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

SEDIMENTARY FACIES ANALYSIS

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

S P E C I A L P U B LI CA TI O N N U M B E R 22 I N T E R N A T I O NA L A S S OCIA T I O N

OF

THE

OF S ED I M E N T O L O GI S T S

Sedimentary Facies Analysis A TRIBUTE TO THE RESEARCH AND TEACHING OF HAROLD G. READING

EDITED BY A. GUY

b

Blackwell Science

PLINT

© 1995

The International Association

of Sedimentologists and published for them by Blackwell Science Ltd Editorial Offices: Osney Mead, Oxford OX2 OEL

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1988, without the prior permission of the copyright owner. First published 1995

Library of Congress Cataloging-in-Publication Data Sedimentary facies analysis: a tribute to

Set by Setrite Typesetters, Hong Kong

the research and teaching

Printed and bound in Great Britain

of Harold G. Reading/

at the Alden Press Limited,

edited by A. Guy Plint.

Oxford and Northampton

p.

em.

(Special publication number 22 of the International Association of Sedimentologists) Includes bibliographical references and index. ISBN 0-86542-898-0 1. Facies (Geology) 2. Sedimentation and deposition. II. Reading, H.G.

[[[.

I. Plint, A. Guy.

Series: Special publication

of the International Association of Sedimentologists; no. 22. QE651.S43 1995 552' .5-dc20

94-30445 CIP

To Harold

Photograph courtesy of Tim Barrett

We offer this collection of papers as a token of our appreciation for your friendship, guidance and inspiration. In remembering your infectious enthusiasm, dedication and sometimes daunting expectations, we realize how deeply we were influenced by your philosophy and attitude; a gift that has, in no small measure, shaped the course of our professional ]jves. Your former students

Contents

IX

Preface

XI

Harold G. Reading

xm

Introduction

Clastic Facies Analysis 3

Alluvial palaeogeography of the Guaritas depositional sequence of southern Brazil Paulo S. G. Paim

17

Sedimentology of a transgressive, estuarine sand complex: the Lower Cretaceous Woburn Sands (Lower Greensand), southern England Howard D. Johnson and Bruce K. Levell

47

An incised valley in the Cardium Formation at Ricinus, Alberta: reinterpretation as an estuary fill Roger G. Walker

75

Gravelly shoreface and beachface deposits Bruce S. Hart and A. Guy Plint

101

The return of 'The Fan That Never Was': Westphalian turbidite systems in the Variscan Culm Basin: Bude Formation (southwest England) Robert V. Burne

137

Depositional controls on iron formation associations in Canada Philip Fralick and Timothy J. Barrett

157

Facies models in volcanic terrains: time's arrow versus time's cycle Geoffrey J. Orton

Tectonics and Sedimentation 197

Coarse-grained lacustrine fan-delta deposits (Pororari Group) of the northwestern South Island, New Zealand: evidence for Mid-Cretaceous rifting Malcolm G. Laird VII

Contents

vm

219

Sedimentation and tectonics of a synrift succession: Upper Jurassic alluvial fans and palaeokarst at the late Cimmerian unconformity, western Cameros Basin, northern Spain Nigel H. Platt

237

The use of geochemical data in determining the provenance and tectonic setting of ancient sedimetary successions: the Kalvag Melange, western Norwegian Caledonides Rodmar Ravnas and Harald Fumes

265

Differential subsidence and preservation potential of shallow-water Tertiary sequences, northern Gulf Coast Basin, USA Marc B. Edwards

Sequence and Seismic Stratigraphy in Facies Analysis 285

Seismic-stratigraphical analysis of large-scale ridge-trough sedimentary structures in the Late Miocene to Early Pliocene of the central North Sea Joe Cartwright

305

Millstone Grit cyclicity revisited, II: sequence stratigraphy and sedimentary responses to changes of relative sea-level Ole J. Martinsen, John D. Collinson and Brian K. Holdsworth

Facies Analysis in Reservoir Sedimentology 331

Productive Middle East clastic oil and gas reservoirs: their depositional settings and origins of their hydrocarbons Ziad R. Beydoun

355

The evolution of Oligo-Miocene fluvial sand-body geometries and the effect on hydro­ carbon trapping: Widuri field, west Java Sea Ray Young, W.E. Harmony and Thomas Budiyento

381

Index

Preface

This book stands out in the series of Special Publi­

Bureau considered that this would be best done

cations of the International Association of Sedimen­

through a Special Publication on a subject in line

tologists.

It

is

an

acknowledgement

of Harold

with Harold's work (obvious topics were clastics,

Reading's commitment to lAS, for whom he has

facies and depositional environments, sedimentation

been Publications Secretary, General Secretary and

and tectonics).

President successively, over the last 30 years.

It is therefore most appropriate that Guy Plint,

Harold has not only been source and inspiration

as Editor chosen for this special publication, has

of many of the lAS policies and activities over this

brought together a collection of original scientific

time, he has also been at the roots of 'facies sedimen­

papers authored by Harold Reading's students, or

tology' as an art in itself, and as a major tool in the

students of theirs. To honour Harold Reading's own

broader field of geology.

scientific scope, the subject chosen is broad: sedi­

More than providing his own personal contribution

mentary facies analysis. The contributions contained

to this branch of the earth sciences, Harold created a

in this Special Publication show to what extent facies

flourishing school of teaching and research. Harold's

sedimentology, as fostered by Harold Reading, is

approach has burgeoned from Parks Road, Oxford,

now established as a necessary basis to any under­

to

standing of sedimentary rocks.

become

international

not

only

through

his

students, but also through 'his book'.

PETER HOMEWOOD

The Bureau of the lAS, taking up a suggestion by

/AS Publications Secretary

Robert Campbell of Blackwell Science, decided to put together a scientific tribute to Harold. The

IX

Harold G. Reading

Harold Reading was born in 1924 and, on leaving

at Rijswijk, arranged for Harold to investigate the

school, joined the Indian Army. This early experience

context of reported turbidites associated with English

left a lasting impression and undoubtedly contributed

Carboniferous deltaics in the Pennines and in south­

to Harold's later concern for international cooper­

west England.

ation. He went up to Oxford in 1948, initially to read

Cooperation with de Raaf and with Roger Walker,

Forestry, but his interests were diverted towards

one of Harold's earliest research students, developed

geology and he graduated in that subject in 1951.

a detailed appreciation of sedimentary structures

As an undergraduate, he visited North Norway to

and their role in understanding processes, and led

investigate the Late Precambrian and Cambrian

to the development of the style of facies analysis

stratigraphy of the Digermul Peninsula. This under­

exemplified by the 1965 classic paper on the Carbon­

graduate expedition not only shed significant new

iferous cycles of North Devon. Thereafter, Harold's

light on the stratigraphy of the area but also sowed

stable of research students grew rapidly as this volume

the seeds of a later rich sedimentological harvest.

amply testifies. Until his retirement, it was unusual

Three years at Durham under K.C. Dunham led to a

for him to have fewer than five or six doctoral

PhD with a project that involved mapping Carbon­

students at any one time, this in addition to a full

iferous Y oredale cycles across an area of bleak

undergraduate teaching programme and responsi­

Pennine moorland. Although the main thrust of the

bilities in college. This formidable work load was

study was stratigraphy and structure, the experience

carried out with great conscientiousness but Harold

of Carboniferous cyclicity was to set a further pointer

still had time to spare for external activities such

for the future.

as his involvement with lAS and JAPEC. During

On completion of his PhD Harold joined Royal

Harold's long career in Oxford, he only spent sus­

Dutch Shell and immediately found himself in the

tained periods away on sabbatical on two occasions,

contrasting field conditions of Venezuela. This multi­

the first in Leyden in the mid-1960s and the second

national, multidisciplinary environment developed

in Canada in 1972. The period in Holland led to

an appreciation of broader geological perspectives

close cooperation with structural geologists working

and the pragmatic, though rigorous, approach to

in the Cantabrians, an important extension of his

problem solving that has characterized Harold's

interests outside Britain.

career.

Of particular significance was a visit to

Harold's earliest students developed his early

Venezuela by Ph. Keunen who was, at that time,

interests, the Carboniferous deltas of Britain, and

actively promoting his pioneering work on turbidity

the tillites, shallow-marine and fluvial sediments of

currents

rigorous

northern Norway. Later, the Lower Palaeozoic of

approach to understanding depositional processes

Ireland and the Carboniferous of northern Spain

struck a chord with Harold, which was to be a

were added. As students were attracted to Oxford

and

their

deposits.

Kuenen's

cornerstone of his approach to sedimentology.

from around the world, the geographical spread

Harold returned to Oxford in 1957 as lecturer in

grew. However, geographical diversification was not

geology, a position that he held until retirement in

an end in itself but largely a result of Harold's

1991. When he took up his post, his teaching responsi­

curiosity about wider controls on sedimentation,

bilities included mapping and palaeontology and

particularly the role of tectonics. He understood

stratigraphy. Sedimentology, as we know it, hardly

very early the implications for sedimentology of

existed. Harold first revived his interests in northern

Plate Tectonics, as exemplified by his pioneering

Norway through a further, largely stratigraphical

paper with Andrew Mitchell. Curiosity about new

expedition to Digermul. Perhaps more importantly,

geological ideas and the need to investigate their

he developed his interest in sedimentary process and

implications for sedimentology and vice versa has

environments through a relationship with Shell.

been a hallmark of Harold's geological thinking. By some standards, Harold has not been a prolific

Maurits de Raaf, then Head of Geological Research

XI

Harold G. Reading

XII

author, although his papers are always thoughtful

President have already been acknowledged by the

and stimulating. Published evidence of Harold's

Association

influence lies mainly in the rigour, originality and

Membership to Harold. It is worth remembering

appreciation of the wider geological perspective that

that it was in no small measure due to Harold's

itself

in

the

granting

of

Honorary

characterize many of the publications of his research

efforts that the Association changed from a largely

students and of second and third generation students.

European organization to one of real international

Harold edited one Special Publication of the lAS

stature. Harold's tireless efforts to meet and encour­

on-strike-slip mobile belts, but his most valued publi­

age sedimentologists of all ages and backgrounds

cation is the textbook Sedimentary Environments and

around the world and his endless patience and

Facies, initially written largely by Harold's former

diplomatic skill have been well rewarded in the

students and rigorously edited to reflect the high

healthy Association that we enjoy today. Harold

standards he espouses. The 3rd edition currently

has additionally been honoured by the Geological

occupies much of his 'retirement'.

Society of London with the award of the Lyell Fund

Although this book is essentially a celebration

and the Prestwich Medal and, most recently, by

of Harold's scientific influence, it is important,

SEPM with the award of its prestigious Twenhofel

especially in a Special Publication of the lAS, to

Medal.

acknowledge

his

enormous

contribution

to

the

development of sedimentology internationally. His unstinting efforts on behalf of the lAS, as Publi­ cations Secretary,

as General Secretary and as

JOHN CoLLINSON Shrewsbury, UK

Introduction

This volume is a very personal compilation. Unlike

from Harold's former graduate students and their

previous lAS Special Publications, it is not centred

students and co-workers, but to impose no constraint

on a specific geological theme, and for that I make

on topic, in order to illustrate the scope of Harold's

no apology. Instead, my intent was to illustrate, and

knowledge, interest and vision. In consequence, the

celebrate, the breadth of interest, energy and inspi­

contents of this book are eclectic. The collection

ration that Harold Reading has brought to the field

of papers serves to highlight the power of facies

of sedimentary geology.

analysis, whether the rocks be volcanogenic, bio­

Few would deny the depth of Harold's influence

genic, siliciclastic,

or even 'catastrophic'

on sedimentology, world-wide. In part, this is due to

olistoliths!),

his publications, in particular the enormously suc­

method fostered by Harold.

cessful

Sedimentary

Environments

(mega­

the scientific

Facies,

It is particularly appropriate that, amongst the

unquestionably the cornerstone for all those who

contributions, Ole Martinsen, John Collinson and

embark on sedimentary facies

and

and of course reflect

analysis! Equally

Brian

important of course, has been his pivotal role in the

Holdsworth

offer

new

interpretations

of

Namurian deltaic rocks in the northern Pennines,

foundation and development of the lAS, a contri­

(upon which Harold cut his sedimentological teeth),

bution acknowledged recently with honourary mem­

but which, judging from referees comments, still

bership of that Association.

provide fuel for heated debate! In similar vein,

His philosophy and attitude has of course travelled

Bob Burne presents a review and discussion of the

with his graduate students, drawn from 13 countries

depositional environment of the enigmatic Bude

on six continents. Because many of these students

Formation

returned home upon completion of their work in

1960s), but which is still subject to sharply divergent

Oxford, and others now work and teach outside

interpretations.

the UK, the approach Harold fostered during their

Roger Walker shows how important it is, both to

(which In

Harold a

studied

salutory

in

lesson

the to

early

us

all,

graduate days has continued to spread. (He may not

separate facts from interpretations, and to ques­

know this, but in a geneological sense, Harold

tion one's cherished interpretation, when he boldly

is now a great-great grandfather to at least one

reinterprets as an incised valley fill, rocks he pro­

young sedimentology student who doubtless is quite

claimed a turbidite channel deposit just nine years

unaware of the history of the supervisory influence

ago! As Editor of this volume, I am indebted to the

that has been passed down!)

following people whose thorough reviews served

Although initially conceived as a thematic volume with contributions to be invited from a panoply of

to clarify the papers, and who made my job that

leading

much easier: Gail M. Ashley, Timothy R. Astin, T.

sedimentologists,

two difficulties quickly

arose: first, just what was to be the theme? As

Christopher

Harold has been involved in so many areas of

Charlie S. Bristow, H. Edward Clifton, Thomas C.

Baldwin,

sedimentary geology, selection of any one topic

Connally,

simply served to highlight gross neglect of another.

Peter G. DeCelles, Frank G. Ethridge, Jill Eyers,

Edward

Janok

Cotter,

P.

Bhattacharya,

William

R.

Dupre,

Secondly, it rapidly became apparent that numerous

Stephen S. Flint, Edward C. Freshney, Robert L.

former students were anxious to pay their own

Gawthorpe, Roland Goldring, Anthony J. Hamblin,

personal tribute, and whose contributions could,

Alan P. Heward, Phillip R. Hill, Richard N. Hiscott,

alone, easily constitute a hefty volume! Of the 34

Richard S.

students whom Harold guided through doctoral

McCabe, Kathleen M. Marsaglia, Franco Massari,

theses between 1961 and 1994, 16 have authored, or

Gerrard V. Middleton, Robert A. Morton, George

co-authored papers in this volume.

Postma, William C. Ross, Alastair H. Ruffell, Bruce

In keeping with the sentiment of this festschrift, I

Hyde, Elana L.

Leithold,

Peter J.

W. Sellwood, Gary A. Smith, Roger G. Walker,

took the decision to limit contributions to those

James

xiii

D. L.

White,

John

A.

Winchester

and

Introduction

XIV

Jonathon Wonham, plus two people who chose to

clandestine spmt of this project,

remain anonymous.

essential intelligence on both Harold and his former

I am very grateful to Susan Sternberg, Edward Wates and Julie Elliott at Blackwell Science who provided guidance at critical phases in the prep­ aration of this book. I also thank Diana Relton (Earth Sciences,

Oxford) who entered into the

graduate students.

A. GuY PuNT London, Ontario

and provided

Clastic Facies Analysis

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

Spec. Pubis int. Ass. Sediment.

(1995) 22, 3-16

Alluvial palaeogeography of the Guaritas depositional sequence of southern Brazil P A U L O S. G. P AlM* Earth Sciences Department , Oxford University, Parks Road, Oxford OXI 3PR, UK

ABSTRACT

The Guaritas sequence is the uppermost stratigraphical level of the Camaqua Basin (southern Brazil) and comprises an alluvial, deltaic and aeolian continental facies association up to 800 m thick. Facies mapping of this unit has revealed a lateral association of tributary fans and trunk braided rivers developed under semi-arid conditions. Two main regions (lobes) of alluvial fan development can be discriminated and the source points of both coincide with synforms in the nearby basement. This depositional system presents a normal down­ fan facies change. An anomalous lateral change of facies within the trunk river system is interpreted as having been inherited from pre-existing alluvial fan deposits. The main alluvial facies comprise trough cross-stratified (74%) and horizontally bedded (7%) sandstones, massive (16%) and tabular cross-stratified (2%) orthoconglomerates, and massive mud­ stones (1%) . Vertical aggradation of three-dimensional subaqueous dunes, followed by an upper flow regime plane-bed phase, characterized the depositional events of the sandy areas of the alluvial system. Diffuse gravel sheets and minor longitudinal and transverse bars were the main geomorphological features of the gravelly alluvial reaches. Fine-grained sediments represent temporarily abandoned areas within the braided channel network.

INTRODUCTION

trending tectonic structure, in southern Brazil (Fig. 1}, and evolved during the latest phases of the Brasilia no orogenic cycle (strike-slip basins of Brito­ Neves & Cordani (1991}}. An extensional or transtensional event at the end of the Brasiliano orogenic cycle, and the consequent formation of intermontane basins, has been pro­ posed as the tectonic setting of the Camaqua Basin during the deposition of the Guaritas sequence (Fragoso-Cesar et al., 1984, 1992; Beckel, 1990, 1992). In the past decade, the Guaritas depositional sequence has received attention from several authors in terms of facies analysis and palaeoenvironmental interpretation (Becker & Fernandes, 1982; Fragoso­ Cesar et al., 1984; Jost, 1984; Lavina et a!., 1985; Beckel, 1990). Generally, these papers have indi­ cated continental sedimentation characterized by

The Guaritas depositional sequence constitutes the uppermost unit of the Camaqua Basin infilling and it is an unconformity-bounded stratigraphical unit: it overlies older deformed molasse strata (angular unconformity) and is covered by Permian sedimen­ tary rocks of the Parana Basin. The Guaritas sequence, about 800 m thick, is almost always flat-lying, although, near to regional faults some extensional reactivation has tilted the Guaritas deposits. The available radiometric dating, summarized in Soliani et al. (1984) and Fragoso­ Cesar et al. (1984), indicates a Cambro-Ordovician age for the deposition of the Guaritas sequence. The Camaqua Basin is located in a NE-SW * Permanent address: UNISINOS- Departamento de Geologia, Av. Unisinos 950, Sao Leopoldo RS, Caixa Postal 275, CEP 93022-000, Brazil.

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

3

P.S.G. Paim

4

Permo Triassic

T �

BRAZIL

>-

� L:...:!l

Upper Vendian to Ordovician

t-':(:�: ., .. . .•

� >:.1 � t:;;:.t1

Mainly alluvial facies

Mainly volcanic rocks

Q o

Mainly eolian facies Pre-Guaritas basement

E:f=3-g

Mainly deltaic facies

LA: j

Permo Triassic

Fig. 2. Three-dimensional view of Camaqua Basin and surrounding area (same region of Fig. 1): (A) topography and (B) sketch of the Guaritas sequence facies.

