Modern and Ancient Lake Sediments
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
Modern and Ancient Lake Sediments EDITED BY ALBERT MATTER AND MAURICE E. TUCKER
Proceedings of a symposium held at the H.C. 0rsted Institute, University of Copenhagen, 12-13 August 1977. Sponsored by the International Association of Sedimentologists and the Societas Internationalis Limnologiae
S PE C I A L P U B L I C AT I O N N U M B E R 2 O F T H E I NTERNAT I O NA L A S S O C I AT I O N O F S E D IMENTO LO G I ST S . P U B L I S H E D B Y B LA C KWELL S C IE N T I F I C P UB L I C A T I O N S O X F O RD L O N D O N E D I N B U R G H M E L B OU RN E
© 1978 The I nternational Association of Sedimentologists Published by Blackwell Scientific Publications Osney Mead, Oxfoord 8 John Street, London WC I 9 Forrest Road, Edinburgh P.O. Box 9, N orth Balwyn. Victoria. Australia All rights reserved. No part of this publication may be reproduced, stored in a retrieval system. or transmitted, in any form or by any means. electronic, mechanical, photocopying. recording or otherwise without the prior permission of the copyright owner.
First published 1 978 British Library Cataloguing in Publication Data Modern and ancient lake sediments.---{International Association of Sedimentologists. Special publications; No. 2) I . Sediments (Geology)-Congresses 2. Lakes-Congresses I. Matter, Albert I I. Tucker, Maurice E. Ill. Series 551.4'82 QE47 1 .2 ISBN 0-632-00234-4 Distributed in the U.S.A. by Halsted Press, a division of John Wiley & Sons, Inc., New York Printed and bound in Great Britain by Burgess & Son Ltd Abingdon, Oxfordshire
Contents
Introduction A lbert Matter and Maurice E. Tucker
7
Saline lakes and their deposits: a sedimentological approach Lawrence A . Hardie, Joseph P. Smoot and Hans P. Eugster
43
Late Pleistocene-Holocene evolution of the Kivu-Tanganyika Basin Peter Stoffers and Robert E. Hecky
55
Holocene carbonate evolution in Lake Balaton (Hungary): a response to climate and impact of man German Muller and Frank Wagner
81
Permian Saar-Nahe Basin and Recent Lake Constance (Germany): two environments of lacustrine algal carbonates Andreas Schafer and Karl R. G. Stapf
107 Origin of the carbonate sediments in the Wilkins Peak Member of the lacustrine Green River Formation (Eocene), Wyoming, USA Joseph P. Smoot
127 Late Neogene sedimentation in the Black Sea Kenneth J. Hsii and Kerry Kelts
145 Turbidites and varves in Lake Brienz (Switzerland): deposition of clastic detritus by density currents Michael Sturm and Albert Matter
167
Lacustrine facies in the Pliocene Ridge Basin Group: Ridge Basin, California Martin
187
H.
Link and Robert
H.
Osborne
Lacustrine sedimentation in an evaporitic environment: the Ludian (Palaeogene) of the Mormoiron Basin, Southeastern France Georges True
203 Triassic lacustrine sediments from South Wales: shore-zone clastics, evaporites and carbonates Maurice E. Tucker
223
Permo-Triassic lacustrine deposits in the Eastern Karoo Basin, Natal, South Africa D. E. van Dijk, D.
K.
Hobday and A . v
J.
Tankard
vi
Contents
239 Subaqueous clastic fissure eruptions and other examples of sedimentary transposition in the lacustrine Horton Bluff Formation (Mississippian), Nova Scotia, Canada Reinhard Hesse and Harold G. Reading
257 A Proterozoic lacustrine interlude from the Zambian Copperbelt Harry Clemmey
277 Economic significance of playa lake deposits C. C.
Reeves Jr.
Spec. Pubis int. Ass. Sediment. ( 1 978) 2, 1-6
Modern and ancient lake sediments: an introduction
A L B E R T M A T T E R andM A U R I C E E . T U C K E R Geol. Jnst. Universitiit Bern, Sahlistrasse 6, CH-3000 Bern, Switzerland and Department of Geology, The University, Newcastle upon Tyne NEJ 7R U,
U. K.
Research on lake sediments really began towards the end of the last century. In Europe, lakes in Switzerland received the attention of scientists from the mid nineteenth century, with early work concerned with water circulation and chemistry, and then later with the lake sediments themselves. Research by Forel ( 1 886- 1 892) demonstrated the presence of channels and levees on the Rhone delta in Lake Geneva. These Forel attributed to the underflow of sediment-laden river-water. In the early part of this century Nipkow ( 1920, 1 928) sampled the bottom sediments of Lake Zurich and, although primarily interested in the flora, described rhythmically laminated sediments ('non-glacial varves') consisting of a summer calcareous layer and a winter organic-rich layer. He also noted thick detrital layers (turbidites) which could be correlated from core to core. Work on true varves forming in proglacial lakes also goes back to the last century, although interest has largely been focused on their chronologie and stratigraphic use (e.g. De Geer, 19 1 2). In North America, classic work was carried out by Russell ( 1 885) and Gilbert ( 1 890) on the Pleistocene lakes Lahontan and Bonneville. The shore features of Lake Lahontan with their evidence for an ancient deep lake were first recognized in 1 858/9 by Henry Englemann, geologist to an expedition crossing the Carson Desert region. Russell ( 1 885) described the whole Lake Lahontan basin, with its beach terraces, bars, spits and tufas, and erected a stratigraphy of the lake deposits. He was able to demonstrate two deep-lake periods with an intervening period of complete lake desiccation. Russell's interpretation of the lake history has generally been proved correct by later work (Morrison, 1 964). It was to be many years, however, before a detailed description was published of the Great Salt Lake, one ofthe remnants of Lake Bonneville. Eardley's ( 1 938) work on the Great Salt Lake sediments is still one of the few detailed accounts of modern lacustrine carbonates with which ancient lacustrine limestones can be compared.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
2
A lbert Matter and Maurice E. Tucker
Studies of lake sediments, like those of Fore! mentioned above, and then later studies of sedimentation in reservoirs, such as Lake Mead, Colorado (Grover & Howard, 1 938; Gould, 195 1), have had an important bearing on the evolution of the concept of turbidity currents. The latter of course are an important sediment transport mechanism in continental-margin situations and in marine basins and troughs. Of all ancient lake sequences, the one which has received most attention, has the most extensive literature and is invariably quoted in comparisons, is the Eocene Green River Formation of Wyoming, Utah and Colorado, U.S.A, representing the deposits of lakes Uinta and Gosiute. Classic work was published by Wilmot Bradley from the 1920's onwards (e.g. 1929, 193 1 , 1 964) following reconnaissance mapping of the area by U.S. Geological Survey officers in the late 1 800s. Bradley envisaged a relatively deep, permanently stratified lake, particularly during deposition of rhythmically laminated oil shales in the basin centre. Recent investigations, however, have shown that a playa-lake model is more applicable (Eugster & Surdam, 1 973; Eugster & Hardie, 1 975 and also Smoot, 1978). A notable feature of the Green River Formation is the great range of evaporite and authigenic minerals present (Milton & Eugster, 1 959; Bradley & Eugster, 1969) many of which are unique to this formation. One of the first rock sequences to be interpreted as lacustrine is the Middle Old Red Sandstone (Devonian) of Caithness, N.E. Scotland. The Caithness Flagstones, deposited in Lake Orcadie, have long been famous for the fossil fish they yield, particularly in the Achanarras Limestone. These rocks were initially described in 1829 by Sedgwick & Murchison, and then later by Geikie ( 1 878) and Crampton & Carruthers ( 1 9 1 4). Apart from the fauna, the lacustrine interpretation was also based on the dark grey, carbonaceous and finely-laminated nature of the flagstones, which contrasts markedly with the red conglomerates, sandstones and siltstones that characterize the continental Devonian elsewhere. It is interesting to note that in the later studies of the Orcadian lake sediments (Rayner, 1963; Donovan, 1 975) comparisons were made with the findings of Nipkow ( 1 920, 1 928) from Lake Zurich. Significant slopes around Lake Orcadie are indicated by slumped and brecciated horizons in the flagstones. Crampton & Carruthers ( 1 9 1 4) noted the absence of beach deposits and used this as an argument in favour of a lacustrine interpretation. Recent work, however, has recognized lacustrine shore-line clastics (Donovan, 1 975), as well as stromatolites (Fannin, 1969; Donovan, 1 973). The latter were fairly accurately described by Crampton & Carruthers but interpreted as subaerial tufas. There has been an increasing interest in lake sediments in the past decades with many papers dealing with modern examples in various parts of the world, but relatively few concerned with ancient lakes (see for example references in Picard & High, 1 972). At the International Sedimentological Conference in Nice ( 1 972), it was realized that a number of people are actively working on lake sediments. It was therefore proposed to hold a special symposium on this subject, under the auspices of the International Association of Sedimentologists. The proceedings of that sympo sium, held in Copenhagen August 1 2- 1 3, 1 977, coinciding with the International Limnological Congress (S.I.L.), constitute this Special Publication. In inviting contributions to this book, the editors have attempted to cover most sedimentological aspects of lakes, from clastic to chemical and biochemical, and from modern to ancient. Aspects of lake sediments not covered in this book include those of an engineering geological nature, such as the silting up of dammed lakes, and those of environmental importance, such as the poisoning of lake sediments and waters through industrial pollution.
Introduction
3
The book begins with a review paper on arid-zone lakes by Hardie, Smoot & Eugster. In Saline lakes and their deposits: a sedimentological approach the physical and chemical conditions of sedimentation are discussed. Various subenvironments are defined and the characteristic features of their sediments presented. It is then shown how the sequence of facies can be used to interpret the history of the lake basin. Low latitude lakes in Africa are described in Stoffers and Reeky's paper on the Late Pleistocene-Holocene evolution of the Kivu- Tanganyika Basin. The variations in diatom assemblages and mineralogy (Mg-content in carbonates, pyrite-siderite) are described from cores collected in Lake Kivu. The variations are related to the stratification history of the lake, which was controlled by climatic changes of the last 14,000 years. Contemporaneous volcanic activity in the form of sublacustrine springs is also considered important. The control of climate (and later man) on sedimentary mineralogy is also demonstrated by Mtiller and Wagner in Holocene carbonate evolution in Lake Balaton (Hungary) : a response to climate and the impact to man.
When Lake Balaton had no outlet in Pre-Roman times, low Mg-calcite was precipitated during periods of high-water level and low rates of evaporation, whereas protodolomite and high Mg-calcite were precipitated during periods of low-water level and high evaporation. Since the construction of an outlet for the lake by the Romans, calcite has been precipitated periodically at times of algal blooms. Continuing with lacustrine carbonates, Schafer and Stapf combine the modern and ancient in their Permian Saar-Nahe Basin and Recent Lake Constance (Germany) : two environments of lacustrine algal carbonates. The oncolites of Lake Constance are of two main types, rough and smooth surfaced forms, which are reflections of water depth and the algal community. Two notable features are the preservation of algal filaments within the precipitated calcite and the various fabrics of the algal balls which can be related to the blue-green algal genera. The Permian oncolites show a greater diversity of shape and size, but are distinguished from the Lake Constance examples by the lack of preserved algal filaments. In Origin of the carbonate sediments in the Wilkins Peak Member of the lacustrine Green River Formation (Eocene), Wyoming, U.S.A. by Smoot the carbonate is chiefly peloidal intraclastic dolomite. Smoot reasons that the carbonate is not a lacustrine precipitate, but was derived from disintegration of surface crusts, tufas and caliches which developed on exposed mudflats, in stream beds and on alluvial fans around the lake. These subaerial dolomites were then transported as clasts into the lake by sheet flows following rainstorms. Deep-sea drilling results are reported by Hsii and Kelts in Late Neogene chemical sedimentation in the Black Sea. The sediments are chiefly calcitic chalks, laminated in the lower part, again like Nipkow's carbonate varves of Lake Zurich. It is thought that the carbonate mineralogy was controlled by salinity, with horizons rich in dolomite being formed in one instance when a shallow salt lake existed, and in another through sabkha diagenesis. Siderite, occurring at two levels, is thought to be a direct precipitate at a time of high dissolved-iron input. Deeper water clastic sediments are described by Sturm and Matter in Turbidites and varves in Lake Brienz (Switzerland) : deposition of clastic detritus by density currents. The turbidity currents are generated through river inflow at high stages and high rates of sediment supply. They deposit graded beds with scoured bases on the central basin plain. During normal river stages when the density of the inflowing river water is lower, interflows and overflows lead to the formation of rhythmically-laminated ('varved') sediments below the thermocline. Homogeneous sediments accumulate on the shore terrace and upper slope above the thermocline.
4
A lbert Matter and Maurice E. Tucker
Link & Osborne describe one of the thickest known lacustrine sequences (over 1 2 k m thick) in Lacustrine facies of the Pliocene Ridge Basin Group: Ridge Basin, California. Practically all possible lacustrine facies are developed, including various shore-zone and deltaic clastic sand bodies, sublacustrine fans and turbidites, evaporites, organic mudstones and stromatolites. The thickness and rapid lateral and vertical facies changes are related to vertical and strike-slip faulting along the margins of the Ridge Basin. Tertiary lacustrine sediments from S.E. France are described by True in the next paper: Lacustrine sedimentation in an evaporitic environment: the Ludian (Palaeogene) of the Mormoiron Basin. As is typical of many lakes, coarse clastics were deposited around the edge of the fault-bounded lake basin, during the Lower Ludian, and organic-rich muds accumulated in the centre. Arid conditions during the Upper Ludian led to evaporites and dolomites being precipitated in the basin centre, along with sepiolite and magnesium smectite-rich muds. The effect of climate on lake sediments is also discussed by Tucker in Triassic lacustrine sediments from South Wales: shore-zone clastics, evaporites and carbonates. The depositional environment of the red homogeneous Keuper Marl has been a subject of some debate but in this paper Tucker describes the lateral, marginal equivalents, which are of lacustrine shore-zone origin. Triassic shore terraces with beach gravels, cliff-lines with screes and wave-cut notches are all preserved, cut into Carboniferous Limestone at the lake margin. During periods of lake shoreline retreat (regression), calcretes developed within the shore-zone sediments and sabkha-type evaporites formed within the exposed lake floor sediments (the Keuper Marl). Van Dijk, Hobday & Tankard describe sedimentology and palaeoecology of Permo- Triassic lacustrine deposits in the Eastern Karoo Basin, Natal, South Africa. The sediments consist of repetitive coarsening-upward cycles representing shore-zone and deltaic sandstones prograding over offshore lacustrine siltstones. Coastal bays were infilled by distributary overbank flooding and crevasse splays. A climatic change resulted in shallow playa lakes with emergence structures and evaporites. Vertebrate remains, fish and plants are common throughout with their prevalence and preservation varying from facies to facies. From the Palaeozoic, Hesse & Reading describe Subaqueous clasticfissure eruptions
and other examples of sedimentary transposition in the lacustrine Horton Bluff Formation (Mississippian), Nova Scotia, Canada. Spectacular sandstone dykes and
collapse structures are related to contemporaneous earthquake activity associated with nearby fault movements. A Proterozoic lacustrine interlude from the Zambian Copperbelt is described by Clemmey. The lacustrine sediments are distinguished from underlying and overlying marine strata largely on the basis of preservation of delicate trace fossils (the oldest in the world) and sedimentary structures. The latter include synsedimentary folds and faults, gas-burst structures, syneresis cracks and various types of rain print. Equivalent structures are noted from a modern ephemeral lake, also in Zambia. Cycles of offshore to nearshore and sabkha sediments (with original evaporites still preserved) are recognised in the Copperbelt rocks and attributed to periods of lake-shoreline progradation. The final paper in this volume by Reeves illustrates the Economic significance of playa lake deposits. The various sediments and minerals of playa lake basins are described and their industrial uses noted. It is likely that playa lakes will take on an even greater importance in the future with the increasing demand for natural resources. With ancient sediments, problems can arise in distinguishing true lake environ-
Introduction
5
ments (enclosed water bodies) from lagoons, estuaries, deltaic interdistributary bays and other paralic water bodies with a permanent or semi-permanent marine connection. In many cases the fauna or flora alone is sufficiently diagnostic of a continental setting (e.g. Van Dijk, Hobday & Tankard, 1 978; Hsii & Kelts, 1 978). When the fossils are not diagnostic or are absent, however, then the sediments themselves, and their facies sequences and relationships must be used. Few, if any, sedimentary structures are restricted to the lake environment (Picard & High, 1 972) but taken together an assemblage of sedimentary structures can be sufficient to indicate an enclosed water body (e.g. Clemmey, 1978). Sedimentary features which are common in lake sediments include wave-formed ripples, syneresis and desiccation cracks and 'varves'. The mineralogy of the sediments and geochemistry may help too, since certain clay minerals and evaporites are restricted to non-marine settings. The characteristic lacustrine facies sequence is a coarsening-upwards unit representing progradation of lake shore-zone coarse clastics over offshore basin-centre silts and muds (e.g. Van Dijk et al., 1978; Clemmey, 1 978). Two other characteristic features are the very rapid lateral facies changes from beach gravels to offshore silts, particularly at coincident lake margins, and the very rapid vertical changes, of lacustrine into subaerial facies, which occur at non-coincident lake margins (Donovan, 1975; Tucker, 1 978). Facies associations on a larger scale, such as relationships to fluviatile or aeolian sediments, will also provide valuable information on environmental interpretations, but in addition on basin geometry and depositional controls (e.g. Link & Osborne, 1978; True, 1 978; and Tucker, 1 978). There is still much to be learnt about Recent lake sediments to facilitate the interpretation of their ancient equivalents. It is likely that re-evaluation of ancient floodplain, aeolian, deltaic or paralic formations will reveal lacustrine sequences hitherto unrecognized. It is hoped that this book will make some contribution towards an understanding of ancient and modern lake sediments. R E F ER E N C E S BRADLEY, W.H. ( 1 929) Algae reefs and oolites of the Green River Formation. Prof Pap. U.S. geo/. Surv. 154,
203-233. BRADLEY, W.H. ( 1 93 1 ) Origin and microfossils of the oil shale of the Green River Formation of Colorado
and Utah. Prof Pap. U.S. geo/. Surv. 168. BRADLEY, W.H. ( 1 964) Geology ofthe Green River Formation and associated Eocene rocks in south western
Wyoming and adjacent parts of Colorado and Utah. Prof Pap. U.S. geo/. Surv. 496-A. BRADLEY, W.H. & EusTER, H.P. ( 1 969) Geochemistry and palaeolimnology of the trona deposits and
associated authigeni c minerals of the G reen River Formation of Wyoming. Prof Pap. U.S. geo/. Surv. 496-8. CLEMMEY, H. ( 1 978) A Proterozoic lacustrine interlude from the Zambian Copperbelt. In: Modern and Ancient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 259-278. CRAMPTON, C.B. & CARRUTHERS, R.G. ( 1 9 1 4) The geology of Caithness. Mem. geo/. Surv. U.K. DE GEER, G. ( 1 9 1 2) A geochronology of the last 1 2,000 years. Int. Geo/. Congr. XI Sess. Stockholm, 241-253. DONOVAN, R.N. ( 1 973) B asin margin deposits of the middle Old Red Sandstone at Dirlot, Caithness. Scott. J. Geol. 9, 203-2 1 1 . DONOVAN, R.N. ( 1 975) Devonian lacustrine limestones at the margin of the Orcadian Basin, Scotland. J. geo/. Soc. 131, 489-5 10. EARDLEY, A.J. ( 1 938) Sediments of Great Salt Lake, Utah. Bull Am. Ass. Petrol. Geo/. 22, 1 305- 1 4 1 1 . EuGSTER, H.P. & HARDIE, L.A. ( 1 975) Sedimentation i n an Ancient Playa-Lake Complex: The Wilkins Peak Member of the Green River Formation of Wyoming. Bull. geo/. Soc. Am. 86, 3 1 9-334.
A lbert Matter and Maurice E. Tucker
6
EuGSTER, H.P. & SURDAM, R.C. ( 1 973) Depositional environment of the Green River Formation of
Wyoming: A preliminary report. Bull. geo/. Soc. Am. 84, 1 1 1 5 - 1 1 20.
FANNIN, N . G.T. ( 1 969) Stromatolites from the middle Old Red Sandstone of Western Orkney. Geo/. Mag.
106, 77-8 8 . FOREL, F.A. ( 1 885) Les ravins sous-lacustre des fleuves glaciaires. C. r. hebd. Seanc. A cad. Sci., Paris, 101, 725-728. FOREL, F.A. ( 1 892) Le Leman, 1. GEIKIE, A. ( 1 87 8) On the Old Red Sandstones of Western Europe. Trans. Roy. Soc. .£din. 28, 345-452. GILBERT, G.K. ( 1 890) Lake Bonneville. Mongrr. U.S. geol. Surv. 1. GROVER, N.C. & HowARD, C.S. ( 1 938) The passage ofturbid water through Lake Mead. Proc. Am. Soc. Civ. Engr,. 103, 720-7 32. · GouLD, H.R. ( 1 95 1 ) Some quantitative aspects of Lake Mead turbidity currents. in: Turbidity Currents and the Transportation of Coarse Sediments to Deep Water: a Symposium. Spec. Pubis Soc. econ. Paleont. Miner. , Tulsa, 2, 34-52. HsD, K.J. & KELTS, K. ( 1 978) Late Neogene chemical sedimentation in the Black Sea. In: Modern and Ancient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 1 29- 1 45 . LINK, M.H. & OSBORNE, R.H. ( 1 978) Lacustrine facies i n the Pliocene Ridge B asin Group: Ridge B asin, California. In: Modern and Ancient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 1 67- 1 89. MILTON, C. & EuGSTER, H.P. ( 1 959) Mineral assemblages of the Green River Formation. In: Researches in Geochemistry (Ed. by P. H. Abelson), pp. 1 1 8- 1 50. John Wiley & Sons, New York. MORRISON, . B . ( 1 964) Lake Lahontan: geology of southern Carson Desert, Nevada. Prof Pap. U.S. geol. Surv. 401. NIP�ow, F. ( 1 920) Vorlaufige Mitteilungen iiber Untersuchungen des Schlammabsatzes im Ziirichsee. Z. Hydro/. l, 1 00- 1 22. NIPKOW, F. ( 1 928) Ober das Verhalten der Skelette planktischer Kieselalgen in geschichtetem Tiefenschlamm des Ziirich- und Baldeggersees. Z. Hydro/. 4, 7 1 - 1 20. PICARD, M.D. & H IG H L.R. ( 1 972) Criteria for recognizing lacustrine rocks. In: Recognition of Ancient Sedimentary Environments (Ed. by J. K. Rigby and W. K. Hamblin). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 16, 108-145. RAYNER, D.H. ( 1 963) The Achanarras Limestone of the middle Old Red Sandstone, Caithness, Scotland. Proc. Yorks. geo/. Soc. 34, 1 1 7- 1 3 8 . RuSSELL, I . C. ( 1 885) Geological history of Lake Lahontan, a Quaternary lake o f northwestern Nevada. Monogr. U.S. geol. Surv. 2. SEDGWICK, A. & MuRCHISON, R.I. ( 1 829) On the Old Conglomerates and other Secondary Deposits of the north coast of Scotland. Proc. geo/. Soc. Lond. I, 1 -77. SMOOT, J.P. ( 1 97 8) Origin of the carbonate sediments in the Wilkins Peak Member of the lacustrine Green River Formation (Eocene), Wyoming, U.S.A. In: Modern and Ancient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 109- 1 27 . TRUC, G. ( 1 978) Lacustrine sedimentation i n an evaporitic environment: the Ludian (Palaeogene) o f the Mormoiron Basin, S.E. France. In: Modern and Ancient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 1 89-203. TUCKER, M.E. ( 1 978) Triassic lacustrine sediments from South Wales in shore-zone clastics, evaporites and carbonates. In: Modern and Ancient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 205-224. VAN DuK, D .E., HOBDAY, O.K. & TANKARD, A.J. ( 1 978) Permo-Triassic lacustrine deposits in the Eastern Karoo Basin, Natal, South Africa. In: Modern andAncient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 225-239. ,
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7-4 1
Saline lakes and their deposits: a sedimentological approach
L A W R E N C E A . H A R D I E , J O S E P H P . S M O O T and HANS P. EUGSTER Department of Earth and Planetary Sciences, The Johns Hopkins University, Baltimore, Maryland, U.S.A.
A B S TRAC T Saline lakes (lakes with >5000 ppm dissolved solutes) are common throughout the arid regions of the world. Their distribution is controlled by tectonic setting and climate, and thus their deposits take on an importance beyond their size and abundance in the geological column. To exploit this aspect of saline lakes an understanding is needed of their sedimentary record. Saline lakes occupy the hydrographically lowest areas of closed drainage basins and are surrounded by a complex of interrelated depositional subenvironments that result mainly from the characteristics of the inflow. Our approach is to identify distinctive subenviron ments, each subject to a distinctive set of hydrological, biological, chemical and sedimentological processes and hence each with a diagnostic set of sedimentary features in their deposits, as follows: (I) alluvia/fan; coarse gravelly wedges composed of braid channel deposits, incised channel fills, sieve deposits and debris flows; (2) sandflat; flat unchannelled sandy apron at base of fan; planar and wavy laminated coarse sand (upper flow regime bedforms); (3) dry mudflats; exposed plain of mudcracked muddy sediment fringing the saline lake, covered with thin saline efflorescent crusts; sediment laminated but disrupted by mudcracks, sheetcracks, and saline mineral growth; (4) ephemeral saline lake consisting of an inner salt pan (thin beds of crystalline salts with mud partings) and an outer saline mudflat ( massive mud crowded with salt crystals that have destroyed layering); (5) perennial saline · lake; bottom sediment of laminated carbonates, gypsum, etc., or, if very saline, thin-bedded halite, etc.; (6) dunefield (aeolian deposits); (7) perennial stream floodplain (braid or meander deposits); (8) ephemeral stream floodplain (braid deposits); (9) springs; travertine and tufa mounds and sheets; ( I 0) shoreline features (deltas, beach ridges, spits, etc.). By recognizing from the sedimentary record which subenvironments were present and how they were arranged in space and time, we can interpret the history of a saline lake basin, either modern or ancient.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
8
L. A. Hardie, J.
P.
