiid worm shells, 410 Spirorbis cf, helicteres, 173 sponges Ireland, 134-5, 137, 140, 141, 142, 258, 259 Sacramento Mountains, 87, 88, 89, 90, 91 Alamogordo Member, 115, 116, 117, 119, 120, 123, 124 Urswick Limestone Formation, 224, 226-7, 230-1, 232 Spotted Limestone, 292 Stacheia, 232 stacheiids (stacheiins), 115, 116, 117, 119, 120, 121, 123, 124, 139 Stacheoides, 134-5, 224, 225, 232, 259 Stainton Quarry, 226-7, 228-9, 230-1,233, 235 'Stemmatias bicristatus', 392 'Stemmatias simplex', 392 Stemmatias-like denticles, 417 'Stemmatiid' denticles, 391,392 Stethacanthidae Lund, 389-92

stethacanthids, 423 Stethacanthus, 417 Stethacanthus altonensis, 390-2, 409 Stethacanthus obtusus, 389-90, 409, 411 Stigmaria, 198, 363, 364 storm beds southwest Spain, 149, 150, 151 Urswick Limestone Formation, 231,233, 236 Streblodus cf. oblongus, 405, 411,413 Strepheoschema fouldensis White, 411 Strepheoschema sp., 410, 411 Streptorhynchus, 430 stromatactis in buildups, 73, 90, 102, 104, 133, 134-5, 138 Ireland, 133, 134-5, 138, 258, 259, 260 stromatolites, 71, 72-3, 74, 77, 140, 246 stromatoporoids, 320 strophomenides, 430 subsidence, effect on exposure surfaces, 294-5, 298 sulphide deposits Iberian Pyrite Belt, 153 Ireland, 4, 6, 8, 10, 14, 15, 23, 31-2 massive, 24, 26, 30, 31-2 sulphur, Derbyshire Platform, 36 Surtseyan eruptions, Shannon Trough, 274-5 Sychnoelasma hawbankense, 377 Sychnoelasma urbanowitschi, 377 Symmoriida Zangerl, 391-94 Symmoriidae Dean, 394 symmoriids, 417, 419 Symmorium, 418, 419, 422, 423-4, 425 Symmorium occidentalis, 394, 412 Syringopora, 149, 151,319 Syringopora reticulata, 209 syringoporoids, 145 Syringothyris cuspidata, 165 Syringothyris Limestone, Solway Basin, 165 Table Top Member, 97, 106 tabulate corals Holy Cross Mountains, 319, 322, 323 southwest Spain, 145, 146, 149, 151 Tabulipora, 135 Taeniolepis trautscholdi, 408, 409, 410 Tatestown, 10, 15, 215 tectonic control on exposure surface development, 295, 298 Holy Cross Mountains, 326 Ireland, 248, 253-61,269, 272, 276-7, 341 of mineralization, 47 North Staffordshire, 353 Solway Basin, 163, 180 Templemary Member, 245, 246, 247 tennantite, Ireland, 32 tephra, silicified, 310-11,312 terra rossa, 283, 290 Tetrataxidae, 225 Tetrataxis, 133, 136-7, 138-9, 140, 322, 378 Thalassinoides, 222 Thirlstane 'Reverse' Fault, 177 Thirlstane Sandstone, 170, 177, 180

461

462

INDEX

thoriferous bitumen nodules, Solway Basin, 51-2, 54-6, 57, 58-9, 62 dating, 59-60, 62 formation, 56 petrography, 54-6 thorite, Solway Basin, 55, 60 Thrinacodus, 418, 421,425 ?Thrinacodus sp., 399-400, 411 thrombolites, 71, 73, 77, 133, 140 tidal point bar deposits, northwest Ireland, 191 Tierra Blanca Member, 84, 86-91, 92, 93, 97, 99, 100 Tikhinella fringae, 320 Tikhinetla measpis, 320 Tournaisian buildups, northwestern Europe, 74, 83, 91-2 clasts, Holy Cross Mountains, 321-3 fish assemblages, East European Platform, 387-411, 412-13 sequence biostratigraphy, Russian Platform, 365-9 Tournayella discoidea, 378 Tournayellidae, 227 trace elements, release from calcite cements, 45-6, 47 Traeth Bychan Limestone, 292 transgressive cycles, Russian Platform, 359-64 transgressive surfaces, 202 transgressive systems, Urswick Limestone Formation, 233, 235 transtension, 1-18 trepostome bryozoa/bryozoans, 67, 247, 258, 259 trilobites Holy Cross Mountains, 320, 321,322 Ireland, 258, 259 Sacramento Mountains, 115, 117, 119 Scotland, 433 southwest Spain, 146, 151 Urswick Limestone Formation, 224, 232 Trim Borehole, 266 Trowbarrow Quarry, 233 Trwyn Dwlban, 297 Tuberendothyra tuberculata Lipina, 367, 368 tufts, 272, 274-76, 303, 306, 310, 345-7, 348-56 Tula Formation, fish assemblages, 390, 402, 403, 405, 411 Tula (Tulsky) Horizon, 362, 363, 380 Tulathyris vogdti, 410 Tumulispora malevkensis Palynozone, Russian Platform, 361 Tunstead Quarry, 283, 285, 287 Tylothyris, 430, 431 Tylothyris laminosa, 432 Tynagh, metal deposits, 3, 12-14, 15 typicus Zone, Lake Valley Formation, 84