7

Palaeogeography of the Guaritas sequence

Table 1. Classification and relative percentage of the sedimentary lithofacies (lithofacies code adapted from Miall (1977, 1978) and Rust (1978)) Rock type Conglomerates (G)

Sandstones (S)

Mudstones (F)

Facies code

Description

Percentage

Gm

Massive or crudely bedded conglomerates (cobbles, pebbles and granules)

l6

Gp

Small- to large-scale tabular cross-stratified conglomerates (granules and pebbles)

Gt

Small- to large-scale trough cross-stratified conglomerates

Gms

Massive, matrix-supported conglomerates (boulders to granules dispersed in a muddy sand matrix)

St

Small- to large-scale trough cross-stratified sandstones (pebbly to very fine-grained)

Sh

Horizontally bedded sandstones (medium to very fine-grained)

Sp

Medium to pebbly sandstone with small- to large­ scale planar cross-stratification

Fm

Massive mudstones with mudcracks

Fl

Laminated to rippled very fine sandstone to siltstone

Table 2. Relative percentage of trough cross-stratification and horizontal lamination in each sandy textural class Texture Pebbly to very coarse grained

Facies St

Sedimentary structures

Percentage

Small scale Medium scale Large scale

10 49 41

Sh Coarse to medium grained

St

0 Small scale Medium scale Large scale

10

Sh Fine to very fine grained

St

Sh

10 43 37

Small scale Medium scale Large scale

15 37 28 20

facies commonly occurs associated with facies Gm (Fig. 3). Facies Gt is rare, finer grained than facies Gm and Gp and characterized by small- to large-scale trough cross-stratification (alternations of small pebbles and gravelly sands). This facies interfingers with facies Gm and grades into facies St (Fig. 4). Matrix-supported conglomerates (facies Gms) are

2

74 7

also rare and occur, locally, near the eastern border of the Camaqua Basin. The main characteristic of this facies is its chaotic arrangement of pebbles, cobbles and, less commonly, boulders floating in a muddy to sandy matrix. Sandstones

Trough cross-stratification (facies St, 74% ), in places disrupted and/or deformed by convolution, and horizontal bedding (facies Sh, 7% ) are the main features of the alluvial sandy deposits (Figs 3 & 4). Planar cross-stratification (facies Sp) is rare. Facies St is characterized by very fine- to very coarse-grained sandstones with trough cross-bedding (Table 1). The cross-strata are predominantly of medium to large scale in all textural classes, but the proportion of small-scale trough cross-stratification increases as sandstones become finer grained (Table 2). This facies is the most common in the fining upward cycles. Convolute bedding is common in trough cross­ stratified sandstones (facies St). Within a single cross-stratified set, all gradations may occur from oversteep foresets, recumbent folding to intense deformation and even complete destruction of the former bedding (facies Sm and Spo of Bromley, 1991). Deformation near the top of the cross-

8

P.S.G. Paim

Fig. 3. Main alluvial lithofacies: facies Gt, St, Sh, Spo and, in the uppermost part of the picture, Fl, Gm and Gp. Bar scale is 2 m long.

stratified set is commonly characterized by downcur­ rent oversteepening of the cross-strata (Figs 3 & 4), and the intensity of convolution increases down the slip-face. Horizontal bedding (facies Sh) does not occur associated with pebbly and very coarse-grained sand­ stones and comprises 10% of the sedimentary struc­ tures of medium- to coarse-grained sandstones and 20% of the fine- to very fine-grained sandstones (Table 2). This facies is often related to the upper­ most parts of the fining upward alluvial cycles (Figs 3 & 4). Planar-tabular cross-stratified sandstones (facies Sp) are not common in the Guaritas sequence alluv­ ial deposits (Table 1). They occur as small- to large­ scale sets in pebbly to medium-grained sandstones and are normally interlayered with facies St. Other facies

Massive mudstones are rare and commonly mud­ cracks are their most conspicuous feature (facies Fm). Very fine-grained sandstones and siltstones (facies Fl) are also, and can be either horizontal (Fig. 3) or, more rarely, cross-laminated (Table 1). Both usually occur in the uppermost parts of the fining upward alluvial cycles. Alluvial facies: summary of general features and interpretations

The textural aspects (Table 1) suggest that the alluv­ ial facies of the Guaritas sequence represent bedload stream deposits in which the bedload was predomi-

nantly sandy and the suspension load, if deposited, was almost completely eroded by subsequent flood events. This type of stream commonly has a braided pattern characterized by low sinuosity and highly mobile channels (Collinson, 1986). The sheet-like geometry of the fining upward cycles enclosed by fifth-order bounding surfaces suggests broad, shallow channels. In terms of the gravelly facies, the dominance of clast-supported conglomerates (Table 1) is indicative of gravel deposition by strong tractive flows, whereas the finer grained material (sand and mud) was still being carried in suspension (Rust & Koster, 1984). Thin beds of facies Gm associated with laterally extensive channels suggest the development of dif­ fuse gravel sheets (Hein & Walker, 1977) by very extensive and shallow sheet-floods (Collinson, 1986). Thicker deposits of facies Gm suggest deeper and less ephemeral flows (Rust, 1978) causing more extensive vertical aggradation of gravel bars with low depositional dips. These deposits commonly have been associated with the development of longi­ tudinal and/or diagonal gravelly bars (Smith, 1970; Rust, 1972, 1978; Miall, 1977, 1978; Rust & Koster, 1984; Collinson, 1986) under high water and sedi­ ment discharge (Hein & Walker, 1977). Conglomerates with planar cross-stratification (facies Gp) has been related to (i) two-dimensional dune migration (transverse and/or linguoid gravel bars of Hein & Walker (1977), Miall (1977) and Middleton & Trujillo (1984)) as well as to (ii) later modifications of longitudinal bars (Smith, 1970; Rust, 1978; Enyon & Walker, 1974) in modern alluvial gravelly reaches. The frequent occurrence of

;;,o

!:)

�

�

-§ � -.:;, �

"'

C)

§

;::. s

"' "'

.E Fig. 4. Detailed view of Fig. 3 (enlargement of its lower part): facies Gt, St, Sh, Spo and thin tabular beds of Gm. Bar scale is 2 m long.

"' "' ;:s '"' "'

'.0

10

P.S.G. Paim

isolated sets of facies Gp within deposits of facies Gm could be explained more easily by the second hypothesis. Trough cross-stratified conglomerates are rare (facies Gt) and have been associated with (i) three­ dimensional dune migration, as observed by Fahnestock & Bradley (1973) and Galay & Neill (1967), and (ii) channel scour-and-fill structures (Miall, 1977; Middleton & Trujillo, 1984). The same criteria previously used to interpret facies Gp can also be applied in this case: the solitary nature of this facies suggests the deposition of gravel in depressions around diffuse gravel sheets. Matrix-supported conglomerates (facies Gms) are also rare and represent mud- and debris-flow deposits commonly associated with an alluvial fan setting (Blackwelder, 1928; Bull, 1963; Hooke, 1967; Rust & Koster, 1984; Collinson, 1986; Blair & MacPherson, 1992). Sandy sediments constitute the majority of the Guaritas alluvial deposits (Table 1) and are exten­ sively dominated by facies St (Table 2). Trough cross-stratified sandstones have been related almost invariably to migration of three-dimensional dunes (e.g. Collinson, 1970; Williams, 1971; Harms et al., 1975; Miall, 1977; Rust, 1978). In braided alluvial settings these bedforms usually have been associated with in-channel deposition (Cant & Walker, 1976, 1978; Cant, 1978; Walker & Cant, 1984). Such repetitive sand deposits commonly are considered as flood-stage bedforms (Williams, 1971) and are larger in deep channels (Cant, 1978). The association of facies St with the lower and middle part of sheet-like fining-upward cycles suggests this facies could be related to flood stage in shallow channels. Subcritical climbing trough cross­ strata (facies St) indicate subaqueous dune aggra­ dation. Sporadic lateral accretion of these bedforms is indicated by inclined planes (first-order bounding surfaces of Miall (1988)) dipping perpendicular to the dune migration direction (Paim, 1994). The absence of third-order surfaces (except the rare occurrence of lateral accretion surfaces) associ­ ated with the subcritical climbing of the trough cross-bedded sets (facies St) suggests rapid depo­ sition of a sandy load, transported by traction plus suspension, without macroform (sensu Jackson, 1975) development. Deformation of trough cross-stratified sandstones is a very conspicuous feature of the Guaritas sandy alluvial facies. Recumbent folding in cross-bedded sandstones commonly has been attributed to shear

stress acting on a liquefied sand bed and caused by current drag (Allen & Banks, 1972; Doe & Dott, 1980; Owen, 1987) or by the movement of large bedforms over an unconsolidated substrate during high-flow stages (Plint, 1983). Horizontal bedding (facies Sh) occurs most often in the finest fraction of the sandy deposits (Table 2). This textural control, associated with the occurrence of parting lineation and scattered small pebbles and granules near the base of the horizontally bedded sets, indicates its origin as an upper flow regime bedform. Deposits with the same characteristics of facies Sh usually have been linked to an upper flow regime phase developed during flood stages on the channel floor (McKee et al., 1967; Williams, 1971; Miall, 1977) or under the influence of high-velocity and low depth flows on the top of sand-flats (Cant and Walker, 1978; Miall, 1977; Collinson, 1986). The common occurrence of this facies (Sh) on the upper­ most parts of the fining upward cycles supports an interpretation involving upper flow regime currents reworking the top of the previous alluvial deposits. Planar-tabular cross-stratified sandstones (facies Sp) are rare. Within alluvial settings this facies commonly has been related to slip-face advance of two-dimensional dunes (transverse -linguoid or lobate bars of Collinson (1970, 1986), Smith (1970), Williams (1971), Asquith & Cramer (1975), Miall (1977), Cant & Walker (1978) and Cant (1978); or sand waves and straight-crested megaripples of Smith (1970), Collinson (1986) and Miall (1978)). Smith (1970) related the origin of the two­ dimensional dunes to the development of 'deltas' in pre-existing channel-floor depressions, whereas Cant & Walker (1978) related them to flow expansion at channel junctions or places where the channels widen. Facies Fm and Fl are not common in the alluvial system of the Guaritas sequence (Table 1). Their rarity and generally lenticular geometry (Fig. 3) are suggestive of waning flood deposits settling on to temporarily abandoned areas of the braided system (Cant, 1978; Cant & Walker, 1978; Miall, 1978). In general, diffuse gravel sheets and longitudinal/ diagonal bars were the main geomorphological elements of the gravelly reaches, whereas sub­ aqueous three-dimensional dunes characterized the sandy portions of the Guaritas alluvial system. The predominance of vertical aggradation of dunes instead of downstream and/or lateral accretion of more stable sandy accumulations (e.g. sand-flats)

11

Palaeogeography of the Guaritas sequence suggests a highly variable hydrological character and predominance of the upper part of lower flow regime conditions within the channels. Debris-flow and sheet-flood deposits suggest the presence of alluvial fans within the alluvial system as well as flashy discharge due to sporadic, but torren­ tial, rainy seasons. The above interpretations together suggest an alluvial drainage developed under semi-arid con­ ditions (large discharge fluctuations) with alter­ nation of flood events and dry seasons. These conclusions are reinforced by the aeolian associ­ ation (Lavina et al., 198S) and by petrographical evidence related to early diagenetic processes (De Ros et a/., 1994).

ALLUVIAL

FACIES:

LATERAL CHANGES

The previous section describes the pattern of alluvial sedimentation in terms of mean regional values and, in this way, reflects the major features of the alluvial deposit. In the following section, spatial variation in some sedimentary features is described and, when possible, interpreted. To achieve this, the mean values, per unit area, of several sedimentary par­ ameters were calculated using the outcrop locations and grid presented in Fig. SA.

Palaeocurrent pattern

The pattern of sediment transport within the entire Camaqua Basin was calculated using the grid and outcrops shown in Fig. SA. In order to eliminate problems associated with the analysis of several types and scales of sedimentary features (Miall, 1977) mean vectors were calculated only from trough cross-stratification. By using only one rank of sedimentary features, difficulties related to vector magnitude were eliminated (Allen, 1963; Miall, 1974). In addition, dunes seem to be associ­ ated with high-stage flow and, consequently, should be good indicators of the true downstream direction (Miall, 1977). The distribution of the palaeocurrent vector means (Fig. SB) indicates two major dispersal compart­ ments within the alluvial system: 1 from the eastern border to the basin axis the sedimentary transport was almost perpendicular to the regional tectonic trend (a general mean vector of

282°, with a correlation coefficient of 0. 86), reflecting a sedimentary input towards the basin axis; 2 from the basin axis to the western border, palaeo­ currents were predominantly parallel to the struc­ tural trend (general mean vector of 211 with a correlation coefficient of 0.96). o,

Pattern of textural dispersion

The alluvial deposits of the Guaritas sequence are composed primarily of sandstones (mainly facies St and Sh), minor conglomerates (mainly facies Gm and Gp) and trace amounts of fine-grained sediments (facies Fm and Fl), as has been described in the previous section. In this paper three types of alluvial deposits are distinguished: sandy (:2: 70% sand­ stones); mixed (70-30% sandstone); and conglom­ eratic ( ::::: 30% sandstone). Figure SC shows the percentage of sandstone (rela­ tive to conglomerate) through the entire basin and illustrates a gradual decrease from sand dominated alluvial deposits along the basin axis (axial alluvial sedimentation), to mixed alluvial deposits toward both basin margins (marginal alluvial sedimen­ tation). Likewise, Fig. SD presents a plan view of the spatial changes of the percentages of coarser grained sediments (conglomerates plus pebbly and very coarse-grained sandstones) relative to finer grained sediments (coarse to very fine sandstones). A pattern quite similar to the former (Fig. SC) can be seen. Clearly, the facies St and Sh are gradually replaced by facies Gm towards both basin borders. In both cases (Figs SC & SD) the only exception to the general pattern of sediment distribution is a NW -SE trending intrusion of coarse material in the southeast region of the basin. Alluvial facies: interpretation of lateral changes

Figure 6 presents an interpretation of the alluvial palaeogeography of the Guaritas sequence based on the lateral variations of the textural and palaeocur­ rent data. This figure was constructed according to the following considerations. The palaeocurrents suggest the coexistence of two distinct alluvial subenvironments (Fig. SB) with almost orthogonal mean sedimentary transport pat­ terns (282° versus 211°). 1 The first dispersal system (282°), developed in the eastern part of the basin, is characterized by the highest palaeocurrent vector dispersion and by palaeo­ flow almost perpendicular to the tectonic trend of

12

P. S.G. Paim B

12 -

Mean vector and number of readings per area Boundary between tributary alluvial fan system and trunk braided river system

50

�

Mean vector and number of readings of both alluvial systems

\

5

A

3 46 ,514 / 44 ti 99

!

26 / 64 /

;;

101

I

--

32

-36

c

§

>90 80-89 70-79

D

1

60 41

\

100 ""'-89 14

"4

deposits

"����

"'



0: u

--

"'

�u

:=C

en"' (f)

Silly Beds

�\�� ''\�

� . ''''"

,,,,,

''-'-'-'->--."."''-':-

c

"'

5

.0"'

�-g �"'

Q)(f)

a. a.

=>

u c

"'

Silver Sands

(f)

:;;

.2 Ui

30-

"'

..: :;;

a. a. =>

"'

u c

"'

(/)

E

40 -

::l .0 0

Heterolithic Sands

3:

--

"'

u c "' (f) c

5

.0 0

3: :;;

-'

�� ------�-

-s==: �L

-

�-

�

.. ...

----� ---� !"- .,_.

-- -- ----

High energy, ebb dominated channel! shoal complex (estuary mouth or open marine, sea strait environment)

(:==.:r> =

==M /Low energy, estuary abandonment! transgressive deposits

J High energy, ebb dominated, estuary mouth channel/shoal complex (= ebb-tidal delta environment)

�-

c

"li

� -� ��

=

= =M

= = = =

Low-moderate energy, estuary shoal deposits

=

�

�--.1!....,

"' u c

50 -

"'

� ..}.\.._ �

(f) c 3:

-

e

en

�-

�

�� ��

..... � -:::'{J! � � � /// !))})

3:

0

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·-··--�

Orange Sands 60-

.__

�

Moderate-high energy, flood dominated, channel-fill sands intercalated with tidal shoal deposits

�--��---""!.

? 70 -

Large-scale cross-bedding Trough cross-bedding Ripple cross-lamination Intraclasts

Basal transgressive deposits with phosphate nodules and reworked faunas

��[�

UPPER JURASSIC

J:?/-::-1 B. Q D

== =

. EJ c=J 1:-�;1 E;J '

Wavy bedding

� Strongly bioturbated

Flaser bedding

6:11 Shells and shell debris

Low-angle erosion surfaces

Q Plant debris

Concretions/nodules

-& Occasional burrows

H. D. Johnson and B.K. Levell

22

Table 1. Summary of the l ithofacies and reservoir characteristics of the main units within the Woburn Sands Interval

Lithology

Gault Clay

Grey fossiliferous claystones

Transition Series

Iron-cemented pebbly sands (basal beds); glauconitic & phosphatic fine-coarse, partly argillaceous Transgressive lag deposits. sands. In-situ lenses of richly fossiliferous limestone (Shenley Lmst.). Reworked clasts of iron-cemented sst. (Carstone) & Shenley Lmst. Rapid lateral lithological variations.

Muddy shelf.

Three main types of cross-bedding: I

Red Sands

Med.-v.coarse sand Mod.-poorly sorted. Ferruginous with up to 20% bv detrital iron oxide (red colouration) up to 2.5% bv heavy minerals 100% sand.

Silty Beds (/) Q z N

::t: \:::)

0' �

a

Cl

;: ., ;: "'­ t:l;)

;>;: !:""'

"' "' "' :::::

Fig. 12. Large-scale, sand-wave-type cross-bedding in the Silver Sands (Munday ' s Hill Quarry) . Individual sets are up to 4 m thick and are internally complex, ranging from avalanche foresets to low-angle surfaces (6-8°) separated by upslope- and downslope-dipping cross-bedding (see Fig. 13 for details) . Note also the sharp, planar contact between the Silver Sands and Silty Beds and the erosional contact between the Silver and Red Sands. The dark area along this contact represents a ridge of strong iron cementation ('Carstone rib ' ) . Finally the sequence is capped by the marine Gault Clay.

Transgressive estuarine sand complex

33

Fig. 13. Internal characteristics of the sand-wave-type cross-bedding in the Silver Sands in the Munday 's Hill- Double Arches area (large-scale structures illustrated in Fig. 12). (A) Tabular cross-bedding with simple, angle-of-repose foresets. Note small oppositely-dipping set (by lens cap) separating the main , thick sets and a prominent reactivation surface in the upper thick set (Double Arches Quarry) . (B) Relatively small tabular cross-beds, with each displaying an upward decrease in grain size (Munday's Hill Quarry) . (C) The low-angle surfaces dipping to the left represent an ebb­ directed (to southwest) sand-wave lee face . Internally these coarse to granule grade sands display upslope dipping cross­ bedding which represents flood-directed megaripples (Munday ' s Hill Quarry ) . (D) Herringbone patterns developed in relatively small sets of trough cross-bedding (Munday's Hill Quarry ) .