Smoot and H. P. Eugster
I N TRO D U C T I O N
Closed drainage basins with saline lakes occupying the hydrographically lowest areas are common throughout the arid regions of the world today. By saline lake we mean a lake that normally contains water carrying more than 5000 ppm of dissolved solutes, the upper salinity tolerance of most freshwater aquatic organisms (Beadle, 1 974, p. 264). Such saline lakes may be perennial (like Great Salt Lake, Utah) or ephemeral (like the Etosha Pan, southern Africa), may be hundreds of metres deep (like the Dead Sea, Middle East) or seldom covered by more than a few centimetres of brine (like Saline Valley, California), rnay expand to thousands of square kilometres at high stand (like Lake Eyre, South Australia) or never become larger than a fraction of a square kilometre (like the Basque Lakes, British Columbia), and may occur at elevations below sea level (like the Death Valley salt pan, California) or thousands of metres above sea level (like Lake Uyuni, Bolivia). All these saline lakes have in common an environmental setting in which annual evaporation exceeds annual inflow, and so their distribution and characteristic features are primarily controlled by the climate of the basins. Tectonism, however, can create orographic (rain-shadow) deserts far north or south of the arid sub-tropical global belts, so control of the distribution of many saline lakes is first and foremost a tectonic one. It is these climatic and tectonic controls that make saline lakes and their deposits take on a geologic importance well beyond their size and abundance in the geologic record. The processes that operate in saline lakes are hydrological, chemical, biological and sedimentological and are all intimately interdependent. A great deal is now known about the chemical aspects of saline lake brines (see, for example, Hardie & Eugster, 1 970; Eugster & Hardie, in press), something is known of the hydrology of closed lake systems (Langbein, 1 96 1 ; Jones, 1 965; Reeves, 1 968, pp. 1 34-1 53), a little is known about the unusual biota of saline lakes (Beadle, 1 94, pp. 259-282), but very little work has been done on the sedimentology ofsaline lake basins. While we will discuss each of these aspects, it is the last one, the sedimentology, that we want to emphasize in this paper. This aspect carries the greatest importance in ancient saline lake deposits because it is the sediments, both clastic and chemical, that preserve the most legible record of the hydrological and physical environment, a record without which the chemical processes could not be properly interpreted. Our presentation here of the sedimentary characteristics of saline lake basins represents only a tentative beginning and is far from comprehensive because many of the crucial sedimentological observations are missing. We hope, however, that this will be enough to encourage other workers to look more carefully at the details of the depositional features, particularly the facies patterns, of saline lake basins both modern and ancient.
H Y D R O L O G I C A L A S P E C T S O F S AL I N E L A K E S
We are concerned here only with a brief summary of the general aspects of the hydrology of saline lake basins, in particular what determines whether a lake will be saline or not, perennial or ephemeral, and what kinds of inflow systems move water, dissolved solutes and detrital sediment within the basin. The conditions which must be met for a saline lake to form are ( 1 ) evaporation must
Saline lakes and their deposits
9
exceed inflow, and (2) the basin should be hydrologically closed, or at the least, outflow must be very restricted. Arid (desert) or semi-arid (steppe) climates, where annual evaporation rate is greater than annual inflow rate, occur in a number of settings: (a) the dry, high-pressure, 'horse latitude' global belts of the subtropics, (b) orographic (rain-shadow) deserts and steppes, independent of latitude, (c) mid-latitude, mid continent deserts and steppes far from sources of ocean moisture (these dry areas, like the Gobi, Turkestan and Tarim deserts and adj acent vast steppes, may at least in part be orographic), and (d) the dry, high-pressure polar 'deserts'. Closed drainage basins and lakes with no outlet have many origins (see Hutchinson, 1957), some of the more common being (a) tectonic basins, particularly block-fault and rift valleys, (b) wind deflation hollows, (c) interdunal depressions in a windblown dune field, (d) volcanic craters, etc., (e) valleys dammed by lava flows or by landslides, (f) cut-off stream meanders, and (g) stream floodbasins isolated by channel levees. Perhaps the most favourable settings for saline lake development are rain-shadow tectonic basins like the block-fault desert basins of the western United States, combining arid climate with hydrological closure. In these basins the enclosing mountains (horsts) act as effective traps for precipitation keeping the basin floors (graben) arid, yet, at the same time, providing enough inflow for significant solute accumulation in the saline lake. Death Valley, California, is an excellent example of this type of setting, where rainfall on the valley floor is < 6 em per year while in the upfaulted Panamint Range more than 2500 m above the valley floor the annual precipitation averages more than 35 em (Hunt et a!. , 1 966, pp. 5-7). The nature of the inflow to a saline lake-whether the inflow is perennial or intermittent, whether it is high or low compared to evaporation rate-will determine if the lake will be perennial or ephemeral. Stable perennial lake conditions will be favoured by perennial inflow and a low evaporation-to-inflow ratio, while ephemeral lake conditions are brought on by intermittent inflow and a high evaporation-to inflow ratio. For example, although the Dead Sea is undergoing evaporation at a rate of 160 em/year resulting in a loss of 1· 58 km3 /year of water, it remains a deep perennial lake (400 m) with an annual water level change of no more than 50 em. This balance is maintained mainly by the perennial inflow of the Jordan River which provides an estimated 1 ·25 km3 of water annually (Neev & Emery, 1967, pp. 72-73). Lake Eyre, Australia, on the other hand, with an annual evaporation rate of about 2 1 5 em/year has no perennial stream input and remains dry except when a rare flashflood inundates the basin with freshwater runoff that is channelled to the lake by a massive ephemeral braid stream system (see Bonython & Mason, 1 953). Inflow to saline lakes can take the form of (a) perennial streams that feed directly into the lake, (b) ephemeral streams, which flow only occasionally with storm runoff or perhaps seasonally with meltwaters from the spring thaw in the mountains, (c) unchannelled sheetflow during storms, (d) perennial or ephemeral springs at the lake edge, and (e) perennial or ephemeral groundwater input. Saline Valley, California (Hardie, 1 968; Eugster & Hardie, 1 978), is an example of an ephemeral saline lake that is recharged only by sheet�ow storm-runoff, springs and groundwater. The Inyo Mountains, over 3000 m above the Saline Valley floor, trap enough precipitation to maintain small perennial mountain streams. These mountain streams sink into the porous alluvial fan gravels as soon as they debouch from the mountain canyons, and the water then flows down to the valley floor as a perennial groundwater body. In places this groundwater surfaces at the dry playa edge as isolated perennial springs but
10
L. A . Hardie,
J.
P. Smoot and H. P.
Eugster
the bulk of this subsurface inflow evaporates from the vadose zone to produce a highly concentrated brine that soaks the sediments of the playa throughout the year, and allows salts to continuously accumulate. Like Lake Eyre, only after sporadic flashfloods does the Saline Valley lake contain a surface water body, but this water reaches the lake by direct sheetflow runoff from the fringing alluvial fans rather than by ephemeral stream channels. It is only during these flashfloods that detrital sediment is transported and deposited in the Saline Valley basin. For detailed information on the hydrology of specific saline lakes the reader is referred to works such as those of Jones ( 1 965) and Phillips & Van Denburgh ( 1 97 1), and for a general discussion o f hydrological cycles in closed basin lakes the study of Langbein ( 1 96 1 ) is recommended.
C H E M I C A L E V O L U T I O N O F S A L I N E LA K E B R I N E S
Chemical aspects of saline lake environments have been discussed by Jones ( 1 966), Jones & Van Denburgh ( 1 966) and Hardie & Eugster ( 1 970), among others, and have been summarized at some length by Eugster & Hardie ( 1 978). These latter authors considered the acquisition of solutes by the inflow waters, the modification of water compositions by early precipitation of alkaline earth carbonates and gypsum, the development of concentrated brines from which the most soluble saline minerals crystallize, and finally the diagenetic reactions between sediments and occluded brines. In this report we will give a brief synopsis of these aspects, and leave the reader to consult the above cited works for detailed information. Saline lakes are known to have a wide range of composl.tions, dominated by the solutes Si02, Ca, Mg, Na, K, HC03, C03, S04 and Cl. Na is by far the most abundant cation, while the anion concentrations are quite variable. The major brine types are (a) Na-C03-Cl-S04, (b) Na-Cl-S04, (c) Na-Mg-Cl-S04 and (d) Ca-Mg-Na-Cl. Hardie & Eugster ( 1970) have identified the factors which determine the composition of a particular salt lake. They found that the final composition of the brines is inherited from the very earliest stages of water evolution, the weathering reactions which occur in the watershed. Since these reactions are dominated by the nature and compositions of the minerals involved, it is basically the bedrock lithology which controls water compositions. Bedrock weathering reactions are of several kinds: (a) congruent dissolution of non-silicates like halite, gypsum, calcite, dolomite; (b) congruent dis solution and hydrolysis of non-aluminous silicates like olivines; (c) incongruent dissolution and hydrolysis of alumino-silicates like feldspars to produce clays; and (d) oxidation of metal sulphides to produce metal oxides and sulphate ions, e.g. pyrite gives geothite + SO/-. It should be noted that carbonic acid (atmospheric C02 dissolved in rainwater) weathering is a maj or chemical process in arid closed basins, as is demonstrated by the dominance of HC03- in closed basin inflow waters (see, for example, analyses in Jones, 1 965, and Hardie, 1 968) and by the abundance of clay minerals in the fine sediments washed into desert basins (Droste, 1 9 6 1 ). In this regard the weathering of a plagioclase-rich igneous or metamorphic rock can contribute as much Ca2 + and HC03- to the inflow waters as dissolution of limestones. As the inflow waters with their solutes newly derived from the weathering reactions in the highlands move downslope toward the closed lake, they are subject to evaporative concentration. This includes direct evaporation of surface waters,
Saline lakes and their deposits
11
subsurface evaporation of groundwater in the vadose zone, as well as evapotranspira tion. C02-degassing may also take place as the C02-charged groundwaters surface and exchange with the atmosphere. These evaporation and degassing processes lead to ever increasing solute concentrations until supersaturation with the least soluble chemical precipitates is reached. All the inflow waters we have tested (Hardie & Eugster, 1 970), whatever their source bedrock type, show saturation first with alkaline earth carbonates, calcite or aragonite or Mg-calcite (Nesbitt, 1 974). The early precipitation of Ca, Mg and COl has a profound effect on the subsequent fate of the brine chemistry. If HCOl >> Ca + Mg at the point of initial alkaline earth carbonate precipitation, then the subsequent evaporative concentration will produce a water rich in COl + HCOl and depleted in Ca + Mg, a sodium carbonate brine. For water initially with Ca + Mg >> HCOl, a brine enriched in alkaline earths and depleted in COl + HCOl will result. The amount of carbonate formed depends on the initial HCOl/Ca + Mg ratio: if that ratio is small or very large, little carbonate can form. Conversely, for molar ratios near unity, carbonate production can be extensive. In that case, low-magnesian calcites are followed by high-magnesian calcites and eventually protodolomite and in some cases even magnesite, because Mg is enriched
II
Co+Mg» HC03 Co »Mg
Co rich HC0 3 poor
1
HC03?: Co + Mg Mg?: Co
Mg CI+S04
pptn.
c:;
No- S04-CI
brine
br ine
Bristol
Dry L.
HC03 rich Co poor
>Co > HC03
�
4rich Co poor
\l
Co-No- CI
C03 »Co+Mg Co»Mg
M g c ol c i te ppt n.
low
gypsum Co r\ so4 poor
� II I
in f I o w
unde rsaturated
Na-C03-S04-CI
high Mg cal cite pptn.
(±
pr otodolo mite )
aragonite pptn.
or
Co+Mg »>HC03
Soline Volley Death Vo I ley
"t
Co »
gypsum pptn.
\i
S04 S04»>
'); �
HC03
Mg » Co
brine
�
L. M ogod i >Co+ Mg
- No- Cl- s o4 -(C03l
a
Deep Springs
No- Mg-(Co)-CI
No- Mg-S04-CI
br ine
brine
Dead
Sea
B o sque
I
L.
L.
Fig. I. A schematic flow sheet for brine evolution as evaporative concentration progresses (direction of arrows).
12
L . A. Hardie, J.
P.
Smoot and H. P. Eugster
preferentially in the residual waters (Fiichtbauer & Hardie, 1 976). * These effects are shown schematically in simplified form in Fig. l (a more complete scheme is given by Eugster & Hardie, in press, fig. 5). Upon further evaporative concentration many waters next become saturated with respect to gypsum. This is still an early precipitation product, and waters at first precipitation of gypsum will normally have an ionic strength of less than l (Hardie & Eugster, 1970, p. 279). Like calcite, gypsum precipitation is an important branching point, with the subsequent path of the water depending upon the Ca/S04 ratio; that is, waters may now become depleted in Ca or in S04, producing either a Na-Cl-S04 or a Ca-Na-Mg-Cl brine (Fig. 1 ). The nature and fate of the precipitation products will be discussed later, but it is clear that a zonal arrangement may result, with the less soluble carbonates on the outside of the evaporating basin, in contact with more dilute waters, and gypsum closer to the centre (Hunt et a!., 1 966; Hardie, 1 968). The next minerals to form upon evaporation are usually quite soluble and their saturation is not reached until the solute load is increased manyfold. During this stage of evaporative concentration an additional process for increasing the solute load becomes important: dissolution of efflorescent crusts by rain and ephemeral runoff. Such crusts are formed at or near the surface and they are the product of complete evaporation to dryness of surface water or groundwater drawn to the surface by evaporative pumping (Hsii & Siegenthaler, 1 969). Dissolution, on the other hand, is fractional, with the most soluble salts being removed first. Most of the alkaline earth carbonates contained in the crusts are not redissolved and thus are permanently lost from the evolving brines. Such losses, as well as that of silica, have been discussed by Eugster ( 1 970), Jones, Eugster & Rettig ( 1 977) and Eugster & Jones ( 1 978). Other important losses were observed for potassium, removed through ion exchange reactions, and sulphate, involved in bacterial reduction. The most extensive gains in efflorescent crust dissolution are achieved by Na, Cl and some of the less abundant elements, such as Br, B, F. By now the solute load of the brines may have increased more than a thousand-fold over that of the inflow waters and the stage is set for final mineral precipitation. This can occur in one of two modes, (a) at the surface from the open lake brine, or (b) within the sediment from occluded brine. The nature of the minerals precipitated from such brines is of course dictated by the evolutionary track the brines have followed previously. Common products are mirabilite (Na2 S04. 10H20), thenardite (Na2 S04), halite (NaCl), trona (Na2 C03.NaHC03.2H2 0), burkeite (Na2C03.2Na2 S04), bloedite (Na2 S04.MgS04.4H2 0), epsomite (MgS04.7H2 0). Common salts formed within the sediment by reaction of the earlier-formed precipitates with occluded brine are gaylussite (Na2 C03.CaC03.5H20), pirssonite (Na2 C03.CaC03.2H20), glauberite (Ca S04.Na2 S04), nahcolite (NaHC03). Interaction of occluded brine with detrital sediment may also produce less soluble minerals such as authigenic silicates. The best-known examples are the zeolites formed by volcanic glass reacting with alkaline brines (for a recent summary see Eugster & Hardie, 1 978). Many ofthese deposits exhibit mineral zonation, presumably recording concentration gradients in the subsurface brine. Common successions, from dilute perimeter to the saline centre, are: montmorillonite-zeolite (phillipsite, *For kinetic reasons aragonite may form instead of calcite, particularly in dilute surface lakewaters with initially high Mg/Ca ratios (Fiichtbauer & Hardie, in preparation).
Saline lakes and their deposits
13
clinoptilolite, erionite, mordenite)-analcime-K feldspar. A wide variety o f other minerals formed in a similar manner has been reported, the most spectacular list being assembled for the Eocene Green River Formation (Milton, 1 97 1 ).
B I O L O G I C A L P R O C E S S E S I N S AL I N E L A K E S
Tolerance of some organisms to high salinity (over 50o/00) and/or high alkalinity is extraordinary, for example, the unusual Tilapia fish, the brine shrimp A rtemia, several species of rotifers, copepods, nematodes, insect larvae, worms, blue-green algae, bacteria and halophyte higher plants (see Beadle, 1 974, pp. 259-282 for a summary of saline lake biota). These salHolerant organisms may have an important input to the sedimentary record of saline lakes, particularly in connection with carbonate precipitation, bioturbation of bottom sediment, and deposition and decomposition of organic matter. In non-saline temperate zone perennial lakes, summer blooms of planktonic algae are thought to be responsible for precipitation of calcite. Photosynthesis removes C02 from the water, increasing pH and hence CO/- activity, which may lead to supersaturation with respect to CaC03• This is the model used by Nipkow ( 1 920) to explain the annual CaC03 laminae in Lake Zurich. Whether a similar model applies to saline lakes is not clear. In the Dead Sea aragonite precipitation is not seasonal but continuous throughout the year and appears to be inorganically precipitated (Neev & Emery, 1 967). So, too, in Deep Springs Lake, California, the alkaline earth carbonates apparently are inorganic precipitates (Jones, 1 965; Peterson, Von der Borch & Bien, 1 966; Clayton, Jones & Berner, 1 968). Blue-green algae certainly are associated with precipitated tufas and travertines in inflow stream channels and around spring orifices in saline lake basins (e.g. Dunn, 1 953; Scholl, 1 960; Slack, 1 967). Again, however, it is not clear whether the metabolic processes of the algae have induced carbonate precipitation or whether the algae are simply entombed by inorganically precipitated carbonate. The tufas and travertines are discussed in more detail under 'spring subenvironment.' Bioturbation of the sediments of saline lake basins by organisms is particularly obvious in deposits surrounding the lake. For example, in alluvial fans, sandflats, dunes and mudflats the depositional layering may be disrupted or destroyed by halophyte plant roots (cf. dikaka of Glennie & Evamy, 1 968; Glennie, 1 970, pp. 1 1 3- 1 1 7), by insect burrows, by oligochaete worm burrows, by bird feeding pits and by terrestrial animal burrows. In perennial saline lakes the bottom sediments may be burrowed and pelleted by worms and brine shrimp. In lake brines that contained dissolved SO/ - , sulphate-reducing bacteria (Baas Becking & Kaplan, 1 956; Goldhaber et a!. , 1 977) in bottom sediments can play an important diagenetic role: they induce anaerobic conditions that favour dissolution of gypsum and precipitation of iron sulphides and phosphate (as discussed below under 'perennial lake subenvironment'). Blue-green algal mats and coccoid ooze may accumulate on the bottoms of saline lakes. These algal deposits will be the prime candidates for diagenetic conversion to 'kerogen', the hydrocarbon complex that makes the organic components of the 'oil shale' of the Eocene Green River Formation (Bradley, 1 973). Before the saline lakes
14
L. A . Hardie, J. P. Smoot and H.
P.
Eugster
become too saline, diatoms may thrive and diatom ooze may collect on the bottom, effectively removing Si02 from the lake waters and storing it as a sediment deposit (see, for instance, the Dead Sea, Neev & Emery, 1967, p. 8 1). Finally, phreatophyte plants (Robinson, 1 958) that typically occur as widely spaced but abundant growths on fan toes, sandfiats and mudflats where the water table is shallow, act to increase the dissolved solute concentration of groundwater by evapotranspiration. Hunt ( 1 966) has excellent descriptions of such phreatophytes (as well as xerophytes) in Death Valley, California.
S U B E N V IR O N M E N T S O F S A L I N E L A K E D E P O S I T I O N A L COMPLEXES
Saline lakes are surrounded by a complex of genetically interrelated depositional subenvironments* that result mainly from the characteristics of the inflow to the lake. So, from a sedimentological point of view, it is necessary to consider the closed saline lake basin as an integrated whole with the saline lake as simply one part of a complex system of subenvironments. It is our experience that no two modern saline lake basins are quite the same, and so it is difficult to present a typical model for saline lake deposition. Our approach instead is to isolate the individual depositional subenviron ments that are associated with modern saline lakes, and to describe the diagnostic features of their sediments (sedimentary structures, textures, geometry, etc.) and the processes responsible for these features. With this approach the individual subenvironments, like building blocks, can be assembled as necessary to provide an overall sedimentological description of any particular saline lake depositional complex. Which subenvironments are actually present and how they are arranged in both space and time represent sensitive responses to the particular processes that are operating, or have operated, in a particular saline lake basin. This kind of approach has been outlined for carbonate tidal-fiat deposits by Hardie ( 1977, pp. 1 88- 1 8 9) and has great utility in reconstructing ancient depositional environments, provided the appropriate criteria are available from modern deposits. With this approach in mind we have outlined below the significant features of the different major sedimentary subenvironments that we have recognized in modern saline lake basins. The subenvironments we have considered are, (I) alluvial fan, (2) sandfiat, (3) mudflat, (4) ephemeral saline lake (saline mudflat and salt pan), (5) perennial saline lake, (6) dune field, (7) perennial stream floodplain, (8) ephemeral stream floodplain, ( 9) springs, ( 1 0) shoreline features of saline lakes. We will first discuss the characteristic features of each subenvironment in turn and then we will very briefly present a few examples of typical combinations of subenvironments found in modern saline lake basins, such as those of Death Valley (Fig. 4), Saline Valley (Figs Sa and b) and Deep Springs Valley (Fig. 6) in California; Lake Eyre basin (Fig. 7) in Australia; the Dead Sea basin in the Middle East; and the Great Salt Lake basin (Fig. 8) in Utah. By 'subenvironment' we mean any part of the surface of the basin that has a distinctive physiography and on which a distinctive set of physical, chemical and biological processes operate (see Hardie, 1 977, p. 3 ) The rock record of a subenvironment is a 'subfacies' (Reinhardt & Hardie, 1 976, p. 1 6). •
.
Saline lakes and their deposits
15
Alluvial fan and sandflat subenvironments
A large number of saline lake basins are formed as a result of block-faulting and rifting. The resulting high relief in such horst-and graben systems (Fig. 3c) leads to the development of coarse gravelly alluvial wedges that surround the saline lake of the fiat valley floor (see, for example, figs 3, 4, 5 , 28, 57, 58, 6 1 and 75 in Hunt & Mabey, 1 966). These gravelly wedges are coalescing alluvia/fans, which in turn are cone-shaped piles of very coarse sediment built out radially onto a fiat valley floor from a stream point source in the mountains (see Bull, 1 972; Cooke & Warren, 1973, p. 1 74). We have found that alluvial fans in saline lake basins commonly grade downslope into a fringing sandy apron that has features different from the alluvial fan as well as from stream floodplains. This distinctive sandy apron to alluvial fans Hardie ( 1973) has been called the sandjlat subenvironment. Alluvial fans are concave in radial profiles and convex in cross-fan profiles (Bull, 1 972, p. 63) with average slopes generally less than 1 0° (Cooke & Warren, 1 973, fig. 3.5); in the rugged Basin-and-Range province of the western U.S. fan slopes are commonly greater than 4°. Relief from apex to toe of the fan may be as much as 500 m over a distance of 20 km as in Baja California at the northwest corner of the Gulf of California. If faulting is long-lived then very thick (but rather narrow) fan deposits measured in thousands of metres can accumulate (see e.g. Bull, 1 972, p. 80 and fig. 16). Alluvial fan surfaces are dissected by a characteristic radial pattern of braided ephemeral stream channels that become shallower and less distinct toward the toe of the fan. This pattern reflects the essential depositional events that build alluvial fans: catastrophic storm-flooding in the mountains produces a surge of water which gushes down the canyon feeder streams and spreads out onto the fan (McGee, 1897, pp. 99- 1 05). These flood events produce four maj or kinds of deposits on fans (cf. Bull, 1972, pp. 66-7 1 ; Blissenbach, 1 954, pp. 1 78- 1 79): ( 1) shallow braid channel deposits; (2) fills of deeper, incised channels; (3) sieve deposits; and (4) debris flow deposits. Braid channel deposits (sheetflood sediments of Bull, 1 972, pp. 66-68) are seen on the fan surface as coarse gravel bars isolated by a braid system of very shallow channels ( < 1 m deep), or 'washes', floored by coarse sand, grit and fine gravel (Fig. 2a, see also figs 77 and 78 in Hunt & Mabey, 1 966). In vertical sections these deposits, like typical braid stream sediments (see Doeglas, 1 962, Williams & Rust, 1 969) are a series of cross-cutting lenses (Spearing, 1 975, fig. 1 ). Thickness is typically less than 1 m. The gravel bars consist of lenses of framework-supported boulders and pebbles (Fig. 2b), moderately sorted, and, in some cases, imbricated. Sand and grit may fill the framework voids by infiltration. The interbar 'wash' sediments are recognized by the planar-parallel horizontal lamination and low angle inclined laminated gritty sand. Maximum grain size in the gravel fraction decreases dramatically down-fan (Krumbein, 1 942). It appears that at maximum flood velocity the gravel lenses are produced as large lenticular longitudinal bars. As the flood wanes, the water shallows and the stream competence decreases so that fine gravel, grit and sand are deposited in the interbar channels as upper flow regime bedforms (plane beds, antidunes and 'washed out' dunes, Williams, 197 1 , p. 8; Simons, Richardson & Nordin, 1 965, p. 37). Incised channelfills are narrow gravelly lenses elongated downstream and generally much thicker than the braid channel deposits, up to several metres. They represent fills of sinuous channels cut earlier by the same flood at maximum discharge or by previous .floods. These fills commonly show a fining-upwards in grain size (see Glennie, 1 970,
16
L . A . Hardie, J. P. Smoot and
H.