Ungdarella Ireland, 241,243, 244, 246, 248, 249, 375 Ardagh buildup, 130, 133, 134-5, 136-7, 139, 140, 142 Urswick Limestone Formation, 225, 228-9, 230-1, 232 Upa Horizon, fish assemblages, 385, 388, 390, 394, 399, 400, 402, 403, 408, 409, 410, 411 Upa Horizon, 361,362 Upinsky-Karakubsky Horizon boundary, 368

Upper Calcarenite Member, 26 Upper Dark Limestones, Navan, 208, 210 Upper Wavy Laminated Unit, 28, 29, 33 Uralodiscus, 379 Urals see East European Platform; South Urals; Volga-Urals uranite radiometric dating method, 54, 60 uranium, Solway Basin, 55 Uriconian rocks, Apedale, 347, 348, 352, 356 Urswick Limestone Formation, 221-2, 287, 292 Asbian carbonates, 222-36 vadose cements, 284, 290-1 Vallatisporites pusillites, 361 Valvulinella, 322 Valvulinella sp., 257 Variscan basement, Germany, 303-12 vegetation, on exposure surfaces, 288-90, 292-3 Venev(-sky) Horizon, 363, 364, 380 vertisols, 290 Vicinesphaera, 368 Virgen de la Pea, 156 Visran bioherm frameworks, Ireland, 140-1 buildups bacterial, 67, 68, 69, 72, 74, 75 Ireland, 127-42 northwestern Europe, 83, 91-2 carbonates and shales, northwest Ireland, 195-6 clasts, Holy Cross Mountains, 322 coral reefs, southwest Spain, 141, 145-51 fish assemblages, East European Platform, 385-411 polymictic debrites, Holy Cross Mountains, 318-23, 325-7 sequence biostratigraphy, Russian Platform, 365-9 Viseidiscus, 379 Vissariotaxis, 135, 225, 230-1,235, 336, 380 Vissariotaxis compressa, 241,243 Vissariotaxis sp., 256, 257 volcanic rocks, 272,274-76, 303,306, 310, 347-8, 353-4 Volcanic-Siliceous Complex (VSC), 153-4, 155 volcanism, Shannon Trough, 274-6 Volga-Urals, sequence biostratigraphy, 365-7, 368 Voronezh Anteclise, 365 cyclic stratigraphy, 359-64 fish assemblages, 411,413 Wales, exposure surfaces, 283, 284, 285 Wall Hill Sandstone Group, 52, 53, 165, 166 Walterstown Block, 270 Walterstown Fault, 269, 270 Warnantella, 243 Warton Crag, 226-7, 228, 229, 230, 231-2 water depth indicators, 227, 235 Waulsortian buildups, 24, 26, 28, 29, 31, 32, 33, 66, 67, 73, 74, 75, 76, 77, 100, 208, 209-10, 217, 337 Europe, 92, 93, 111, 112 Ireland, 92, 260-1,268 New Mexico, 83-95, 97-108, 111, 112-13, 121-4, 140, 266 calcite, 30 phases, depth interpretations, 123-4

INDEX wave-dominated delta environments, Iberian Pyrite Belt, 157-8 wave-resistant structures, 140, 141, 145 Weberiphillipsia, 431 Welsh Marches, Uriconian rocks, 348 Wem-Red Rock Fault System, 355, 356 west Cumbria Shelf, 282 Western Anticline, Apedale, 345, 346, 347 western Ireland brachiopod facies, 427-35 see also Ireland Westphalian Zone, Germany, 307, 312 White Limestone, 292 Woodtown boreholes, 266, 376, 377 Woodtown Member, 266 woody material, 441-45

463

Xenacanthidae Fritsch, 39 xenacanthids, Russia, 412, 425 Xenacanthiformes Berg, 398-9 Xenacanthus ?nebraskensis Johnson, 399, 412, 420, 421, 425 Xenacanthus luedersensis Berman, 425 Yoredale cyclothems, Ireland, 239 zaphrentids, 258, 259, 435 Zareby Beds, 316, 317, 325 zinc, Derbyshire Platform, 39, 42, 43, 45, 47, 48 zinc-lead(-silver) deposits, Ireland, 1, 3, 10, 11, 23-34, 207, 217, 264 Zolotoy Verkh Ravine, fish assemblages, 394, 400 Zoophycus, 107, 201

Recent Advanc es in Lower Carboniferous Geology edited by P. Strogen, I. D. Somerville and G. LI. Jones (University College Dublin)

Rocks of Lower Carboniferous age are widely developed across Europe. Apart from their instrinsic geological interest, they are hosts to major Zn-Pb-Cu-Ba deposits in Ireland and Au-FeS2 deposits in the Iberian Pyrite Belt. Further, the Upper Palaeozoic rocks of Euro I~e are increasingly becoming the targetof oil and gas exploration. The wealth of data on Lower Carboniferous rocks,while not guaranteeing success, will be an invaluable tool for exploration. This book brings together in one volume advances over the last decade in several specialist subdisciplines of geology. It contains papers on carbonate and clastic sedimentology, palaeontology, palaeoecology, stratigraphy and biostratigraphy. • • • •

472 pages 283 illustrations international field of contributors index

Cover illustration: Basinal calciturbidites and shales of the late Vis~an Loughshinny Formation, Loughshinny Harbour, County Dublin. These northwards-verging chevron folds are thought to have originated as slump folds moving down the back-tilted footwall block of intrabasinal synsedimentary faults; later burial and Variscide deformation has tightened them. (Photograph by I. D. Somerville.)

ISBN 1 - 8 9 7 7 9 9 - 5 8 - 6

II 111ILl1

9 781897