In the extreme west of the area, west of Heath and Reach (Sheepcott quarry; Fig. 2) , a heterolithic facies of decimetre-scale cross-bedded sands with persistent clay drapes overlies the Orange Sands. These beds, formally assigned to the Lower Woburn Sands, could be a lateral facies variant of the Silver Sands. In general, the Silver Sands are weakly bio­ turbated, apart from the upper few metres in some pits (e.g. New Trees) , where clay-lined Ophiomorpha-type burrows occur. Interpretation

The textural and mineralogical maturity of the Silver Sands, the uniformly large size of cross-bedding and

the lack or scarcity of clay drapes and burrows suggest deposition in a higher energy environment than both the Heterolithic and Orange Sands. None the less, the relatively infrequent evidence of revers­ ing palaeocurrents still suggests tidal deposition. The extensive subhorizontal erosion surface at the base of the unit and the evidence of a bedform complex up to 1 5 m thick, at least in the Heath and Reach area, suggest deposition in a high-energy, current-dominated bar system after a period of wide­ spread erosion. There is no clear channel-fill facies and it is possible that deposition occurred in an environment in which interbar depressions were created by the accretion of shoals rather than the active cutting of channels. In more detail, the cross-bedding indicates large ,

34

H. D. Johnson and B.K. Levell

possibly sand-wave-type bedforms with super­ imposed megaripples over the majority of the out­ crop area, and microdelta-like forms infilling scour hollows in the north. Taken together with a similar change in cross-bedding style in the younger Red Sands (see below ) , this change could be related to a southward increase in water depth across the study area (Fig. 2 ) , which allowed the construction of larger bedforms in the south of the area. In comparison with the shoal complex of the Heterolithic Sands, current energy was apparently much higher and there was less time for the settling of suspended clay. The textural and mineralogical maturity of the sands also indicates a more intensive phase of reworking prior to final deposition. Silty Beds

Description

The Silty Beds (Lamplugh & Walker, 1903; Lamplugh , 1922) comprises a thin interval (mainly 1 - 2 m ; maximum 3 .6 m) of grey-green, glauconite­ and carbonaceous-rich clays, silts and argillaceous fine-grained sands with rare thin beds of coarse- to very coarse-grained, well-sorted sands. The fine­ grained sediments are strongly (90- 100 % ) bio­ turbated (Fig. 14) and have in places suffered soft-sediment deformation. The coarse beds are 50- 1 00 mm thick , have sharp bases and flat or rippled tops, and are internally cross-laminated or plane-laminated. The unit sharply overlies a slightly undulating surface on top of the Silver Sands (Fig. 12). This undulose surface was considered by Lamplugh ( 1922) and Bentley ( 1970) to represent the preserved morphology of shoal or sand-wave crests in the underlying Silver Sands. The Silty Beds are sharply overlain by either the Transition Series or the Red Sands (Fig. 4). Interpretation

This interval represents a period of slow sedimen­ tation in a marine environment in which both current and wave energy were apparently too weak to cause significant erosion of the unconsolidated sands on the sea-floor. The coarser grained sand layers are interpreted as lags, resulting from reworking and winnowing of the Silver Sands (cf. Levell, 1980). The Silty Beds are interpreted to represent a phase of low energy and low sediment influx (an estuarine abandonment phase) relative to the high-energy,

14. Silty Beds, overlying the Silver Sands with a sharp, planar contact (next to trowel ) . The mottled texture in the Silty Beds is the .result of extreme bioturbation. A low-angle laminated sand bed with an erosional base (storm layer) is present in the upper part of the sequence (Bryant's Lane Quarry) .

Fig.

subtidal sand shoal and channel complexes described from the underlying deposits. The precise depositional environment of this unit is debatable. The lack of marine fauna and abundant carbonaceous material could indicate restricted marine conditions, possibly within a deeper, pro­ tected part of an estuary ( Lamplugh, 1922). The facies could also be interpreted as tidal flat deposits (Eyers, 1992b) , although the absence of tidal channels and the characteristic fining upward sequence capped by in situ marsh deposits (e.g. Evans, 1965) fails to support this. Alternatively, the abrupt cessation of high-energy conditions marked by the sharp base of this unit, and the presence of reworked and glauconite-rich horizons within it, is more characteristic of a sudden deepening, rather

Transgressive estuarine sand complex

than shallowing in water depth. Hence , a more offshore shallow marine environmen t is a further possibility . The Silty Beds probably represen t a relatively long time interval, which marks the end of the underlying Orange- Heteroli thic and Silver Sands depositional succession. This i n terpreta tion implies that the overlying Red Sands form a genetically separate sand body. Red Sands .

Description

The unit comprises moderately to poorly sorted, medium- to very coarse-grained sands (occasionally gravelly), which are distinguished from all older units of the Woburn Sands by their high conte n t (up to 20% bulk volume) of detri tal iron oxide in the form of rust-coloured ooliths of goeth i te and angular ir9nstone chips. It is also rich in heavy minerals ( Bentley, ( 1970) reports up to 2 . 5 % bulk volume) . I n addition, diagene tic iron oxide occurs a s a local cement around mud clasts and on certain foresets. The Red Sands erosionally truncate all earlier deposits. In the Shenley Hill area they were shown by Bentley ( 1970) to occupy E- W trending shallow scours up to 6 m deep and 100 m wide. South of this area the Silty Beds and the Silver Sands are progressively cu t out by erosion beneath the sou th­ ward thickening Red Sands until, south of Leighton Buzzard, the sand pits expose only this last unit (Fig. 4), which is presumed to overlie Orange and/ or Heterolithic Sands. The erosion surface at the base of the Red Sands is i n terpreted as a sequence boundary by Ruffell and Wach (in press). Three main types of cross-bedding characterize the Red Sands. Type I consists of thick (up to 5 . 5 m) sets of avalanche cross-bedding that overlie prominent erosional scours up to 3 m deep and more than 100 m wide . The overall set geome try thus resembles giant­ scale trough cross-bedding. The uppermost se t of the unit is overlain by a coset of successively smaller se ts with a particularly i n triguing geometry (Fig. 15). The thick set of avalanche foresets at the base dip at 25 -30° and abut sharply against a basal erosion surface locally lined with granules and pebbles. Major reactivation surfaces periodically separate bundles of foreset laminae with an average thickness of 130 mm (range 50-220 mm) . Each of

35

these tabular or wedge-shaped bundles is in turn subdivided by some times i n tersecting minor dis­ continuity planes that define the individual foresets. The major reactivation surfaces are convex-up, flattening toward the top of the set, where they are separated by cosets or sol i tary sets of decimetre­ scale cross-bedding. At the top of the sets, where the reactivation surfaces flatten out, there are some up-dip thinning sets 0.05 - 0.2 m thick of oppositely dipping cross-bedding. This geometry closely resembles, on a large scale, the ebb and flood shields described from several modern tidal environments (e.g. Klein , 1970) . Type ll is characterized by wedge-shaped cross-sets (0.3- 2 m thick) with sou th- or southwest-dipping foresets, which are superimposed on 2-4-m-thick sets of low-angle southward-dipping erosion surfaces (dips 4-8°) . Individual cross-sets taper either up or down the inclined surfaces. Small sets (less than 0.5 m thick) occasionally thicken down-dip into single large-scale ( c. 2-3 m thick) avalanche sets. Type II cross-bedding passes laterally i n to type I in both down- and up-palaeocurren t direction . Type III comprises 0. 1 -0.4-m-thick sets o f trough and tabular cross-bedding bounded by subhorizontal erosion surfaces, and forms the bulk of the Red Sands in the channelized facies of the northern area.

Additional information on the complex and varied geometry of these sedimen tary structures has been obtained recently from ground-penetrating radar studies ( Bristow, in press) . The Red Sands are extensively bioturbated (Fig. 16). The most widespread burrow type is a c. 5-mm­ diameter colour mottling in which pale-coloured burrow fills contrast with the darker iron-rich sands surrounding them (Fig. 16A) . This structure results from horizontal burrowing and has caused only minimal disruption of the primary stratification. Such bioturbation is apparently similar to that termed 'cryptobioturbation' by Howard & Frey ( 1975) . This structure was found by them in the G eorgia (USA) estuaries and a ttributed by them to amphipod crustaceans. The colour mottling could be due either to a slight grain-size fractionation ( the iron pellets fall mostly in the fine- to very fine-grained sand classes ( Bentley, 1970) or to some oxidation/ reduction difference caused by organic slime. The second and most distinctive burrow type is the nested inverted cone or funnel-shaped struc-

36

H. D. Johnson and B. K. Levell

Fig. 15. Large-scale (c. 3-4 m thick ) , ebb-directed (to south) , avalanche foresets and associated top set deposits.,.(A) Foreset packets are separated by major reactivation surfaces. The latter are ascribed to stronger than normal flood-currents because they are overlain by packets of oppositely dipping, wedge-shaped cross-beds. The large-scale structure is interpreted as an ebb-dominated bedform which filled an erosional hollow. (B) Close-up of central part of (a) showing oppositely-dipping (to northeast), flood-directed sets, possibly analogous to flood-shield bedforms. Locality: Pratts' Lane Quarry.

ture (Fig. 1 6 B) . The V-shaped laminae (in two dimensions) of these structures could have formed either by sediment collapse i n to the underlying horizontal tubes, which are sometimes associated with these structures, or could have formed as an escape structure by upward movement of an animal (e.g. large pelecypod or crustacean). The escape s tructure interpretation may be supported by the association of this sort of burrow with large erosion surfaces overlain by migrating bedforms. Perhaps the animals preferred the relatively slow deposition areas of channel floors but were none the less able to escape by efficient vertical burrowing when buried beneath advancing bedforms.

A third burrow type consists of horizontal 40-mm­ diameter subhorizontal tubes with miniscus-shaped spreiten. The size, shape and back-filling spreiten resemble the 'press structures' produced by £chino­ cardium cordatum ( Reineck & Singh , 1973). If these burrows are echinoid burrows, they would indicate fully marine conditions. Interpretation

The Red Sands accumulated in a high-energy marine or marginal marine environment dominated by southward flowing, but sometimes reversing, cur­ rents (Fig. 8) . There is no evidence of distinct

Transgressive estuarine sand complex

37

Fig. 16. Biogenic sedimentary structures in the Red Sands. (A) Typical mottled appearance of the Red Sands resulting from horizontal burrows. Note the alternation of non-bioturbated laminae and the retention of structure within the bioturbated layers. Locality: Pratt's Lane. (B) Two V-shaped burrows occurring within a large-scale cross­ bedded sand body at the contact between two different sediment types (coarser, massive sands above) . The lower cross-bedded sands also display the mottled texture detailed in (a) . Locality: Grovebury South.

channel-fill or shoal deposits, and the whole uni t seems to represent rather a blanke t-like field of megaripples and megarippled sand waves (cf. J . R.L. Allen, 1980) , which increased systematically in height southwards. The sedimentary s tructures closely resemble tidal sand waves found in modern subtidal environments (e.g. Berne et a!. , 199 1 ) . There i s no evidence of emergence o r s ignificant wave activity , but the b ioturbation suggests similar marine or marginal marine cond itions to the under­ lying units. The lack of clay drapes indicates a more or less constantly active water column .

Transition Series

Description

The Transition Series comprises a d istinctive thin (c. 1 -2 m thick) , and laterally extensive deposit at the junction between the previously described units and the overlying Gault. I t has been studied inten­ sively by palaeontologists because of its rich fauna (for more extensive descriptions see Lamplugh ( 1922 ) , Wright & Wright ( 1947 ) , Hancock ( 1958) , Casey ( 1961 ) and Owen ( 1972)) . Lithologically the Transition Series is laterally variable , consisting mainly of partly cemented pebbly

38

H. D. Johnson and B.K. Levell

sands (at the base) and argillaceous, glauconitic and phosphatic fine- to coarse-grained sands and sandy clays. A distinctive and unusual limestone, the Shenley Limestone, occurs as discontinuous in situ bands, up to c. 0.6 m thick , and as reworked frag­ ments in the lower part of this unit. This limestone, which sharply overlies the underlying deposits, contains a unique fauna, including rich and dis­ tinctive brachiopods and a unique assemblage of echinoids and Crustacea. The whole Transition Series is highly fossiliferous, with abundant ammon­ ites, bivalves, belemnites, gastropods and oysters. There is abundant evidence of reworking and dis­ continuous sedimentation in the form of reworked faunas and several condensed horizons. The Tran­ sition Series forms a laterally extensive drape over the underlying Woburn Sands and extends laterally over Jurassic sediments (Owen, 1972) . Interpretation

The Transition Series represents a composite trans­ gressive , shallow-marine lag, which forms a type of abandonment deposit on top of the underlying, mainly high-energy, subtidal deposits. The rework­ ing of limestone beds, concretions and iron­ cemented sandstones ( ' Carstone'), mixing of faunal zones and condensed intervals indicate slow depo­ sition and strong current activity. The abundant and highly varied fauna supports a shallow marine, probably offshore, depositional environment, with the coarser grained layers representing winnowed lag deposits.

PALAEOGEOGRAPHICAL EVOLUTION

Reconstruction of the palaeogeographical evolution of the five erosionally bounded units depends not only on the process interpretations given above but also on the overall palaeogeographical setting and the stratigraphical relationships between the units. The first point to establish is the direction of the transgression . Several factors lead us to propose a northward transgression . 1 The affinities of the fauna described by Casey ( 196 1 ) from the lower part of the complex are with faunas from the southern Wealden Basin, rather than from the northern Southern North Sea Basin. 2 In tidal embayments it is common for ebb-directed sets to be preferentially preserved, either because ebb flows are often enhanced by freshwater run-off

or because these flows reach maximum velocities at times of lower water, and are thus preferentially preserved in the topographically lower parts of the deposit. The predominance of southwestward­ directed palaeocurrents leads to the proposal of open sea in the southwest. 3 The increase in set size southwards in the Red Sands (and possibly in the Silver Sands) , coupled with a southward increase in the thickness of the Red Sands, could be explained by greater water depth in the south. Water depth is, of course , of especial importance in the preservation of sets up to 5 . 5 m thick. In modern tidal basins (e.g. estuaries and embayments) the largest bedforms frequently occur in the deeper areas (see Fig. 17). 4 The geometry of the deposit is one of southward­ dipping facies units (Fig. 4) . This is most readily explained by a northward transgression . Two factors lead us to suggest that the Red Sands should be considered separately from the older sub­ divisions of the Woburn Sands. 1 They rest erosively on the Silty Beds, which appears to be a slowly deposited unit representing the cessation of sand supply to the underlying Silver Sands shoal system. 2 Their distinctive petrography, rich in detrital iron­ oxide fragments with goethite ooliths, suggests a change in the source material. The most likely possibility is the erosion of presently unexposed iron-rich intervals derived from further updip to the north , which may include the iron-stained Orange Sands, the carstones of the Silver Sands or some other, older or laterally equivalent formation(s) (e.g. the early Cretaceous Carstone of the Norfolk area, Casey & Gallois, 1 973). The only reported occurrence of similar goethite ooliths is from an early Cretaceous oolith-rich sand unit further north (Deepdale Pit), which may be equivalent to the Silver Sands (Eyers, 1992a) . The deposition of the Woburn Sands began with the partly erosional and partly tectonic creation of a trough trending NE-SW through the Bedfordshire area. The basal deposits are gravels and phosphates with rolled, derived faunas and abundant wood material, suggesting transgressive coastal erosion (whether open-marine or on the margins of a brackish embayment is unknown). The overlying Orange and Heterolithic Sands represent subtidal channel and shoal deposits, respectively. The relationship between these two units could be either lateral interfingering or a sharp erosional contact (Fig. 17). l n the first case

Transgressive estuarine sand complex

39

�� Tidal channel fills I . I High-energy tidal shoals C:=J Moderate-energy tidal shoals

}

}

Orange Sands Heterolithic Sands

b) Model 1

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. __ .•.•.. •

::::

HIGH ENERGY CHANNE HIGH ENERGY L SHOAL SHO L - -. . . ::::::::::-::=·. .� : - .:0 - _ - =-== = . '\ . ;- _

MODERATE ENERGY SHOAL

== -==;::;===;;:::_:::: ----== _ :: :; :

B

0

?

c) Model 2

A

5m B

17. (a) Distribution of the Orange and Heterolithic Sands (Lower Woburn Sands or Brown Sands) below the Silver Sands erosion surface, and (b) two possible models depicting possible lateral relationships (model 1 preferred) .

Fig.

40

H. D. Johnson and B.K. Levell

(Fig. 17b) , the Heterolithic Sands would be situated towards the centre of the estuary or embayment, with the Orange Sands being a marginal , flood­ dominated , channel complex. Such a palaeo­ geography with marginal channels hugging the edges of an embayment or estuary would be consistent with the generally rotary circulation models for estuaries, where tidal currents are deflected by Coriolis force to the margins (e.g. Schubel , 197 1 ) . Nevertheless, a problem remains with this inter­ pretation in that it is difficult to understand why the Heterolithic Sands shoal complex accreted steadily to a thickness of 25 m , without being dissected by a migrating channel system. However, the similarity between the Heterolithic and Orange Sands leads us to infer a laterally interfingering relationship (Fig. 17b). Two features place these two units in the relatively landward and protected parts of the embayment or estuary, rather than on the open shelf. 1 The clear presence of distinct channel and shoal complexes is characteristic of the estuary proper, rather than the seaward extensions. Furthermore the flood dominance of the Orange Sands channels indicates mutually evasive ebb and flood channels. 2 The well-developed clay drapes indicate periods of slack water in a generally current-dominated regime. Whether these are diurnal drapes or not, periods of tidal slack water are more likely to occur in systems with rectilinear tides. Such tides are more common in the inner and more restricted parts of estuaries, whereas in estuary mouths and on the open shelf, tides tend to be rotary (Terwindt, 1973). The major erosion surface at the base of the Silver Sands indicates a shift in depositional setting, and is explained most readily by the lateral shift of a major, wide, estuary mouth channel system. The overlying Silver Sands, with the lack of a distinct channel-fill facies and the large-scale bedform complexes up to 15 m high , suggest, in the context of a continuing northward transgression, deposition in estuary mouth subtidal shoals. Modern analogues might be the outer reaches of Chesapeake Bay ( Ludwick 1970, 1974; Colman eta/. , 1988), or Delaware Bay ( Knebel et al . , 1988) , or the seaward portions of ebb-tidal deltas (Greer, 1975 ; Hayes, 1975 ) . A s mentioned above , the Silty Beds represent the cessation of sand supply to the Silver Sands shoal, followed by a period of slow deposition. Reasons for the abandonment of this shoal could be autocyclic (e.g. a shift in the shoal pattern related to the movement of main tidal channels) or allocyclic (e.g.

sea-level rise leading to the area becoming an open­ marine shelf rather than an estuary mouth. However, alternative interpretations consider this interval to represent tidal flat deposits and hence a relative shallowing of sea-level of the facies succession has been proposed (Eyers, 1992b) . In gross facies terms (e.g. the style of cross­ bedding, grain size and lack of clay drapes) the Red Sands closely resemble the Silver Sands. However, there is a major difference in their petrography. The Red Sands could, therefore , represent a similar outer estuarine shoal complex but formed at a time when the up-dip tidal channels were eroding different material. They would therefore represent the shift of tidal channels to debouch again in the Leighton Buzzard Area. An alternative and more dramatic interpretation would relate the Red Sands to the final breaching of the London-Brabant land mass and thus to continued coastal erosion at the margin of the transgressing embayment. As discussed above, the erosion surface at the base of the Red Sands has been interpreted as evidence for a sea-level low­ stand (sequence boundary) punctuating the overall late Aptian transgression. This explanation would presumably equate the depositional settings of the Silver Sands and Red Sands, with the Silty Beds representing either the most distal (offshore) or most proximal (tidal flat) depositional settings. On the grounds of the facies differences we have described , and of simplicity, we prefer to view the succession as one of continued overall transgression and attribute the baSal erosion surface of the Red Sands to changes in tidal dynamics, possibly associ­ ated with the breaching of the Lori don - Brabant land mass. The Transition Series marks the final abandon­ ment of the high-energy, tide-dominated sanely shoal complexes. Sedimentation was slow, discontinuous and interspersed with periods of physical reworking. Subsequent basin deepening resulted in the wide­ spread deposition of shallow-marine/shelf muds (the Gault) .

DEPOSITIONAL MODEL FOR A TRANSGRESSI V E ESTUARINE - EMBAYMENT SYSTEM

A generalized depositional model can be constructed for a transgressive estuarine -embayment system based on observations from both the Woburn Sands and from known facies and bedform distributions

Transgressive estuarine sand complex

within modern meso- and macrotidal estuaries and embayments (Fig. 18). This model also can be compared with estuarine facies models developed mainly from a synthesis of modern environments and from conceptual considerations (Dalrymple et al., 1 99 1 ) . The essential elements o f the inner estuarine environment are as follows. 1 Discrete , mutually evasive, ebb and flood tidal channels and interchannel tidal shoals. 2 Tidal channel sands will be relatively coarse-grained and characterized by large-scale cross-bedding. 3 Tidal shoals will consist of more heterolithic sands, generally with smaller scale structures, stronger bio­ turbation and more frequent and laterally extensive clay layers than the higher energy channel-fill deposits.

b)

41

4

Reservoir quality and heterogeneity will be strongly influenced by the contrast between channel and shoal deposits. In general, the channel fills may be expected to form discrete zones of higher per­ meability (Fig. 18b), whereas the tidal shoals will comprise variable but generally lower quality and more heterogeneous reservoirs ( c.f. Orange and Heterolithic Sands). The essential features of the outer estuarine environ­ ment are as follows. 1 Minor lithological differentiation between tidal channel and shoal deposits. 2 Well-sorted sands with fewer clay intercalations and less bioturbation compared with the inner estuarine sands. 3 Large-scale cross-bedding, partly reflecting large estuary mouth bars and sand waves, and also

Tidal shoals

'

3

a)

Basic Characteristics

Woburn S. Equiv.