P. Eugster
Fig. 2. (a) View of the surface of an alluvial fan, B aja California, showing a gravel bar (right) and a gritty,
sandy, braid channel or wash (left). Trenching tool for scale. This is a typical view of the mid-fan zone. (b) Cross-section through an alluvial fan deposit, Death Valley, California, showing typical flat lense-shaped bedding of gravels and pebbly grits and sands. (c) Cross-section through sandflat deposit, Baja California, showing wavy and inclined bedding (antidune bedding) and scour-and-fill structures. The surface of the sandflat was covered by a thin efflorescent halite crust (chips at the surface are crust fragments broken during trenching). Ballpoint pen for scale. (d) Polygonally cracked surface of layered halite of salt pan, Saline Valley, California. Pick handle for scale (approx. I m long). A fluffy efflorescent halite crust is forming at the surface of the pressure-cracks as brine is drawn up the cracks by evaporation. This salt pan is surrounded by the efflorescent crust covered saline mudflat shown in Fig. 3c.
fig. 18; Picard & High, 1973, pp. 1 98- 1 99), and overall sorting is not as good as that of the gravel bars of the braid system (Bull, 1964). Mudflow deposits are also found as channel fills. Sieve deposits on fans were first recognized and described by Hooke ( 1 967, pp. 453-456). These unusual deposits consist of longitudinal tongues of relatively well sorted open-framework gravel (see fig. 5 in Bull, 1 972) that sieve out coarse gravel by allowing floodwaters to sink down into the sediment, carrying with them the finer sediment and leaving only boulders on the surface. Debris flows are uniformly thick (up to 3 m) lobes of very poorly sorted 'pebbly mudstones' or 'sandy mudstones' that have resulted from downslope flow of a dense, viscous, wet, muddy sediment in which coarser grains are suspended. Bull ( 1 964, table 1 7) showed that debris flows are 40-90% mud. Debris flows that are made up of sandy mud, rather than gravelly mud, are usually identified as 'mudflo ws'. Profiles across old mudflows in Death Valley (Hunt & Mabey, 1 966, fig. 48) show well defined
Saline lakes and their deposits
17
channel-levee systems. Bull ( 1972, p. 70) noted that if the debris flow is very fluid then graded bedding and imbricate structure can be seen in the gravel component. For very viscous flows, the internal texture is uniform, although vertical orientation of platy boulders may be seen. On the basis of the surface features of fans and the distribution of the above four kinds of fan deposits we subdivide the fan subenvironment into three zones, ( I ) fan apex, (2) mid-fan, and (3) fan toe (see also Spearing, 1975). At the fan apex the radiating channels are well defined, few in number and commonly incised (e.g. fig. 1 in Bull, 1 972). The sediments are typically channel fills and sieve deposits. In the mid-fan zone the shallow braided channel-gravel bar system dominates so that the sedimentary record is mainly boulder-gravel lenses interbedded with planar horizontal to inclined laminated sands and grits (Fig. 2b). At the fan toe the braid channels are still the main dispersal systems but the channels are extremely shallow and the gravel bars are mainly in the pebble size range. Sand and grit dominate overall in these fan toe sediments. At the distal end of the fan toe where it passes into the sandfiat subenvironment the sand in the channels may be rippled (mainly linguoid current ripples), and, in some cases, the waning currents may produce a 'powering down' vertical sequence of poorly laminated gritty sand overlain by horizontally laminated medium sand and capped by a ripple cross-laminated fine sand (Hardie, 1 973), a sequence much like the Bouma sequence usually attributed to waning turbidity currents in deep water (Bouma, 1 964). The distal end of the fan toe will grade into the sandjlat subenvironment where the braid channels lose their identity and the floodwaters disperse as unchannelled, unconfined sheetfioods across a narrow fiat ( < < 1 slope) sand plain (Hardie, 1 973). Sand is the main component of the sediment and characteristically occurs as planar parallel horizontal laminae and wavy laminated beds (Fig. 2c). These structures are produced by plane bed and antidune flow regime bedform migration: even though the flow competence is only in the sand range the flow is upper regime because the water depths of the flood sheets are very small (perhaps only a few centimetres). Shallow ponding of water on the sand fiat as the adjoining saline lake expands with the flood may result in wind wave reworking of the surface of the sandfiat, producing a variety of shallow-water wave-ripple bedforms (e.g. interference ripples, fiat-top ripples) on top of the sheet-flood laminated layer. Deposition by each flood event may range from a package of laminated sand tens of centimetres thick to a thin sheet one or two centimetres thick (Fig. 3d). Both the sandfiat and fan toe surfaces are normally subaerially exposed and so are subject to significant reworking and redeposition of sand by the wind, a point emphasized by Glennie ( 1970, pp. 29-56) for 'wadi' sediments. Indeed, the sandfiat surface and sediment may take on features very similar to the backshore of a beach, with cut-and-fill structures, low-angle inclined bedding and heavy mineral lag laminae. Small wind-blown dunes, with avalanche cross-bedding, slump structures, etc., may march across the sandfiat and up the fan toe where they may coalesce into a large dune field. Significant post-depositional diagenetic processes operate on alluvial fans and sandfiats. Perennial springs that surface along fault-zones at the fan apex will give rise to short travertine-lined channel bottoms (Slack, 1967) along which the spring waters flow before sinking into the porous fan sediment. Travertine and tufa may also form from springs that emerge along the fan toe edge (Jones, 1 965, pp. 34-35; and Hunt & •
18
L. A . Hardie, J. P. Smoot and H. P. Eugster
Mabey, 1 966) (see 'spring subenvironment'). Evaporative pumping (Hsii & Siegenthaler, 1 969) from the groundwater body below the fan surface, particularly in the braid channel 'wash' sediments, rnay result in the formation of calcite pore cements and caliche crusts, coatings, root moulds and nodules within the fan sediment (Bull, 1 972, p. 65; Lattman, 1973; Glennie, 1 970, pp. 33-36; Hardie, 1 968, p. 1 288; Blissenbach, 1 954, p. 185; Hunt et a!. , 1 966, p. 1 4). At the distal end of the fan toe and beneath the sandflat in the vadose zone the pore and vug filling cements are more likely to be gypsum and/or high-Mg calcite (perhaps even 'protodolomite'). If gypsum does form in these zones then dehydration to anhydrite and/or bassanite may occur (Hardie, 1 967). The groundwater from which these chemical precipitates formed could be recharged by the aperiodic surface flooding (ephemeral recharge) and from perennial springs and mountain streams (perennial recharge, as for example, in Saline Valley, California, see Hardie, 1 968). Phreatophyte and xerophyte bushes and trees with large root systems, like creosote bush (Larrea tridentata), paloverde (Cercidium sp.) and mesquite (Prosopis sp.) of the western U.S. desert basins (see Hunt, 1 966), will not only disrupt the layering in alluvial fan and sandflat sediments, but also act by evapotranspiration to concentrate both ephemeral and perennial groundwater (phreatophytes are particularly effective on fan toes where perennial groundwater is shallow but still relatively dilute; see Hunt, 1 966, fig. 5). On death and decay of the roots, the network of root holes provide important conduits for groundwater movement and hence can also become plugged with calcareous caliche or gypsum. Bushes may also act as baffles for wind blown sand and build small pyramidal mounds of sand on the fan toe and sandflat. Finally, pebbles and boulders exposed for long periods on fan surfaces can become coated with 'desert varnish' (Hunt et a!., 1 966, p. 6; Lustig, 1 965, p. 1 34; Hunt & Mabey, 1 966, pp. 90-92; Glennie, 1 970, pp. 1 9-20), a film of iron-manganese oxides thought to be produced by in situ weathering. Some authors, such as Bull ( 1 972, p. 80), believe that desert varnish can be used as a criterion for recognizing ancient alluvial fans.
Dry mudflat subenvironment
Saline lakes, particularly ephemeral ones, are fringed by a subaerially exposed plain of fine-grained sediment, a supralittoral mudflat. This mudflat subenvironment would fall within the definitions of 'playas' and 'inland sabkhas' of other workers (e.g. Cooke & Warren, 1 973, pp. 2 1 5 and 2 17; Glennie, 1 970, p. 60; Neal, 1975, p. l). We have distinguished between a 'saline mudflat', saturated with brine from which a mass of salt crystals have grown to completely destroy the depositional sedimentary structures, and a 'dry mudflat' in which the depositional sedimentary structures are well preserved. The 'saline mudflat' has strong affinities with ephemeral salt pan deposition and so is dealt with under the 'ephemeral saline lake' subenvironment. Here we will discuss only the dry mudflat subenvironment. The surface of the dry mudflat subenvironment is characterized by polygonal mudcracks and thin saline crusts. The mudcracks are generally narrow (a millimetre or two wide) and irregular to wavy, isolating polygons typically 5-25 em across (Fig. 3a; see also Cooke & Warren, 1 973, figs 2 . 1 4 and 2. 1 6). The mudcracks may extend to depths of a few millimetres to tens of centimetres. The surface crusts are of two kinds: ( I ) very thin (a millimetre or two thick), hard, dense crusts made of micritic alkaline
Saline lakes and their deposits
19
3. (a) Close-up view of mudcracked surface of dry mudflat, B aj a California. Pencil for scale. Note the secondary mudcracks within the primary polygons. Gypsum is crystallizing within the mudcracks. (b) Vertical slab of plastic-impregnated core through dry mudflat sediments, Baja California, showing characteristic millimetre-lamination. Pencil point for scale. Gypsum (white euhedral shapes) has crystallized within sheetcracks, mudcracks and fenestral pores. (c)View ofblocky efflorescent crust covering saline mudflat of Saline Valley, California. This porous crust, about 30 em thick, is so hard it must be broken with a pick in order to get through to the underlying gypsum- and glauberite-studded, brine-soaked soft mud. Note the very steep faulted valley walls that rise 3000 m above the mudflat. At the base of these precipitous mountains can be seen coalescing alluvial fans which are some 300 m high. (d)Trench dug into a saline mudflat bordering a sandflat, B aj a California. B allpoint pen and buried sample core-can for scale. The sandy-silt sediments are crowded with anhydrite nodules and lenses of nodules, and partly dehydrated gypsum crystals, the growth of which has destroyed most of the layering. The surface of the shallow brine groundwater body is exposed at the bottom of the trench. The efflorescent crust that once covered this saline mudflat was dissolved during a massive storm which washed a thin layer of wavy-laminated sand over the mudflat (top of trench), prograding the distal end of the sandflat over the mudflat. Since this photo was taken an efflorescent halite crust has once again covered the surface.
Fig.
earth carbonate (commonly high-Mg calcite or 'protodolomite' as in Salt Flat Graben, Texas); these brittle crusts may crack into small (millimetre or centimetre scale) platy fragments and be reworked by wind or flood water into pockets of fiat-chip gravel, grit or sand intraclasts; and (2) puffy, porous crystalline crusts up to several centimetres thick of soluble saline minerals such as halite, thermonatrite, trona or thenardite; these
20
L. A. Hardie,
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Smoot and H. P. Eugster
salt crusts are dissolved with each succeeding stoim flooding event and so are not preserved in the underlying sediment, but their effects may be seen as disruption and destruction of the upper part of each depositional unit. The mudcracks are desiccation cracks while the crusts are surface efflorescences chemically precipitated from subsurface brines drawn up through the vadose zone to the surface, probably by evaporative pumping (Hsii & Siegenthaler, 1 96 9). Where the subsurface brines are relatively dilute, as typically occurs near the outer edge of the mudflat, the crusts are likely to be the carbonate micrite type, but where the subsurface brines are quite concentrated and hence have evolved beyond the alkaline earth carbonate stage (Hardie & Eugster, 1 970; Eugster & Hardie, 1 978), the crusts are of the soluble salt type. Other surface features of the mudflat are patches of sand (both clastic grains and intraclastic peloids) that are starved wind ripples or longitudinal trains blown across the mudflat from the adjoining sandfl.at and alluvial fan or derived from erosion of the mudcrack polygons on the mudflat itself. In some places at the outer edge of the mudflat, small sandy rills may protrude from the sandfl.at onto the mud surface. Also at the outer edge of the mudflat where it meets the sandfl.at, springs may surface and deposit mounds of travertine and spring tufa (see under 'spring subenvironment'). The sedimentary features within the sediments of the mudflat subenvironment of saline lakes are not well known, and this represents a major gap in our documentation of saline lake sedimentation. We present here some of our cursory findings as well as apply data from analogous settings, e.g. marginal marine supratidal saline mudflats such as those of the extensive playa-like sabkha at the northwest corner of the Gulf of California (Thompson, 1 96 8 ; Hardie, 1 973). The predominant sedimentary structures that should be expected in the sediments of the mudflat subenvironment are: ( I ) millimetre-scale lamination which consists o f either (a) lenticular laminae o fvery fine sand and coarse silt alternating with more continuous laminae of fine silt and clay (Fig. 3 b) or (b) graded laminae and very thin beds; (2) disruption of the laminae by shallow and/or deep mudcracks, commonly filled or partly filled with fine sand and silt (including peloids); deep cracks may be filled with several different generations of sand and silt; (3) disruption of the laminae by horizontal sheet-cracks that propagate from mudcracks along bedding planes (Fig. 3b); (4) crystals of gypsum, glauberite, mirabilite, etc. may fill or partly fill mudcracks and sheet-cracks which have acted as conduits for subsurface brine movement beneath the mudflat (Fig. 3b); (5) disruption of the laminae by displacive growth of saline minerals like gypsum in the vadose zone; (6) destruction of layering by efflorescent crust growth of soluble saline minerals like halite at the mudflat surface; (7) thin dense alkaline earth carbonate micrite crusts interlayered with the unlithified sediment laminae; and (8) thin lenses of mudchip or micrite crust intraclast sands and grits. We speculate that there might be three basic types of mudflat, each produced by a different kind of depositional process and resulting in a different kind oflayering, ( 1) sheetwashed mudflats, (2) ponded water mudflats, and (3) exposed old perennial lake bottom mudflats. Deposition on sheetwashed mudflats would occur when thin sheets of sediment-charged stormwaters stream off the sandfl.ats and across the mudflats en route to the central saline lake. The essential process here would be traction deposition from moderate velocity, but very shallow, unchannelled flow (hence fine-grained sediment but upper flow regime bedforms). The resulting layering probably would be graded, with traction load, fiat lenticular sandy or silty laminae capped by a fine mud
Saline lakes and their deposits
21
(clay-sized sediment) drape deposited when the flow waned. These mud drapes, as well as the finer-grained lenticular lamination, would become mudcracked (centimetre scale) on drying out; such small shallow mudcracks should be beautifully preserved when the next flood washes sandy sediment over the bottom and into the old surface cracks. Deposition in ponded (standing) water on mudflats occurs during storm flooding when the saline lake temporarily expands outward over the fringing mudflat subenvironment. Under these conditions the sediment-charged sheetwash off the sandfiats should rapidly decelerate as it enters the expanded lake and quickly deposit its load. We envisage that deposition would take the form of a graded thin bed or thick lamina deposited (a) from the sheetwash that entered the shallow standing lake water as a waning turbid underflow (a 'turbidite'), or (b) from the mixed inflow-lake water body as a simple 'settle-out'. After deposition these graded layers could be reworked by wind waves to produce either coarse silt-fine sand lenses and muddy drapes or, if the layer is thick enough, simple surface rippling. Finally, as the lake waters recede by evaporation, the mudflats are exposed once more and mudcracking will disrupt the layering. The last mudflat type, the old lake-bottom, is not an active depositional mudflat but is a response to a major long-term climatic change that caused a perennial lake to dry up. Reeves ( 1 968, p. 1 20) identified many ofthe playas ofthe Great Basin of the western U.S. as being of this type. The mudflats to the east of the Bonneville Salt Flats in Utah may be an example of the exposed perennial lake-bottom type of mudflat. At present the surface is covered with polygonal mudcracks (tens of centimetres across) and a thin efflorescent halite salt crust, but the features in the sediments below suggest perennial lake deposition rather than mudflat deposition. The evidence is as follows: the sediments are finely laminated alkaline earth carbonate micrites with no internal mudcracks to signal exposure between depositional events; only the uppermost 20-30 em, which is a highly churned soil-like zone, shows evidence of recent exposure to mudcracking, salt-crust deformation, oxidation, insect burrowing and, in places, rooting by halophytes. Ephemeral saline lake subenvironment
By ephemeral saline lake we mean a shallow body of water, normally a concentrated brine, that at least once every few years dries up, leaving in the low central area an exposed layer of salt(s) that precipitated out as the brine evaporated. This subenvironment has been described as a salina, alkali lake or playa lake when wet and as a playa, dry lake, alkali fiat, salt fiat, salt pan or inland sabkha when dry (see Reeves, 1968, p. 87; Cooke & Warren, 1 973, p. 2 1 5 ; Glennie, 1 970, p. 60; Neal, 1 975, p. 1). Modern ephemeral saline lakes may cover at maxiumum stand an area as large as 8000 km2 (Lake Eyre, Australia) or as small as 0·008 km2 (Basque Lakes, British Columbia). Maximum depth of water may be several metres (e.g. 4 m in Lake Eyre during the 1 949- 1 952 flood, Johns & Ludbrook, 1 963, p. 23). The important aspect of the ephemeral saline lake subenvironment is that recharge is by surface runoff during infrequent catastrophic storms and by springs. In the long periods between fioodings the standing storm waters of the newly expanded lake will slowly shrink and become saline brines by evaporative concentration, ultimately reaching saturation with respect to salts such as halite or trona before drying up completely or almost completely. This cycle of catastrophic expansion (with freshening) and gradual contraction (with increasing salinity) of the lake (see
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L. A. Hardie,
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P. Smoot and H. P. Eugster
Langbein, 1 96 1 ) leads to a sub-division of the subenvironment into two parts based on the features of the bottom sediment: ( 1) a salt pan, underlain by layered salts (Fig. 2d), in the lowest part of the lake area; and (2) a saline mudflat, underlain by muddy clastic sediment crowded with crystals of salt minerals, surrounding the salt pan. Deposition of clastic sediment in the saline lake subenvironment appears to take place mainly as 'settle-out' of silt-clay sized grains from suspension when the turbulence of the flood waters subsides. This leaves a lamina or thin bed of mud as a storm layer over the expanded lake bottom. The thickness of this settle-out layer will depend on the concentration of suspended sediment and the depth of water, but would rarely exceed a few to tens of millimetres because mud-sized grains are not a major product of weathering in arid zones and because the floodwaters are seldom very deep. After settle-out of the storm-layer, wind induced waves may cause rippling of the surface of the layer or even reworking of the entire layer into silt-clay lenticular lamination. Blue-green algal spores washed in by the storm may bloom into a pervasive lake-bottom mat, but because the mat develops after deposition there will be no algal influence on the sedimentation. During the flood, the freshwater runoff streaming over the efflorescent crust covered mudflats surrounding the lake will preferentially dissolve the most soluble saline minerals from the crusts, producing a chemically 'simple' brine dominated by one or two major solute species, e.g. NaCl, Na2 S04 or Na2C03, a brine different from one produced by evaporative concentration of dilute inflow waters and sequential precipitation of carbonates and sulphates (as outlined by Hardie & Eugster, 1970) . This ponded brine will, over the ensuing weeks and months, slowly evaporate until halite or trona (or other soluble salt appropriate to the brine chemistry) precipitates. This stage of precipitation of salt out of the ponded surface brine may not occur until the lake area has been reduced considerably, for example, Lake Eyre expanded to 8000 km2 in 1950 but after evaporation to dryness by 1 952, a layer ofhalite a few centimetres thick had been deposited over only 1 / 1 0 of this area (from data in Bonython, 1956). A similar value for lake/salt pan ratio was found for Saline Valley, California (Hardie, 1968), which is only one hundredth the size of Lake Eyre. At Lake Magadi, on the other hand, the lake/salt pan ratio is much larger; the shorelines are steep in this graben lake and so annual flooding does not drastically expand the lake area (Eugster, 1 970). The thickness of the salt layer deposited by complete evaporation of the ponded brine will generally be in the order of centimetres, e.g. a sheet of brine l m deep and saturated with respect to halite will yield a halite crystal layer about 1 5 -20 em thick. The textural features of the salt layer reflect nucleation of seed crystals at the brine-air interface and subsequent epitaxial overgrowth on the foundered seeds. Shearman ( 1 970) and Eugster & Hardie ( 1 978) have described the major features of halite layers in salt pans while Eugster ( 1 970) has discussed trona crystallization. Interesting post depositional features include polygonal cracking (Fig. 2d) and overthrusting of the dry salt layer (see descriptions in Eugster & Hardie, 1 978), growth of salt crystals in the underlying mud layer, and bacterial reduction of S04 to produce H 2 S and iron sulphides in the mud layer making it black and anaerobic (Reeves, 1968, p. 78; Baas Reeking & Kaplan, 1 956). Overall, then, in the salt pan a single storm will result in the deposition of a couplet of a thin mud layer (millimetre scale, black, iron sulphide-rich, crowded with salt crystals) overlain by a thicker crystalline salt layer (centimetre scale) (see Bonython, 1 956, plate VIII, fig. b). Repeated storms will superimpose one couplet upon another
Saline lakes and their deposits
23
making a salt pan facies that could reach tens or even hundreds of metres in thickness. For example, more than 300 m of salt-mud interbeds underlie the Death Valley salt pan (Hunt & Mabey, 1 966, table 1 9), while more than 40 m of trona underlie Lake Magadi (Baker, 1 958). Immediately surrounding the salt pan the newly deposited mud is left exposed as the saline lake shrinks, and may become extensively mudcracked. Also, as the lake shrinks the brine will soak into the mud and persist as a subsurface brine body or join an existing perennial groundwater brine, the upper level of which will be controlled by the salt pan brine level. Evaporative pumping (Hsu & Siegenthaler, 1 969) will cause continued evaporation of the subsurface brine in the upper vadose zone and lead to widespread intrasediment precipitation of salts as displacive and/or poikilotic crystal growths. At the very surface an efflorescent salt crust will quickly form (Fig. 3c; see also Eugster & Hardie, 1 978). Where the brine body is close to the surface, massive intrasediment growth of salts will destroy any layering so that only a structureless mud full of salt crystals will result (Fig. 3d). This zone of efflorescent salt-encrusted, brine soaked, massive saline mud fringing the salt pan we have called the 'saline mudflat'.
Perennial saline lake subenvironment
By perennial saline lake we mean a surface body of brine that persists for many years (tens, hundreds or even thousands) without drying up. It may be shallow or deep (metres to hundreds of metres) but if more than several metres deep it is usually stratified (meromictic). Examples are Great Salt Lake, Utah (southern basin, about 1 2 m deep) and the Dead Sea on the Israel-Jordan border (northern basin 400 m deep). Perennial saline lakes require a substantial perennial inflow, normally a large river or rivers (e.g. Jordan River flows into the Dead Sea; the Weber, Bear and Jordan rivers empty into Great Salt Lake). The perennial stream inflow not only keeps the lake constantly supplied with water so that it does not dry up by evaporation, but also supplies both dissolved solutes and clastic sediment. Considerable amounts of dissolved solutes can be provided by perennial springs around the perimeter ofthe lake (e.g. the many fault-line brine springs around the Dead Sea; Bentor, 1 96 1 ) while aperiodic flash floods can bring in clastic sediment from all parts of the drainage basin as ephemeral sheetflood inflow. The primary process in perennial saline lakes is evaporation from the lake surface. In the arid climate necessary for a saline lake to persist, evaporation is essentially continuous and leads to ( 1 ) concentration of the surface brine, and (2) nucleation and growth of saline minerals in this surface brine. Both the newly concentrated brine and the saline minerals precipitated from it will sink down toward the bottom of the lake, and the less dense, less concentrated inflow will 'float' in over the brine to evaporate and sink down in turn. This continuous repetition of inflow � evaporation � saline mineral precipitation � sinking of brine and settling of chemical sediment, is the essential feature of stratified perennial lakes in which evaporation exceeds inflow. Which saline minerals will actually precipitate from the surface waters will depend, apart from kinetic factors, upon the evaporation/inflow ratio, the chemical composition of the inflow, and the stage in the history of the lake (see Hardie & Eugster, 1970; Eugster & Hardie, 1978). For example, for a non-alkaline brine, the saline minerals could be alkaline earth carbonates, or alkaline earth carbonates +
24
L. A. Hardie, J. P. Smoot and H. P. Eugster
gypsum, or alkaline earth carbonates + gypsum + halite. If evaporation only slightly exceeds inflow then it is probable that only alkaline earth carbonate saturation will be reached. On the other hand in order to precipitate a halite-bearing assemblage, not only must the evaporation/inflow ratio be very high but the bottom brine must be dense enough to support the concentrated surface brine at the high density a halite precipitating brine achieves, otherwise the surface brine will sink before halite saturation is reached. Hence, the mineralogy of the chemical sediments of a perennial saline lake clearly is a sensitive record of the relative evaporation/inflow ratio as well as the 'maturity' of the lake. This leads to a simple geochemical evolution model for perennial saline lakes deep enough to be stratified. Let us consider the changes that might take place in a non-alkaline perennial lake in an arid climate as inflow decreases because of changes in rainfall in the enclosing highlands. If inflow initially was greater than evaporation then the lake would be deep and quite dilute. At this stage the lake would support a diverse freshwater fauna and flora in spite of the aridity, an aridity that would be evidenced only in the fringing mudflats and isolated 'lagoons' where saline mineral precipitates would be abundant. A good example of this stage is Lake Chad in Africa (Maglione, 1 974). Now, if inflow rate was decreased so that it became less than the evaporation rate, lake expansion and deepening would stop, reverse, and evaporative concentration of the surface waters would lead to chemical precipitation of alkaline earth carbonates (aragonite, low-Mg calcite or high-Mg calcite depending on the Mg/Ca ratio and ionic strength, see Fiichtbauer & Hardie, 1 976). Initially, the evaporating surface waters, because of their increased density, would sink before saturation with respect to saline minerals could be reached. The sinking surface brine may either mix with the bottom waters increasing the overall brine density, or simply displace the bottom waters forcing them toward the surface where they could evaporate. Eventually the whole brine body would be dense enough that the residence time of the surface waters was long enough to reach supersaturation with respect to an appropriate alkaline earth carbonate mineral. * As the tiny crystals of precipitated carbonate (see, for example, Neev & Emery, 1 967, pp. 88-90 for size data on micron scale aragonite needles precipitated from the Dead Sea surface waters) slowly settled through the mixed or displaced bottom waters, most would dissolve, further increasing the salinity and density of the lake brine body. An extended period might ensue before the entire brine was saturated throughout with respect to the carbonate precipitate and net accumulation of carbonate sediment on the lake bottom would occur. Walker Lake in Nevada is a good example of this stage of development, but the carbonate precipitate is the unusual monohydrocalcite (CaC03.H2 0) (R. Spencer, personal communication, 1 976). Continued evaporative concentration of the surface waters will slowly increase the overall concentration of the brine body because the residence time of the surface waters becomes progressively longer as the brine density progressively increases. Ultimately, in a manner analogous to the carbonate saturation stage, the lake brine becomes saturated with respect to gypsum and a stage of gypsum + carbonate co precipitation and accumulation sets in. The Dead Sea is at this stage today (Neev & Emery, 1 967), although it appears to have reached this point from the opposite *S upersaturation of the surface waters may be reached before sinking of the brine because the surface waters are likely to be warmer and hence fractionally lighter than the undersaturated but cooler lower waters.
Saline lakes and their deposits
25
direction, that is, increased inflow rate since early Holocene times has replaced a halite stage with a gypsum + aragonite stage. As long as the inflow rate of our hypothetical perennial saline lake remains much smaller than the evaporation rate, the lake will eventually reach saturation with respect to halite in the same manner as described for the carbonate and gypsum accumulation stages. The solubility of halite, however, is an order of magnitude greater than that of gypsum (in relative mass terms) and so a long period of time will be required to reach the halite stage. And, of course, the level of the lake will have fallen considerably, perhaps it would have no more than 1 / 1 00 the volume of the lake at carbonate saturation stage. If the lake basin sides are not too steep, this drop in lake level could expose a large part of the lake-bottom underlain by carbonate or carbonate + gypsum. Great Salt Lake, Utah, appears to be at this stage today, with wide fringing mudflats that represent old perennial lake-bottom. At this stage, upon further evaporation, the fate of the perennial lake will follow one of two possible paths depending upon the depth of the brine saturated with respect to halite (or some other salt, say trona) and the inflow-evaporation balance. If the brine body was shallow, say no more than a few or a few tens of metres deep, the lake may dry up completely and deposit an appropriate layer of salts. This happened at Owens Lake, California, in response to inflow cutoff by man and it nearly happened at Great Salt Lake during the drought of 1 930- 1 935 when a thick halite layer was deposited over the lake bottom. If the drought had continued, Great Salt Lake would quickly have joined the many examples of ephemeral salt lakes. Accumulation of salt could have continued as long as springs brought in the necessary solutes, as they do now at Lake Magadi. Instead, inflow increased again and Great Salt Lake remained a perennial lake, but now considerably diluted. In fact, much of the solute load of the dense bottom brine may have been derived by dissolution of the bottom salt crust. The transition from perennial lake to ephemeral salt lake is documented many times in the geologic record, such as in the transition from the High Magadi beds to present Lake Magadi (Baker, 1958; Eugster, 1 970), the transition from Bottom Mud to Lower Salt and from Parting Mud to Upper Salt at Searles Lake (Smith, 197 8), from oil shales to trona beds in the Wilkins Peak Member of the Green River Formation (Bradley & Eugster, 1 969) and the transition back is equally common. In contrast, if the saline brine body is very deep, say hundreds ofmetres, the lake will not readily turn into an ephemeral salt lake. Instead, a massive evaporite deposit must eventually accumulate from such a perennial salt lake. This is what would happen to the Dead Sea, if the Jordan River were cut off. We do not know of a contemporary example where this is taking place and we therefore can only suggest a reasonable model. This model of halite or trona precipitation from a deep lake is simply an extension of the model of carbonate and gypsum precipitation from a perennial lake discussed earlier. Evaporation takes place at the surface, producing a concentrated surface brine, that when its density becomes great enough, will flow to the bottom of the lake, displacing lighter bottom water. This process continues until the whole brine body is saturated with, say, halite, when massive halite precipitation would set in. The lake could remain such a perennial halite-depositing lake for thousands of years or more, provided the proper and delicate evaporation/inflow balance is maintained. The model we have outlined here is similar in principle to that suggested by Schmalz ( 1 969) for deep water marine evaporites. It is not restricted to monomineralic deposition and applies equally well to complex brines.
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L. A . Hardie, J. P. Smoot and H. P. Eugster
The vertical sequence of chemical sediments that would record the stages in the evolution of the hypothetical non-alkaline lake outlined above would simply be a halite + minor gypsum + trace of carbonate package overlying a gypsum + minor carbonate package which in turn overlies a carbonate package. The thickness of each package will depend very much on the chemical composition of the inflow, the evaporation/inflow ratio, the original lake volume, and the length of time involved, but probably would be in the order of metres, tens of metres or even hundreds of metres. Of course if a climatic reversal occurs and inflow increases at any point, then the sequence would record it faithfully by a reversal of stage, as for example in the Dead Sead where cores through the bottom show a thin late Holocene carbonate package (80 em) overlying an early Holocene-late Pleistocene halite package (Neev & Emery, 1 967). The sedimentary structures typical of chemical sediments of perennial saline lakes deep enough to be stratified are relatively uncomplicated. The major feature is areally widespread (laterally continuous) saline mineral laminae and thin beds with clastic sediment (mud) partings. Many authors identify each saline mineral layer-clastic layer 'couplet' as a 'varve' (e.g. Reeves, 1 968, fig. 50). In both Searles Lake (Smith, 1 978) and the Dead Sea (Neev & Emery, 1 967, pp. 84-85), however, the carbonate laminae are not annual but are rather irregular, only one lamina deposited per several years (the Dead Sea, 1 5-20 laminae in 70 years). The varve interpretation arises because saline lake deposits have been compared to temperate zone non-saline lake deposits (e.g. Bradley, 1929). In saline lakes, however, evaporation is normally continuous so that chemical precipitation is continuous throughout each year, as Neev & Emery ( 1 967, pp. 82-85, p. 93) showed for co-precipitation of aragonite and gypsum in the Dead Sea. Perennial saline lake layering, then, is not due to seasonal pulses of chemical precipitation but rather seems to be due to storm-flood influxes of clastic sediment (mainly mud) that punctuate the continuous rain of chemical sediment (compare with Calvert's, 1 966, explanation for the clastic mud-diatom ooze couplet lamination in the sediment of the Gulf of California). The clastic influxes could be seasonal but are more likely to be quite irregularly spaced, as for example is typical of the Great Basin of the western U.S. (e.g. Hardie, 1 968) or the Lake Eyre, Australia, basin (e.g. Bonython, 1956). The typical thicknesses of the chemical layers (i.e. thickness between clastic partings) depend upon the solubility of the saline mineral, the net annual evaporation (typically 1 -2 m) and the average interval between storm floodings (perhaps 1-5 years). Roughly, we would expect carbonate laminae to be fractions of a millimetre, gypsum laminae to be several millimetres, and halite to occur in thin beds of several centimetres. The most probable mechanism of deposition is simple 'settle-out' to produce uniformly thick layering. The clastic sediment is also likely to be deposited by 'settle-out' as fine mud laminae, but an occasional massive storm over the whole basin could introduce unchannelled sheet-floods (carrying coarse sediment from the surrounding mudflats) as a dense underflow from which a graded clastic layer could be deposited. Other maj or sedimentary features besides the layering are diagenetic in origin. The fine-grained chemical sediments soaking in the bottom brine might recrystallize to a more stable coarse crystalline mosaic, destroying in the process the 'settle-out' fabric. Where organic matter has accumulated in the bottom sediment, the interstitial brines may undergo bacterial reduction of sulphate ions (Bass-Becking & Kaplan, 1 956). If gypsum, or other sulphate mineral, is present this bacterial reduction can result in
Saline lakes and their deposits
27
dissolution of the mineral, such as Neev & Emery ( 1 967, p. 94) believe is responsible for the loss of gypsum in the bottom sediments of the Dead Sea. The sulphide-rich anaerobic conditions produced in the sediment by the bacterial reduction of sulphate will also allow ferric oxides and oxyhydroxides brought in with or on the clastic grains to dissolve, raising the ferrous ion concentration. This can lead to precipitation of black iron sulphides like mackinawite and greigite which ultimately will convert to pyrite or marcasite (Berner, 1 970) and/or precipitation of iron phosphate (Bray, B ricker & Troup, 1 973). At the same time bubbles of H 2 S or CH4 gas generated by the bacterial reduction process may disrupt or deform the layering in the sediment, leaving a fenestral fabric (cf. Forstner, Muller & Reineck, 1 968). Very shallow perennial saline lakes that are always thoroughly mixed by wind action would probably give a sedimentary record like ephemeral saline lakes and perhaps should be included with such ephemeral lakes.
Windblown dune field subenvironment
The sandy alluvial fan toes, the sandflats and the inflow stream floodplains in saline lake basins are dry and exposed most of the time and without a pervasive vegetation cover. So, whenever protecting efflorescent salt crusts are absent, the surface sediments of these subenvironments are highly vulnerable to wind erosion and deposition (see, for example, Glennie, 1 970, pp. 45-47). Dune fields and deflation flats and hollows are the major products of this wind action. In deep, narrow elongate closed basins, such as the graben valleys ofthe Great Basin of the western U.S., small sand dune fields (a few square kilometres in area, with dunes perhaps no more than 1 0 m high) become established (essentially non-migratory) at the upsloping ends of the basin, fringing the playa flats. This takes place because the winds are forced by the topography to blow up or down the long axis of the valley. In Death Valley (Hunt & Mabey, 1 966, plate 1 ) and Saline Valley (Lombardi, 1 963, plates 1 and 2) in California, the dune fields cover less than 5% and 1 0%, respectively, of the depositional basin area, and rest on sandflat and fan toe substrates. In wide shallow closed basins like those of Lake Eyre, Australia, or the Etosha Pan, Southwest Africa, the dune fields cover most of the basin interior. These massive dunefields mainly take the form of extensive longitudinal dune ridges with bare 'claypans', deflation pavements, caliche- and silcrete-covered pavements, and isolated small ephemeral saline lakes in the interdunal troughs (see King, 1 956, fig. 1 and plates 3, 4, 5; Twidale, 1 972, fig. 1 and pp. 6 1-65; Goudie, 1 973, p. 1 1 7; Glennie, 1 970, enclosure 1 ). Dune sand is not always derived by wind erosion of alluvial sand fringing a playa or salt lake, but may be derived by erosion of the playa or dried saline lake surface. In these cases the dunes may be composed of saline minerals like gypsum originally precipitated within the playa or lake mud (McKee, 1966; Glennie, 1 970, p. 1 36) or perhaps even mud intraclasts (Huffman & Price, 1 949). The diagnostic sedimentary features of aeolian sediments have been well described by many authors, particularly McKee & Tibbitts ( 1 964), McKee ( 1 966), Sharp ( 1 966), Glennie ( 1970), Bigarella ( 1 972) and Hunter ( 1 977) and we refer the reader to these works.
28
L. A. Hardie, J. P. Smoot and H. P. Eugster
Perennial stream floodplain subenvironment
Inflow to perennial saline lakes is invariably dominated by one or more perennial streams. For example, almost 80% of the water reaching the Dead Sea annually is brought in by the Jordan River which rises from spring-fed Lake Tiberias and flows 105 km down the narrow Jordan Valley before reaching the Dead Sea (see Neev & Emery, 1 967, pp. 72-73). To maintain Great Salt Lake, Utah, 82% of the present annual inflow is supplied by the perennial Weber, Jordan and Bear Rivers (Hahl & Langford, 1 964) that rise in the towering Wasatch Range which makes the eastern buttress of the basin. The depositional elements of perennial stream floodplains, both meander and braid streams, are among the best known of all modern sedimentary environments, and so will not be repeated here. We refer the reader to works such as those of Doeglas ( 1 962), Harms, Mackenzie & McCubbin ( 1 963), Allen ( 1 965), Harms & Fahnestock ( 1 965), Coleman ( 1 969), Williams & Rust ( 1 969), Allen ( 1 970, pp. 1 1 8- 1 48), McGowan & Garner ( 1 970) and Reineck & Singh ( 1 975, pp. 225-263), where point-bar, scroll-bar, levee, crevasse-splay, flood basin, ox-bow lake, braid-bar, braid-channel, etc. deposits are described and diagnostic vertical sequences outlined. We would add only that in saline lake basins, perennial stream floodplains would be subject to arid conditions. So the normally subaerially exposed subenvironments like levees and upper point-bars may have significant caliche and/or silcrete deposits and intrasediment growth of saline minerals such as gypsum, while the floodbasins and ox-bow lakes may become ephemeral saline lakes and/or desiccated mudflats.
Ephemeral stream floodplain subenvironment
Ephemeral streams are sporadically-flooding river systems that are usually dry. Such systems are distinguished from alluvial fan-sandflat complexes by their longitudinal extent away from the source highlands. They are a major subenviron ment component of wide shallow closed basins like Lake Eyre, Australia (Bonython & Mason, 1 953) and the Etosha Pan, South-west Africa (Gevers, 1 930), or of sandy deserts like those of Libya or the Arabian Peninsula where the ephemeral streams are known as 'wadis' (Glennie, 1 970, pp. 29-56, p. 198). Ephemeral stream floodplains are simply an expanse of braid channels and braid bars (Bonython & Mason, 1 953, plate 1 ; Twidale, 1 972, fig. 17; Glennie, 1 970, pp. 30-32; Williams, 1 97 1). Grain size of the sediment decreases away from the highlands (Williams, 197 1 , p. 2; Picard & High, 1 973) so that gravel bars and gritty channels may predominate near the source areas and fine downstream to sandy material. The floodplains are commonly partly covered with windblown dunes that have derived their sand from the dry stream beds. The surface sediment may be coated with caliche crusts, silcretes or even gypsum (Bonython & Mason, 1 953; Bonython, 1 956; Gevers, 1 930; Goudie, 1 973), or may be cemented by alkaline earth carbonate vadose cements, particularly the braid channel bottoms (Glennie, 1 970, pp. 33-36). The diagnostic sedimentary structures of ephemeral stream floodplain deposits have been described by Williams ( 1 97 1 ), Karcz ( 1 972) and Frostick & Reid ( 1 977), and have been discussed and illustrated at some length by Glennie ( 1 970, pp. 1 1 - 1 4, pp. 29-56) and Picard & High ( 1 973). Because ephemeral streams are such important features of saline lake basins, both in floodplains and in alluvial fans, we will briefly summarize
Saline lakes and their deposits
29
their sedimentary record. In typical braid stream fashion ephemeral stream deposits consist of cross-cutting lenses of coarse braid bar sediments and finer channel sediments (Doeglas, 1 962; Williams & Rust, 1 969) that show inclined planar lamination and thin bedding, antidune wavy bedding, megaripple cross bedding, horizontal planar lamination, ripple cross lamination, mud-cracked clay drapes, mud chip pockets, and steep-edged scour-and-fill. These lenses occur as fining-upward sequences that reflect a waning flow-regime. Near the source area highlands the deposits are mainly channel-fill conglomerates that show inclined bedding and grits with horizontal to low inclination lamination. Trough cross-bedded and planar to wavy laminated sands and grits and ripple cross-laminated sands dominate the middle of the floodplain. At the distal ends near the saline lake finer-grained sand with horizontal planar to wavy lamination and ripple cross lamination is the norm. The crucial evidence for an ephemeral, as opposed to a perennial, braid stream is the presence of ( 1 ) mud-cracked mud drapes (Glennie, 1 970, pp. 49-55) at the top of thin (tens of centimetres) waning flow sequences that are dominated by relatively fine grained upper flow regime structures, and (2) mud-chips (peloids) scattered through the sands or as grit, breccia or conglomerate pockets and lenses. Glennie ( 1 970, p. 1 2) stressed the presence of rippled and horizontally-laminated aeolian sand beds separating the water-laid sequences. Picard & High ( 1 973) suggested terraced channel sides formed by receding flood waters are characteristic of ephemeral streams. Finally it is the presence of caliche, silcrete and gypsum crusts, alkaline earth carbonate vadose cements and saline mineral intrasediment growths (e.g. gypsum crystals, anhydrite nodules) that record the essentially arid climate ofthese ephemeral stream deposits in saline lake basins.
Springs and spring-fed ponds
Springs, although small features, are important in saline lake basins because they not only are perennial suppliers of water and particularly of dissolved solutes to many lakes (see, for example, Bentor, 196 1 ; Jones, 1 965; Hardie, 1968; Eugster, 1 970) but they are settings for primary production of chemical and biochemical sediment such as travertine and tufa. Springs, which are surface outlets of groundwater, usually discharge in one of two situations: ( 1 ) along faults, commonly at the apices of alluvial fans, at saline lake margins and even beneath lakes (Scholl, 1 960; Jones, 1965; Hunt et a!., 1 966, p. 29; Hardie, 1 968; Neev & Emery, 1967); and (2) at the intersection of a porous sediment (aquifer) and impermeable layer (aquiclude) such as where alluvial fan debris has built over mud flat or lake bottom sediments (Jones, 1965, p. 1 5 ; Hunt et a!., 1 966, p. 29; Hahl, 1968). There are many other spring settings, of course, but these two modes are apparently the most common in saline lake basins. Once the springs discharge, the outflow may quickly seep into a porous substrate after flowing a few metres (a situation commonly found in alluvial fans, e.g. Jones, 1 965, p. 14; Hardie, 1 968), they may feed perennial streams (Slack, 1 967; Dunham, 1972, p. 1-5 8), they may form ponds, marshes or 'lagoons' (Baker, 1 958; Jones, 1 965, p. 1 5 ; Hunt et a!., 1966, pp. 32-36, fig. I I ; Eugster & Jones, 1 968), or they may mix directly with saline lake waters as they exude into them (Scholl, 1960; Scholl & Taft, 1 964). The compositions of spring waters vary considerably reflecting the history of the
30
L. A. Hardie, J. P. Smoot and H. P. Eugster
groundwater. Some examples of this variation are: ( l ) springs are dilute where groundwater is near its source such as the apices of alluvial fans or the proximal portions of ephemeral streams (Jones, 1 965, table 7; Hardie, 1 968, table 2). These springs can reach supersaturation with alkaline earth carbonates, particularly low-Mg calcite, by 'degassing' col from the spring water on encountering the atmosphere, or by evaporative concentration of the waters as they flow on the surface; (2) groundwaters may be concentrated by evaporation in the vadose zone of porous sediments (such as the alluvial fan) resulting in spring waters which are supersaturated with respect to high-Mg calcite, protodolomite or even gypsum (Hunt et al. , 1 966, figs 39, 44, 45 and 47); (3) solutes may be added to ground waters by seepage of storm runoff that has dissolved surface saline crusts. This groundwater can emerge near the lake as brine springs, such as described by Jones ( 1 965, table 7, p. 30), Hunt et al. ( 1 966, tables 50 and 52), Hardie ( 1 968, table 2) and Eugster ( 1 970); (4) springs may be fed from deep circulating groundwater that dissolves old evaporites (see Bentor, 196 1 on the Dead Sea brine springs); (5) hot groundwaters may react with bedrock to produce brines with unusual compositions (see Muffler & White, 1 969 on the Salton Sea, California). The most obvious and probably most diagnostic depositional products of springs are travertines and tufas. These structures are composed of alkaline earth carbonates, usually low-Mg calcite (Slack, 1967; Irion & Muller, 1 968) and occasionally high-Mg calcite (Barnes & O'Neil, 197 1 ) . Travertine and tufa form low mounds, sheets, coated grains, and pore-filling cements at spring orifices, along outflow channels and margins of ponds and marshes (see for instance Hunt et al., 1 966, p. 1 6; Slack, 1 967; Dunham, 1972, p. I-58 and fig. I-5 8) and apparently form 'pinnacles ' where springwaters exude into lakes (Scholl, 1 960; Scholl & Taft, 1964). Travertines are laminated, commonly with alternating layers of colour-banded micrite and fenestral, palisade-structured micrite (Smoot & Hardie, in preparation; Irion & MUller, 1 968, figs 3, 4 and 5; Dunham, 1 972, figs I-59-I-63; Monty, 1976, fig. 7). Tufas are generally less well laminated, showing predominately a fenestral fabric, like the porous layers of travertines (Scholl, 1960; Scholl & Taft, 1964; Golubic, 1969; Monty, 1 976, figs 28 and 29) . The laminar layers probably form by precipitation of carbonate minerals from films of water while the porous layers and tufas are produced by precipitation around plants such as algae or moss, so that the structures produced are 'chemical stromatolites'. The role of the plants is probably as a static substrate, but they may aid precipitation by providing a preferred nucleation surface or increasing supersatura tion by photosynthetic uptake of C01 (see Weed, 1 889, or discussion in Hardie, 1 977, pp. 1 70- 175). The porous tufas and fenestral travertines are very friable and can break down into carbonate sediments (peloids from granule to mud size) which Smoot ( 1976, 1978) emphasizes as a major source of non-skeletal carbonate sediment production in arid basins. McGannon ( 1 975) presented a similar model in describing a Pleistocene, cross-bedded, fining-upward fluviatile deposit composed of tufa fragments, ooids and pisoids that could be traced back to a spring source. Other chemical deposits reportedly made by springs include: ( l ) 'chemical deltas' which result from springs mixing with saline lake water. Deposits that have been interpreted in this way include the oolitic sands of northern Pyramid Lake, Nevada (Surdam & Wolfbauer, 1 975, p. 343) and the protodolomite muds(?) ofMound Playa, Texas (Reeves, 1 968, pp. 64-65, figs 5 1 , 53 and 54); (2) Hunt et al., ( 1 966, pp. 56-59) describe spring-fed marshes floored with precipitated masses of crystalline gypsum as well as crusts containing glauberite, thenardite, halite and trona; (3) hot springs
Saline lakes and their deposits
31
surfacing along faults in Lake Magadi are precipitating alumino-silicate gels (Eugster
& Jones, 1 968); and (4) silica sinters are forming around hot springs in Yellowstone National Park, Wyoming (Weed, 1 889; Walter, 1 976; Walter, Bauld & Brock, 1 976)
and similar sinters have been reported in Pleistocene Lake Lahontan marginal sediments (Morrison, 1 964, fig. 1 2). Finally, spring ponds and marshes may support, in addition to a rich flora (algae, fungi, saltgrasses, rushes, pickleweed, cattail, mesquite, tamarask, palm, etc., see Hunt, 1 966), brineshrimp (Jones, 1 965, p. 17) and desert fish (Hunt et a!., 1966, p. 35; Beadle, 1 974). The remains of the vegetation can accumulate at the bottom of these spring ponds, marshes and 'lagoons' to produce organic-rich layers (Jones, 1965, p. 1 5) which could become buried by storm runoff sediment layers or chemically precipitated saline mineral layers. Shoreline features
Shoreline features such as deltas, beaches, beach ridges, spits, bars and platform and mound build-ups have not been widely reported from saline lakes. This may well be due to the fact that most of these features are products of strong currents and/or energetic wave action found only in large relatively deep perennial lakes, a setting that is uncommon in arid closed basins. During the flooding of Lake Eyre, Australia, in 1 949-1 950 the dune sands and ephemeral stream sediments were worked into spits, longshore bars, beaches and beach ridges along the temporary lake shore (see Bonython & Mason, 1 953, plate 3 and King, 1 956, fig. 4). Presumably these features would look like their non-saline temperate perennial lake (Gilbert, 1890, pp. 23-89) and marine counterparts (see Reineck & Singh, 1 973, pp. 280-349). Deltas in freshwater perennial lakes have been examined (e.g. Houbolt & Jonker, 1 968; Forstner et a!., 1 968) but to our knowledge similar detailed studies on saline lake deltas have not been published. Two unusual kinds of 'delta' in saline basins have been reported: ( 1) Krinsley ( 1970, see reprint in Neal, 1 975) has mentioned a 'fan delta' from the Sabzevar Basin, Iran, which seems to be a lobate deposit (characteristics not described) that shed directly from an alluvial fan into a temporary lake; and (2) a 'chemical delta' appears to be forming at the mouth of the narrow channel between the Gulf of Karaboghaz and the Caspian Sea (Teodorovich, 1 96 1 ; Dickey, 1 968) due to precipitation of alkaline earth carbonates as the inflowing Caspian Sea water concentrates by evaporation on entering Karaboghaz. Broad submerged shoreline platforms and isolated mounds built up of marly sediment have been reported from freshwater 'marl lakes' (Hooper, 1956; Wilson, 1936, 1 938; Moxham & Eckhart, 1956; Wetzel, 1 970) but not as far as we know, from arid climate saline lakes. This may simply be an oversight. Subaqueous algal mounds and ooid sand shoals have been described from Great Salt Lake (Eardley, 1 938; Carozzi, 1 962) but it is not clear whether these features are being actively produced today or under what conditions they formed (see Halley, 1976). S O M E E X AM P L E S O F S AL I N E LA K E D E P O S I T I O N AL COMPLEXES
No two saline lake basins are exactly the same (cf. Twidale, 1972, p. 2 1 2) because mode of origin, tectonic setting, bedrock types and patterns, regional and local climate,
32
L. A . Hardie,
J.
P. Smoot and H. P. Eugster
v v v vC' v v v v v v v 0.>-.>- v v v v v v
� be d r o c k ( pre - Quater nary l � a l l u v i a l fan li"':'"{i\;'d s a n d f l a t
� m u d f lat �jt,§ spr i n11 pond - marsh lv v vi ephemeral sa l i ne lake
4 . Surface subenvironment map o f the northern part o f Death Valley, California. Alluvial fan and ephemeral saline lake subenvironments dominate Death Valley. T, Tertiary volcanic and sedimentary bedrock; P, Palaeozoic sedimentary bedrock; pC, Precambrian igneous and metamorphic bedrock.
Fig.
and past history vary from one basin to the next. So which depositional subenvironments are, or have been, present and how they are arranged in space and time will be different from one basin to the next. Certain associations of subenvironments of modern saline lake basins are, however, recurrent and allow us to present a few 'typical' examples of saline lake depositional complexes. Our presentation is tentative only, because our approach has not yet been used in the field as a specific mapping procedure. (1) Alluvial fan-ephemeral saline lake complex.
Deep, narrow rift and block-fault graben-valleys enclosed by towering mountains
Saline lakes and their deposits
33
(a )
� � w / l � -
S A L I N E V A L L E Y, CALI F O R N I A
... � � c:::J ..
b e d rock a l l uv i a l f a n sa n d f l a t
v
dune field
s p r i n g p ond- m a r s h so l i n e m u d f l a t sa l t p a n s p r i n g t r a ve r t i n e
( b) 1000
0 o A
South
Km
SALINE VALLEY, CA LIFO R N I A
A' North
Fig. 5 . ( a ) Surface subenvironment map o f Saline Valley, California. (b) Cross-section along a north-south
line in Saline Valley, California (A'-A in Fig. Sa).
34
L. A . Hardie,
J. P.
Smoot and H. P. Eugster
1 18°00'
+ + + + + + + +
D E E P S P R I N G S VALLEY, CAL I FOR N IA
� b e d rock � a l l u v i a l fan r:'7:'7:1 s a n d f l a t w i t h LJJJJ sma l l d u n es
� sa l i ne m u d f l a t !IIIJ so I t p a n li'ill s p r i n g - marsh
Fig. 6. Surface subenvironment map of Deep Springs Valley, California. Alluvial fan subenvironment
dominates here, extending 15 km north of the map area.
are typically dominated by huge alluvial fans that slope down to touch a central, narrow, salt-encrusted ephemeral saline lake that may be ringed by a narrow dry mudflat (see Eugster & Hardie, 1978, fig. I). A sandflat strip normally connects the fans with the mudflat-saline lake zone. Small perennial streams rise in the rocky canyons of the mountains but quickly sink into the porous alluvial fan gravels so that inflow to the lake is entirely by perennial and ephemeral groundwater and ephemeral storm runoff. There are many examples of this basic theme to be found in the Great Basin of the western United States (see Eugster & Hardie, 1978, fig. 3); we will report briefly on three such basins in California, Death Valley, Saline Valley and Deep Springs Valley, to show the essential elements of, as well as the local variations in, this kind of environmental complex. All three basins lie between 36. and 3r N latitude, are narrow block-fault valleys with a minimum relief of 1 000-2000 m and represent mid latitude orographic desert basins.
Saline lakes and their deposits 1 30"
1 34°
138°
35
142°
22"
26°
30°
- - - - drai naQe divide
)!//
d u ne f i e l d
�
f l ood p l a i n
'j l/
� e phemeral stream ......_ eph emera l
...... sa l i ne l a ke
{
j;:::_:·: .<·:l sa l i ne mu d f l a t
� � so 1 1 pan
Fig. 7. Very generalized surface subenvironment map of the Lake Eyre basin, South Australia. Inset shows
the ephemeral saline lake area.