Fossiliferous clays

Gault Clay

Reworked, fossiliferous, partly cemented sands

TRANSGRESSIVE LAG

Transition Series

Well-sorted sands. Large-scale, complex cross-bedding. Ebb-dominated palaeocurrent directions. Mainly minor bioturbuation.

OUTER ESTUARINE/ EMBAYMENT TIDAL SHOALS

Silver Sands and Red Sands

Pre-transgressive deposits

� l

Upper Jurassic Clays

18. A model for a transgressive estuarine- embayment depositional system (see Fig. 3 for legend). (a) Idealized vertical section through a transgressive estuarine-embayment complex based on the Woburn Sands. (b) Block diagram illustrating some of the essential differences in sand body characteristics between the inner estuarine/embaymer.t and outer estuarine/embayment environments. (c) Idealized section through a transgressive estuarine- embayment indicating the potential stratigraphic trap geometry.

Fig.

H. D. Johnson and B.K. Levell

42

the infilling 0f deep ebb-dominated channels and interbar troughs. 4 Possibly an overall seaward i ncrease in the size of cross-bedding in response to deeper-water conditions. 5 Higher quality and more homogeneous reservoir characteristics compared with the inner estuarine sands as a result of both stronger tidal currents and increased wave activity. Shale layers are likely to be very thin and discontinuous, or absent. The vertical sequence through this type of trans­ gressive sand complex would tend to have lower reservoir quality sands in the lower part and higher reservoir quality sands towards the top (Fig. 1 8a) . The transgressive nature of the sequence would lead ultimately to the blanketing of the overall lenticular sand-body complex by shelf muds and could thereby, form a stratigraphical trap (Fig. 1 8c) .

DISCU S S ION

The difficulty of postulating a generalized facies model for this type of succession is that these deposits are influenced by a particularly wide range of physical processes, both autocyclic and allocyclic; they may receive sediment supplied from different sources (landward and/or seaward; Schubel, 197 1 ) and their preservation potential i s linked closely to the rate and sense of relative sea-level fluctuations. These variables form the basis of Dalrymple et al. 's ( 199 1 ) classification of estuarine facies, in which they argue that these deposits are diagnostic of transgressive conditions. In this scheme the Woburn Sands would be classified as a tide-dominated estuary in which fluvial processes were negligible . Compar­ able modern environments in terms of physical processes, facies and sedimentary structures would include the southwest Netherlands area (e.g. Haringvliet and Oosterschelde Estuaries) (Oomkens & Terwindt, 1960; Terwindt, 1973; van den Berg et al. , 1980; Yang & Nio, 1989 ) , Severn Estuary/ Bristol Channel area, (Hamilton, 1979; Harris & Collins, 1985 ; J . R . L. Allen, 1 990) , Ossabaw Sound (Greer, 1975 ) , Jade Estuary/German B ight ( Reineck, 1963; Reineck & Singh, 1973) and Broad Sound, Australia (Cook & Mayo, 1977). The stratigraphical relationships recorded in the Woburn Sands and summarized in the depositonal model allows comparison with other stratigraphical models associated with tide-dominated estuaries, incised valleys and transgressive systems tracts , and

allows consideration of their sequence stratigraphical implications (e.g. van Wagoner et at. , 1990) . The major third-order lowstand and maximum flooding events would be recorded respectively by the basal lag deposit of the Woburn Sands and by the condensed horizons in the Lower Gault Clay. We have seen no direct (i.e. preserved) evidence of fluvial incision at the base of the Woburn Sands, which comprises a typical shallow-marine lag deposit. The non-preservation of predicted lowstand fluvial deposits among shallow marine sandstones successions is commonly argued to be the result of either (i) minor lowstand fluvial deposition, due to a predominance of fluvial bypassing during incision, and/or (ii) extensive shallow marine reworking during subsequent relative sea-level rise. In many cases this combination results in the apparent merging of lowstand and initial marine flooding sur­ faces, and removal of all lowstand fluvial deposits (e.g. Bergman & Walker, 1987; Plint & Walker, 1987; Plint & Norris, 1 99 1 ; Thorne & Swift, 199 1 ; Walker & Bergman, 1992). The greatest amount of such reworking is likely to occur in tide-dominated settings, particularly in confined areas such as estuaries or embayments, due to high-energy tidal currents and their ability to incise deeply into the substrate. Holocene incised valleys on the inner continental shelf of southeastern USA , for example, show a hierarchy of erosional events that reflect not only earlier fluvial incision but later deep tidal scour (c. 15 -30 m ) , particularly at the confluence of tidal streams, in tidal channels associated with tidal flats, in tidal inlets and in channels associated with ebb tidal delta shoals (Oertel et a/. , 199 1 ) . Preservation potential o f transgressive estuarine successions is relatively high due to their relatively protected position within palaeovalleys (Swift et a/. , 1980; Demarest & Kraft , 1987). The degree of preservation is related most closely to the depth of incision of the basal erosion surface , or sequence­ bounding fluvial incision, plus any additional down·· cutting by tidal scour associated with the initial flooding surface and , to a lesser extent, the degree of erosion associated with the ravinement surface (Dalrymple, 1992). The latter is equivalent to the ravinement/shoreface erosion surface of wave­ dominated shorelines , but in tide-dominated systems it is characterized by tidal current erosion (e.g. the bedload parting zones of tidal seas; Stride, 1982 ) . Complete preservation of an idealized tide­ dominated transgressive stratigraphy, which seems to be extremely rare, would predict tidal shelf sand

Transgressive estuarine sand complex

deposits (e.g. tidal sand ridges and tidal sheet sands with sand waves and associated deposits; Stride, 1982; Nummedal & Swift, 1987; Dalrymple, 1992) overlying the ravinement surface . All the main sand bodies in the Woburn Sands appear to occur below the transgressive marine erosion surface and hence are assigned to the estuarine part of the sequence. Shelf sand wave deposits may occur further to the south and southeast within the Folkestone Beds (Narayan, 1 97 1 ; J . R. L . Allen, 1 982; Bridges, 1982). The main difference between tidal shelf sands and tidal estuarine sands is that the latter will be domi­ nated by complex interfingering of tidal-channel, tidal-shoal and, in places, tidal-flat deposits (cf. Maguregui & Tyler, 1 99 1 ) . Thus in the Woburn Sands, concave-upward erosion surfaces and rela­ tively rapid lateral thickness changes are common. In more basinward settings individual sand bodies are likely to have tabular geometries, with individual sand and shale layers having much higher continuity (cf. Surlyk and Noe-Nygaard, 1 99 1 ) . Distinguishing between these different types of shallow-marine sand body is particularly important in evaluating the nature of such deposits in the subsurface because of their different geometries and lateral extent (e.g. Banerjee, 199 1 ; Brownridge & Moslow, 199 1 ; Jenette et a!. , 1992).

ACKNOW L E DGEMENTS

We would like to extend our sincerest thanks to Harold Reading for giving us our initial introduction to the Woburn Sands and for making us aware of its sedimentological significance. The basis for this work, its methodology and evaluation stems directly from the lessons learnt by the authors under Harold's previous supervision. We are extremely grateful for this and for his ongoing interest and influence in our work. The authors would also like to thank the Konin­ klijke/Shell Exploratie and Produktie Laboratorium, Riswijk, The Netherlands, for supporting both the project and its publication. We also acknowledge Shell Research B . V. and Shell Internationale Pet­ roleum Maatschappij B . V . for their permission to publish this paper. We are also indebted to Guy Plint for his patience and encouragement during the preparation of this paper. The assistance and openness of the two referees, Jill Eyers and Alistair Ruffell , is also gratefully acknowledged. Finally, Ian Glenister is thanked for redrafting the figures.

43 REFERENCES

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45

Aktua-palaeontologie nach Studien in der Nordsee, p. 666. W. Kramer, Frankfurt. ) . ScHUBEL, J . R . ( 1971 ) Sources o f sediments to estuaries. In: The Estuarine Environment, Estuaries and Estuarine Sedimentation (Eel. Schubel, J . R . ) , pp . V, 1 - 19. American Geological Institute, Washington. SCHWARZACHER, W. ( 1953) Cross-bedding and grain size in the Lower Cretaceous sands of E. Anglia. Geol. Mag. , 90, 322-330. SHEPHARD-THORN , E . R . , HARRIS, P . M . , HIGHLEY, D . E . & THORNTON, M . H . ( 1 986) An Outline Study of the Lower Greensand of Parts of South-east England. Report for the Department of the Environment, British Geological Survey, Keyworth. SPEARING, D.R. ( 1975) Shannon Sandstone, Wyoming. In: Depositonal Environments as Interpreted from Primary Sedimentary Structures and Stratification Sequences. Society of Economic Paleontologists and Mineralogists, Short Course No. 2, pp. 104- 1 14. STRIDE, A . H . ( 1982) Offshore Tidal Sands: Process and Deposits. Chapman & Hall, London, 213 pp. SuRLYK, F. & NoE-NYGAARD, N . ( 1991 ) Sand bank and clune facies architecture of a wide intracratonic seaway: Late Jurassic- Early Cretaceous Raukelv Formation, Jameson Land, East Greenland. I n : The Three­ Dimensional Facies Architecture of Terrigenous Clastic Sediments and Its Implication for Hydrocarbon Discovery and Recove1y (Eels Miall, A . D . & Tyler, N . ) . Soc. econ. Paleontol. Mineral . , Concepts Sediment. Paleont . , Tulsa, 3 , 261 - 276. SwiFT, D . J . P . & THOR N E , J .A . ( 1 99 1 ) Sedimentation on continental margins, 1: a general model for shelf sedimentation. I n : Shelf Sand and Sandstone Bodies: Geometry, Facies and Sequence Stratigraphy (Eels Swift, D . J . P . , Oertel, G.F. , Tillman, R.W. & Thorne, J . A . ) . Spec. Pubis i n t . Assoc. Sediment . , No . 1 4 , p p . 3 - 3 1 . Blackwell Scientific Publications, Oxford. SWIFT, D . J . P . , MOIR, R . & FREELAN D , G . L. ( 1980) Quaternary rivers on the New Jersey shelf: relation of seafloor to buried valleys. Geology, 8, 276-280. TERWINDT, J . H . J . ( 1 973) Sand movement in the in- and offshore tidal area of the southwestern part of The Netherlands. Geol. Mag. , 52, 69-77. THORNE, J . A . & Swwr, D . J . P. ( 1991) Sedimentation on continental margins, VI: a regime model for depositional sequences, their component systems tracts, and bounding surfaces. I n : ShelfSand and Sandstone Bodies: Geometry, Facies and Sequence Stratigraphy (Eds Swift, D . J . P . , Oertel , G . F . , Tillman, R . W . & Thorne, J . A . ) . Spec Pubis int. Assoc. Seclimento l . , No. 14, pp. 189-255. Blackwell Scientific Publications, Oxford. VAN DEN BERG, J . M . , GoESTEN , M . J . B . G . & S M U LDERS, F. ( 1980) Dynamics and sequential analysis of a mesoticlal shoal and intershoal channel complex in the eastern Scheidt (southwestern Netherlands). Sedimem. Geol. , 26, 263 -279. VAN WAGONER, J .C . , MITC H U M , R . M . , CAMPIO N , K . M . & RAHMA N IAN , V . D. ( 1 990) Siliciclastic Sequence Stra­ tigraphy in Well Logs, Cores, and Outcrops: Concepts for High-resolution Correlation of Time and Facies. American Association of Petroleum Geologists, Tulsa, Methods in Exploration, 7, 55 pp. ViSSER, M . J . ( 1 980) Neap-spring cycles reflected in

46

H. D. Johnson and B. K. Levell

Holocene subtidal large-scale bedform deposits: a preliminary note. Geology, 8, 543-546. WALKER, R.G. ( 1 984) Shelf and shallow marine sands. I n : Facies Models ( E d . Walker, R . G . ) . Geosci. Can . , Reprint Ser. 1 , 1 4 1 - 170. WALKER, R.G. & BERGMAN , K .M. ( 1992) Late Cretaceous clastics: cores and outcrops. In: Field Trip Guidebook: Trip No. 6. Canadian Society of Petroleum Geologists/ American Association of Petroleum Geologists, Calgary, 77 pp. WALKER, R . G . & P uNT, A . G . ( 1992) Wave- and storm­ dominated shallow marine systems. In: Facies Models: Response to Sea Level Change (Ed. Walker, R.G. & James, N . P . ) , pp. 2 19-238. Geological Association of Canada, St Johns. WoNHAM, J . P . ( 1993) Sedimentology and sequence stra­ tigraphy of tidal sandstone bodies: implications for reservoir characterisation. Unpublished PhD thesis, University of Liverpool.

C.W. & WRIGHT, E . ( 1947) The stratigraphy of the Albian Beds at Leighton Buzzard. Geol. Mag. , 84, 1 6 1 - 168. WvArr, R.J . , LAKE, R . D . , MoORLOCK, B . S . P . & SHEPHARD­ THOR N , E . R . ( 1986) The Leighton Buzzard Area: Geo­ logical Report on 1: I 0 000 Sheets SP82NE, Southeast; SP92NW, Southwest; SP93NW, Northeast, Southwest, Southeast. British Geological Survey, Keyworth . YANG, C.-S. & N i o , S.-D. ( 1989) An ebb-tide delta depo­ sitional model - a comparison between the modern Eastern Schelde tidal basin (southwest Netherlands) and the Lower Eocene Roda Sandstone in the southern Pyrenees (Spain ) . Sediment. Geol. , 64, 175- 196. ZIEGLER, P . A . ( 1 988) Evolution of the Arctic- North Atlantic and the Western Tethys. Mem . Am. Assoc. petrol. Geo l . , 43, 198 pp. ZIEGLER, P.A. ( 1990) Geological Atlas of Western and Central Europe. Shell Internationale Petroleum Maatschaappij B . V . , The Hague, 239 pp.

WRJGHT,

Spec. Pubis int. Ass. Sediment. (1995) 22, 47-74

An incised valley in the Cardium Formation at Ricinus, Alberta: reinterpretation as an estuary fill R O G ER G . W A L K E R Department of Geology, McMaster University, Hamilton, Ontario L8S 4Ml, Canada

ABSTRACT

The Ricinus incised valley (Turonian-Coniacian Cardium Formation) is at least 50 km long, 10 km wide and up to about 30 m deep. It trends NW-SE, parallel to tectonic trends and shoreline trends in the Cardium Formation. I t cuts into shallow-marine mudstones and sandstones (Raven River Allomember of the Cardium), and is truncated by a transgressive surface of erosion (E5). The valley fill forms a small oil reservoir in the subsurface, but the field is disrupted by thrust faults, and the western margin is poorly constrained because of thrusting and Jack of wells. The valley fill consists of interbedded sandstones and mudstones, with both brackish and fully marine trace fossil assemblages. Some of the sandstones are structureless, with individual beds apparently over 4 m in thickness; others are parallel-laminated with beds up to 3 m thick. Cross-bedding in sets thicker than 5 em is extremely rare. The mudstones contain some thin silty and sandy laminations, but are commonly extensively bioturbated with a fully marine trace fauna (robust Helminthopsis, Zoophycos, Planolites, Anconichnus, Skolithos, Ophiomorpha, Thalassinoides, Teichichnus, Asterosoma, Terebellina and Rosselia). The mudstones can be traced as a sheet about 4 m thick along the entire length of the valley; they appear to pinch out against the eastern wall of the valley. Chert-pebble conglomerates are scattered throughout some wells, but tend to be concentrated near the northern end of the valley. Here, they arc associated with sandstones and with fully marine mudstones. The valley was formerly interpreted to have been cut and filled by turbidity currents, but this interpretation is difficult to reconcile with the depositional environments of the rest of the Cardium Alloformation. It is suggested here that the valley was cut by a river during a falling stage of relative sea­ level, and was filled transgrcssivcly as an estuary. There is absolutely no evidence for tidal currents, and little indication of waves. The very thick and structureless sandstones suggest pulses of sediment swept into the estuary from the bay head, possibly as density underflows. There is little or no evidence for the reworking of these deposits within the estuary. There arc no central basin mudstones (no turbidity maximum), although this could be due to incomplete preservation of the entire length of the estuary. The geometry of the estuary and the facies distribution of the fill suggest that the open sea Jay to the north. Chert pebbles moving in the longshore drift system were introduced into the estuary from the north by storms; they were not introduced from the fluvial (southern ) end of the estuary. The orientation of the incised valley, parallel to depositional strike for over 50 km, remains a maj or problem.

INTRODUCTION

storm-dominated shallow-marine sandstones (Walker & Eyles, 1 988) , and transgressive incised shoreface deposits ( Bergman & Walker, 1987 , 1988; Pattison & Walker, 1 992). The main purpose of this paper is to re-examine the allostratigraphical (NACSN , 1983) position of the sand body (Figs 1 & 2 ) , and to reinterpret the incised valley fill as estuari ne. Ricinus is a small oil field, discovered in 1969.

The Ricinus sand body in the Cardium Formation of Alberta (Figs 1 & 2) is over 50 km long, at least 10 km wide, and up to about 30 m thick (Walker, 1985 ) . I originally interpreted it as a channel fill, with channel cutting and filling by turbidity currents. This interpretation has always been difficult to reconcile with the sedimentology of the other facies in the Cardium Formation, which consist largely of

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

47

48

R.G. Walker

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Cardium fields associated with the E4 (fields outlined) and E5 (black) bounding discontinuities. Note location of Ricinus close to and partly straddling the edge of the deformed belt (thrust symbol, inset lower left). The main map shows locations of wells and cores. All of the wells immediately east of Ricinus are shown, but only a few Caroline wells have been used. Note also the locations of cross-section A-A' through to G-G'. Number at right indicates Townships, and number along top indicates Ranges west of the 5th Meridian (114°).

Fig. 2. Stratigraphy of the Cardium Alloformation in the Ricinus area, drawn roughly to scale vertically. Bounding discontinuities E2, E3, E6 and E6.5 cannot be recognized here. Sandier upward successions A-D are assinged to the Hornbeck Allomembcr (note that succession D is commonly truncated by E4). E5 RSE (regressive surface of erosion) defines the base of the Ricinus Allomember, and E5 TSE (transgressive surface of erosion) defines the top. There is some evidence that E5 TSE cuts downward progressively southwestward. IT and RT refer to initial and resumed transgressions, as discussed in the text.

Cardium Formation incised valley

Initial oil-in-place was estimated as 2.88 x 106m3 (18. 1 million bbl) , with initial recoverable oil esti­ mated at 2.27 x 106m3 ( 14.3 million bbl) (Geological Survey of Canada, 198 1 a , b). However, the signifi­ cance of the field goes well beyond its minor econ­ omic potential. It is the only incised valley known in the Cardium, which is surprising in a formation dominated by eight regionally extensive erosion surfaces, each controlled by major fluctuations of relative sea-level. The incised valley interpretation is complicated by the fact that the valley is straight, at least 50 km long, and is parallel to regional strike (Fig. 1 ) . The new interpretations are based upon all o f the data available as of April 1993 , as illustrated in seven new cross-sections presented here (see Figs 3-9 and Appendix 1 ) . There are 172 on-field wells, of which 71 are cored (Fig. 1 ) . All measurements are given in metric units, except where specific reference is made to wells where the well logs and cored intervals were originally recorded in feet.

INTRODUCTION TO THE CARDIUM ALLOFORMA TION

The Cardium Alloformation (NACSN , 1983) is Upper Cretaceous (Turonian - Coniacian), and occurs in the Alberta foreland basin. This basin is part of the Western Interior Seaway, which was open from the Boreal Ocean to the Gulf of Mexico during the Turonian . Sediment was supplied from the active cordillera in the west. The Cardium is about 100 m thick, and is dominated by open-marine mudstones and storm­ influenced sandstones with hummocky cross­ stratification (HCS). There is one prograding shoreface succession (Kakwa Allomember; Plint & Walker, 1987) overlain by non-marine deposits (Musreau Allomember), and several examples of incised transgressive shoreface deposits. The relationships of these various deposits have been documented by mapping eight bounding discon­ tinuities, at least five of which can be traced basin­ wide (Piint et a/. , 1986; Wadsworth & Walker, 199 1 ; Walker & Eyles, 199 1). These were originally termed erosion surfaces and lettered El- E6, E6. 5 and E7. The regional trend of the various shoreface sand bodies is consistently NW-SE (Fig. 1 ) , parallel to the active cordillera to the west of the foreland basin.

49

Structural problems at Ricinus

The exact morphologies of the Ricinus valley and sand body are difficult to reconstruct because Ricinus lies at the eastern limit of the Foothills deformed belt (Figs 1 , 10 & 1 1 ) . The sand body is broken by several SW-NE trending faults, interpreted to have formed in response to different distances of thrusting toward the northeast. The fault positions have been located mainly by offsets in the northeastern margin of the field, which can be fairly well constrained by the close juxtaposition of wells with 'Ricinus' well­ log responses versus 'off-field' log responses (Figs 1 & 3-7). Within Ricinus, the sand body is repeated by thrusting in many wells (see Fig. 5 ) . In a few wells the dips exceed 45° and are locally vertical. The former turbidite interpretation

The interpretation of Ricinus as a channel cut and filled by turbidity currents (Walker, 1985) was based on the presence of very thick structure less sandstones and the presence of many beds characterized by a Tabc and Toe internal structure (Bouma, 1962). In the early 1980s, the possibility of turbidity currents in the Western Interior Seaway was suggested by the need to explain how a very large number of sharp­ based HCS beds had been emplaced well below fair­ weather wave base. Thinking was also influenced by the occurrence of well-defined turbidites overlain by HCS beds in the Jurassic Passage Beds in the Banff area. Here , both the turbidites and HCS beds have identical palaeocurrent orientations (Hamblin & Walker, 1979), suggesting emplacement of the HCS beds by turbidity currents, but above storm wave base. Turbidite ideas were first applied to the Cardium before the presence and significance of sea-level changes in the Cardium had been documented (Piint et a/., 1986), and before much of the important recent work on estuarine deposits had been pub­ lished (Dalrymple et a/. , 1992). The main purpose of this paper is to suggest an alternative to turbidity currents for cutting and filling the incised valley at Ricinus. Allostratigraphy

The most useful subdivision of the Cardium is allostratigraphical (NACSN , 1983) rather than lithostratigraphical (Fig. 2). An allostratigraphical unit 'is a mappable stratiform body of sedimentary

50

R.G. Walker

rock that is defined and identified on the basis of its bounding discontinuities ... [one may thus define] as single units discontinuity-bounded deposits charac­ terized by lithic heterogeneity' (NACSN , 1983 , p. 865). I regard the Cardium Alloformation as a formally defined stratigraphical unit, and have appended some thoughts on allostratigraphy in Appendix 2. The Cardium Alloformation has been subdivided into 14 allomembers, with eight well­ defined bounding discontinuities (E1 through to E7; Plint et al. , 1986) . Throughout this paper, informal use of the term Cardium (without a modifier) implies the Cardium Alloformation. The term Cardium Formation will be used only to designate the litho­ stratigraphical unit.

regionally extensive sheet-like sandier upward successions of the Hornbeck Allomember (Fig. 2). These successions are labelled A through to D in Figs 2-8 and collectively form the datum in these cross-sections. Succession C is cored in wells 7-9-36-8W5 and 10-17-34-7W5 (Fig. 4; all wells lie west of the 5th Meridian , and the W5 will be omitted from all subsequent well numbers cited). Succession C consists of bioturbated sandy mudstones that gradually become sandier upward; the top consists of a 1-cm-thick layer of chert grains about 0.5 mm (7-9-36-8) or 3 mm (10-17-34-7) in diameter. This sandier upward succession, capped by a gritty layer, is similar to those described by Walker & Eyles ( 1 988) from the Raven River Allomember; suc­ cessions A, B and D have not been cored but are probably similar to succession C. The sand bodies that form reservoirs at Garrington, Crossfield and Caroline (E4 fields in Fig. 1; Pattison & Walker, 1992) are assigned to the Burnstick Allomember. They rest in NW-SE trending steps on the E4 erosion surface (Fig. 2) , and are inter­ preted as transgressive shorefaces. Each incised shoreface is underlain by a surface of initial trans­ gression (IT in Fig. 2) , and is truncated by a surface of resumed transgression (RT in Fig. 2). An RT surface can pass southwestward into an IT surface if and when the transgression pauses and another still­ stand shoreface becomes incised. An example can be seen in well 7-33-36- 10, section A (Fig. 3 ) , where

OUTLINE OF CARDIUM STRATIGRAPHY AND SEDIMENTOLOGY

The base of the Cardium Alloformation is taken at bounding discontinuity El (Figs 2 -6). In the Ricinus area E1 has never been cored, and it is probably a correlative conformity (see Appendix 2) rather than a surface with demonstrable erosion . Its log identi­ fication is compatible with the E 11og identifier used by Plint et al. ( 1 986) and Wadsworth & Walker ( 199 1 ) . Erosion surfaces E2 and E3 cannot be ident­ ified in this area. Surface E4 has demonstrable erosion (Figs 3, 5 & 6), and cuts into a series of

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51

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infiltrate? Comparison of the size of clasts inferred to have been rolled on the bed with the size of the sand in the matrix is sometimes used as a starting point (Harms et al., 1982) . The application of this approach to shoreface conglomerates is difficult because: (i) shallow marine settings are typified by combined flows (especially during significant sedi­ ment transport events such as storms) and criteria for the threshold of movement under such conditions are not easily established (especially for gravels), and (ii) this approach is probably not valid for high sediment concentrations or transport rates such as might typify the surf zone. Textural evidence can sometimes provide clues as to the origin of the matrix. For example, in the Cardium Formation at Bay Tree, a sandy matrix is common in clast-supported shoreface conglomerates but is typically absent from adjacent finer (granule or fine pebble) conglomerates (Fig. SA) . This may suggest that the sand was able to infiltrate the larger pore throats of the coarser deposits but not through the more restricted openings between smaller clasts.

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Fig. 7. Longshore trends in D10 in shoreface conglomerates of Cardium Formation at: (A) Bay Tree and (B) Elephant Ridge (Mount Niles), showing a weak, southeasterly fining trend, which, based on regional palaeocurrent data (Hart, 1990; Hart et a/., 1990) is the dominant longshore transport direction. Each 'observation' at Bay Tree corresponds to a D10 value calculated from individual beds in each section. The D10 values at Elephant Ridge represent averaged values for each measured section. The location of sampling points is shown in Fig. 4A & B .

mode (see next section) must also be considered. Carter & Orford (1991) emphasized that episodic supply can induce textural changes that are inde­ pendent of process regime. The origin of a sandy matrix in a conglomerate can be problematic. Was the sand transported and deposited with the gravel , or is the sand a later

Many studies have suggested that form indices can be useful for distinguishing between 'beach' and 'fluvial' pebbles (Cailleux, 1945; Dobkins & Folk , 1970). The study by Dobkins & Folk ( 1970) suggested that the 'maximum projection sphericity' and 'oblate- prolate index' were useful discrimi-­ nators for distinguishing between beach and river clasts and several studies have since investigated the usefulness of these criteria (partial summary in Hart 199 1 ) . Hart (199 1 ) , using pebble shape data from Bay Tree and Chesil Beach , concluded that no form parameter derived to date can be used to unambigu­ ously distinguish shoreface or even beachface gravels and conglomerates from fluvial deposits. Thus, although pebble-shape segregation does develop in fluvial and beachface environments, and pebble zonation does develop on individual beaches (Orford, 1975), shape criteria alone cannot be con­ sidered diagnostic of depositional environment . Although Howard ( 1 992) was able to show statistical differences between shape indices of ancient fluvial and beach gravels, considerable overlap was present between fields and none of the 'cri tical values' derived from studies of modern gravels provided a successful means of differentiating beachface conglomerates.

Gravelly shorefaces

85

Fig.

8. Clast-supported pebbly shoreface conglomerate. (A) Bay Tree outcrop; note open framework portions. (B ) Core ( 10-18-63-5W6) note crude stratification and possible imbrication to right.

B EDFORMS AND SEDIMENTARY STRUCTURES

In most ancient deposits, sedimentary structures provide the key information for palaeoenvironmen­ tal reconstruction. Knowledge of the types of bed­ forms and morphological elements in analogous modern environments is therefore critical . Unfortu­ nately little is known about the variety of bedforms that may form on gravelly shorefaces. Shoreface

Gravel wave ripples (Fig. 9A & B ) , formed by oscillatory water motion are excellent indicators of subaqueous deposition at depths above storm wave base but can be formed at considerable depths, well beyond the limit of the shoreface. Leckie (1988) reviewed existing literature on the generation and environmental significance of 'coarse-grained' rip­ ples. Although most examples appear to be from shelf settings, gravel wave ripples have been reported from modern (Shipp , 1984; Hart & Plint, 1 989} and ancient (Leithold & Bourgeois, 1984; Massari & Parea, 1988} gravelly shorefaces. Crestlines tend to

be nearly shore-parallel due to wave refraction (Fig. lOA}, and both onshore and offshore migration (suggested by asymmetric profiles) have been reported. Gravel wave ripple form sets may be preserved , especially when the bedform is draped by finer-grained sediment. Cross-bedding may be produced by migration of the bedform under the influence of asymmetric wave motion or wave motion with weak superimposed unidirectional currents. Bi- and polymodal fabrics with modes oriented in onshore and offshore directions may indicate the presence of shore-parallel symmetrical gravel wave ripples (ct. Leckie & Walker, 1982) even if form sets are not observed (Hart & Plint, 1989 ) . Alternatively, migrating gravel wave ripples may produce fabrics with dominant dips in the direction of ripple migration (Hart & Plint, 1989; Fig. lOB) . In the absence of preserved form sets, it may be difficult to distinguish landward dips generated by onshore migration of gravel wave ripples from those produced in gravels by strong unidirectional offshore flows (cf. Cheel & Leckie, 1992) . Nearshore bars of various types are also formed by wave processes. These bars develop on gently dipping shorelines which tend to be sand-dominated

86

B.S. Hart and A. G. Plint

Fig.

9. Coarse-grained wave ripples developed in poorly sorted coarse sandy conglomerate; Cardium Formation on Elephant Ridge. (A) Symmetrical crestlines and (B) interference ripples. Scale bar in centimetres.

(Carter, 1988). Large, broadly convex upward 'macroforms' apparently representing nearshore bars have been reported from shoreface successions comprising sandstones, pebbly sandstones and con­ glomerates by Leithold & Bourgeois (1984), Massari & Parea (1988) and Massari eta!. (1986). Such forms are found in the upper shoreface (Davidson-Arnott & Greenwood , 1976) and may or may not migrate. Migratory bars centimetres to decimetres in height may move onshore and (particularly in tidal settings) weld themselves to the beachface, typically produc­ ing onshore oriented tabular to sigmoidal cross-beds

capped by planar stratified sandstones or pebbly sandstones (Fig. llA & B ) . Troughs associated with longshore bars have been recognized (Leithold & Bourgeois, 1984; Massari & Parea, 1988) as broadly lenticular conglomerates with concave upward bases cut in sandstones. Such troughs should be elongate and nearly shore-parallel. Ross & Long (1989) cored a modern nearshore bar system in a mixed gravel and sand setting and found that gravel was not restricted to the troughs, but could also be found on the bar crest. The presence of longshore bars and troughs from ancient gravelly

Gravelly shorefaces

87

8

A

Mean x-bed orientation

Mean crestline orientation

Crestline orientation

Fig.

10. Summary of palaeocurrent information from Elephant Ridge section of Cardium Formation (Fig. 4B) . (A) Coarse-grained wave ripple crestlines (a) and decimetre-scale cross-bedded conglomerates (b) . Wave ripples indicate NW-SE trending shoreline, in agreement with other data from the Cardium Formation in ths area (Hart, 1990) . (B) Pebble fabric from a near-symmetrical coarse-grained wave ripple at Elephant Ridge suggesting onshore migration of the wave ripple form.

shoreface deposits suggests that steep, 'reflective' profiles (typical of modern transgressive gravelly shorefaces) were not developed at the time of depo­ sition (Wright & Short, 1984). Broad, pebble-lined scours that are oriented per­ pendicular, or nearly so, to the palaeoshoreline in shoreface sandstones may be interpreted as rip-current channels (e.g. Fig. 6C; Leithold & Bourgeois, 1984; Massari & Parea, 1988; Gruszcyzynski et a/., 1993 ) . Shell debris may form a significant component of such lags, facilitating the recognition of depositional environment. Cross-bedded sandstones and conglomerates are common in shoreface deposits, and may be oriented onshore, offshore or alongshore (Dupre et al., 1980; Leithold & Bourgeois, 1984; Massari & Parea, 1988). These structures may be the product of asymmetrical wave motion, and rip, longshore, tidal and (close to river mouths) fluvial currents . At the Bay Tree section of the Cardium Formation, cross-bedded conglomerates several decimeters thick (with cross­ bedding defined by size or textural variations) are present. This large-scale cross-bedding (Fig. 12A) was apparently produced by the migration of dune­ like forms at least 60 em high. Palaeocurrent evi­ dence indicates that these bedforms and similar cross-sets in other outcrops of the Cardium migrated in a shore-parallel direction (Figs lOA, 12B & 13A ) . Such bedforms have not been documented from modern gravelly shorefaces, although B ujalesky & Gonzalez-Bonorino (199 1 ) reported asymmetric 'washed-out (gravel) dunes' up to 10 em high with

wavelengths typically about 5 m and crests oriented at high angles to the shoreline from the lower beach­ face (intertidal) . These bedforms appeared to have been produced originally by longshore currents at high tide. At Bay Tree (Fig. 4A) over 50% of the conglomer­ ate is organized into crudely bedded, amalgamated units, 1-3 m thick. Centimetre to decimetre thick, massive clast-supported pebble conglomerate beds are the primary constituent of the amalgamated units, but open framework , imbricated, matrix­ supported conglomerates and thin sandstone beds (generally < lO cm thick) are also present. Shallow scours, depositional thinning and lateral facies changes limit the lateral extent of individual beds to a few metres. Despite the internal scouring, the overall appearance of the conglomerate, revealed especially by the sandstone beds, is of a pronounced horizontal stratification in the strike sections at Bay Tree (Fig. 14) . Presumably, 'clinostratification', such as that illustrated by Massari & Parea ( 1 988), would be observable in shore-normal sections. Most of these shoreface conglomerates can be interpreted as the products of gravel bedload sheets such as described from gravelly fluvial systems (Hein & Walker, 1977; Whiting et a/., 1988). The lateral dimensions of these 'bedforms' were probably of the order of several metres, explaining the lateral impersistence of individual conglomerate beds in the amalgamated deposits. The possible existence of this type of sediment transport has been suggested by Mathews ( 1980), who found that tracers on a

88

B.S. Hart and A. G. ?lint

Fig.

1 1. Migratory bars of upper shoreface and beachface. (A) Poorly sorted pebbly sand , Island View Beach, British Columbia ( B ) Sigmoidal, landward-directed sandstone cross-set in upper portion of Cardium Formation shoreface conglomerates at Bay Tree, Alberta.

Gravelly shorefaces

89

Fig.

12. Decimetre-scale cross­ bedding in Cardium Formation shoreface conglomerates. (A) Bay Tree (scale= 15 em) and (B) Elephant Ridge (pocket knife for scale) . In both cases, migration direction was alongshore, to the southeast.

modern gravelly foreshore moved alongshore on the lower beachface (i.e. in the surf zone at high tide) as discrete 'slugs' of sediment representing volumes of 1-10m\ rather than as a uniform gravel sheet. These 'slugs' moved 10- SOm alongshore during periods of high wave energy, but were reworked during fair-weather conditions. One may speculate that, during strong longshore flows, the bedload sheets may have grown in height and developed slipfaces, such as has been observed in modern gravelly streams (Whiting et at., 1988) . Because of the amalgamated (and hence indistinguishable) con­ tacts between individual beds ('flow units') , the use

of quantitative techniques such as plots of maximum clast size versus bed thickness or Markov chain analyses become impractical. Pebble fabrics from the shoreface portion of the conglomerate at Bay Tree are typically polymodal (Fig. 13B ) . In most cases, the principal mode is towards the south or southeast (alongshore) . Other modes are observed at about 90° and 180° to the principal mode in nearly all cases. All of these samples represent crudely stratified to cross-bedded clast-supported pebble conglomerates generally from the middle to lower portions of the section. The dominant dip direction is approximately that of

(A�

90

�

B.S. Hart and A. G. Plint (B)

(C)

(E)

(F)

�

�

� (G)

~

(Hl

(I)

� �

Fig.

13. Palaeocurrent data from Bay Tree section. (A) Cross-bedded sandstones and conglomerates. (B- I ) Pebble fabrics from conglomerates. Pebble fabrics are for sites labelled in Fig. 4A: (B) BTI-11 (beachface conglomerates); (C) BTI- 15 (plunge-step conglomerates); (D) BT4-2 (?wave rippled conglomerate) ; (E) BTI-5; (F) BT3-7; (G) BT3- 1 2; (H) BT5-9; ( I ) BT4- 13. See text for further description.

the dominant cross-bed orientation in the sand­ stones. This suggests a downcurrent dip of clasts lying on the lee side of gravel bedforms transitional

between bedload sheets and dunes. Massari & Parea (1988) also reported the presence of polymodal fabrics from shoreface conglomerates but did not elaborate on the potential significance of their obser­ vations. They also indicated that rose diagrams of clast long axes from upper shoreface deposits could be isotropic, polymodal, bimodal or unimodal. Clifton (1973) suggested that pebble segregation and the lenticularity of conglomerates, although not necessarily diagnostic, were useful criteria for distinguishing between wave-worked and fluvial gravels. Wave-worked gravels tend to be more 'lat­ erally regular' than lenticular, owing to the uniform­ ity with which waves act over large areas. He also suggested that winnowing of a deposit by waves could segregate sand- and gravel-sized material . Sub­ sequent studies (Leithold & Bourgeois 1984; Massari & Parea 1988; Hart & Plint 1989) have tended to confirm Clifton's textural and stratigraphical obser­ vations. It should be reiterated, however, that rip­ current deposits and longshore trough conglomerates may both have lenticular geometries, providing that they were buried below the influence of wave rework­ ing. Lenticular, channelized conglomerates may also be present in river mouth settings (Kleinspehn et al., 1984; Fig. 4B). Planar cross-bedded conglomerates several deci­ metres thick at the top of some sections at Bay Tree (Fig. 15) probably record the progradation of a gravel plunge step at the toe of the beach (low­ tide level) -a feature common on coarse-grained beaches (Forbes & Taylor, 1987; Massari & Parea,

Fig.

14. Pronou nced horizontal stratification in shoreface conglomerates at Bay Tree. Stratification emphasized by sandstone interbeds.