A simplified areal distribution of surface subenvironments for a portion of the Death Valley basin is shown in Fig. 4, which is based on the excellent geological maps of Hunt et a/. ( 1 966, plate 1 ) and Hunt & Mabey ( 1 966, plate 1 ) and on our own observations. This should be compared with Fig. 5, which is a subenvironment map and cross-section of Saline Valley, California (Lombardi, 1 963; Hardie, 1 968; Eugster & Hardie, 1 978). Saline Valley is similar in most respects to Death Valley but lacks a dry mudflat subenvironment, mainly because the perennial groundwater inflow is very concentrated by the time it reaches the mudflat, where massive intrasediment crystallization of gypsum, glauberite and halite has destroyed all layering. Deep Springs Valley (Jones, 1 965; Eugster & Hardie, 1 978) has an assemblage of surface subenvironments (Fig. 6) like that of Saline Valley but the brine chemistry is different so that alkaline earth carbonate muds, including protodolomite (Peterson et a/., 1 963, 1 966; Clayton, Jones & Berner, 1 968), have precipitated out in the saline mudflat area, and sodium carbonate salts dominate in the salt pan.
(2) Ephemeral stream floodplain-dune field-ephemeral saline lake complex
In wide, shallow basins in arid sub-tropical high pressure belts ephemeral stream floodplains with widespread windblown dune fields are the dominant depositional subenvironments. In such basins catastrophic storms spaced tens of years apart wash
36
L. A. Hardie, J. P. Smoot and H. P. Eugster 1 1 2" 00'
1 13"00'
41 . oo '
r+t b e d r o ck lL_:t] ( m o u n t a i ns ) � alluvi a l fan L:2i] sand flat complex
�
r--1 pe renn i a l saline L.__j lake ( 1969 shorel ine ) � perennial stream floodplain � d e l t a com p l e x
m u d f l a t ( m a in l y o l d l a ke - bot t o m )
Fig. 8. Generalized surface subenvironment map of the Great Salt Lake basin, Utah.
masses of sediment for many kilometres over the normally dry braided floodplain, and may leave a huge area inundated with ponded floodwaters. This shallow ephemeral lake will slowly evaporate over the ensuing months (or even years), precipitating a crystalline layer of saline minerals before drying up completely to leave a salt pan surfaced by a hard salt crust and fringed by a saline mudflat. The type examples of this kind of saline lake basin are the Lake Eyre basin in South Australia (Bonython & Mason, 1 953; Bonython, 1 956; King, 1 956; Twidale, 1 972) and the Etosha Pan in Southwest Africa (Gevers, 1 930; Goudie 1 973, p. 1 27). Figure 7 shows the distribution of subenvironments in the Lake Eyre basin, as far as we can discern them from the published maps and descriptions.
(3) Perennial stream floodplain-perennial lake complex
Two well known examples of very concentrated perennial saline lakes with perennial streams flowing directly into the lake are the Dead Sea on the lsrael-Jordan border (Neev & Emery, 1 967) and Great Salt Lake, Utah (Eardley, 1 938). They are
Saline lakes and their deposits
37
quite different and so they will serve to show some of the many possible variations on the perennial saline lake complex theme. The Dead Sea occupies the low point in the Jordan Valley rift and is a deep, narrow stratified saline lake (North basin, 400 m deep) with alluvial fans reaching right to the water's edge (see Neev & Emery, 1 967, figs 3 and 4A). Beach ridges, built on drowned fan toes during a higher lake stage in the late Pleistocene (Lisan stage), are preserved in places above the present lake shore (Neev & Emery, 1 967, fig. 4). The southern end of the lake is a shallow arm (South Basin, < 1 0 m deep), partly dammed by man-made structures. The Jordan River floodplain which stretches northward of the lake along the rift for over 1 00 km, makes the northern shore of the Dead Sea, so that the entire lake-floodplain complex is a long, narrow, deep trough filling. Great Salt Lake, by contrast with the Dead Sea, is a shallow (max. depth 1 2 · 5 m), wide stratified lake surrounded by dry mudflats (old lake bottom type) which pass into a narrow alluvial fan wedge at the basin edge. Small floodplains, grading into even smaller deltas, make up the eastern shoreline of the lake where the perennial Bear, Weber and Jordan Rivers debouch from the massive Wasatch Mountains. Figure 8 summarizes the basic subenvironment distribution. Both the Dead Sea and Great Salt Lake at present are perennial lakes because the inflow streams are perennial and of relatively large discharge. The sources of the perennial stream waters, however, are quite different in the two cases. Great Salt Lake is a mid-latitude (40-42° N) orographic desert basin with a plentiful precipitation inflow source in the Wasatch Mountain catchment ( 1 5 0 em precipitation annually). The Dead Sea, on the other hand, is a subtropical high pressure 'horse latitude' (3 1 -33° N) desert basin with almost all the inflow water ultimately supplied by deep circulating groundwaters of uncertain origin and age (Bentor, 1 96 1 ; Jones et al. , 1 976) that rise along the rift valley faults.
A P P L I C ATION O F THE A P P R O A C H TO MODERN AND ANCI ENT SALINE LAKE B A S I N S
Mapping depositional subenvironment surfaces of modern active saline lake basins, using our approach as outlined above, should provide an organized format for attack on the problem of identifying and understanding the processes that operate in such settings and the record these processes leave in the sediment. When the approach is extended to three dimensions using outcrops, borehole data, etc., to build a facies block diagram, then we have a powerful method of attack on the problem of the vital history of the basin, with its implications for climatic and tectonic changes. Finally, such studies carried out on existing saline lake basins would supply us with a set of subfacies and facies models for identifying and interpreting ancient saline lake basin deposits.
ACKNOWLEDGMENT
The work was supported in part by NSF grant G A 3 1 076.
38
L . A . Hardie, J. P. Smoot and H. P. Eugster REFERENCES
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Saline lakes and their deposits
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40
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Saline lakes and their deposits
41
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Spec. Pubis int. Ass. Sediment. ( 1 978)
2,
43-55
Late Pleistocene-Holocene evolution of the Kivu-Tanganyika Basin
P E T E R S T O F F E R S and R . E . H E C K Y Jnstitut fur Sedimentforschung, Postfach 1 03020, 6900 Heidelberg 1, Germany, and Freshwater Institute, 501 University Crescent, Winnipeg, Manitoba R3T 2N6, Canada
ABSTRACT The available stratigraphic records of Lakes Kivu and Tanganyika are rich in microfossils, especially diatoms, and distinctive mineral assemblages which can be interpreted in terms of the salinities, redox conditions and trophic status of the lakes. Prior to 1 3,000 years B . P. both lakes were closed basins because of the much drier climate at that time. With the onset of more moist conditions in the Early Holocene, Lake Kivu overflowed into Tanganyika via the Ruzizi River. Since then, both lakes have responded to the changing climate and volcanic events in the Kivu Basin. The crenogenic meromixis which is the unique limnological feature of modern Lake Kivu was set up at least twice in the past 5000 years and was obliterated between approximately 4000 years B.P. and 2000 years B.P. Volcanic events in the Kivu Basin were almost certainly involved in the initiation of meromixis at 5000 years B.P. and were likely involved at 2000 years B.P. A cooler and drier climate between 4000 and 2000 years likely caused the breakdown of the meromictic condition. These events in the Kivu basin are responsible for patterns of carbonate deposition observed in Lake Tanganyika. The climatic sequence of dry and pluvial periods inferred for the Kivu-Tanganyika basin is consistent with palaeoclimatic reconstructions from other places in East Africa for the last 14,000 years.
INTRODUCTION
Lakes Kivu and Tanganyika are situated in the great western rift zone of East Africa. They have attracted limnologists and zoologists since their discovery in the mid-nineteenth century. This is due to their tropical location, their unusual hydrography and unique ecology (Damas, 1 937; Capart, 1 95 2; Kufferath, 1 952; Verbeke, 1 957; Coulter, 1 963). Little is known, however, about the sedimentary history. Recently, the discovery of bedded metalliferous deposits that have formed or are forming in the vicinity of active rifts as in the Red Sea (Degens & Ross, 1 969) has focused the attention of geologists on the sediments of the large East African rift lakes (Degens & Kulbicki, 1 973; MUller & Ffustner, 1 973; Degens & Ross, 1 976). Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
Peter Stoffers and R. E. Hecky
44
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Three expeditions conducted by scientists from Woods Hole Oceanographic Institution, Woods Hole, Mass. during 1 970- 1 972 to Lakes Kivu and Tanganyika have recovered a number of cores revealing a wealth of information (Degens, Von Herzen & Wong, 197 1 ; Degens et a!., 1 973; Reeky & Degens, 1 973; Stoffers &
Kivu- Tanganyika Basin
45
Fishbeck, 1 974; Stoffers, 1 975). The core locations are shown in Fig. 1 . In this paper, we discuss our data on mineralogy and diatoms obtained from these cores in view of the interaction of climate, volcanism and sedimentation for the Late Pleistocene Holocene time. To provide some background information on the two lakes. a briefsynopsis of their limnology and water chemistry is given in Tables 1 and 2. The distribution of major cations versus depth in Lake Kivu water and a temperature profile from the deep northern basin are given in Fig. 2a and b. Below 70 m of water the temperature and salinity increase with depth in Lake Kivu.
Table I . Selected morphometric and hydrological parameters of Lakes KiYu and Tanganyika. values
generated from data of Capart ( ! 952), Verbeke ( ! 95 7 ). De gens ( 1 973)
Area (km')
A" AL
=
=
(m)
a/. ( 1 97 1 . 1 973 ) and Hecky & De gens
KiYU
Tanganyika
Catchment
7. 1 40
Lake
2.060
23 1 .000 (exclusive of Kivu) 32.600
240 485 5 83 0·33-0-4 3·2 * 1 10 190
Z Mean depth zm, maximum depth Volume (km3) v Volume storage-area relationship t>. VI t>.A L Outflow (km3 /year) Residence time water (years) Refill time (runoff alone) (years)
Depth
et
=
=
570 1 .470 1 8.880 0·66 2·7 430 1 .000
*Ten year average discharge supplied by Amenagement Hydro-Electrique des Chutes de Mururu for its dam on the upper Ruzizi River.
Table 2. Chemical composition of Lakes Kivu and Tanganyika and major tributaries to Lake Tanganyika.
Source of data: a, b and c from Degens et a/. ( ! 97 3 ), d from Beauchamp ( 1 939). n.d., not determined Na (mg/1)
K (mg/1)
Ca (mg/1)
Mg (mg/1)
Cl (mg/1)
so, (mg/1)
co, (mmol/kg)
a Kivu 1 2 1 ·6 b R uzizi (near B ukavu) 1 1 7·0 c Tanganyika 66·3 d Malagarisi 1 6-4
97-4 1 00·2 34·2 2·4
4·8 4·6 8·2 12·9
87 88 4 1 ·5 9· 1
55 n.d. 21 1 5 ·5
23·8 n.d. 3-4 2· 1
1 2·5 1 3· 1 5·64 1 · 55 (mEq/ 1 )
The density increase produced b y higher salinities just offsets the density decrease inherent with the higher temperatures which results in a stable thermohaline density structure. This meromixis greatly reduces the exchange of water between the upper 70 m of water which circulates at least seasonally and the anoxic water of greater depth. Lake Tanganyika is thermally stratified with a perennial thermocline at approximately 1 00 m. Vertical temperatures and major ion distributions can be found
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(_____
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� � �
!::> ;:s !::>..
� �
I
I
. ... .. ...
. '" ·
.
• • .
.
..
K
·: :.
..
··
300
.
.
200 I
.. .
Temperature (•c)
-
Ca
Q) c
23
50
-·
Q.
(b)
400
•
-
.s:.
400
300 I
-
. •. .
300
e
200 I
!
.
200
400
100 I
.
.. ,..
Na
�.'1
).. .
450
Fig. 2. (a) D istribution of major cations (Ca, Mg, K, Na) in Lake Kivu waters from several stations, after Degens et al. ( 1 973). (b) Temperature profile versus depth for Lake
Kivu (deep northern basin) after Degens et al. ( 1 973).
� (")
�
Kivu- Tanganyika Basin
47
in Degens et a!. ( 1 9 7 1 ). The Ruzizi River presently flows from Lake Kivu to Lake Tanganyika and is the major contributor to the water and salt budget of Lake Tanganyika (cf. Table 1 ). The impact of the Ruzizi on Lake Tanganyika is clearly reflected in the chemistry of the surface waters (Table 2) especially the proportions of potassium and magnesium.
S EDIMENTS Lake Tanganyika
Eleven cores were studied. B ased on sediment colour, diatoms and mineralogical composition, four stratigraphic units can be recognized (Fig. 3). Sedimentation rates are in the order of 30 - 50 em per 1 000 years in the deep basins and approximately ten times less on the shallow sill which separates the two deep basins (Degens et a!., 1 9 7 1 ) . The longest core studied was 1 0 · 74 m ; i t was taken b y the Department o f Zoology of Duke University, Durham, N.C., in the southern part of the lake. This core (core TL 2) is dark grey and stratigraphically uniform; the average sedimentation rate is 49 em per 1 000 years (Livingstone, 1 965). 0 Sill Lake Tongonyi ko Sediment stratigraphy
5 .c::
a.
500
"' "0
Northern Basin
[:::J I Corbonotes/5Y2 / I - N 6 EJ 2 Koo linite/N4 [:::J 3 Smectite/595/1 I!]]J 4 Smectite/N6
Nontronite layer
Ar- Aragonite Ioyer MQC - - Mg - Colcite Ioyer Carbonate content>IO 0/o
Strati graphic units
0
oooo
1 00 Di s t a n c e
( kml
200
250
Fig. 3. Gross stratigraphy of selected cores from Lake Tanganyika based on characteristic minerals and
sediment colour.
Unit 1 (the youngest) is characterized by alternating light (diatoms) and dark layers (clay) forming a varve-type pattern. High-magnesium calcite (9- 1 1 mol per cent MgC03 in the northern basin, 7-9 mol per cent in the southern basin) is present as well as three thin aragonite layers which only occur in the cores from the northern basin. Kaolinite is by far the dominant clay mineral. The abundant diatoms all belong to the
48
Peter Stoffers and R. E. Hecky
genus Nitzschia. Organic carbon content is high ( > 4%) and pyrite in the form of framboids is found frequently. Unit 2 is medium grey in colour. Lamination is not as pronounced as in the first unit. The dominant components are kaolinitic clays and diatoms (Stephanodiscus, Melosira). Organic carbon content is high (6- 1 2%) and pyrite layers are frequent. No carbonates were detected. Concerning their texture, these two upper units are mainly clayey silts although sand layers are present occasionally which are mostly of turbiditic origin. Units 3 and 4 are only found in the cores fmm the sill area. Unit 3 is bluish-grey in colour whereas Unit 4 is light grey. Both are composed of carbonate-free fine-grained mud. The clays are predominantly smectite . Organic carbon content is low ( 1-2%) reaching its lowest values in Unit 4. No pyrite w as found in these units. Microfossils are rare.
Lake Kivu
Ten con :s from the various basins of Lake Kivu were studied. Details on the mineralogy and chemistry of the sediments can be found in Reeky & Degens ( 1 973), Degens & K ulbicki ( 1 973) and Stoffers ( 1 975). In Fig. 4 a stratigraphic correlation of some repres �ntative cores based on the occurrence of characteristic minerals is attempted. T he 14C dates are from Reeky & Degens ( 1 973). It seems that during the coring operat: on a substantial part of the upper sediment section was lost in some cores (A. Driscoll, 1 Jersonal communication). Especially at Station 4, superpenetration of 1 he corer was I toticed which is confirmed by stratigraphic considerations, as well as by ; tge dating. A c 'ate of 1 1 , 200 years B .P. was obtained from the top of this core. The base 1 ;ave an age of 1 3,560 years B.P. which is the oldest sediment recovered from the lake. ! :edimentation rates for the upper parts of the cores can only be given as an ; tpproximation They are greater than 30 em per 1 000 years based on core 9 and 1 0, .vhere not too Tiuch sediment seems to be missing. In the lower parts of the cores, sedimentation t ates are much higher, ranging between 1 and 5 m per 1 000 years. Core 1 4, take1 t in 3 1 0 m water depth in the western par� of the lake, is particularly noteworthy. In , he lower part of the core, well-rounded poorly sorted pebbles prin ;ipally of ml tam orphic rocks-and shell fragments, are the dominant constitu ents No sign of ! :rading was observed. F ·om Fig. 4 it c an be seen that the oldest sediments deposited in Core 1, 1 2 and 1 3 are s iderite-rich se iiments In Core 1 3 and 1 2 these sediments are coarse grained with a . 'and content of me re than 25%. Especially in Core 1 3, these sediments have a crumbly tt. xtu -e similar to a soil. During the deposition of the siderite-rich sediments, finely lan.'in 1ted organic-rich muds irregularly interrupted by maj or diatomite intercalations ( > 0· 5 em) were dc.oosited at Station 4 which is situated in the deepest area of Lake Kivu. , \lso at Statio.1. 4 are found infrequent layers of manganosiderite, one of which wa s ap Jroximately 1 em thick and nearly pure. The siderite-rich sediments are topped by a se< liment sequence alternating between carbonate-free pyrite-rich sediments and ara. �onitic muds. It t the lower parts of the cores, the pyrite-rich muds are dark grey, texturally stiff and struc tureless. At Stations 1 and 9, these sediments are mainly clays (kaolinite) whereas at St� tions 1 0, 1 2 and 1 3 silts dominate. Towards the top of the cores, the pyrite-rich muds .ue dark brown in colour. These muds often have a gel-like consistency with a high m·ganic carbon content ( > 1 0%). Water content is very high (90-95%). Only in ·
49
Kivu- Tanganyika Basin K 15
K 13
K1
K 12 - -? - -
LOST
LOST
LOST
LOST
I
-�
K 4 ? LOST
K 10
'
?
440m
86m
?
STRATIGRAPHY
OF
LAKE KIVU SEDIMENTS
�o�o;o:o�o�o�o�o;� Aragonite � :��ydro (:·:·:·:·:·:·:·:·:·:·:·:·1: ��t;da�f��i�e c=J �����n�ich h C=:J �!�;':,!::�
! 1t0
90m
( W.ter Depth ) 600
700
13 1\0 ! llO
800 l
_____J 402m
Fig. 4. Gross stratigraphy of selected cores from Lake Kivu based on characteristic minerals.
Cores 9 and 10 the carbonate intervals start with high-magnesium calcite ( - 1 2 mol per cent MgC03) and pass up into aragonite. Aragonite muds are found in the lower part of Station 15 which is located in the Bukavu basin. This basin is separated from the main lake by a shallow sill (-30 m). Core 15 is characterized by the presence ofmonohydrocalcite, high-magnesium calcite and protodolomite. The monohydrocalcite intervals are laminated with brownish yellow 0·5-3 mm thick layers of monohydrocalcite alternating with brownish-green layers of diatoms. Below this sequence are layers of black sapropelic sediments containing hardly any carbonates. o1RO values for the aragonites present in Station 1 5 fall within the narrow range of + 2 · 2-3· 1 per mil relative to PDB. Also the o13C data are very constant ranging from + 6 · 3 to 7 ·0. The isotopic composition of the monohydrocalcite/protodolomite shows lighter o1RQ and o13C values falling into the range of + 1 · 1 - 1 · 5 and 2 · 5-5 · 7, respectively.
50
Peter Stoffers and R. E. Hecky
Diatoms
Characterization of the biostratigraphy of Lake Kivu is based on siliceous microfossils present in the northern basin cores, 1 0 and 4. These two cores are considered, on the basis of their stratigraphy and their radiocarbon chronology, to give a nearly complete record of sedimentary events through the last 1 4,000 years. The methods of sample preparation and counting are described in Reeky & Kilham ( 1 973). Sampling frequency was at least once every lO em. Ecological interpretations are based on Richardson ( 1 968, 1 969). From the bottom of the superpenetrated Core 4 to 1 30 em in Core 10 only three species comprise well over 90% ofthe flora at any sample level. These three are Stephanodiscus astraea v. minutula, Nitzschia fonticola, and Nitzschia spiculum. Among the other diatoms only Nitzschia acicularis achieves abundances over 5% for short periods of time. The relative abundance of Stephanodiscus astraea declines almost linearly from about 90% at the base of Core 4 to zero at 1 30 em in Core 10 (Fig. 5). Nitzschia fonticola also disappears at this depth interval which forms a sharp contact between laminated aragonitic mud below and dark brown highly organic pyrite-rich mud above. At this interval, a thin layer of volcanic ash is also present. In a space of 5 mm the diatoms change from Nitzschia spiculum to purely Stephanodiscus astraea to relatively short and wide species of Nitzschia such as N. Palea and N. acomodata which are present in the dark brown material. Even more striking are the changes in absolute abundances, as diatoms are scarce in the brown material, and minute siliceous scales of the chrysophyte Paraphysomonas vestita became the most abundant microfossil. In the upper sediment section of Core 1 0, Nitzschia spiculum is joined by N. bacata and N. mediocris as new elements in the diatom plankton. Another new genus is a Chaetoceros sp. This genus is rarely recorded from inland waters, as it is typically marine.
[ B.P.J yrs. X
Frequency % 0
50
100
10 3
2
4
.. "' <( (.) ;!
6
8
10
12
Fig. 5. Dominant diatoms of Lake Kivu sediments found in Cores 4 and 10 from the central part of the lake.
Kivu- Tanganyika Basin
51
DISCU S S ION
The superpenetrated Core 4, located in the deep northern basin of Lake Kivu, covers the time interval from about 1 3,700 to 1 1 ,200 years B .P. At 1 3 ,700 B.P. Nitzschia fonticola is prominent. The sediments are silty with high A l 203 contents suggesting a low lake level. The prominence of N. fonticola at this time indicates that the waters are relatively rich in silica and a nearby tributary or at least near-shore conditions are implied. The well-rounded poorly sorted pebbles and shell fragments in the lower part of Core 1 4 are thought to represent a former beach deposit (Hecky & Degens, 1 973). Support for this interpretation comes from seismic profiles showing that at depths of less than 300 m, which are found mostly in the southern part of the lake, there is only a thin veneer of sediments overlying the structural basement. The beach deposit has not been directly dated but it is thought to reasonably correlate with the bottom of Core 4 (Hecky & Degens, 1 973). During this time period Kivu was a highly concentrated closed lake. Palaeohydrological reconstruction (Hecky & Degens, 1 973) suggests that a combination of mean annual temperatures 3 ·c lower than the present and rainfall approximately one half the modern level would be adequate to maintain the lake at minus 300 m. Fluctuating lake levels are suggested by the changing diatom population of N. fonticola, N. spiculum and S. astraea. But generally a rising lake level is indicated. At 1 2,000 years B.P. N. spiculum became prominent in Core 4 implying a rather stable stratification. The deep water and sediments became more reducing. Sedimentary pyrite and organic carbon concentrations increased at Station 4. At the same time, coarse grained siderite-rich sediments were deposited in the shallower parts of the lake (Cores 1 and 1 2) suggesting weakly reducing waters. The soil-like sediments found in the lower section of Core 1 3 which was taken in 86 m of water indicate that the lake stood much lower than its present level, and it was probably lower than the coring site before then. Initiation of sedimentation at Station 1 3 occurred at about 1 2,000 years B.P., at which time the lake level exceeded minus 86 m. The rising lake level brings a corresponding rise in the interface between weakly and strongly reducing sediments and changes the sedimentary regime at the shallower stations. A shift from relatively coarse-grained siderite-rich to fine-grained pyrite-rich sediments can be observed. The association of these conditions with the evidence for volcanism (high Ti02 content, Degens & Kulbicki, 1 973) during this period could indicate sublacustrine volcanic activity and the introduction of saline spring water at depth (crenogenic meromixis, Hutchinson, 1 957). * Equally plausible is a climatic cause. The climate was becoming warmer and more moist during this time, and a closed saline lake was being freshened. Prolonged intervals of warm, wet years might favour the establishment of a relatively warm dilute surface layer of water overlying cooler saltier water. The resulting sharp, stable thermocline would reduce mixing and Nitzschia spiculum would be favoured as nutrients would accumulate below the thermocline and be unavailable for algae in surface waters. At about 1 1 ,500- 1 1 ,000 years B.P. aragonite deposition started in the shallower areas of the lake (Cores 1 , 1 2 and 1 3). In the deeper regions ofthe lake (Cores 9 and 10) high-magnesium calcite occurred around 10,500 years B.P. followed by aragonite *Crenogenic meromixis is a stable stratification induced by submerged saline springs that deliver dense water to the lake usually at deep portions of the basin.
52
Peter Stoffers and R. E. Hecky
sedimentation. The difference in the carbonate mineralogy between the shallow and deep water cores may indicate the much higher photosynthetic activity in the near shore water. The sudden occurrence of carbonates indicates a change in water chemistry. The lake probably established a deep perennial thermocline with no salinity stratification similar to the modern Lake Tanganyika. The shift in sedimentation from aragonite to monohydrocalcite in the Bukavu basin around 1 0,000 years B.P. seems to mark the opening of the main lake to the south. The low 8180 values of the monohydrocalcite as compared to the aragonite data also suggest a change from a closed to an open system. The return to aragonite sedimentation for a short time may indicate some fluctuating water levels. Aragonite deposition ceased around 9300 years B.P. at Station 1 5 giving way to monohydrocal cite, protodolomite and high-magnesium calcite sedimentation. The high organic carbon content associated with these sediments also points to a higher lake level. From 9000 to 5000 years B .P. S. astraea declines in abundance to be replaced by N. fonticola and subsequently by N. spiculum. This reflects a general warming trend in the climate. The thermocline is becoming more stable as the temperature and density differential of surface and deep water increases. Reduced mixing and lower productivity result. It must be emphasized that shifts between S. astraea and N. fonticola reflect only the availability of silica, and do not necessarily imply any changes in the total production of the phytoplankton or even of the diatoms. A relative increase in N. spiculum, however, does mean that overall phytoplankton production has decreased as it reflects low mixing rates. The crenogenic meromixis which is the most distinctive feature of the modern Lake Kivu is first evident in the available record about 5000 years B.P. when carbonate deposition suddenly ceased and the diatom flora dramatically changed. The history of the lake since then is dominated by the creation and destruction of the meromictic condition. In the modern lake, the density differential between surface water and 3 water below 350 m is not large, being approximately 0·002 g/cm (Schmitz & Kufferath, 1 955). Thus the modern lake is delicately poised, and rather small changes in density effected by either a combination of cooling and/or concentrating surface waters or heating and/or diluting the deep water could obliterate the meromictic condition. Two conditions are required for the maintenance of crenogenic meromixis: ( 1) a deep source of denser spring water must exist, and (2) surface waters must be constantly diluted by runoff and loss of salts at the outflow. Lake Kivu had the Ruzizi outlet since 9500 years B . P. Therefore, the imposition of crenogenic meromixis seems dependent on the first condition. Either spring activity greatly increased at that time, or the deep spring's have always been present but some catastrophic event introduced a source of dense water at depth. This denser water could have imparted a stability which spring activity caused to persist for nearly 1 000 years. Damas ( 1 937) suggested that hot lava entering the lake would yield warm salt-rich waters which would collect in the deep basin and such an occurrence has been witnessed in modern times. Similarly, the eruption of the sublacustrine volcanoes (Verbeke, 1 957) found in Lake Kivu could yield a source of hot saline water. Peaks of Ti02 occur at each shift from aragonite diatom-rich mud to aragonite-poor sediments suggesting some relationship with volcanism (Hecky & Degens, 1973; Degens & Kulbicki, 1 973). The breakdown of the meromixis is probably related to climate which in the period 5000-2000 years was becoming cooler and drier. These conditions would lessen the density differential and reduce the work required to effect total mixing (Hutchinson,
Kivu- Tanganyika Basin
53
1 957). Loss of the Ruzizi outlet would certainly lead to destruction of the meromictic condition. At 2000 years B.P. meromixis was again set up and it seems to have persisted to the present. The modern climate is certainly warmer and wetter than that of 30002000 years B .P. (Richardson & Richardson, 1972). The microfossil data described above are consistent with a catastrophic origin for the meromixis at 5000 years B . P. Within a vertical distance of5 mm in the core, there is a shift from N. spiculum dominance to nearly pure S. astraea to diatom deficient sediments. This shift can be interpreted as a change from relatively unproductive nutrient-poor conditions to extreme eutrophy to very oligotrophic conditions, and this all occurred in less than 1 5 years, if the mean sedimentation rate is applied. Chaetoceros sp. appears only after 5000 years B.P. indicating that the lake was more saline than previously. This would be expected with the introduction of saline water to the lake. The second initiation of meromixis is not as dramatically apparent in the microfossil stratigraphy, and the more moist climate which led to a dilution of surface waters and perhaps the renewal of the Ruzizi outlet may have been sufficient to re establish a stable chemocline. However, there are also high Ti0 2 concentrations and the cessation of carbonate sedimentation was abrupt, so volcanism may have been involved again. The stable haloclines established during the last 5000 years resulted in a greatly reduced mixing, acidic deep waters, and extremely low algal productivity in surface waters. Carbonate deposition ceased below the halocline. Under the low nutrient conditions the long thin Nitzschia were expected. N. spiculum was joined by N. bacata and N. mediocris. The ecological relationship of these three is unclear as little is known about any of them. N. mediocris and N. bacata have their greatest relative abundances in the cores where calcium carbonate is low or absent indicating that stratification was most stable and of a thermohaline character. N. spiculum has a maximum for the past 5000 years over a time when calcium carbonate was being deposited and perhaps only a deep thermocline was present. The major climatically and volcanically-induced alterations in Lake Kivu had profound effects on Lake Tanganyika as both lakes are presently joined by the Ruzizi river. The discharge of Lake Tanganyika via the Lukuga river at present is much smaller than the amount of incoming Ruzizi water (Table 1 ), and closure of the Kivu system would be adequate to close Lake Tanganyika. As a consequence of Lake Tanganyika's morphology, the lake level would be drastically lowered. For the time period before 1 3 ,000 years B . P. when Lake Kivu was a closed alkaline lake, a low lake stand of minus 600 m is assumed for Lake Tanganyika by Reeky & Degens ( 1 973) roughly corresponding to the minus 550 m sonic reflecting layer reported by Capart ( 1949). This would imply that Lake Tanganyika was split into two isolated basins with the sill area being exposed to erosion. However, textural, mineralogical and geochemical studies of the cores from the sill area do not show any signs of erosion or soil-formation for the corresponding time interval. Subaqueous erosion while water levels were rising might eliminate traces of lower levels. The very low sedimentation rates on the sill may reflect discontinuous sedimentation. The predominance of smectite over kaolinite found in the lower section of the sill cores may indicate different weathering conditions before 1 3,000 years B .P. As the climate was dry smectite formation was favoured. Highest abundance of kaolinite occurred between about 1 0,000-5000 years B.P. when the general evidence for East Africa pointed to a pluvial period.
54
Peter Stoffers and R. E. Hecky
The linkage between Lake Kivu and Lake Tanganyika is demonstrated best by studying the carbonates in the Tanganyika cores. Prior to about 4000 years B.P. no carbonates could be detected. The carbonate deposition in the upper sediment sequence is related to the incoming Kivu water, but with an appropriate delay which reflects the large residence time of water and salts in Lake Tanganyika. Carbonate precipitation from surface water would not start until the Ruzizi-derived calcium began to increase in the surface water of Lake Tanganyika. The influence of the Kivu water with its high magnesium/calcium ratio is mirrored in the lateral distribution of the individual carbonate minerals. Aragonite and high-magnesium calcite ( > 9 mol per cent MgC03) is only found in the cores from the northern basin whereas only high magnesium calcite ( < 9 mol per cent MgC03) is present in the southern basin. No carbonates could be detected in Core TL2 taken at the extreme southern end of the lake. Fluctuations in the carbonate content reflect the changing amount of water supplied by the Ruzizi to Lake Tanganyika. The stratigraphic record for the two lakes deduced from mineralogical and diatom studies has provided an insight into the interaction of climate and geology in determining the hydrochemistry and sedimentation during the Late Pleistocene Holocene period. The climatic interpretation obtained here is more or less correlative with other East African data over the available time period. (cf. Livingstone, 1 975; Stoffers & Holdship, 1 975; Richardson & Richardson, 1972).
REFERENC ES BEAUCHAMP, R.S.A. ( 1 939) Hydrology of Lake Tanganyika. Int. Rev. Hydrobio/. 39, 3 1 6-353. CAPART, A. ( 1 952) Le milieu geographique et geophysique. Exploration hydrobiologique du Lac
Tanganyika ( 1 946- 1 947). Inst. Roy. Sci. Nat. Be/g. 1, 3-27. CouLTER, G.W. ( 1 963) Hydrogeological changes in relation to biological production in southern Lake
Tanganyika. Limno/. Oceanogr. 8, 463-477. DAMAS, H. ( 1 937) La stratification thermique et chimique des lacs Kivu, Edouard, et Ndalaga (Congo
Beige). Verh. Int. Limno/. 8, 5 1 -68. DEGENS, E.T. & KULBICKI, G. ( 1 973) Data file on metal distribution in East African rift sediments. Tech.
Rep. Woods Hole Oceanogr. Ins/. 73- 15, 1-280. DEGENS, E.T. & Ross, D.A. (Ed.) ( 1 969) Hot Brines and Recent Heavy Metal Deposits in the Red Sea.
Springer, New York. DEGENS, E.T. & Ross, D.A. ( 1 976) Strata-bound metalliferous deposits found in or near active spreading
centers. In: Ores in sediments, sedimentary and volcanic rocks (Ed. by K. H. Wolf), pp. 1 65-2 1 1 . Elsevier, Amsterdam. DEGENS, E.T., VoN HERZEN, R.P. & WONG, H.K. ( 1 9 7 1 ) Lake Tanganyika: water chemistry, sediments, geological structure. Naturwissenschaften, 58, 229-24 1 . DEGENS, E.T., VON HERZEN, R.P., WONG, H.K. & J A N NASCH, H.W. ( 1 973) Lake Kivu: structure, chemistry, and biology of an East African rift lake. Geo/. Rdsch. 61, 245-277. HECKY, R.E. & DEGENS, E.T. ( 1 973) Late Pleistocene-Holocene chemical stratigraphy and paleolimnology of the Rift Valley Lakes of Central Africa. Tech. Rep. Woods Hole Oceanogr. Ins/. WHOI 73--28. HECKY, R.E. & K t L HAM , P. ( 1 973) Diatoms in alkaline, saline lakes: ecology and geochemical implications. Limno/. Oceanogr. 18, 53-72. HuTCHINSON, G . E. ( 1 957) A Treatise on Limnology, V ol. I. Wiley and Sons, New York. K UFFERATH, J. ( 1 952) Le milieu biochimique. Exploration hydrobiologique du Lac Tanganyika ( 1 946-47). Inst. Roy. Sci. Nat. Be/g. 1, 3 1 -47 LiVINGSTONE, D .A. ( 1 965) Sedimentation and the history of water level change in Lake Tanganyika. Limnol. Oceanogr. 10, 607-6 tO. LIVINGSTONE, D.A. ( 1 975) Late Quaternary climatic change in Africa. Ann. Rev. eco/. Systematics, 6, 249-280.
Kivu- Tanganyika Basin
55
MOLLER, G . & FORSTNER, U. ( 1 973) Recent iron ore formation i n Lake: Malawi, Africa. Mineral. Deposita, 8, 278-290. RICHARDSON, J . L. ( 1 968) Diatoms and lake typology in East and Central Africa. Int. Revue Ges. Hydrobiol. 53, 299-338. RICHARDSON, J . L. ( 1 969) Characteristic planktonic diatoms of the lakes of tropical Africa. Int. Revue Ges. Hydrobiol. 54, 1 75- 1 76. RICHARDSON, J . L. & RICHARDSON, A. E. ( 1 972) History of an African rift lake and its climatic implication. Ecol. Monogr. 42,499-534. SCHMITZ, D.M. & KuFFERATH, J. ( 1 955) Problemes poses Ia presence de gaz dissous dans les eaux profondes du Lac Kivu. Bull. Seances Acad. roy. Sci. Coloniales, N.S. 1, 326-356. STOFFERS, P. & FISCH BECK, R. ( 1 974) Monohydrocalcite in the sediments of Lake Kivu. Sedimentology, 21, 1 63- 1 70. STOFFERS, P. & HOLDSHIP, ST ( 1 975) Diagenesis of sediments in an alkaline lake: Lake Manyara, Tanzania. IXth International Congress of Sedimentology, Nice, theme, 7, 2 1 1 -2 1 7. STOFFERS, P. ( 1 975) Sedimentologische, geochemische und paliioklimatische Untersuchungen an ostafrikani schen Riftseen. Habilitationsschrift Universitiit Heidelberg, 1 1 7 pp. VERBEKE, J . ( 1 957) Recherches ecologiques sur Ia faune des grands lacs de !'est du Congo Beige. Exploration Hydrobiologique des Lacs Kivu, Edouard et Albert, Vol. 3, Fasc. I , 1 77 pp.
Spec. Pubis int. Ass. Sediment. ( 1 97 8) 2, 57-8 1
Holocene carbonate evolution in Lake Balaton (Hungary): a response to climate and impact of man*
G ER M A N M U L L E R and F R A N K W A G N E R
Institut fur Sedimentforschung, Universitiit Heidelberg, P. O. Box 1 0 30 20 D-6900 Heidelberg, West Germany
AB STRACT
Mineralogy, geochemistry and oxygen isotope composition of Lake Balaton carbonate sediments reflect fluctuations in the composition of the lake water, which were strongly influenced by both climate and man during the past 8000 years. During the 'Pre-Roman Stage' (about 7500-2000 y. B.P.) when the lake had no outflow (closed basin), calcite with low Mg, Sr and Na concentrations was precipitated at high water levels during periods with a relatively low rate of evaporation. High magnesian calcite with up to 20 mol per cent MgC03 and elevated Sr and N a concentrations and protodolomite formed at low water levels during periods of high evaporation from solutions with higher Mg/Ca and Mg + Ca/Sr ratios and elevated Na concentration. These conclusions are also strongly supported by oxygen isotope data of the autochthonous carbonates. Lithium, associated with clay minerals correlates positively with Mg, Sr and Na. In addition to the vertical fluctuations in carbonate composition within each core, pronounced lateral changes are found between the different cores: From core A (closest to the main inflow, the Zala River) to core F (farthest distance from the Zala River) the concentrations of Mg, Sr and Na incorporated in carbonates increase more or less steadily. The interstitial waters in the cores show a similar development: The Mg/Ca ratios and the Na concentrations increase generally from core A to core F. Within each core the highest Mg/Ca ratios were found to occur in the lower half of the core where they are still close to the zone where the highest Mg concentrations of carbonate minerals are found. Two periods with evaporation maxima can be traced along the long axis of the lake: one towards the end of the Atlanticum (about 5000 y . B.P.), another towards the end of the Subboreal (about 3000 y. B . P.). After an artificial outflow was built by the Romans about 2000 y. B.P., the lake changed from a closed to an open basin with only minor fluctuations in water level and hydrochemistry
Since then, high magnesian calcite with a more or less
constant rate of MgCO,-, Sr and Na incorporation has been precipitating during periods of algal blooms. *Dedicated to Professor Dr Martin Schwarzbach, Ki:iln on the occasion of his 70th birthday.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
58
German MUller and Frank Wagner FOREWORD
Already in his first semester, a student of geology will learn about the importance of some sediments as climatic indicators. Tillites derived from glaciers are witness of cool climate, whereas coral reefs are believed to have been grown in warm seas. He will also learn of the paradox that raindrop imprints are more an indicator of a warm and arid climate than of a climate of abundant rainfall. During the past decades, methods used in the reconstruction of palaeoclimates have become more sophisticated. The application of geochemical and isotope-geochemical methods and the addition of clay minerals, zeolites and other sedimentary materials have enlarged our knowledge of how climatic changes can be recognized in sediments. The present study is an attempt to apply such methods to homogeneous lacustrine carbonate muds with the aim of reconstructing the historical development of a lake and its sediments and to develop a stratigraphy based mainly on the chemistry and mineralogy of the autochthonous carbonates. Since carbonate chemistry is strongly related to the chemistry of the water from which precipitation took place, the carbonates strongly reflect the chemistry ofthe lake water at the time of their deposition. As the water level and the hydrochemistry of a lake (especially of a closed basin), is a response to the ratio evaporation : inflow + precipitation, climatic changes might be expressed in the carbonates precipitated in the lake. This is not the first attempt to reconstruct palaeoclimatic conditions from carbonate composition. In a study of numerous lakes from different climatic zones, Muller, Irion & Forstner ( 1 972) established the general relationship between carbonate mineralogy and Mg/Ca ratio of the lake or interstitial water from which precipitation or transformation took place. Stoffers ( 1 975) applied changes in carbonate mineralogy of sediments from Lake Victoria (Africa) to reconstruct palaeo Mg/Ca ratios. Stoffers & Reeky ( 1 97 8) used mineralogical and oxygen isotope changes in sediments of the Kivu-Tanganyika Basin to establish palaeo-salinities. Rothe, Hoefs & Sonne ( 1 974) observed an enrichment in 180 in Tertiary carbonate sediments from the Mainz Basin (Germany) by repeated evaporation of a closed basin which was paralleled by increasing concentration of Sr. Preliminary results of a combined oxygen isotope and trace element study of Precambrian carbonate sediments from the Bambui Group, B razil (in preparation), clearly indicate that even in very old and diagenetically altered carbonates, these methods might be useful to show changes in water composition which could then be related to general climatic conditions during the time of carbonate deposition. INTRODUCTION
Lake B alaton is one of the best studied lakes in the world. Towards the end of the last century, Louis Loczy ( 1 849- 1 920) organized a research team for the investigation of the lake and its surroundings. Between 1 897 and 1 920 a Balaton Monograph was published in Hungarian (A Balaton Tudomanyos Tanulmanozasanak Eredmenyey ) and German (Resultate der wissenschaftlichen Erforschung des Balatonsees) by the Hungarian Geographical Society with the contributions of sixty scientists 'dealing with the problems of geology, geomorphology, archaeology, ethnology, anthropology and history of the region' (Ronai, 1 969) - an excellent example of inter-disciplinary cooperation.
Holocene carbonate evolution in Lake Balaton
59
The investigation of the lake bottom by underwater coring (thirteen drillings with max. depths of 23· l m) was part of the programme and first results were published in the Appendix ( 19 1 1) to Vol. I, 1 (which appeared in 1 9 16) of the Balaton Mongraph: Uber die Sande des Balatonbodens (G. Melczer), Der Grund des Balatonsees, seine mechanische und chemische Zusammensetzung (P. Treitz) and Die chemische Zusammensetzung des Schlammes und des Untergrundes vom Balaton-Baden (K. Emszt). Treitz assumed two sources for the fine-grained sediment sequence ofthe lake: wind-transported clastic material and chemically precipitated lime-a conclusion which is confirmed by our studies. An investigation of the surface sediments of Lake Balaton (MUller, 1 969, 1 970) revealed that sedimentation in this lake is mainly governed by the precipitation of high-magnesian calcite. According to our knowledge, this was the first report of Recent high-magnesian calcite formation in a fresh water environment. Since then, this mineral has been found in many other lakes and other non-marine environments (Muller et al., 1 972). Analyses of a 1 · 1 5 m sediment core taken in the Lake Balaton surface sediment study showed that the rate of Mg incorporation in the high magnesian calcite and other chemical parameters varied with depth. As a consequence, a series of sediment cores comprising the full Holocene sedimentary sequence was taken along the long axis of the lake. The results of the investigations of these cores are presented in this report.
S ETTIN G
A compilation of data concerning the limnology of Lake Balaton was presented by Entz & Sebestyen ( 1946); the period 1 946-1960 is covered in a compilation by Sebestyen ( 1 962). Both reports are printed in German and Hungarian. An Excursion guidebook Study Tours (in English) prepared for the International Symposium on Palaeolimnology at the B iological Research Institute of the Hungarian Academy of Sciences at Tihany, 28-3 1 August 1 967, sums up the most important data on the lake. Geology
Lake Balaton, situated in the centre of the western region of Hungary called Transdanubia (Fig. 1 ), is the largest and shallowest lake in Central Europe. It stretches in a WSW-ENE direction along the border of the foothills of the Hungarian Central Mountains, which are built up of Palaeozoic and Mesozoic sedimentary rocks, partly covered by Tertiary and Quaternary sedimentary rocks (including basalts of Late Tertiary age). Clastic sedimentary rocks of Upper Pliocene age (Pannonian), mostly covered by Quaternary loess and soils, border the southern and eastern side of the lake. The lake itself is in an area of Pannonian sediments which are overlain by a thin cover of Late Pleistocene/Holocene sediments (Ronai, 1 969). Pollen analyses carried out on the lake sediments by Z6lyomi ( 1 953) clearly revealed 'that the present lake as a whole was born at the beginning of the Holocene' (Ronai, 1 969). Climate
The Lake Balaton area has a particular meso-climate within the general climatic conditions of Hungary. It belongs to the climatic belt with prevailing Westerlies and
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I. Geology of the catchment area of Lake Balaton (after Geological map of Hungary, 1 :300,000). I, Quaternary sediments; 2, Pliocene (Pannonian) sediments;
3, Basalts of Pliocene age; 4, Triassic sedimentary rocks, mainly limestones and dolostones; and 5, Permian sandstone. Wind directions at Si6fok, average of July 1 90 1 - 1 950 (after Szestay, 1967) .
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Holocene carbonate evolution in Lake Balaton
61
continental-type temperature conditions. Annual average temperatures range between 10 and 1 1 'C. Monthly air temperatures of the coldest month (January) are slightly lower than O'C, the July average (warmest month) is 2 1 ·6'C. The water temperatures of the shallow lake follow the seasonal variations in air temperature with only a slight time lag. They are, however, always several degrees higher than the corresponding air temperatures (see below). Hydrology and water budget
Szestay ( 1 967) presented the following data (long-term averages) for the water balance of Lake Balaton: drainage area: 5200 km2; water surface area: 600 km2; mean and maximum depth: 3·3 m and 1 2 m, respectively; mean water temperatures (t) in summer months as compared with mean air temperatures ( 1) t
T
V 19·5 16·2
VI 23·2 1 9·4
VII VIII IX 25·0 23·8 19·8'C 2 1 ·6 20·7 17·0'C
mean annual water level fluctuation: 45 em; mean annual precipitation: 630 mm; mean 3 annual inflow: 860 mm ( 1 7 m /sec); mean annual outflow from the lake controlled by 620 mm. The average annual water balance is the gates at Si6fok: 12 m3/sec characterized by the equation: =
Precipitation + Inflow 630 + 860
=
Evaporation + Outflow 870 + 620mm
The average water exchange period amounts to only 2·2 years. The river Zala delivers more than half the amount ( 1 0 m3 /sec) of the inflowing waters. The Sio canal, originally a natural outflow regulated at first by the Romans, is the outlet of the lake, regulated by the locks at Si6fok. If under present-day conditions the outflow were stopped, a closed system would be established with a surface area of 1730 km2 (enlargement by a factor of 2·88). Hydrochemistry
The water of Lake Balaton is a well-buffered alkaline water (alkalinity 4·6; pH 8·6) of the HC03- - Mg2+ - Ca2+ type. Thirty-four perennial and twenty-six temporal inflows carry water of various chemical compositions into the lake. If the average chemical composition of Lake Balaton tributaries and of outflowing Lake Balaton water is compared (Table 1 , based on measurements between 1950 and 1 959, Entz, 1959), it becomes evident that (a) the outflow contains less dissolved salts than the inflow; and (b) all ions - with the exception of Ca2+ and HC03-- are enriched in the outflow. Two factors are responsible for this change in hydrochemistry: (a) the rate of evaporation within the lake basin is greater than the rate of precipitation (see the previous paragraph) which leads to a general increase of most of the ions in the outflow; and (b) due to large scale precipitation of high-magnesian calcite (Muller, 1969, 1 970) as a result of the photosynthesis of the phytoplankton, large amounts of Ca2+ and HC03- (plus minor amounts of Mg2 + ) are extracted from the lake water. Both factors combined lead to the observed enrichment of most of the ions but at the same time to a depletion ofCa2+ and HC03- in the lake and to a general decrease ofthe total electrolyte concentrations in the outflow.
62
German Muller and Frank Wagner
Table 1. Average chemical composition of water from rivers flowing into Lake B alaton and Lake Balaton water itself (after Entz, 1 959)
K+ Na+ Ca2 +
M g2 +
HCO,C1SO,' -
Inflowing water from rivers (mg/1)
Lake B alaton water (mg/1)
5·6 18·1 80·6 36· 1 336·8 9·0 46· 1
6· 1 27·0 32.0 47.8 280·0 1 2·0 55·0
Fig. 2. Scanning electron photomicrographs of contemporaneous high-Mg calcite (upper pictures)
and a mixture of high-Mg calcite (upper picture) and a mixture of high-Mg calcite and cation-disordered protodolomite (lower pictures), from core E.
Holocene carbonate evolution in Lake Balaton
63
Chemical investigation of the lake water along the axis of the lake during two periods (23-25 June 1 95 9 and 28-29 June 1 960) by Paszto ( 1 963) clearly show that from the western to the eastern part of the lake, the Ca2 + concentrations decrease and the Mg2 + concentrations increase. As a result, the Mg2 + /Ca2 + atomic ratio (hereafter abbreviated Mg/Ca ratio) increases considerably from West to East (Fig. 2). In the 1 959 period the increase from W to E was from about 1 ·5 to 2·6, during the 1 960 period from about 1 ·0 to 2·0. It must be emphasized that the Mg/Ca ratios encountered during the two sampling periods were random and are in no way averages or even maximum values; they represent only the momentary situation in the lake water during the periods of sampling. The average MgC03 concentration found in the calcite lattice of the sediments permits the assumption that the Mg/Ca ratio of the lake water at the time of precipitation of the calcite was generally higher (Fig. 1 1).
P HYTO PLANKTON AND S UB A Q U E O U S PLANTS: I MPORTANT CARBONATE PRODUCERS
In the course of phytoplankton studies performed in 1 967 (Tamas, 1 972), 1 47 phytoplankton species (including one species of aquatic mycophyta) belonging to six phyla could be identified. About half of the species are diatoms (Chrysophyta). High magnesian calcite precipitated during phytoplankton blooms causes a turbidity while in suspension, the so-called 'blondness' of the Balaton water (Felfoldi et al. , 1 969). As compared with the large amount of carbonates precipitated due to the metabolic processes of the phytoplankton, macrophyta 'produce' only negligible carbonate quantities. According to T6th ( 1 967) the macro-vegetation of the water is relatively poor in species. The bulk consists of Potamogeton perfoliatus and, to a lesser extent, of Myriophyllum spicatum. The carbonate crusts which develop on the upper surface of Potamogeton leaves consist of aragonite (Bidl6, 1960; Muller, 1 97 1). Other carbonate producers: molluscs
Unionida shells (mainly A nodonta and Unio species) cons1stmg of aragonite, commonly occur on the lake bottom. In 1 932, the 'zebra-mussel' Dreissena polymorpha appeared and rapidly spread over the entire lake, and by 1 934 the number of individuals had surpassed that of the U nionida. On a mud surface of 2 1 · 1 8 mZ, a total of26,568 Dreissena shells were found during a sampling campaign (P6nyi et al., 1 974) . Since Dreissena has its optimal distribution in water depths between 2 and 5 m, the whole of Lake Balaton offers very favourable conditions as far as water depth is concerned. However, compared with the amounts ofhigh-magnesian calcite produced during algal blooms, aragonite derived from mollusc shells is minimal, comprising less than 1% of the biogenic carbonates.
R E C ENT C A R B O NATE S EDIM ENTATION
In 1 968, a total of 1 7 5 surface samples (twenty-five traverses, each with seven samples) were collected from Lake Balaton. The results of the mineralogical
64
German Muller and Frank Wagner
investigations can be summarized as follows (Muller, 1 969, 1970). Recent sedimenta tion in the lake is governed chiefly by two processes which contribute in about equal proportions: (a) clastic material (quartz, feldspar, calcite, dolomite, phyllosilicates, heavy minerals) introduced by wind and rivers is distributed on the bed by currents; and (b) during the warmer seasons considerable fine-grained high-magnesian calcite is precipitated from the lake due to the extensive growth of phytoplankton. The precipitates are equant grains or aggregates of irregular shape. (Fig. 2) About half the grains (chiefly aggregates) are 2 to 6·3 J.Lm in size, and the rest smaller. The average MgC03 concentrations ofthe precipitates range between 6 and 8·5 mol per cent (Fig. 3) and increase from west to east along the axis of the basin. This increase reflects the general increase of the Mg/Ca ratio in the water of the lake, from west to east, during the algal blooms. Thus, the rate ofMg incorporation into the calcite lattice seems mainly to be determined by the Mg/Ca ratio of the water from which precipitation takes place. Figure 4 shows the carbonate content of the surficial lake sediments (after Muller, 1 969). The finest grained muds have the highest carbonate content (up to 69%); carbonate-poor sediments are coarse-grained and consist mainly of allochthonous minerals.