Gravelly shorefaces

91

Fig. 15. Planar tabular conglomerate cross-set at Bay Tree, i nferred to have been produced by the offshore progradation of a plunge step at the base of the beachface. From top of section BT2 ( see Fig. 4A) .

1988; Hart & Plint, 1989) . Somewhat larger steps are reported from some gravelly shorefaces (Kirk, 1980; Carter, 1988; Massari & Parea, 1988; Postma & Nemec, 1990) . A pebble fabric from the plunge­ step deposits at Bay Tree shows offshore dipping pebbles (Fig. 13B) . B ased on (unpublished) obser­ vations at Chesil Beach, progradation of the plunge step seems most likely to occur either following high tide or after a storm surge when the water table in the beach is still high. At such times, backwash dislodges pebbles from the beachface and transports them seaward. These clasts then avalanche over the plunge step causing it to prograde. Body and trace fossils provide the clearest evidence of marine deposition (Clifton, 1973; Kleinspehn et a/., 1984; Leithold & Bourgeois, 1984) . Unfortunately, neither are common in high­ energy, conglomerate-dominated shoreface deposits (Leithold & Bourgeois, 1984; Massari & Parea, 1988; Hart & Plint, 1989 ) . Trace fossils are rare in conglomeratic shoreface deposits of the Cardium Formation, although traces are present locally in thin interbedded sandstones (Fig. 16) . Although gravels are generally inhospitable to infaunal macro­ benthos, we have observed (using SCUBA) crab burrows in gravels on the Chesil Beach shoreface. However, the preservation potential of these bur­ rows seems very low owing to the lack of a contrast­ ing sediment fill. In general , the abundance of trace fossils may be related directly to the percentage of

Fig. 16. Burrowing ( Diplocraterion) in sandstone within shoreface conglomerates, Cardium Formation (core from 16-19-5 1 - 1 0W6, scale bar= 5 em).

92

B.S. Hart and A. G. Plint

sand, partly because this forms a more hospitable substrate, and partly because sand is better able to preserve traces of burrowing. Beachface

The structure of beachface gravels and conglomer­ ates is well established (Maejima, 1982; Nemec & Steel, 1984; Forbes & Taylor, 1987; Massari & Parea, 1988; Hart & Plint, 1 99 1 ) . Well-stratified, laterally continuous gravels and interbedded sands that dip gently (typically S0-20°) in a seaward direction are

considered diagnostic of gravelly beaches (Fig. 17 A & B ) . In examples from the Cardium Formation, individual gravel ·laminae/beds can be normally or inversely graded. Sand tends to be well-segregated from gravel, at least in high-energy systems , and commonly forms 'sand runs' in the intertidal zone (Fig. 18A) . Beach­ face dips are generally inversely related to the per­ centage of sand in the system and directly related to grain size. Shape sorting and seaward-dipping pebble imbrication (Fig. 18B) may be well developed. Landward-dipping stratification and imbrication may

Fig.

17. Beachface deposits. (A) Gently dipping interstratified beachface conglomerates and sandstones, Bay Tree. Note low­ angle truncation surface. (B) Gently dipping stratification in beachface conglomerates, Cardium Formation, Cutpick Hill, Alberta.

Gravelly shorefaces

93

Fig.

18. Modern beachface gravels at Chesil Beach, England. (A) Good segregation of sand- and pebble-sized constituents on beachface. (B) Offshore (to right) imbrication of beachface gravels. Shape sorting is not well developed at Chesil Beach.

be present in cases, and are apparently the result of berm accretion (Maejima, 1982; Forbes & Taylor, 1987 ) . Maejima ( 1 982) noted that broad , concave­ upward scour surfaces may represent the effects of beach cusp formation. Massari & Parea ( 1988) found that beachface conglomerates were internally trunc­ ated in places by low-angle erosion surfaces inter­ preted to represent storm-wave planation of the beachface (Fig. 17A). Enigmatic 'high-angle scours' were described by Leithold & Bourgeois ( 1 984) and Bourgeois & Leithold ( 1 984) from gravelly sandstone 'shoreface' deposits; similar structures have been observed in

outcrops of the Cardium Formation (Fig. 19). The scours may have steep walls (in places stepped or undercut) that scour into hummocky/swaley cross­ stratified and planar laminated sandstone. Multiple phases of fill can be inferred for some units, and sandstone 'intraclasts' , woody debris and shell frag­ ments are observed in places. Leithold & Bourgeois ( 1 984) suggested very rapid cut-and-fill (nearly instantaneous) at shallow depths just seaward of the breaker zone. An alternative interpretation is that the scours are the product of small streams incised into the beachface. Intuitively, it seems unlikely that high-angle scours can be

94

B.S. Hart and A. G. Plint

Fig.

19 Conglomerate filling �teep­ sided scour cut into planar­ laminated, fine-grained sandstone ?beachface deposits of Cardium Formation along Murray River, British Columbia. The scour may have been cut subaerially by a small stream crossing the beach, or may be the product of subtidal scour by rip-currents (Chiocci & Clifton, 1991) . See text for discussion.

maintained in non-cohesive sandy substrates in a subaqueous setting, especially in the presence of intense wave activity. However, Gruszcyzynski (pers. comm . , 1993) reported seeing (using SCUBA) vertical-sided, rip-current channels cut into sandy sediment, and speculated that bacterial or algal binding may have rendered the sediment cohesive. Alternatively, high-angle walls (locally overhanging) can be cut in loose sand, and maintained in subaerial settings by capillary forces (Fig. 3A) . Cohesion may be promoted by salt cementation (in a spray zone j ust above the limit of the swash) and by organic films. The example shown in Fig. 19 is from planar stratified sandstone, interpreted as beachface , which directly overlies swaley cross-stratified (shoreface) sandstone, the stratigraphical position of which, and relation to relative sea-level changes, is well con­ strained (Hart & Plint, 1993a) . Facies descriptions of wash over deposits have been provided by Carter & Orford ( 198 1 ) and Forbes & Taylor (1987) . Massari & Parea ( 1988) have described washover deposits from Messinian and Pleistocene gravelly. shorelines.

FACIES SUCCESSIONS

Stochastic variations in storminess, sediment supply, relative sea-level and other factors probably account for much more of the variation in sedimentary struc­ tures and sediment textures than is commonly appreciated . The morphology of gravelly shorefaces

and beachfaces can be very sensitive to textural factors such as the amount of sand in the· system. The effects can be non-linear: process affects mor­ phology which in turn affects process. In this respect , it may be overly simplistic to attribute facies vari­ ations at any given site to short-term changes in energy level (e.g. storm -non-storm) , sea-level , etc. Modern gravelly beachfaces show much intersite variation and , in places, much temporal variation at any given site. Facies models based on 'snapshots' of any particular beachface or shoreface are therefore clearly of limited applicability. It is equally clear that sedimentary facies and textural or pebble shape criteria derived from the beachface must not be applied when attempting to identify shoreface conglomerates. Several previous authors have presented sum­ maries of criteria for distinguishing between fluvial and shoreface/beachface gravels (Clifton, 1973; Leckie & Walker, 1982; Ethridge & Wescott, 1984; Nemec & Steel, 1984; Reddering & Illenberger, 1988). Based on these studies, and our own studies of shoreface and beachface gravels and conglomer­ ates, an attempt is made in Table 1 to outline the expected character of preservable gravelly shoreface and beachface deposits as a function of several vari­ ables. The factors that appear to be of greatest significance are: (i) change in relative sea-level (direction and magnitude) , (ii) textural character-­ istics of the sediment, including proportion of sandl in the system , (iii) proximity to sediment source (river mouth) ; and (iv) orientation of the shoreline

95

Gravelly shorefaces Table 1.

Summary of expected preservable characteristics of gravelly shoreface and beachface deposits Geometry and stratigraphical position

Transgressive

Processes

Facies

Thin lags continuous with flooding surfaces Topographically 'trapped' shoreface deposits

Barrier rollover Rafting

Thin lags Wave ripples and storm beds

Lobate (plan) Lenticular (cross-section) Correlate with other progradational deposits

Rivers/waves/tides Slope failures

Lenticular beds Mixed palaeocurrents Mix of fluvial and marine facies

Mixed sand and gravel

Tabular (cross-section) ?Shore-parallel elongation Correlate with other progradational deposits

Waves/tides - shoreface Swash/backwash - beachface

Longshore bars/troughs Rip-current scours Palaeocurrents reflect processes Segregation of sand and gravel increases with wave energy

Gravel - drift aligned

Tabular (cross-section) Shore-parallel elongation Correlate with other progradational deposits

Longshore currents dominant

Weak shore-normal zonation Bedload sheets/dunes Palaeocurrents shore parallel

Gravel - swash aligned

Tabular (cross-section) ?Beach ridges Shore-parallel elongation Correlate with other progradational deposits

Onshore/offshore motions

Pronounced shore­ normal zonation Wave ripples Onshore/offshore palaeocurrents

Regressive River mouth

with respect to dominant wave-approach direction. It is expected that in most instances transgressive gravelly shorelines will leave little permanent record of their existence. Thin pebbly lags representing abandonment of isolated clasts during shoreface retreat, and possibly organic rafting seaward of the shoreface (see Leithold, 1989; Woodborne et a!., 1989; Liu & Gastaldo, 1992), are the expected litho­ logical expression . Thicker successions might be preserved where the shoreface has 'backed up' against pre-existing relief ('bevels', cliffs, etc . ; see Bergman & Walker, 1987, 1988; Hart & Plint, 1993b) . Stratigraphically, these deposits will correlate with marine flooding surfaces (Plint et a!. , 1986) . Progradational packages are likely to represent regional shoreline regression, although they may also develop in areas of high sediment supply during relative sea level rise (e.g. modern fan-deltas) . Ideally the shoreface conglomerates will overlie shelf deposits and , in turn, be overlain by beachface conglomerates. The latter tend to be quite distinctive

both in outcrop and in core and the relative strati­ graphical position can be used to help infer a shore­ face origin for the underlying deposits. Non-marine facies may or may not be present above the beach facies. Roughly tabular geometries are expected for indi­ vidual progradational packages, except close to river mouths where lenticular geometries may occur (Fig. 4B) . These river-mouth deposits will reflect the interaction of fluvial, tidal and wave-induced cur­ rents. Rapid deposition in the transition zone will generate poorly sorted deposits and , potentially, instability related gravity flows (Kleinspehn et a!., 1984 ) . Channelized and burrowed deposits may be found in close proximity. Away from river mouths, the influence of marine processes will become more apparent. In 'mixed' settings with appreciable amounts of sand, segre­ gation of sand and gravel constituents should be proportional to the wave energy level of the environ­ ment. As a general rule , the potential for finding

96

B.S. Hart and A. G. Plint

marine trace fossils increases with the amount of sand in the deposit. Longshore bar systems and rip­ current deposits may be distinguishable (Leithold & Bourgeois, 1984; Massari & Parea, 1988; Gruszcynski et al., 1993). Studies of modern gravelly barriers have led to a distinction between swash-aligned and drift-aligned barrier systems, each with distinctive process­ morphology characteristics (Carter & Orford, 199 1 ) . Swash-aligned barriers are oriented nearly perpen­ dicular to the direction of wave approach and are dominated by onshore-offshore motions. Drift­ aligned barriers are those oriented at an angle to the direction of wave approach, and are dominated by longshore transport. Shore-parallel facies zonation is best developed in swash-aligned beaches and are 'smeared' in drift-aligned systems (Carter & Orford, 199 1 ) . Gravelly shoreface deposits tend to appear massive, with crudely developed stratification and sedimentary structures revealed by subtle textural variations (Bergman & Walker, 1987 , 1988; Hart & Plint, 1989, 199 1 ; Plint & Hart, 1988). By careful study of palaeocurrent information, both in the conglomerates and in interbedded sandstones, it should be possible to recognize the dominant direc­ tion of wave approach (Hart & Plint, 1989 ) .

SUMMARY

At present, sufficient criteria exist to permit the identification of gravelly shoreface deposits, at least in outcrop. However, few truly 'diagnostic' criteria exist, with faunal or trace faunal evidence being possible exceptions. The inadequacy of previously proposed diagnostic criteria (such as shape para­ meters) reflects the range of processes that can affect the character of shoreface conglomerates. In outcrop, the stratification style, sedimentary structures and palaeocurrent information can be powerful indicators of depositional process. In the subsurface , conglomerate body morphology, orien­ tation and stratigraphical relation to other units may aid interpretation. In both outcrop and subsurface , detailed facies analysis should make it possible to distinguish subtypes of gravelly shoreface deposits by clarifying the nature and relative importance of specific depositional processes (e.g. wave-induced and gravity driven currents, etc . ) . More studies are needed t o document the range of sedimentary processes and structures present on modern gravelly shorefaces. To do so will require

the development of innovative sampling and moni­ toring techniques. Difficulties encountered with establishing the threshold of grain movement and bedload transport rates under combined flows typical of sandy shoreface settings are doubly present on gravelly shorefaces, where clast morphology, bulk textural parameters and fabric play important roles. The establishment of phase diagrams for bed con­ figurations remains problematic even in sandy shore­ face settings, and extremely little documentation is available for gravelly shoreface settings. Empirical investigations of ancient deposits may provide, for some time to come, the best insight into the processes operative on modern gravelly shorefaces.

ACKNOWLEDGEMENTS

The work on shoreface conglomerates of the Car­ dium Formation which forms the nucleus of this paper was completed during the senior author's doctoral studies at the University of Western Ontario. Funding for that study was made possible through research grants (to A . G . P . ) from the Natu­ ral Sciences and Engineering Research Council, the Department of Energy, Mines and Resources and the University of Western Ontario. Additional logis­ tical support was provided by Canadian Hunter Ltd . , Esso Resources Ltd . , Home Oil Ltd. and Unocal Canada Ltd. To all these agencies and companies, we are very grateful . Helpful comments by reviewers E. Leithold and F. Massari are greatly appreciated. We thank the Canadian Society of Petroleum Geol­ ogists for permission to reproduce Figs 4A, SA, 9A and 12A, originally published in Plint & Hart ( 1988). This paper was written during the tenure of a Visiting Fellowship by the senior author at the Geological Survey of Canada's Pacific Geoscience Centre.

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97

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99

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Spec. Pubis int. Ass. Sediment. (1995) 22, 101-135

The return of 'The Fan That Never Was': Westphalian turbidite systems in the Variscan Culm Basin: Bude Formation (southwest England) R OB E R T V . B U R N E Australian Geological Survey Organisation, P O B ox 378, Canberra, ACT 2601, Australia

ABSTRACT

The Westphalian Bude Formation of southwest England was one of the first successions to be interpreted as a subsea fan deposit. This interpretation has been questioned recently, and the alternative depositional environment of a wave-affected shelf has been proposed. The tectonic setting, palaeo­ geographical relationships and facies succession of the formation have been reassessed in an attempt to resolve this controversy. The formation was deposited in the synorogenic Culm Basin. It contains thick sandstone units, but is in other respects similar to the underlying turbidite-bearing Crackington Formation (Namurian-Westphalian) . There is no depositional continuity with the contemporaneous paralic and deltaic deposits of the Bideford Group that occur in a tectonically distinct succession near Westward Ho! Four facies are identified in the Bude Formation; the black shale facies, the muddy siltstone facies, the interbedded sandstone-shale facies and the sandstone-dominated facies. Both thickening upward and thinning upward facies successions occur at some levels, but generally the facies succession shows less predictable, although not random, alternations between the four facies. The suggestion that the Bude Formation represents the deposits of a wave-influenced shelf does not stand up to scrutiny. There is no substantial evidence for shallow water environments in the Formation . Evidence for storm-generated hummocky cross-stratification and combined-flow wave ripples is equivocal. The ichnofacies have no depth connotation and there is no evidence for sedimentological continuity with undoubted deltaic successions. By contrast there is overwhelming evidence for deposition of turbidite systems in an isolated fresh- or brackish-water basin. The black shale facies represent the deposits of fine-grained turbidity currents in an anaerobic or dysaerobic basin. The muddy siltstone facies represent the deposits of aerobic environments characterized by gentle currents, possibly fan-levee environments. The interbedded sandstone-shale facies consist of a variety of turbidites and were deposited in either levee, lobe or interchannel environments. The sandstone-dominated facies were deposited in fan-valley environments, and include channel-fill successions. It is concluded that the Bude Formation was deposited as subsea fans on the northern, inactive margin of a land-locked, foreland basin.

INTRODUCTION

The Westphalian Bude Formation (Owen, 1934; King, 1966, 1967, 1971; Freshney et al., 1979) was deposited in the synorogenic Culm Basin of south­ west England (Sedgewick & Murchison, 1840; Ussher, 1892; Thomas, 1988) (Figs 1 & 2). The formation crops out along the coast of north Cornwall and west Devon (Fig. 2) and it had become the custom to equate it with a contemporaneous succession, the Bideford Group, which crops out on

the coast near Westward Ho! in north Devon (Figs 2 & 3). Reading (1963) drew attention to the funda­ mental difference of opinion about the deposition of these successions. Owen (1950) and Prentice (1960a, b) considered that they were examples of paralic 'Coal Measures', but Ashwin ( 1957, 1958) interpreted them as turbidites. Reading ( 1963, p. 69) contrasted the evidence in the rocks around Westward Ho! for 'successive advances into a basin

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

101

102

R. V. Burne A 52 '

----

8

5 1'

ST

50' 6'

D [] �

Permian- Tertiary Upper Carboniferous Devonian - Lower Carboniferous

3'

• I+++I

Lizard and Start complexes

----&....- Thrust fault -- Normal fault

Granite

�

BCFZ

Bristol Channel fracture zone

SLFZ

Sticklepath - Lustleigh fracture zone

Strike- slip fault

Fig. 1. Location of the Culm Basin. (A) Location of southwest England with respect to the British Isles. (B) Location of the Culm Basin with respect to the other principal tectonic elements of southwest England (after Hartley & Warr, 1990). Suggested Silesian position of southwest England with respect to France produced by pre-deformation restoration along the Bristol Channel-Bray Fault (after Holder & Leveridge, 1986), all other structures shown in present-day positions.

of a low coastal plain, possibly fronted by deltas, alternating with periods of reduced supply when basin conditions were re-established' with the fact that 'No cyclic pattern of sedimentation has been recognized, nor has any unequivocal evidence of shallow water or coastal plain deposition been observed' in the Bude Formation. This implied that 'whilst coastal plains reached into north Devon in Lower Westphalian times, there is no evidence that they extended into Cornwall'.

Subsequent workers confirmed the existence of paralic cycles in the Bideford Group, (comprising the Northam and Abbotsham Formations), around Westward Ho! (Walker, 1964a, b; de Raaf et al. , 1965; Money; 1966; Elliott, 1976). They have also confirmed Reading's statement that (1963, p. 69) 'The precise depositional environment of the Bude Sandstones is not easy to establish'. Sedimentologists had generally recognized that the Bude Formation was, in part, a turbidite succession, but did not

103

Bude Formation subsea fans '40

4"20'

'30

'20

Bideford

'30

Bay

Hartland Point

51"00'

'20

Bude

Bay

'10 50'50'

BUDE

·

.

p;.

.