H O L O C E N E C A R B O N AT E EVOLUTION
In 1 972, a series of sediment cores was collected along the long axis of Lake Balaton, six of which (A-F, Fig. 4) are described here. The coring device consisted of a hollow steel tube (diameter 6 em) which was pressed by hand into the soft bottom of the lake. The cores, which are between 1 90 em and 300 em long, all penetrated the carbonate mud sequence completely and bottomed either in a substratum of sandy gravel, or, as in the case of core B, in peat. Megascopic Appearance
In three of the six cores, C, E and F, a 3 to 5 em sandy layer, rich in gastropods (subfamily Naticaea) and plant roots, was found 5-20 em above the substratum. This layer is comparable with sediments which occur at present in the reed-dominated littoral zone of the lake, between Tihany and Balatonfiired. With the exception of this sand layer rich in snail shells, the cores consist of homogeneous mud without any indications of stratification or any other sedimentary structures. The main variations are colour differences in the mud. The deepest third of a core is usually darker grey than the upper two thirds. An exception is a 5 em pinkish-greyish mud layer ('Rosa Zone') which occurs at about the border between the lower third and middle third of cores C, D, E and F. This 'Rosa Zone' is an excellent marker horizon. The pinkish colour of the 'Rosa Zone' can be explained by an admixture of colloidal haematite particles derived from red-coloured Permian sandstones and siltstones which crop out along parts of the northern shoreline of the lake (Fig. 1 ). An unusual, heavy torrential rainfall is assumed to have eroded larger amounts of fine-grained materials from the red beds (and the red soils derived therefrom) and washed them into the lake. The fact that Permian clastic rocks do not occur along the shoreline adjacent to cores A and B might explain the lack of the 'Rosa Zone' in these two cores.
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pinkish layer snails roots
Holocene carbonate evolution in Lake Balaton
67
Mineralogy of sediments
The same detrital minerals observed in surface sediments are present in the carbonate-rich muds of the cores: quartz, feldspar, clay minerals (mainly montmoril lonite and illite with minor amounts of chlorite and kaolinite), calcite and dolomite. Fine-grained autochthonous carbonate minerals (low-magnesian calcite, high magnesian calcite and a Ca-Mg phase with a dolomite-like composition) first appear immediately above the gastropod horizon or, if this horizon is not developed, some centimetres above the sand or peat substratum. Calcite contains a wide range ( 1 -20 mol per cent) of MgC03 in solid solution. A calcium-magnesium mineral not encountered in the surface sediments having a major X-ray reflection near that of dolomite, is rather abundant in some sections of the cores. It varies in MgC03-composition between 35-45 mol per cent. Cation-order reflections could not be detected with absolute certainty. A weak reflection at the position where the long range 22 1 -order peak of dolomite should appear belongs to the (202) ( 13 1 ) reflection of illite, which is present in all samples. The stability of this Ca Mg carbonate with respect to acid treatment is relatively high. In contrast to the associated high-magnesian calcite, it does not dissolve in dilute acetic acid. In dissolution experiments with C02- rich water the material dissolves congruently. This
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German Muller and Frank Wagner
68
behaviour clearly speaks for a mineral with the properties of protodolomite rather than high-magnesian calcite or 'pseudo-dolomite' (Berner, 197 1) . Gaines's ( 1 977) proposal, t o apply the term protodolomite strictly only t o minerals with order reflections in diffraction patterns and to apply the term 'pseudodolomite' to minerals in which cation-order cannot be unequivocally demonstrated, does not take into account other important physical and chemical properties. Until a better definition for these types of sedimentary dolomite-like Ca-Mg carbonates, which are quite common in Recent non-marine sediments, has been found, the designation 'cation-disordered protodolomite' is chosen to describe a mineral with properties, some of which are typical for either protodolomite or pseudodolomite. Morphologically the cation-disordered protodolomite appears to be identical to high-magnesian calcite and cannot be differentiated by purely optical means, including stereoscan observations (Fig. 2). Fig. 5 shows typical X-ray diffractograms of different grain size classes within a sediment sample from the 'Rosa Zone', core E. The sand fraction ( > 63 ttm) contains mainly detrital calcite; in the medium and coarse silt fractions (6· 3-63�-tm) detrital dolomite is abundant. The clay fraction ( < 2 ttm) consists nearly exclusively of primary high-magnesian calcite and cation-disordered protodolomite. This holds true also in the fine silt fraction (2-6·3 ttm) in which smaller amounts of detrital calcite are also present. Since more than 90% of the carbonates in the muds encountered in this study are made up of fine silt and clay fractions, allochthonous carbonate minerals play only a minor role in the carbonate mineralogy.
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Holocene carbonate evolution in Lake Balaton
69
Chemistry of interstitial solutions
From cores A, B, C, E and F, a total of thirty-six sediment sections with a length of 1 0 em each were collected and centrifuged not later than 2 h after retrieving the cores. From the filtered solution the cations (Ca, Mg, Sr, Na and K) were determined by atomic absorption spectroscopy, the anions (C 1, S04, HC03) by standard methods of water chemistry. The results are listed in Table 2. In Fig. 6 some of the data are presented graphically. For comparative purposes, the various thicknesses of the Holocene carbonate mud sections in the different cores were converted to a common (arbitrary) thickness. Results (a) Calcium. Ca concentrations range between 1 0·8 and 25·5 ppm. As a rule higher values normally occur in the top and basal layers of a core, minimum values in the middle or upper third. With one exception (core F), the maximum concentration is always to be found in the top layer. Ca generally tends to decrease from core A to core F. (b) Magnesium. M g concentrations range between 34·8 and 63 ·0 ppm. With one exception (core B), the minimum value always occurs in the top layer. Maximum values are to be found in the middle or in the lower part of the cores. There are no major lateral concentration differences amongst the different cores. (c) Strontium. Sr concentrations range between 0·05 and 0 · 1 7 ppm. Within a core, Sr has some similarity (but less pronounced) with Ca: higher values are found in the top and basal layers, a minimum occurring in the upper third or middle of a core. No pronounced differences between the different cores can be recognized. (d) Sodium. Na concentrations range between 1 8 · 8 and 50 · 8 ppm. With the exception of core B, the N a concentration increases continuously with depth, thus the minimum concentration is to be found in the top layer, the maximum concentration in the deepest layer. The increase is most pronounced in core E ( 1 9 · 8-50·8 ppm). With depth, a general increase of Na occurs within the upper seven sections then lower values were measured in the two lowest sections. From core A to core E the Na concentrations increase continuously. Core F has somewhat lower concentrations than core E, especially in the lower third of the core. (e) Potassium. K concentrations range between 5 ·9 and 1 5·0 ppm. As with Na, there is a general trend towards an increase in K concentrations with depth; the minimum values are always to be observed in the top layer. The lateral variations are less pronounced as with N a, however, cores E and F show distinctly higher concentrations than the other cores. (j) Chloride. Cl concentrations range from 1 0·6 to 23 · 1 ppm. Chloride has a clear tendency to increase from the top to the bottom layer of a core, the steps of increase are, however, smaller than with Na. Lateral differences are also present: Cores E and F have higher concentrations than cores A through D. Core A exhibits the lowest concentration of all. (g) Sulphate. S04 concentrations are the most variable. They vary between 6· 1 and 56 · 4 ppm, i.e. nearly within an order of magnitude. No general trend of distribution can be seen from the data, neither in a vertical nor in a lateral direction. (h) Bicarbonate. H C03 concentrations range between 297 · 8 and 5 85 · 8 ppm. Within a core low values are to be found both in the top and in the bottom layers with a
70
German Muller and Frank Wagner
maximum occurring near the middle of a core. No clear lateral trends in distribution can be traced. (i) Mg/ Ca (atomic) ratio. The Mg/Ca ratio shows the typical development plotted in Fig. 6. Within each core two minima occur: one in (or close to) the bottom and one in the top layer. Between these minima, the Mg/Ca curve bends towards higher Mg/Ca ratios and reaches its maximum at about the top ofthe lower third or the lower half of a core. Only in core E is the development of the curve more irregular. The broad maximum occurring in the other curves is split into two less extensive maxima by a minimum. From core A to core F, the Mg/Ca curves shift continuously from lower to higher Mg/Ca ratios. Again core E behaves differently: its curve crosses the curves of both core C and F irregularly. UJ Sr/Ca 1 000 (atomic) ratio. The Sr/CalOOO ratio varies between 1 ·60 and 3·97. Within a core no general distribution trend is to be observed. Laterally, a general increase from core A to core F may be recognized. (k) NajK (atomic) ratio. The Na/K ratio varies between 3·66 and 7·25. Within a core no common trend of variation can be observed, but generally, the lower parts of a core exhibit somewhat lower ratios than the upper half (exception: core A). From core A to core C, the ratio shows a slight increase. In cores E and F, the ratios are again in the range of core B . (l) Na/Cl (atomic) ratio. This ratio, which permits the calculation o f the portion of Na connected with Cl to form NaCl, varies between 2·02 and 3·69. Within a core, the ratio generally increases with increasing depth. No general trend is to be observed laterally. The results obtained may be summarized as follows. ( 1 ) Vertical distribution (from top to bottom); (a) Na, K, and Cl and the Na/Cl ratio decreases; (b) Mg increases; (c) Ca, Sr, HC03 and the Mg/Ca ratio have minima in the top and bottom layers with a broad maximum occurring towards the middle of the core; (d) Sr and the Sr/CalOOO ratio do not seem to be dependent upon the location in the core. (2) Lateral distribution (from core A to core F): (a) Mg, Sr, HC03 and the Na/Cl ratio do not change significantly; (b) Ca decreases; (c) the Mg/Ca ratio, the Sr/Cal OOO ratio, and Na, K, and Cl increase.
D EV E LO P M E NT O F C A RB O N A T E M I N E R A L O G Y , C HEMI STRY, G EO C H E M I S TRY AND I SOTOPE G EOCHEMISTRY
Core E
The results of mineralogical, chemical and geochemical investigations of core E, including isotope data from Linz ( 1 976) are presented in Tables 3 and 4 and depicted in Fig. 7. It should be mentioned, that the chemical data were obtained from the total sediment, whereas the isotope measurements refer to the < 2 /lm fraction of the sediment only, in which non-detrital carbonates are about the only carbonate minerals. Detritic carbonates occur only in trace amounts (Fig. 5).The original chemical MgC03- and Sr determinations were recalculated on the basis of 1 00% carbonates. Calcite precipitation commences above the gastropod layer with a 4 mol per cent MgC03-calcite which is poor in Sr ( < 500 ppm) and low in 8180 (-3°/00). Within the
LITHOLOGY
TIME
MINERALOGY
'lbMgC03 MOLE% CARBONATES • 100% MgC03 IN Mg CALCITE 3 10 20 30 40 19% 7 15 11 90%
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German Muller and Frank Wagner
72
Table 3. Mineralogical and chemical properties of sediments of core E, Lake Balaton
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54 51 53 52 59 68 70 71 73 74 74 73 72 72 72 70 66 67 69 71 71 70 68 68 68 69 70 69 68 63 60 57 54 52 53 56 56 63 63 64 61 56 60 61 64 57 57 53
39 35 37 37 45 57 60 61 64 65 66 64 63 63 63 60 55 56 59 63 61 60 57 57 57 59 60 59 57 51 47 43 39 36 37 41 41 51 51 38 25 20 23 24 33 43 43 37
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8 8 9 9 9 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 9 10 10 12 II II 12 13 14 12 II 13 12 15 20 18 19 15 13 13 14 12
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0.94 0.9 1 1 .08 1 .07 0.97 1.01 1 .03 0.93 0.91 0.96 0.97 1 .07 1 .00 1 .09 1 . 02 1 .03 0.95 1 .00 1 .00 1 .05 1. 1 1 1 .05 1.17 1.11 1 . 12 1 . 09 1.11 1 .29 1 .24 1 .26 1 .22 1.15 1 .23 1 .29 1 .3 6 1 .35 1 . 17 1 . 38 1 .42 1 .44 1 .59 1 .47 1 .66 1 .39 1 .53 1 .2 8 1 .28 1 .20
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4 6 6 4 6 3 3 4 4 4 4 3 6 4 3 6 3 4 4 5 4 3 5 6 5 4 6 7 6 4 6 6 7 7 7 8 6 7 5 8 8 7 7 7 8 7 7 7
Holocene carbonate evolution in Lake Balaton
73
Table 3 continued
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41 47 40 47 48 45 45 47 44 49 53 52 53 49 48 34 23 22 25 35 36 26 29 25 24 36 32 36 33 35 36 30 27 34 36 33 44 32 26 23 25 31 40 62 28 0 0
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II 10 II 10 10 II II 10 II 10 9 9 9 10 10 10 9 10 10 II 9 II II 12 12 10 10 10 10 10 10 II II 10 10 II II 13 14 14 14 14 II 7 14 18 18
4 3 4 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 4 4 3 3 3 3 3 3 3 3 3 3 3 3 4 4 5 5 4 4 2 4 5 6
13 13 12 12 II 13 13 14 12 10 II 12 13 14 19 17 16 13 15 15 12 18 13 14 12 13 13 12 13 II 12 12 8 8 9 10 7 4 4 4 4 4 4 4 0 0 0
22 22 23 18 18 21 19 22 22 22 20 21 22 24 28 33 35 35 30 26 27 28 27 27 27 24 25 27 25 25 25 24 26 22 21 21 18 14 15 14 14 II 9 4 9 18 23
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1 90 1 80 1 60 210 1 50 1 90 1 80 1 90 1 90 1 90 200 210 200 210 230 290 380 300 290 240 250 230 220 240 220 200 250 240 240 250 220 250 270 230 250 230 220 200 1 90 1 80 1 60 240 270 630 340 1 60 100
7 6 6 6 6 4 4 5 9 6 6 6 6 7 7 7 10 10 7 8 9 12 12 12 12 II 12 12 12 12 II II 12 12 II 8 12 6 10 10 10 8 7 2 7 8 8
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German Muller and Frank Wagner
74
Table 4. 8"0 values in autochthonous carbonates of the < 2 fLm fraction of sediments from cores E and F
(only sample at 1 7 9 em). After Linz, 1 976 Depth within sediment (em) E E E E E E E E E E F
0-1 31 53 84 86 101 1 13 130 133 151 1 79
8"0 in °/00 PDB -2.13 - 1 .88 - 1.19 + 0.36 + 0.22 - 0.38 - 0.98 + 0.77 - 0. 1 5 - 1 .02 - 3 .05
± 0.08 ± 0. 1 5 ± 0. 1 4 ± 0. 1 2 ± 0.26 ± 0. 1 1 ± 0. 1 4 ± 0. 1 8 ± 0. 1 2 ± 0. 1 9 ± 0. 1 5
profile, the different parameters fluctuate systematically and a positive correlation exists between mineralogical, geochemical and isotope data. Higher MgC03 concentrations in the calcite lattice are connected with the occurrence of cation disordered protodolomite, high Sr concentrations and relatively high 8180 values. Amongst the other elements investigated (K, Na, Li, Mn, Pb, Zn, Cu, Ag), the acid soluble amounts of Na (incorporated into the carbonate lattice and not bound to chloride) and Li (adsorbed on the surface of clay minerals) show a positive correlation with the MgC03 concentration of the sediments which in turn depends strongly on the MgC03 contents of the non-detrital carbonates (Fig. 7). Figure 8 shows the relationships MgC03/Na and Li/Na in the sediments of core E. There is no correlation between MgC03 and HCl-soluble K. Two well-defined Mg, Sr, Li, Na, and 180 maxima occur: one at the top of the lower third of the core - in the pinkish-greyish mud ('Rosa Zone') at a depth of about 1 30 em (hereafter called 'Mg 1') - and the other at a depth of about 80 em (hereafter called 'Mg 2'). The fluctuation of the MgC03 content of calcite ceases at about 50 em below the sediment surface. The uppermost sediment section exhibits a very constant rate of 8 mol per cent MgC03 in solid solution in the calcite lattice. A trend towards slightly higher MgC03 concentrations (hereafter called 'Mg 3') occurs in the top 1 0 em of the core. Other cores
The strongly positive correlation between chemical and mineralogical parameters which was observed in core E holds true for all other cores examined. A comparison of the different cores within the lake basin can therefore be made using only two main properties: (a) the composition of the high-magnesian calcite, and (b) the occurrence of cation-disordered protodolomite. Fig. 9 depicts the results of such a comparison, which is summarized as follows. (a) MgC03 incorporation into the calcite lattice: the amount of MgC03 in solid solution increases generally from core A to core F. Within a core, three (cores D, E, F), two (cores B, C) or one (core A) distinct Mg maxima occur, of which the lowermost - Mg 1 - can be traced over the full extension of the lake basin. The second maximum, Mg 2, is to be found in all cores with the exception of core A. The MgC03 contents are generally lower than those of the first maximum. A third
75
Holocene carbonate evolution in Lake Balaton % M g C03 26 22 18
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Fig. 8. MgCOjNa and Li/Na relationships in sediments of core E.
maximum, Mg 3, with MgC03 concentrations much lower than those of the other two maxima, is clearly present in cores D, E, and F. (b) Occurrence of cation-disordered protodolomite. This mineral is definitely connected with MgC03 maxima Mg 1 and Mg 2. In the easterly cores D, E, and F, two protodolomite horizons, D 1 and D 2 are associated with Mg 1 and Mg 2, whereas in core C, only the lower protodolomite zone D 1 is to be found. Cores A and B do not contain protodolomite. Age relationships
Palynological studies carried out by Z6lyomi ( 1 953) on several sedimentary cores from Lake Balaton revealed that the beginning of the authigenic carbonate sedimentation can be sharply defined as the beginning of the Atlanticum. The strata underlying the carbonate mud sequence are characterized by typical Boreal pollen spectra. The carbonate mud sequence of the lake thus comprises 7000-8000 years. If Z6lyomi's absolute time scale based on results of the pollen zonation encountered in the core 'Balaton V' is transferred to core E of our investigations, the time scale as shown in Fig. 7 can be applied.
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Fig. 9. Rate of MgCO, incorporation in primary calcite and zones of cation-disordered protodolomite (Mg l-Mg3) in cores A-F.
Holocene carbonate evolution in Lake Balaton
77
Two absolute datings carried out on carbonate muds ( < 2 Jlm fraction, practically free of detrital calcite and dolomite Fig. 5) of the 'Rosa Zone', Mg 1 ( 1 26 - 133 em) and the Mg 2 maximum (82-88 em) revealed ages of6640 and 4500 years, respectively. The 'true' ages may be considerably younger, since a 'hard water effect', in freshwater carbonates is capable of falsifying the ages by + 1300-3000 years. If 1 500 years are assumed for such a 'hard water effect', the radiocarbon ages would fit perfectly into the time scale depicted in Fig. 7 . There is no doubt, however, that the Mg 1 maximum developed during the Atlanticum and the Mg 2 maximum during the Sub-Boreal.
DISCUS SION OF RESULTS AND CONCLUSIONS
If the major events of carbonate evolution illustrated in Fig. 9 are redrawn on a time scale independent of the thickness of the different cores, the model of Fig. 10 results. In addition to the vertical chemical and mineralogical differentiation of the sequence, a pronounced lateral differentiation exists as well. From core A (closest to the major inflow of the lake, the Zala River) to core F (in the easternmost part of the lake) a gradient of increasing Mg, Na, Sr, and Li concentration in the sediments is observed. The results of the examination of the interstitial water chemistry lead to a similar conclusion. With the exception of sulphate, whose present irregular distribution is the result of the activity of sulphate-reducing bacteria, all other ions examined in this study (and some ratios derived therefrom) show a pronounced vertical and lateral distribution pattern which can generally be related to the vertical and lateral chemical differentiations observed within the sedimentary carbonates. It should be kept in mind, however, that the present pore solutions do not represent the original composition of the lake water at the time when it was included into the newly formed sediment as interstitial water. Mineral-water reactions, diffusion and, to a much greater extent, an upward movement of the pore solution due to sediment compaction must have led to changes in the original chemistry of the solutions. Thus one cannot expect that short-term fluctuations in lake water chemistry - which are well reflected in the sediment's geochemistry - will be preserved in the pore water chemistry. The same lateral gradient as observed in the Holocene carbonate sediments and their interstitial solutions, although less pronounced, exists in the MgC03 contents of the contemporaneous primary carbonate deposits of the lake (Fig. 3), and it seems only logical to assume the same mechanisms for the Holocene carbonate sediments that are in effect today: (a) an increase of the concentration of all ions (with the exception of CaH and HC03 - ) along the long axis of the lake as a result of evaporative dominance over inflow, and (b) the precipitation of calcite (with some MgC03 incorporation) depletes the water in Ca and HC03- and at the same time increases the Mg/Ca ratio. If such changes occur in an open basin (such as the lake represents today), a magnifying effect may be expected at times when the lake had no outlet (closed basin) as was the case before the Romans built an artificial outflow. At least three old terraces ( 140- 1 50 em, 1 80-250 em, and 300-400 em) above the present water level (Keresmaros, 1 939) were witness to this period which hereafter is designated as the 'pre-roman stage' in contrast to the 'post-roman stage' of the lake development during which the lake can be considered to be an open system.
78
German Muller and Frank Wagner 8
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C
D
E
F
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Increase of Mg , S r , N a , L i
Water
Mg/Ca ratio
i n Sediments
Fig. 10. Lateral occurrence of the Mg maxima (Mg l -Mg3) and cation-disordered protodolomite horizons
(D 1 - 02) in Lake Balaton sediments.
During the very dry year of 1 949 the water level of the lake fell to a point where no outflow was reported at the gates of Si6fok. During this very dry period the total electrolyte concentration of the lake water increased considerably (Sebestyen, Entz & Feldoldy, 195 1 ; Sebestyen, 1 962). We assume that the weak Mg 3 maximum in the uppermost 1 0 em of the sediment represents a period in which the outflow was much less than it is today or was even zero during the warmer seasons. The Mg/Ca ratio of the lake or interstitial water not only determines the level of MgC03 incorporation into the calcite lattice thus forming high-magnesian calcites but is also essential for the diagenetic transformation of high-magnesian calcite into protodolomite. In experiments, Winland ( 1 969) found the distribution coefficient for partitioning magnesium between calcite and an aqueous phase at 20'C to be about 0·02 if the precipitate was in equilibrium with the final solution and the Mg/Ca ratio of the solution was kept as near to constant as possible during precipitation (homogeneous or Henderson-Kracek distribution). In a recent study, Fuchtbauer & Hardie ( 1 976) found the following mean homogeneous distribution coefficients for precipitated magnesian calcites at thre� different temperatures: Temp.
50'C
28'C
13 ·5'C
K'
0·042 ± 0·005
0·029 ± 0·004
0·02 1 ± 0·003
The coefficients apply for high-magnesian calcite composition up to about 20 mol per cent MgC03 and solution Mg/Ca mole ratios below about 7-8.
Holocene carbonate evolution in Lake Balaton
79
The graphical representation of the homogeneous distribution coefficient in Fig. 1 1 summarizes the interrelationship between Mg and Ca in solution and in precipitated carbonates. The above distribution coefficients compare quite well with the high magnesian calcite composition and the Mg/Ca ratio of the interstitial solutions of Lake Balaton sediments if a temperature between 28' and 30'C can be assumed (a value which in fact is often exceeded on summer days). The maximum MgC03 concentrations (20 mol per cent) coincide with maximum Mg/Ca ratios in: the pore solutions (8·5).
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8
28 °C
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15 10 MOLE % M g C03 IN CALCITE
20
Fig. 1 1 . Graphical representation of the relationship between •Mg/'Ca in solution and level of MgC03 incorporation in calcite at different temperatures based on mean homogeneous distribution coefficients experimentally determined by Fiichtbauer & Hardie ( 1 976).
According to Muller et a!. ( 1 972) a minimum Mg/Ca ratio of about 7 is required to form protodolomite from high-magnesian calcite. We therefore assume that protodolomite began to develop when the interstitial solution surpassed this critical ratio. The fact that inorganic aragonite does not occur within the Holocene authigenic carbonates is a sign that the Mg/Ca ratio did not reach a value of about 1 2. In lakes, aragonite only precipitates inorganically when this ratio is exceeded (Muller et a!., 1 972). An example where dolomitization is taking place in contemporaneous sediments is the Neusiedler See, a steppe lake a few hundred kilometres NW of Lake Balaton near the Hungarian-Austrian border (Schroll & Wieden, 1 94 1 ; Blohm, 1974). The Mg/Ca ratio in an aquatic system, which in a humid climate is low, can be increased by preferentially extracting Ca from the solution through CaC03 precipitation. If river water is continuously added to a closed lake basin and CaC03 precipitation takes place, the Mg/Ca ratio of the water and of the precipitating calcite is governed by the ratio of evaporation to inflow plus meteoric precipitation. This ratio also determines the total concentration of other ions in solution, and those having a geochemical
80
German Muller and Frank Wagner
affinity for Ca-Mg carbonates (Sr: Kinsman, 1 969; Na: Muller & Billings, 1 965; Kitano, Okomura & Idogaki, 1 975; White, 1 977) or for clay minerals (Li), which tend to be enriched in the precipitating carbonates or with detrital particles respectively. Since an increase of evaporation without a corresponding increase of precipitation results in a shrinking of the water body, extreme levels of MgC03 incorporation into the calcite lattice (together with Sr, Na and Li maxima) are indicative of lower water levels. The fluctuations in the Mg, Sr, Na and Li concentrations in the Lake Balaton sediments are thus an expression of the ratio of evaporation to precipitation, i.e. ofthe relative 'dryness' of the Balaton area during the past 7000-8000 years. According to our results, the periods around 5000 and 3000 years B .P. were the driest. The high 8180 values in the non-detrital carbonates as well as the other findings support this assumption. During evaporation 1 80 is enriched in the water phase while at the same time the water vapour is depleted in this isotope. High rates of evaporation therefore lead to an enrichment of 180 in the water and also in the carbonates precipitating therefrom. The pollen spectrum from the Balaton sediments (Zolyomi, 1 953) does not reflect the two 'dryness' maxima derived from the sediment data but shows a continuous transition from the warm, semi-humid climate of the Atlanticum to the still warm, but rather dry climate ofthe Sub-Boreal period and finally to the present-day conditions of the Subatlanticum period. In our opinion, this discrepancy does not argue against the sedimentary method but only reflects the slowness of the vegetation to adapt to short term climatic changes and shows that vegetation is much less sensitive than water chemistry and carbonates.
ACKNOWLEDGMENTS
We wish to thank our colleagues from the Biological Institute of the Hungarian Academy of Sciences, Tihany, for providing their research vessel and for technical assistance. Thanks also to Drs H. P. Eugster and K. Kelts for critically reading the manuscript and for their many useful comments. G. P. Wee assisted in the English version of the manuscript and U. Kastner prepared the illustrations. This project was financially supported by the Deutsche Forschungsgemeinschaft.
REFERENCES BERNER, R .A. ( 1 97 1 ) Principles of Chemical Sedimentology. McGraw-Hill, New York.