'00

•

Post- Silesian strata

WESTWARD HO!­ BIDEFORD FAULT BLOCK

m t-=-� �

Greencliff Beds Bideford group & Westward Ho! Formation

CULM BASIN

D [·.:.·: ·_.:J

Bude Formation

]

--Fault

SLFZ

SILESIAN

�

Crackington Formation 1

Sticklepath - Lustleigh fracture zone

Town

D

Village

30

National grid number tor Great Britain

Pre- Silesian strata

Fig, 2, Geological map of north Cornwall and west Devon. (Modified after Burne & Moore, 1971; Thomas, 1988.)

appear to be a 'true flysch' (Goldring & Seilacher, 1971). It contained some unusual characteristics, such as massive sandstones, which de Raaf (pers. comm., 1966) had informally termed 'Budes' and interpreted as some form of unusually fine-grained 'fluxoturbidite' (Unrug, 1963, 1965). A sedimentological facies analysis of the middle Bude Formation (Fig. 4) was undertaken to identify features definitive, rather than indicative, of depo­ sitional processes and environments (Burne, 1969a, b; 1970, 1971, 1973, 1976; Burne & Moore, 1971). It was concluded that the Bude Formation had been deposited in a land-locked basin, akin to the present Black Sea or Caspian Sea, that had been isolated by Hercynian continental collision. The

basin was generally filled with fresh or brackish water, although rare marine incursions occurred, probably the result of glacio-eustatic oscillations. The mature, fine-grained sandstones were derived from a land mass of older sedimentary rocks to the north and were deposited as turbidites from bottom­ hugging underflows. The turbidites included con­ ventional Bouma sequences, traction carpet deposits with current-reworked tops, irregularly laminated sandstones, and 'slurried' beds produced by the impact of dense turbidity currents on a soft substrate. The major 'slumped' beds had the same origin as the 'slurried' beds, but some sandstone dykes and volcanoes were formed by load-induced post­ depositional water-escape. The depositional

104

R. V. Burne R

HARTLAND TO BOSCASTLE

a WESTWARD HOI

z "' :J "' I 0.. f­ rJ) UJ

Amaliae Marine Band

s:

Usteri Marine Band Marine

• Q D � �

Greencliff Beds

Bideford Group

Westward Ho! Formation

Bude Formation

Crackington Formation

INDEX SHALES

ccs

Clovelly Court Shale

DPS

Deer Park Shale

ESS

Embury Shale

GAS

Gull Rock Shale

HQS

Hartland Quay Shale

SMS

Sandy Mouth Shale

SPS

Saturday's Pit Shale

TCS

Toms Cove Shale

WGS

Warren Gutter Shale

Brigantian •oom

environment was that of a submarine fan, with open­ fan, fan-levee, active fan-channel, and inactive fan­ channel facies being distinguished. Despite the apparently random nature of much of the succession, it was recognized that predictable associations occurred in parts (Burne, 1969b, Fig. 20) (Fig. 5), including thickening upward successions, interpreted as being due to the advance of a fan-channel/levee complex over an open fan, and thinning upward successions, interpreted as being the result of fan­ channel fill and abandonment.

Fig. 3. Stratigraphical correlation between the Culm Basin and Westward Ho! Silesian sequences. (Modified from King, 1967 ; Thomas, 1988.)

Melvin (1976, 1986) reached similar conclusions from a study of higher levels in the Bude Formation (Fig. 4). He suggested that deposition was from turbidity currents in a relatively shallow basin and noted that sediment coarser than fine-grained sand was absent, and the organization of the bundles of sandstones was complex. He suggested that sediment was supplied from a delta to the north by resedimen·· tation of delta-front sands, and by slumping of delta·· front muds and silts to form a lower prodelta turbidite fan. Melvin could not identify well-developed

105

Bude Formation subsea fans

B § � D G []] � �.

UJ 0 ::;) "'

Nodular shale Shale Siltstone Mudstone Slumped bed Thick-bedded and massive sandstone characteristics of BudeFormation Shales with thin sandstone Medium to thinly bedded sandstones with subordinate shales and siltstones

INDEX SHALES

600

WGS Warren Gutter Shale 'Anthracoceras' aegiranum horizon

SMS

Sandy Mouth Shale

SPS

Saturday's Pit Shale (key shale W)

TCS

Tom's Cove Shale (key shale Q)

LS

Longpeak Shale

HQS

Hartland Quay Shale

GRS

Gull Rock Shale

Gastrioceras amaliae horizon

Gastrioceras /isteri horizon

ES

Embury Shale Gastrioceras subcrenatum

400

200

horizon

Fig. 4. Representative lithological section of the Crackington Formation and Bude Formation based on exposures between Duckpool (SS 201 1 15) and Embury Beach (SS 214 195) (modified after Freshney et a/., 1979). King (1967) originally defined the base of the Bude Formation at a level 265 m below the base of the Tom's Cove Shale and the top of the formation at a level 150 m above the same datum, a section here referred to as the middle Bude Formation.

thickening upwards and thinning upward successions as described by Burne (1969b) from lower in the Bude Formation. However, he presented a facies model of various fan, fan-levee and fan-channel relationships that accounted for the facies suc­ cessions observed in measured sections. Apart from different degrees of organization in the succession, there is general correspondence between the fan interpretation suggested by Burne (1969b) and that proposed by Melvin (1986). Higgs ( 1987) strongly criticized Melvin's inter­ pretation in a vigorous discussion provocatively titled 'The Fan That Never Was?'. Higgs (1983, 1984, 1986a, b; 1991) tentatively identified wave­ influenced structures in the Bude Formation, and pursued the significance of this interpretation to its logical conclusion, that the 'Bude Formation was deposited largely, if not entirely, above storm wave base, as indicated by the fact that wave-influenced

structures occur throughout the succession, spaced no more than a few metres apart.' This paper attempts to resolve the controversy that has resulted from these contrasting inter­ pretations by reconsidering the tectonic setting, palaeogeographic relationships and facies analysis of the Bude Formation.

THE TECTONIC SETTING OF THE CULM BASIN

The convergence and collision of the Laurussian and Gondwanan continental masses between the Devonian and the Early Permian (Burne, 1969b, 1973; Ziegler, 1990) generated a deformation front that migrated northwards with time (Besly, 1988) to form the Variscan orogenic belt. During the Middle Devonian and Early Carboniferous, volcanicity and

R. V. Burne

106

COMPOSITE SEQUENCE

Possible interpretation

Mudstone

Open fan or inter

dominated

-channel sediment

sequence

1-:: . 'Thinning upward'

--- ---

/

-�::--......_

�-=--

::-�--

:� ----------- ----:-:-,

�

sequence

---==- � -==-

Sandstone dominated

l

Sediment in confines of partly active channel

Sediment in confines ----�-�-

of active large

sequence

channels

-

'Thickening upward'

.. ·.:::.:.:_"·

Sediment of channel levee

sequence

Mudstone

Open fan or inter­

dominated

channel sediment

sequence

16/04117

I

Bed scale Approx. 1m (bu1 variable)

Bude Formation subsea fans

subsidence accompanied a period of extension in southwest England. Regional inversion of this Devonian to Early Carboniferous basin occurred in the early Namurian. Loading of its northern margin by the resulting thrust nappe development (Hecht, 1992) formed the Culm Basin that lay north of the developing Variscan Orogen and south of the cratonic Wales-Brabant Massif (Hartley & Warr, 1990) (Fig. 1). Evidence for a thin skinned thrust­ and-nappe regime implies that the orogen did not attain 'Alpine' proportions (Isaac et a/. , 1982) . Hartley & Warr (1990) suggest that the Culm Basin maintained its present geographical relationship with South Wales throughout the Variscan deformation. However, others suggest that the northern boundary of the basin corresponded to an area west of present­ day Le Havre, France (Fig. 1) (Holder & Leveridge, 1986), and subsequent transcurrent faulting along the Bristol Channel and Bray Fault has displaced the basin to its present position, south of the Bristol Channel. In the early Namurian the Culm Basin was prob­ ably part of a continuous deep-water flysch basin, the Rhenohercynian Basin, that extended from Cornwall to Upper Silesia (Ziegler, 1990) . Sedimen­ tation began to exceed subsidence in parts of this basin during the Namurian, and the Culm Basin became one of a number of basins isolated during the last stage of continental collision (Franke & Engel, 1988). In the Culm Basin evidence for this change may be seen in the turbidites of the Crackington Formation. Marine goniatites, which had been scattered throughout the shales lower in the succession, became restricted in the uppermost Namurian to marine bands deposited by marine incursions into an otherwise brackish-water basin (Freshney et a/., 1979). By the early Westphalian most Rhenohercynian basins had shallowed and established paralic conditions providing the classic localities for coal-measure cyclothems (Reading,

107

197 1) . The occurrence of correlatable marine bands throughout these basins indicates that they were separated by very low topographical relief, with the marine bands resulting from short-lived marine transgressions that entered the basins from the Moscow Depression (Ziegler, 1990). The trans­ gressions probably reflect high sea-levels in a series of short-term glacio-eustatic fluctuations, the last of which occurred in Westphalian D time. It was only in the Culm Basin that deep-water conditions persisted into the Westphalian (Ziegler, 1990). Deposition took place on the northern, non­ orogenic flank of the basin, as shown by the lack of palaeocurrents from the south (Fig. 6) (Ashwin, 1957; Burne, 1969b; Freshney eta/. , 1979; Melvin, 1986; Higgs, 1991) and by the compositionally mature, fine-grained nature of the sediments (Fig. 7) (Burne, 1969b; Freshney eta/. ; 1979; Melvin, 1986; Haslam & Scrivener, 199 1). Namurian processes deposited thin-bedded turbidites from currents flow­ ing eastward along the axis of the basin, forming the Crackington Formation (Freshney et a/. , 1979) . Deposition of the Bude Formation began in the Early Westphalian. Although 50% of the beds in the Bude Formation resemble those of the underlying Crackington Formation, it is distinguished by its content of massive sandstones (Fig. 4). The currents that deposited the Bude Formation generally flowed towards the south and southwest, although easterly flowing currents occurred at some horizons (Fig. 6) (Burne, 1969b, Freshney et a/. , 1979, Melvin, 1986). No post-Bude-Formation sediments are preserved in the Culm Basin.

TECTONICALLY

JUXTAPOSED

SUCCESSIONS IN THE CULM BASIN

Freshney & Taylor ( 1972) proposed a simple stra-

Fig. 5. (Opposite.) Composite 'sequence' (i.e. succession in present usage) for the Bude Formation reproduced from Burne ( 1969b, Fig. 20) showing the original interpretation, including the first recognition of 'thickening' and 'thinning' upward 'sequences' in subsea fan deposits. The composite succession includes only one representative of the various bed types· found in each part of the succession , and the scale shown relates to the thickness of these individual beds. Complete successions are actually 10-60 m thick. In the usage of the present paper the mudstone-dominated sequence corresponds to the black shale facies, now interpreted as fine-grained turbidites in an anaerobic or dysaerobic basin; the thickening upward sequence corresponds to part of the interbedded sandstone- shale facies, now interpreted as turbidites deposited in either levee , lobe or interchannel environments; the sandstone-dominated sequence corresponds to the sandstone-dominated facies, now interpreted as the deposits of fan-valley environments, including channel-fill successions; and the thinning upward sequence corresponds to part of the interbedded sandstone-shale facies now interpreted as the deposits of aerobic environments characterized by gentle currents, possibly fan levees.

R. V. Burne

108 A

Tool Marks

Directions

Senses Only

8

Current Scours

Directions

C

Senses Only

�

Cross Laminations

N

I

10 % Directions

N

=

15

20

Number of readings

Fig. 6. Structurally corrected palaeocurrents recorded from sole marks and cross laminations in the middle Bude Formation (data of Burne, 1969b).

tigraphy for the Culm Basin (Figs 3 & 4) in which all the dominantly turbidite sandstone-mudstone successions of Namurian and basal Westphalian rocks were classified as Crackington Formation, whereas the Bude Formation contained the younger

Westphalian deposits which 'while they still contain sequences of turbidites, have within them substantial numbers of massive sandstones and other associated facies characteristic of the rocks described by King (1966)' (Freshney & Taylor, 1972, p. 467). An ar­ bitrary junction between the two formations was set at the top of the Hartland Quay Shale (Freshney et al., 1979), a marker bed containing goniatites of the Gastrioceras amaliae marine band. The Bude Formation includes the youngest Westphalian rocks exposed in southwest England, the Warren Gutter Shale, which contain a fauna correlated with the 'Anthracoceras' aegiranum marine band (Freshney et al., 1979). The sedimentological contrast between the Crackington and Bude Formations and the con­ temporaneous paralic sediments exposed near Westward Ho! led Reading (1965) to suggest that the Culm Basin contained a number of sedimen­ tologically distinct though synchronous strati­ graphical successions separated by low-angle thrust faults. Burne & Moore (1971) concluded that two distinct stratigraphical successions could be recognized; a northern, essentially deltaic succession (the Westward Ho! Formation, Northam Formation, Abbotsham Formation and the Greencliff Beds), and a contrasting southern succession, consisting entirely of basin facies, comprising the Crackington Formation and the Bude Formation of Freshney et al. (1979) (Figs 2 & 3). The two successions are separated by a major E- W trending normal fault, with downthrow several hundred metres to the north (Freshney & Taylor, 1972). This fault is exposed on the coast (Burne & Moore, 1971, pp. 293-294) and can be traced inland eastward to the Mole Valley (National Grid Reference SS 675 260), 3 km east of South Molton (Fig. 1). The existence of tectonically juxtaposed suc·· cessions in the Culm Trough is also indicated by contrasting levels of organic maturation shown by vitrinite reflectance measurements from the two suc­ cessions (Cornford et al., 1987). Samples from the

Fig. 7. (Opposite.) Petrology and grain-size distributions of the various bed types in the Bude Formation. (A) Comparison of sandstone composition of the Bude Formation (data of Burne, 1969b; Freshney et al., 1979; Melvin, 1986); the Crackington Formation (data of Melvin, 1986); and the present-day Mississippi subsea fan (data of Roberts & Thayer, 1985). (B) Matrix composition of various Bude Formation bed types. Note high matrix content of 'slumped' and 'slurried'' beds, low matrix content of laminated and structureless sandstones, and intermediate matrix content of graded silty sandstones (data of Burne, 1969b) . (C) Plot of mean grain size () against sorting ( oG ) of Bude Formation sandstone grain·· size distributions, with and without matrix included (data of Burne, 1969b). Fields recognized by Melvin (1986) for the Bude Formation and Crackington Formation sandstones are outlined (note that Melvin measured o1 ) .

109

Bude Formation subsea fans QUARTZ

A

� D E2J]

FELDSPAR

8

Bude Formation sandstones Crackington Formation sandstones Mississippi Fan sands

ROCK FRAGMENTS QUARTZ

m; LIJ � E:;l

D

FELDSPAR & ROCK FRAGMENTS

c

Cross laminated beds Structureless & irregularly laminated sandstones Graded silty sandstones "Slurried" & "slumped" beds

MATRIX

3.0

2.0

MELVIN (1986) Bude Formation data

0G 1.0

}

MELVIN (1986) Crackington Formation data

0

Matrix Excluded

6.

Matrix Included

BURNE (1969b) Bude Formation data

110

R. V. Burne

section between Westward Ho! and Abbotsham form a statistically different group compared with those from stratigraphically equivalent rocks from the area between Hartland Quay (National Grid Reference SS 223 248) and Crackington Haven (SX 143 968). This suggests that the two areas have suffered differ­ ent thermal histories, i.e. different burial histories. 'These two areas must either now be in juxtaposition due to post burial tectonic events (e.g. thrusting) or they are not stratigraphic equivalents, the Bideford Formation (i.e. Bideford Group) being younger and hence less deeply buried than the Bude and Crackington Formations' (Cornford et al., 1987, p. 463). The stratigraphical equivalence of the sec­ tions was established by Edmonds et at. (1975, 1979) who identified the Gastrioceras amaliae marine band 550 m above the base of the Bideford Group. The Hartland Quay Shale, although of a contrasting facies, also contains this horizon and has been defined by Freshney et al. (1979) as underlying the base of the Bude Formation (Fig. 3). Edmonds et at. (1979) suggested that depositional continuity existed between these successions and concluded that the Northam and Abbotsham Formations (for which they proposed the name Bideford Formation, but for which the term Bideford Group is here retained) lie conformably between the Crackington Formation and the Bude Formation. Freshney et at. (1979) and Melvin (1986) both based their interpretation of the palaeogeography and facies of the Bude Formation on implied continuity between the deltas of the Bideford group and the depositional environments of the Bude Formation. According to Freshney (pers. comm., 1994) 'there is no physical evidence of thrusts in the Bideford area, but it is possible that strike-slip, east-west trending faults may define areas of contrasting thermal and burial history'. Although Freshney (pers. comm., 1994) points out that 'wedge-bedded sandstones characteristic of the Bideford Group can be traced eastward beyond Umberleigh (National Grid Reference SS 610 237) and appear to pass further east into a number of massive Bude-type sandstones', it has yet to be established whether this transition represents depo­ sitional continuity or tectonic juxtaposition. How­ ever, Cornford et at. (1987) concluded that the fluvial-deltaic Bideford Group is allochthonous and therefore the palaeogeography of the basin should not be constrained by the apparent associ­ ation between the Bideford Group and the deeper water Bude Formation. This conclusion supports the interpretation of Burne & Moore (1971) that

the succession in the area of Westward Ho! and Bideford is tectonically juxtaposed to the Culm Basin succession (Figs 1 & 2).

FACIES ANALYSIS OF THE BUDE FORMATION

Four facies are recognized in the Bude Formation, although these are, to a large extent, intergra­ dational. The term facies is used here to denote the assemblage of beds deposited within a depositional system. A depositional system comprises a sedimen­ tary environment together with the processes that may operate within it. Beds are formed by individual depositional processes. Facies descriptions

Black shale facies

Black shales (Fig. 8) are the most persistent facies in both the Bude Formation and the underlying Crackington Formation and form valuable lithofacies for stratigraphical correlation (King, 1967; Freshney

Fig. 8. Black shale facies showing regular millimetre-scale graded mudstone laminae with interbedded grey, muddy siltstone laminations, some with thin ripple cross­ laminations. Scale in centimetres (SS 1985 1313).

Bude Formation subsea fans

& Taylor, 1972). The facies may attain a thickness of 20 m (Fig. 4), and is dominated by uniform, dark mudstone composed almost entirely of sharp-based, graded laminae 5-10 mm thick. Thin sharp-based sets of ripple cross-laminated silt occur at the lower boundary of some of these laminations. Sole marks on these silts comprise prod marks or, rarely, obstacle scours around burrow entrances. Isolated thicker beds can occur in this facies. The clay minerals of the shales consist dominantly of illite (Burne, 1969b). Undeformed shales also contain kaolinite, whereas in deformed shales this is replaced by chlorite, and grey shales contain more kaolinite (Freshney et al., 1979). Siderite bands and ankerite nodules are found in some shales. The grading of the laminae is due to an upward increase in the content of organic carbon. Finely comminuted plant debris is common on parting surfaces. Freshney et al. (1979) record jarosite-natrojarosite efflor­ escences in sulphurous shales. Higgs (1991) found that carbon/sulphur (CIS) ratios (Berner & Raiswell, 1983, 1984) in four out of five black-shale samples were less than 10. The shales contain little evidence of biological activity apart from the common occurrence of Planolites (Fig. 9) (King, 1967). Tops of interbedded sandstones may contain Diplocraterion parallelum, an'd, more rarely Arenicolites, Teichichnus, Phyeodes

Fig. 9. Replacements of Planolites burrows preserved in a siderite vein in the black shale facies. A 4-cm-long live limpet for scale (SS 1983 0313).

111

and Skolithos (Higgs, 1991). Freshney & Taylor (1972) found that in the lower parts of the Crack­ ington Formation a marine fauna of goniatites occurs throughout the shales, whereas from the Namurian G1 zone the goniatites occur only in isolated bands within some black shales. The Goniatites, which are truly marine, occur in bands between layers of nodules that contain fossils of fish that were tolerant of low-salinity environments. Goniatites have been found only at two horizons in the Bude Formation (Fig. 4), the Sandy Mouth Shale and the Warren Gutter Shale. In the Warren Gutter Shale, goniatites are present only as spat (Freshney et at. , 1979). Goniatites have not been found in the middle Bude Formation, although fish remains and coprolites occur at two horizons (King, 1967). Cornuboniscus budensis and Elonichthys aitkeni, both palaeoniscid bony fishes, the acanthodian Acanthodes wardi and an eocarid crustacean Crangopsis huxleyi occur in the Saturday's Pit Shale, whereas the Coelocanth Rhabdoderma elegans occurs in the Tom's Cove Shale (Fig. 4). Muddy siltstone facies

This facies consists of grey to dark grey laminated muddy siltstones and thin, sharp-based silty sand­ stones that show parallel laminations and ripple drift cross-lamination. In places the facies is disrupted by intrastratal folding and small-scale synsedimentary faults. Scouring occurs locally (Fig. 10). The muddy siltstones are composed of carbonaceous material, clay minerals, and silt-sized quartz particles. The grading of the 1-10-mm-scale laminae reflects a decrease in the size of the quartz grains and an increase in the proportion of clay (Fig. 11). The silstones are composed of quartz, illite, kaolinite, chlorite, siderite and minor feldspar (Merriman, quoted in Higgs, 1991). Higgs (1991) found that organic carbon varied from 1.8% to 2.9%, plant fragments were conspicuous, burrows are absent, and C/S ratios (Berner & Raiswell, 1983, 1984) range between 22.8 and 39.0. King (1967) described Kouphnichnium (King, 1965; Goldring & Seilacher, 1971), i.e. xiphosurid feeding trails from this facies. The animals moved over silt-covered surfaces, prodding down into lower silty laminae. Current sole marks were formed before the trackways were emplaced, and the tracks are confined to one level within the graded laminated units, indicating that little sediment was deposited during the time in which the track was being made

112

R. V. Burne

Fig. 11. Photomicrograph of graded muddy siltstone lamination in the muddy siltstone facies. Carbonaceous material appears dark (SS 1 995 0400).

Fig. 10. Muddy siltstone facies showing alternations of laminated muddy siltstone and ripple cross-laminated units. Note irregular scouring and deposition of cross­ laminated beds at the level of the hammer (circled). Section youngs to the right (SS 201 077).

(Goldring & Seilacher, 1971). Other tracks initially regarded by King (1965) as xiphosurid mating-traces (Fig. 12), were reinterpreted by Higgs (1988) as Undichna or trails left by fish dragging their fins along the soft bottom during feeding. Interbedded sandstone-shale facies

This facies comprises interbedded 'event' beds (Einsele & Seilacher, 1982) and black shales. The relative abundance of various types of event beds in this facies shows no consistency of grouping. At some horizons beds of one of the following types dominate the succession. In other cases a predictable succession of event beds occurs, whereas elsewhere the facies consists of a less predictable succession of beds. The following types of event beds can be distinguished.

Fine-grained graded beds range from 1 to 12 em in thickness. Beds have sharp bases, in places bearing sole marks, and comprise laterally continuous, normally graded layers of siltstone and muddy silt­ stone (Fig. 13). The sediment is poorly sorted with median grain size 3.5

.-., c == c -UJ o u (j)

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

142

P. Fralick and T.J. Barrett

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

Fig. 3.

Depositional controls on iron formation

143

Grainstone (reworked chemical sediment)

Fine-grained chemical sediment

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

Fig_ 4_

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

Outer shelf and slope: Middle Member, Baby Formation

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

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

P. Fralick and T.J. Barrett

144 M

60

Fe-C, Fe-0, Fe-S

50

ss

40 SS, Fe-C

30

f,

Fe-S

Fe-C

20

_

Overburden

ss

10

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

0

SS, Fe-0, Fe-S, Si

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

Fig. 5.

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

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

Depositional controls on iron formation

145

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

Fig. 6.

146

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

Channelized turbidity flows

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

Rainout of chemical sediment

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

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

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

Depositional controls on iron formation

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

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

B

A M

M

50

6

40

147

•

IFLAa interlaminated magnetite, siltstone, and slate

II

IFLAa interlaminated magnetite, jasper, siltstone, and slate

4

D lliJ 0

30

20

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

Fig. 8.

10

0

2

0

� -

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

148

P. Fralick and T.J. Barrett

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

Fig. 9.

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

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

Depositional controls on iron formation

Debris-flow

149 Fine-grained chemical sediment

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

Fig. 10.

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

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

3

�

Felsic porphyry intrusive

�-

lnterlaminated pyritic-carbonaceous slate and pyrite

• �

Massive pyrite Chert

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

150

P. Fralick and T.J. Barrett

Fig. 12.

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

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

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

Depositional controls on iron formation

15 1

Sand lobes

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

sediment and clay

Intermediate and mafic volcanics

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

Thick lava-plain successions are common in Superior

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

A

B

Layered siltstone and shale Massive siltstone

E2J � 1-

Layered pyrite Pyrite and layered chert Pyrrhotite and disrupted chert

f+:+l �

Felsic dike Intermediate volcanics

Fig. 14.

Fralick

\A/\ �

II

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

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

et al.,

152

P. Fralick and T.J. Barrett

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

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

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

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

Fig. 15.

Depositional controls on iron formation

153

Off-axis volcanism

proximal chemical sediment

sediment and clay Fig. 16.

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

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

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

DISCUSSION

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

154

P. Fralick and T.J. Barrett

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

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

Depositional controls on iron formation ACKNOW LEDGEMENTS

DEVANEY, J .R. & WILLIAMS,

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

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in

the

Kirkland

Lake

area,

Ontario .

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

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

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

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

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

ABSTRACT

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

INTRODUCTION

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

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

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

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G.J. Orton

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

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

TERMINOLOGY

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

Facies models in volcanic terrains

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

GEOLOGICAL FRAMEWORK

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

159

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

160

G.J. Orton

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

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

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

161

Facies models in volcanic terrains 68

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

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

FACIES ASSOCIATIONS

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

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

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

162

G.J. Orton

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163

Facies models in volcanic terrains Facies interpretation 65

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164

G.J. Orton

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

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

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166

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Fig. 6. Contact between (A) fine­

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

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

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

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

Facies models in volcanic terrains

167 Facies interpretation

Scm 5cm

Stacked 'channel' fills

Small scale trough Cross-stratification

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Low-angle indistinct stratification

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

Fig. 21. Map showing generalized

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

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

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

215

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

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

216

M .G. Laird

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

ACKNOWLEDGEMENTS

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

REFERENCES

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

BALLANCE,

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

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

217

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

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

22, 219-236

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

ABSTRACT

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

INTRODUCTION AND GEOLOGICAL SETTING

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

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

et a!.,

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

219

220

N.H. Platt

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

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

LITHOFACIES STRATIGRAPHY

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

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

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

221

Alluvial fans, palaeokarst and tectonics

A

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

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

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

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

The balance between sediment supply and subsid-

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

275

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

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

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

276

M.B. Edwards

Updip

Down dip 2

3

4

Down dip

Updip

Datum

2

4

3

Datum Top

Top J

J

Paleostructure

2000' A

610m

Base J B

230'

405'

500'

2000'

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

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

PRESERVATION POTENTIAL THRESHOLDS

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

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

CONCLUSIONS

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

277

Dff i erent ial subsd i ence

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FORESHORE Bay/

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B Unstable setting

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Fig. 12.

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

Topography

B

healed by sedimentation

Topography controls sedimentation

Fig. 13.

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

Amalgamated channel complex

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M. B. Edwards

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

279

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

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

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

ACKNOWLEDGEMENTS

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

M. B. Edwards

280

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

REFERENCES BEBOUT,

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

EDWARDS,

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

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

281

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

WEISE,

Sequence and Seismic Stratigraphy in Facies Analysis

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

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

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

ABSTRACT

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

INTRODUCTION

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

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

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

285

286

J.

Cartwright

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

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

Cenozoic Basin evolution

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

GEOLOGICAL SETTING

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sandstone isopach with the oil­ water contact marked. (B) Depth structure map on the top 36-1 sandstone showing that the oil­ water contact is 60ft above structural closure (shaded). The structural spill point is marked by a triangle.

less extensive. Net sandstone thickness from well penetrations varies from 0 to 67ft (0-20m) and a sand body is defined (or predicted from thickness trends) that is elongated roughly N-S and has an average width of approximately 11 000 ft (3354 m) with a zero-edge to the west and an inferred one to the east. The abrupt change to a westwards depo­ sitional strike at the north end of the sand body is defined by well penetrations to the north and west outside of Widuri field.

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364

R. Young, W.E. Harmony and T. Budiyento

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the 35-2 sandstone. The spill point has -remained in approximately the same position but is now at 3680 ft SS ( 1122 m SS) and this is coincident with the oil-water contact for this reservoir. The trapping mechanism is therefore again clearly structural, but this time the structure is full to spill. 35-1 sandstone

The Widuri E-1 core (Fig. 11) is lacking the basal

Fig. 10. 35-2 reservoir. (A) Net sandstone isopach with the oil­ water contact marked . (B) Depth structure map on the top 35-2 sandstone showing that the structural closure (shaded) is coincident with the oil-water contact. The structural spill point is marked by a triangle.

6 ft (1. 8 m ) of the 35-1 sandstone but adequately demonstrates the sedimentary character of the reservoir. The 35- 1 is an erosively based sandstone, which is trough cross-bedded in 1-2-ft (0. 3-0. 6 m) sets, and which averages medium grain size. Intervals of coarse grain size are common towards the base of the sandstone and fine/very fine grain size common towards the top, but the sandstone does not fine upwards systematically and several distinct units with abrupt contacts are evident (see Fig. 11). This

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multistorey character is readily recognizable in all other 35-1 cores and the units are separated occasionally by coals (see Widuri A-1, Fig. 11). The 35-1 sandstone is the most variable of the Widuri reservoirs (see cross-section A-A', Fig. 11). The log character is predominantly gamma-ray increasing upwards (fining upwards) with a gra­ dational top, but is also commonly blocky with a sharp top. The most striking feature of the sand­ stone, however, is the rapid lateral thickness changes within an overall range of 0-64 ft (0-19.5 m).

Fortunately, these thickness vanatwns could be mapped and predicted seismically prior to any development drilling. The 35-1 interval is extremely well constrained between two readily correlatable markers - Coal 'A' at the top and usually a double coal at the base (see cross-section A-A', Fig. 11). Acoustic impedance contrasts associated with these coal markers are responsible for two seismic reflec­ tion which envelope the 35-1 sandstone and can be traced field-wide and beyond. The isochron between these two seismic markers, when linked with

366

R. Young, W.E. Harmony and T. Budiyento

well penetrations, proved to be a very effective predictor of net sandstone distribution (see Young et al., 1991). The net sandstone isopach (Fig. 12A) demon­ strates an anastomosing pattern and the 35-1 sandstone is interpreted to be a multistorey and multilateral fluvial system, possibly representing anastomosing rivers. Individual channels can be mapped with more confidence and over a larger area than those of the 35-2 and 36-1 (e.g. the sand body

trending southwest from the D platform is a single­ storey channel), and sand body widths vary from approximately 3000ft (915 m) in the west to 6500ft (1982 m) in the east. Finer grain size, thinner and narrower sand bodies, and more order to the sequence suggest a more mature fluvial system than the 35-2. Reservoir quality is again excellent, with log porosity averaging 29% and test permeabilities ranging from 10 to 20 D. The depth structure map on top of the 35-1 sand-

+

+

Cl100' 5000 FT

a

Fig. 12. 35-1 reservoir. (A) Net sandstone isopach with the oil­ water contact marked. (B) Depth structure map on the top 35-1 sandstone showing that the oil­ water contact is 20ft below structural closure (shaded).

Fluvial sand-body geometries

stone (Fig. 12B) defines a structural spill point or closing contour at 3660ft SS (1116 m SS). The oil-water contact for this reservoir, however, is the same as for the 36-1 and 35-2 at 3680 ft SS ( 1122 m SS). There is therefore 20ft (6 m) more oil column than can be accounted for by structural closure. Returning to Fig. 12A, it is clear that the structural contour of the oil-water contact does not close to the northwest but that closure is effected by the sand body pinch-out to the northwest. There is therefore a stratigraphical component to trapping and a combination stratigraphical-structural trap results. 34-2 sandstone

The 34-2 sandstone rests directly on top of Coal 'A' and the sedimentary character is demonstrated in core from the B-8 well (Fig. 13). The base is erosive on top of Coal 'A' and the thin, very coarse grained, basal lag contains coal clasts. The overlying sand­ stone consists of predominantly medium- to coarse­ grained sandstone in the basal and middle section before passing into fine- and very fine-grained sand­ stones in the upper part, and finally into rootleted siltstone and coal at the top. Trough cross-bedding in 1 ft (0. 3 m) sets is the dominant sedimentary struc­ ture. Maximum penetrated thickness for the 34-2 sandstone is 54ft (16.5 m) and in the thicker sections the log character tends to be blocky, with a gamma­ ray increasing (fining upwards) signature at the very top (see Fig. 13). In the thinner sections, however, the log character is mostly gamma-ray increasing (fining upwards). Immediately above the top coal is the first indication of fully marine conditions in the Talang Akar Formation. Calcareous mudstone grades up into muddy limestone containing skeletal debris of bivalves, large foraminifera, solitary corals and echinoid fragments. This limestone can be recognized throughout most of the field. The 34-2 net sandstone isopach (Fig. 14A) out­ lines a sinuous sand body with an average width of 4000ft ( 1220m). The shape and character suggest that the sand body is fluvial but the exact type of fluvial system is less clear. Several distinct sedi­ mentary units can be recognized within the 34-2 at B-8 (Fig. 13), but these are not as discrete as those of previous sandstones. Similarly, evidence from other full-hole core, dense sidewall cores in most of the other wells, and log character analysis suggests a broad systematic fining within a single-storey system (see also cross-section A- A', Fig. 13). This implies

36 7

a relatively mature fluvial system but a meandering river model (Allen, 1965) is discounted on the lack of evidence for lateral accretion and on the geometry of the reservoir which is not, so much a sheet as a ribbon of sandstone that tapers towards both sides (see Fig. 14A). Similarly, equations relating sand body thickness to channel depth and channel belt/ sand body width (see Leeder, 1973; Collinson, 1978; Lorenz et al., 1985; Fielding & Crane, 198 7), although having a wide margin of error, predict a much wider sand body if the 34-2 were a high sinuosity channel deposit. It is tempting to use a Mahakam River (East Kalimantan, island of Borneo) type model (Allen et al., 1979; Young & Atkinson, 1993) as the analogy for the 34-2 reservoir. In this case, the river channel does not meander significantly but does have a meandering thalweg. On the inside of every thalweg meander loop is a non-emergent lateral bar which is attached to the channel side and progrades down­ stream. Echo-sounding and dredge sampling suggest that the Mahakam River immediately above the delta has a width, sinuosity, maximum sand thick­ ness, grain size and bedform distribution that is remarkably similar to the 34-2 sandstone. However, irrespective of the exact mechanism of channel fill, the 34-2 reservoir is interpreted to be a single storey, fluvial sandstone. Proximity to the sea is suggested by the maturity of the fluvial system and by the overlying transgressive limestone. However, evi­ dence for marine conditions within the sandstone is absent. Reservoir quality is excellent, with average log porosity of 29% and test permeabilities of 10-30D. The depth structure map on top of the 34-2 sand­ stone is shown in Fig. 14B and the structural spill point/closing contour is at a depth of 3630ft SS (1107 m SS). The oil-water contact for the reservoir is still 3680 ft SS (1122 m SS) and there is again a discrepancy (50ft or 15m) between oil column and structural closure. It is clear on Fig. 14A, however, that the trapping mechanism is the same as for the 35-1. The structural contour of the oil-water contact does not close to the northwest but the sand body pinches out in that direction. The trapping mechan­ ism again combines stratigraphy and structure. 34-1 sandstone

The base of the 34-1 sandstone lies just above the limestone at the top of the 34-2 sequence and all but the top few feet of the sandstone was cored in the

368 A 20

R. Young, W.E. Harmony and T. Budiyento

B-2 API

1600' 120

20

B-3 API

1900' 120

20

B-4 API

2200' 120

5300'

Widu ri- 7

20

API

120

A'

Widuri-5

20

API

120

Top l i m estone

Base coal 'A'

DEPTH LOCATION MAP

GAMMA RAY API100 0

(FT)

Core log 34-2 sandstone Widuri B-8 Muddy l i mestone and calcareous m u dstone Bivalve and fora m i n ifera fragments Root leted si ltstone and coal

Trough cross- bedded sandstone g rading from med i u m to coarse at the base to fine-very fine at the top

4490

Si MCg ,...,... Sh�Cl SA ND I

NO RECOVERY 4500

4510

4520

Coal 'A'

Fig. 13. 34-2 reservoir: vertical and lateral sedimentary character..

B-8 well (see Fig. 15). The sandstone is erosively based and consists of mostly fine- and very fine­ grained, very well-sorted sandstone, which broadly fines upwards giving a predominantly gamma-ray increasing upwards log signature. Cross-bedding is sporadically present throughout but, towards the top, parallel lamination, ripple cross-lamination and ripple/wavy bedding predominate. Carbonaceous

laminations are common, and abundant mudstone clasts in the lower to middle part of the sandstone result in increased gamma-ray readings (Fig. 15). Towards the top of the sandstone (4382. 5-4391f t MJ?) a heavily bioturbated and interbedded·­ interlaminated unit of mudstone, siltstone and sandstone contains predominantly Thalassinoides but also Teichichnus type burrows, which are indica-

Fluvial sand-body geometries

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Fig. 14. 34-2 reservoir : (A) Net

sandstone isopach with the oil­ water contact marked. (B) Depth structure map on the top 34-2 reservoir showing that the oil­ water contact is 50ft below structural closure (shaded).

tive of marine influence. In Widuri-1 the 34-1 was also cored and clay drapes, a common fei!ture of tidal sediments, are abundant in the upper part of the sandstone, where marine acritarchs also have been found. The net sandstone isopach (Fig. 16A) defines a sand body with a maximum penetrated thickness of 49ft ( 15 m) and an average width of 2000 ft ( 610 m). The detailed shape of the contours in Fig. 16A is not

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the result of extrapolated well data. Prior to any development drilling it was recognized that an acoustic impedance contrast between the base of the sandstone and the underlying mudstone was respon­ sible for a seismic peak (see Blue seismic horizon, Fig. 6, for example) which, when mapped, had a restricted areal extent. Seismic modelling suggested that the sandstone thickness could be estimated from the amplitude of the reflection through the use

370 A 20

R. Young, W.E. Harmony and T. Budiyento

D-11 API

2150' 120 20

A-12 API

1600' 120 20

Widu ri-1 API

5000' 120 20

C-6 API

DEPTH

LOCATION MAP GAMMA RAY API

(FT)

3850' 20 120

B-8 API

A' 120

Core log 34-1 sandstone Widuri B-8

Biotur bated mudstone, si ltstone and sandstone

Thalassinoides/ Teichichnus type bu rrows

Fine and very fine grained sandstone

Calcareous mu dstone with shell fragments

Fig. 15. 34-1 reservoir: vertical and lateral sedimentary character.

of a tuning curve (Young et al., 1991), and the net sandstone isopach map was constructed based on reflection amplitude. This map proved to be very accurate and has changed little after development drilling. The shape and dimensions of the sandstone, together with the log and sedimentary character, suggest a single-storey channel which was reworked by tides in the upper part. Within Widuri field, the interval that is laterally equivalent to the channel sandstone consists of mudstone and siltstone with rare, thin sandstones (see cross-section A-A',

Fig. 15). When traced southwards away from the sediment source area, however, Young & Atkinson (1993) have shown that the entire productive interval at Widuri becomes more marine and is dominated mostly by 6-24 ft thick, stacked, coarsening upwards units, which are interpreted to be distributary mouth bars and various other marginal marine bars. Dia­ chronism of the whole interval is clear and the overall interpretation is a deltaic succession domi­ nated by tidal and fluvial processes (Young & Atkinson, 1993). The 34-1 is therefore interpreted to be a distribu-

Fluvial sand-body geometries

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