BIDL6, G. ( 1 960) Balatoni Aragonit-Kivalas. Bull. Hung. geo/. Soc. 90, 224-225. BLOHM, M. ( 1 974) Sedimentpetrographische Untersuchungen am Neusiedler See/Osterreich. Inaug. Diss. Univ. Heidelberg. ENTZ, H. & SEBESTYEN, 0. ( 1 964) Das Leben des B alatonsees. Magyar Bioi. Kut. Munk. (Publ. Hung. Bioi. Res. Inst.) 16, 1 79-4 1 1 . ENTZ, B . ( l 959) Chemische Charakterisierung der Gewasser in der Umgebung des Balatonsees (Plattensee) und chemische Verhaltnisse des B alatonwassers. Annal. Bioi. Tihany, 26, 1 3 1-20 1 . FELFOLDY, L., M USZKALAY, L., RAKOCZI, L . & SZEDZTAY, K . ( 1 960) Origin and movement o f sediment in Lake Bala ton. Mitt. in/. Verein. Limnol. 17, 282-29 1 . FOCHTBAUER, H . & HARDIE, L.A. ( 1 976) Experimentally determined homogeneous distribution coefficients for precipitated magnesian calcites: Application to marine carbonate cements. A bs. Prog. geol. Soc. Am. Meetings, 8, 877 .
Holocene carbonate evolution in Lake Balaton
81
GAINES, A.M. ( 1 977) Protodolomite redefined. J. sedim. Petrol. 47, 543-546.
KERESMAROS, J. ( 1 939) A keszthelyi halomgerine balatoni szinl6. Foldrajzi Kozlemenyek, 67. KINSMAN, D.J.J. ( 1 969) Interpretation of Sr concentrations in carbonate minerals and rocks. J. sedim.
Petrol. 39, 486-508. KITANO, Y., OKOMURA, M. & IDOGAKI, M. ( 1 975) Incorporation of sodium, chloride and sulphate with
calcium carbonate. Geochem. J. 9, 75-84. LINZ, E. ( 1 976) Zur Geochemie stabiler C- und 0. Isotope in nichtmarinen Karbonaten und Karbonatgestein en. lnaug. Diss. Univ. Heidelberg. MOLLER, G. & BILLINGS, G.K. ( 1 965) Beziehungen zwischen dem Natrium-Gehalt biogener Karbonate und der SaliniHit im Bildungsraum. Abst. 55. Jahresvers. Strasbourg 4. -6. Miirz, 1 965, I I . MOLLER, G . ( 1 969) Sedimentbildung im Plattensee/Ungarn. Natunvissenschaften, 56, 606-6 1 5 . MOLLER, G. ( 1 970) High-magnesian calcite and protodolomite i n Lake B alaton (Hungary) sediments. Nature, 226, 749-750. M OLLER, G. ( 1 97 1 ) Aragonite inorganic precipitation in a freshwater lake. Nature, 229, 1 8 . MULLER, G . , IRION, G . & FORSTNER, U . ( 1 972) Formation and diagenesis o finorganic Ca-Mg-carbonates in the lacustrine environment. Naturwissenschaften, 59, 1 5 8- 1 64. PAszT6, P. ( 1 963) A B alaton vizmin6segem!k vizsgalata. VITUKI Rept, 1 1 , 1 - 1 25 (in Hungarian). PONY!, J.E. TUSNADI, G ., V ANGER, E. & RICHNOVSZKY, A ( 1 974) Investigation with computer ICL system 4 on the morphometry and composition of the population of Dreissena shells from the upper sediment layer of Lake B alaton. A nnal. Bioi. Tihany, 41, 2 1 7-234. RONAl, A. ( 1 969) The geology of Lake B alaton and surroundings. Mitt. int. Verein. Limnol. 17, 275-28 1 . RoTHE, P., HOEFS, J. & SONNE, V . ( 1 974) The isotopic composition o f Tertiary carbonates from the Mainz Basin: an example of isotopic fractionations in 'closed basins'. Sedimentology, 21, 373-395. ScHROLL, E. & WIEDEN, P. ( 1 97 1 ) Eine rezente Bildung von Dolomit im Schlamm des Neusiedlersees. Tscherm. miner. petrogr. Mitt. 7, 286-289. SEBESTYEN 0., ENTZ, B. & FELDOLDY, L. ( 1 9 5 1 ) Alacsong vizallassal kapcsolatos biologia jelensegekr61 a Balatonon 1 949 6szen. SEBESTYEN, 0. ( ! 962) Ergebnisse der B alaton-Forschung der letzten fiinfzehn Jahre 1 946-1 960. Annal. Bioi. Tihany, 29, 2 1 7-266. STOFFERS, P. ( 1 975) Die Rekonstruktion paliioklimatischer Verhiiltnisse am Beispiel ostafrikanischer Seen. Ruperta Carola (Heidelberg), 55, 8 1 -86. STOFFERS, P. & H EC K Y, R.E. ( 1 978) Late Pleistocene-Holocene evolution ofthe Kivu-Tanganyika Basin. In: Modern and A ncient Lake Sediments (Ed. by A. Matter and M. E Tucker). Spec. Pubis int. Ass. Sediment. 2, 43-55. SzESTAY, K. ( 1 967) Some hydrologic data of Lake Balaton. In: Study Tours, int. Symp. Paleolimnol. Bioi. Res. Institute Hungarian A cad. Sciences, Tihany, 26-28. TAMAS, G. ( 1 972) Horizontal phytoplankton studies in Lake Balaton based on scooped samples and filtrates taken in 1 967. Annal. Biol. Tihany, 39, 1 5 1 - 1 88. TOTH, L. ( 1 967) Higher water plants. In: Study Tours, int. Symp. Paleolimnol. Biol. Res. Institute Hungarian A cad. Sciences, Tihany, 40-4 1 . WHITE A. F . ( 1 977) Sodium and potassium coprecipitation i n aragonite. Geochim. Cosmochim. A cta, 41, 6 1 3-625. WINLAND, H.D. ( 1 969) Stability of calcium carbonate polymorphs in warm, shallow seawater. J. sedim. Petrol. 39, 1 579- 1 5 87. ZoLYOM!, B . ( 1 953) Die Entwicklungsgeschichte der Vegetation U ngarns seit dem letzten Interglazial. Acta Bioi. Hung. IV, 367-4 1 3 .
Spec. Pubis int. Ass. Sediment. ( 1 978) 2, 83- 107
Permian Saar-Nahe Basin and Recent Lake Constance (Germany): two environments of lacustrine algal carbonates
A N D R E A S S C H A F E R and K A R L R . G . S T A P F Geologisches lnstitut der Universitat, Nussal!ee 8, D-5300 Bonn, and Institut fur Geowissenschaften der Universitat, Saarstrasse 21, D-6500 Mainz, West Germany
ABSTRACT Algal carbonates of a Permian and Recent lacustrine environment are described and their occurrence discussed. The stromatolites of the Permian Saar-Nahe Basin are diverse in shape, forming small domes, mats and oncolites. Their scale ranges from millimetres to metres. Embedded in limestone, mudstone and muddy sandstone, the algal carbonates are locally restricted to distinct horizons. The algal origin is demonstrated by the gross fabrics since there is no evidence for algal filaments within the stromatolite microstructures. In Recent Lake Constance algal carbonates exhibit fabrics which can be related to different blue-green algae. These carbonates mainly occur in the U ntersee of Lake Constance and are restricted to its shallow-water areas. The algal nodules, 0·5-30 em in size, possess both smooth and spongy surfaces. The smooth surfaces are formed by Schizothrix, the spongy surfaces by Phormidium, Ca/othrix and/or Dichothrix; occasionally colonies of Rivu/aria are intercalated. The smooth nodules are found close to the shoreline of the lake. The spongy nodules occur in loose carbonate sands which are influenced by the Rhine river current. Both types are covered by water for most of the year.
INTRODUCTION
Although most algal carbonates have been described and classified from marine environments (Walter, 1 976; Fliigel, 1 977; Wray, 1 977), Kalkowsky's early descrip tions ( 1 908) are based on lacustrine stromatolites and oncolites. The paucity of ancient lacustrine stromatolites (Golubic, 1 973), largely reflects a lack of recognition and
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
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preservation of lacustrine facies in the geological record (Picard & High, 1972). The oldest known lacustrine stromatolites are from the Precambrian of the Canadian Shield (Hoffman, 1 976). Other occurrences of lacustrine stromatolites include the British Old Red Sandstone, Devonian (Fannin. 1 969). the Permian Saar-Nahe Basin, Germany (Stapf, 1 973; Wagner & Lamprecht. 1 974, and this paper), Upper Cretaceous, Eocene and Oligocene of southern France (Freytet & Plaziat, 1965; Colin & Vachard, 1977; Donsimoni & Giot, 1 977), the Tertiary of the Rhinegraben (Reis, 1923; Rutte, l 954a) and Bavaria (Rutte, 1954b) and the Eocene Green River Formation of U.S.A. (Bradley, 1929; Surdam & Wray, 1976). In addition, in this volume, lacustrine algal carbonates are described or mentioned in papers by Smoot ( 1 978), True ( 1 978), Link & Osborne ( 1 978) and Tucker ( 1 978). Recent lacustrine algal limestones have been described from various localities including lakes north of the Austrian Alps (Kann, 1 94 1 : Schneider, 1 977), Lake Constance, southwestern Germany (Baumann, 1 9 1 1 ; Schottle, 1 969), the Green Lake, New York (Eggleston & Dean, 1 976) and the freshwater marshes of Florida and the Bahamas (Monty & Hardie, 1 976). Algal limestones forming in rivers have been described by Fritsch ( 1 950) from Britain, and by Golubic & Fisher ( 1 975) from Pennsylvania, U.S.A. In this paper, fossil stromatolites and oncolites from the Permian lacustrine beds of the Saar-Nahe Basin are described and compared with Recent oncolites from Lake Constance.
SAAR-NAHE BASIN Geological context and sedimentation
At the end of Lower Carboniferous the palaeogeography of Central Europe consisted of a series of basins and mountain ranges, generally trending southwest northeast. Weathering and erosion of highland areas provided detritus which through alluvial, fluvial and lacustrine processes constituted the continental deposits ofUpper Carboniferous and Lower Permian. The Saar-Nahe Basin (Fig. I) was part of an extensive trough-system along the southern border of the Rheinisches Schiefergebirge. It was filled while continuously subsiding, so that a sequence more than 9000 m thick accumulated. Today the Saar-Nahe Basin measures about 90 km in length and 40 km in width, and is bounded by Devonian strata along its northern edge. To the south it is covered by the Triassic Buntsandstein and towards the east by Tertiary sediments of the Rhinegraben and Mainz Basin. Sedimentation in the Saar-Nahe Basin began in the Upper Carboniferous with fluviatile and lacustrine beds, of sandstones, lenticular and bedded conglomerates and mudstones (Falke & Kneuper, 1 972). Associated lacustrine coal measures indicate rich organic productivity. The Upper Carboniferous sedimentation pattern continued into the Permian (Lower Rotliegend) which is here defined by the occurrence of the fern Callipteris conferta Sternberg (Falke, 1 976). Falke ( 1954) recognised that the sequence consisted of fining-upward fluvio-lacustrine cycles which have been used as the basis of the stratigraphy (Fig. 2). Each cycle begins with red conglomeratic sandstone which is immediately succeeded by finer sandstones and muddy sandstones. This part of the cycles represents a fluviatile influx into the basin, of sediment derived from the south.
Permian Saar-Nahe Basin 0 ,----,---, § [2J [;2:] � � ffi Bonn
85
Terltary.Recent Buntsondsletn U-Rot!iegend Mag mottles L - Rothegend Carbontferous oevon•on
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Fig. l. Permian Saar-Nahe Basin and Recent Lake Constance, an ancient and a modern environment of lacustrine algal carbonates. In the Saar-N ahe Basin algal carbonates occur within the Rotliegend sediments of the entire basin and in the Untersee of Lake Constance on colites are present at selected nearshore areas only (stippled).
The cycles then show a gradual transition into mudstones, 'paper-shales', micritic carbonates and coal, which accumulated in lacustrine environments. The carbonates and rhythmically laminated 'paper-shales' are horizons of considerable stratigraphic value, and contain well preserved faunas of fish and amphibians but few benthonic organisms (Boy, 1 977). Fluvio-deltaic sediments rapidly infilled the lake basins and terminated the lacustrine conditions (Rast & Schafer, 1 977). With the beginning of the Upper Rotliegend the pattern of sedimentation changed. Clastics of the N ahe Group are coarse red arkoses of braided stream and alluvial fan origin, locally interbedded with perennial playa lake carbonates. Lavas and tuffs were also extruded during this time (Atzbach & Schwab, 197 1). Recent studies of amphibian footprints (Boy, personal communication) and plant fossils (Boersma, 1 975; Visscher, Huddleton Slater-Offerhaus & Wong, 1 974) indicate an Autunian age for the whole Rotliegend sequence of the Saar-Nahe Basin.
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Carbonate sediments
Within the dominantly clastic sequence of the Saar-Nahe Basin, carbonates are rather rare. They are confined to a few specific horizons and are composed of low-Mg calcite, high-Mg calcite, dolomite, ankerite or siderite. The lacustrine carbonates are chiefly present in the Lower Rotliegend since quiet water conditions were more prenlent at that time. In the Kusel Group there are several carbonate beds which reach 2 m in thickness and in the Lebach Group there is only one carbonate horizon (0·7 5 m thick). In the Nahe Group carbonate horizons of playa-lake origin are around 0·5 m thick and intercalated with coarse alluvial-fan deposits. The carbonates are micritic if low in magnesium, but of a more sparitic texture, if the content of iron and magnesium increases. The content of organic matter is variable and strongly reflects the former oxygen-content of the lake water. In the carbonate rocks of the Kusel and Lebach Groups droplets of bitumen are seen in thin sections and at outcrop. The only organic remains which can be identified in the carbonates are oncolites and stromatolites (Fig. 3). They mostly form distinct layers within the carbonate rocks but in some cases wholly constitute them.
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Algal carbonates
The algal carbonates from the Saar-Nahe Basin have been described by von Gumbel ( 1 894), Reis ( 1903), von Ammon ( 1 9 1 0) and Habicht ( 1 953). In more recent investigations Stapf ( 1 973, 1974) classified the various algal growth forms according to the work of Truswell & Eriksson ( 1 972), while Wagner & Lamprecht ( 1 974) referred to Logan, Rezak & Ginsburg ( 1 964). Truswell & Eriksson ( 1 972) recognised a domical, tabular and an encapsulating type of stromatolite, all of which are found in the Saar Nahe Basin. The domical stromatolites (Fig. 4) form algal mounds, mostly symmetrical in shape and attached to a substrate of oncolites. The diameter of the domes ranges from I to 20 em; their height is up to 10 em. The algal mounds may occur singly or in great numbers. Oncolites are present between mounds and commonly increase in size upwards through the bed. The internal structure of the domical stromatolites is columnar. The columns are some millimetres in diameter and some centimetres in length. They are oriented at various angles, and may branch. The columns themselves contain wavy and crinkled laminae, which in all cases are convex towards their direction of growth. The wavy and crinkled laminae and the interspaces between them are reflections of the outer scarred surface of the entire algal mound. The tabular stromatolites (Fig. 5) form laminated algal mats which may be attached to oncolitic sediment or micritic limestone. Their thickness can reach 5 em, their lateral extent several square metres. Internally the tabular stromatolites are laminated. Columns are present as in the domical mounds, but they are only several millimetres in length. Small oncolites are present within the laminae, and together with the columnar structures give a pustular texture to the surface. The encapsulating stromatolites or oncolites (Fig. 6) are very variable in shape and size. They range from several millimetres in diameter with a spherical shape to a size of 100 x 70 x 30 em with an elongate shape. Small roundish oncolites (0·5-5 mm in diameter) are abundant and have commonly formed around detrital grains. Beds of oncolites 0·5 to 10 em in thickness are commonly inversely graded and may be associated with micritic or detrital limestones, or marly siltstones. Some large elongate oncolites encapsulated logs of wood which have since decomposed. The surface structures of the wood are visible on the inside of the algal crusts. The latter have commonly been compacted or the cavities formerly occupied by the wood filled by calcite and sediment. The small oncolites possess both wavy and crinkled laminae. Internally, the large elongate oncolites consist of columns 5-l 0 mm wide and 3-5 em in length. The columns are radially arranged with respect to the nucleus only on the upper side.
Petrography of the algal carbonates
Microscopic examination of the stromatolites and oncolites shows an intense sparitization of the microstructures. Only the overall organisation of the columns and the arrangement of their laminae are visible. Algal filaments are not preserved and may have been lost through oxictation or diagenetic neomorphism. Carbonate and non-carbonate lithoclasts, pellets and bioclasts are contained within the sparry carbonate matrix of the algal structures. The latter are composed of high-Mg calcite, dolomite or ankerite.
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Environmental implications
Stromatolites and oncolites are mostly found associated with limestones and marly
Fig. 3. Typical algal limestone from the Saar-Nahe Basin showing growth of domical stromatolites on a substrate of oncolites. Biogenic carbonates are light in colour; the surrounding calcareous mudstone is dark, organic-rich with a pitted weathered surface.
Fig. 4. Stromatolite from the Saar-Nahe Basin showing domical mounds resting on a substrate ofoncolites. The internal structure of an algal mound in an early growth stage is seen on the vertical polished section.
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Fig. 5. Tabular stromatolite from the Saar-Nahe Basin which has grown on a micritic limestone. The mat consists of columns with occasional laminations.
Fig. 6. An encapsulating stromatolite or oncolite from the Saar-Nahe Basin. The asymmetry, due to a lack of reworking during growth, is shown by columns on the upper side and planar laminae on the underside.
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mudstones (Fig. 7) and never with coarse clastics or 'paper shales'. Only the large elongate oncolites occur in sandstones. Associated sedimentary structures include parallel lamination, ripple marks and desiccation cracks. indicating a shallow-water to emergent nearshore situation. Reworking of the algal carbonates did not take place. In the Lower Kusel Group, growth of algal mounds was commonly terminated by volcanic ash.
LAKE CON STANCE Environmental setting
With a total area of 539 km2 , Lake Constance (Bodensee) is divided into the larger Obersee with an area of 476 km2 and the smaller Untersee with an area of63 km2 (Fig. 1). The maximum water depth of the Obersee is 252 m, while the Untersee is 45 m in the deep southern basin. Lake Constance is situated 395 m above sea level. The geological history of Lake Constance started during the Pleistocene (Schreiner, 1975). The lake basin was excavated by glaciers out of the Tertiary Alpine Molasse to a depth of about 500 m along some pre-existing NW-SE trending faults in the basement. Later the basin was filled with sediments and separated by glacio-fluvial activity into the separate lakes. The fill of the Obersee basin consists of coarse-grained glacio-fluvial till, overlain by postglacial to Recent fine-grained lacustrine sediments (Miiller & Gees, 1 970). Lake Constance is part of the Rhine River system which rises in the Swiss Alps. A large delta has formed where the Rhine enters the Obersee (Muller, 1 966). Suspended matter settles on the deep lake floor so that the Rhine leaving the Obersee carries little sediment into the Untersee. The Rhine leaves the Untersee at its western end and trends northward to the North Sea. Lake Constance hydrology depends on seasonal variations in water supply which is determined by rainfall and/or water retention in the Alps (for further hydrological and meteorological data see Kiefer, 1 972). Because of its greater depth the Obersee is a monomictic lake (the northernmost of Europe) with complete mixing in late autumn; the Untersee is a dimictic lake and its surface is occasionally frozen during cold winters. The water chemistry ofboth lakes differs only slightly. The Obersee's dissolved solids vary a little depending on the Rhine water discharge, whereas the Untersee is rather constant in chemical composition (HCO- 3 content is about 1 30 mg/ 1 , Ca2 + 45 mg/ 1 , MUller, 1 965). The Obersee's water is oligotrophic, while that of the Untersee is eutrophic (Elster et al. , 1 968). The sediments of the Obersee are mainly derived from the Rhine River which delivers both bed- and suspended load (Miiller, 1 97 1 ; Reineck, 1 974). Carbonate sedimentation has dominated the Untersee in postglacial times (Schafer, 1 972) and during the Atlanticum period lake marl (Seekreide; Schafer, 1 973) was precipitated by subaqueous plants such as Potamogeton and Chara in littoral zones. The Ermatingen Embayment
The Rhine River between the Obersee and the Untersee traverses Quaternary lake marls (with oncolites), local glacial sediments and modern lake deposits. Where the Rhine enters the Untersee, a large sandflat is developed, the area being called the Ermatingen Embayment. Normally it is covered by 0 · 5-1 ·5 m of water, into which the
Permian Saar-Nahe Basin
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Fig. 7. Two stromatolite mounds from the Saar-Nahe B asin which have grown on a thin limestone horizon within a sequence of laminated siltstones. A thin layer of volcanic ash overlies the stromatolites.
Rhine has cut a channel of 5-10 m depth. During the summer the sand flat is densely crowded with subaqueous plants, but during the winter the flat is emergent and free of plants. Sediments on the sand flat consist of muddy fine- to medium-grained sand with up to 75% carbonate content, of which 5-20% is dolomite. This detrital material is derived from the nearby Molasse and from glacial till. Sediment and water plant distributions strongly reflect the current pattern (Lang, 1 969; Schafer, 1 972). The vegetation in Lake Constance and especially that of the Ermatingen Embayment, frequently changes in diversity and density. For instance from 1 970 onwards the water plants Potamogeton pectinatus with minor amounts of Zanichiella palustris were replaced by a dense population of abundant Zanichiella palustris and Potamogeton pusillus. Rich populations of the green algae Spirogyra and Cladophora form thick-threaded masses now replacing former Characees. Algal carbonates of the Untersee
The algal carbonates ('Schneggli') have long been known from Lake Constance, especially from the Untersee. Baumann ( 1 9 1 1 ) mapped their distribution and made an attempt to classify them. Other studies of the algal carbonates include Schmidle ( 1 9 1 0, 1 9 1 1), Pia ( 1 933), Jaag ( 1 938), Kann ( 1 94 1 ) , Golubic ( 1 962, 1 973), Schottle & Muller ( 1 968), MUller ( 1 968), and Schottle ( 1 969). The algal carbonates are most common in the shallow-water zone at the inlet of the Rhine into the Untersee, the Ermatingen Embayment (Fig. 8), but they are also abundant at the outlet ofthe Untersee at Stein am Rhein. It is possible that the oncolite
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Fig. 8. Oncolites ('Schneggli') from the Ermatingen Embayment in the Untersee of Lake Constance in an area of loose sandy sediments and algal carbonate detritus. Different sizes of smooth and hard as well as soft and spongy oncolites are exposed at low water stages in late autumn. Fig. 9. Discoidal soft and spongy oncolite with attached plants and Dreissena-shells. The active blue-green algae are confined to the oncolite's upper surface exposed to daylight in the shallow water of the Untersee of Lake Constance.
Permian Saar-Nahe Basin
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formation at these two localities is related in some way to the Rhine current. Further algal carbonates are found on the eastward trending underwater ridge of the Hori Peninsula, at the southeastern end of the Mettnau Peninsula, at the southeastern part of the northern U ntersee basin and along both sides of the artificial dam which leads to the Reichenau Peninsula. The oncolites of the Ermatingen Embayment
The oncolites are 1 -30 em in diameter and show both a hard and smooth as well as a soft and spongy exterior. Their shape is mostly discoidal (Fig. 9). The oncolites increase in numbers and size towards the Rhine River current and tend to be larger where the current is stronger (up to 1 m/sec at high water levels).
Fig. 10. Cross sections of a soft and spongy oncolite (above) and a smooth and hard oncolite (below) from the U ntersee of Lake Constance, both showing their internal laminated structure. The upper oncolite is strongly asymmetric since the calcite laminae grew on a mussel valve which was clearly in a very stable position. The lower one is more symmetrical about its nucleus, as it was often overturned by water agitation.
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Large oncolites commonly have a slightly convex surface exposed to the open water and a conical subsurface embedded in the sandy substrate. They lie embedded in loose, rippled sediment, with their upcurrent sides free of sediment and their downcurrent sides covered by detritus and/or smaller oncolites. The bottom side of the larger oncolites and the sediment below is always bluish-black in colour as a result of reducing conditions, while the top side of the oncolites and the sediment around has light-yellowish to light-brownish colour (also noted by Kann, 1959, p. 1 74). The upper side of the oncolites may be partly covered by floating green algae, such as Cladophora, and the bivalve Dreissena may be attached. Algae and fabric
The oncolites of Lake Constance are of two different types, one smooth and hard, the other spongy and soft. The former are mostly piled up like pebbles along the shore,
Fig. 1 1 . Calcitized tubes of vertically and horizontally growing filaments, probably made by the blue-green alga Phormidium. Soft and spongy orrcolite, U ntersee of Lake Constance. Height of thin section is l -4 mm.
Permian Saar-Nahe Basin
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while the latter are embedded in the substrate below the water level. The fabric of the oncolites in most cases is laminated (cf. Monty, 1 976) (Fig. 1 0). Spongy and dense laminae, both of micritic calcite crystals, alternate irregularly and may contain detrital grains. Towards the interior of the nodules the laminae are more closely stacked. The nucleus of the oncolites may be a piece of wood, pebble, mineral grain, bivalve, gastropod (which gave rise to the local name 'Schneggli'), or a fragment of broken oncolite. The organization of the stacked laminae results from the growth of algae and the nature of their filaments. Oncolite microstructure can only with difficulty be related to the algae which formed them, since algal filaments decay quickly, soon after the micro-crusts formed around them. Generic names of the modern blue-green algae are given as a guide to the particular algal group which produces a certain structure in the oncolites. Three different types of algal texture can be distinguished: spongy, dense and fan-like. A spongy microstructure is produced by algae of the genus Phormidium. In thin section the algal filaments are seen as tubes which are encrusted with calcite crystals. Vertically-arranged filaments occur together with cross-cutting ones (Fig. l l ). A variable direction of filament growth of these algae is common (Golubic, 1 976) and is similar to that described by Monty ( 1 976, p. 207). Phormidium was recorded from Lake Constance by Baumann ( 1 9 1 1 ), Mattern ( 1 970) and Kann ( 1 973). The structure attributed to this group of algae occurs in both the smooth and hard, and soft and spongy oncolites. Another group of vertically growing algae can be compared with Calothrix and/or Dichothrix (Fig. 1 2). The filaments are tufted and occasionally intercalated with Phormidium, producing fabrics of a bush-like character. Calothrix was described from
Fig. 12. 'Bushes', probably of Calothrix/ Dichothrix-colonies, developed upon a possible Phormidium-1ayer.
Soft and spongy oncolite. Untersee of Lake Constance. Width of thin section is 1-4 mm.
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the Untersee by Schmidle ( 1 9 1 0), Baumann ( 1 9 1 1 ), Pia ( 1 933) and Jaag ( 1 93 8); Kann ( 1 94 1 ) noted Calothrix together with Dichothrix. Small detrital grains encrusted with precipitated carbonate may occur between the loosely-packed and diverging filaments of both Phormidium and Calothrix and/or Dichothrix (Fig. 1 3 ). Following decay of the algal filaments, hollow tubes are left within the micritic carbonate (Fig. 1 4).
Fig. 13. Scanning electron microscope (SEM) view of a vertical cut through part of an oncolite showing
carbonate encrustations around diverging filaments of blue-green algae of the type Phormidium-Calothrix / Dichothrix. An algal filament is seen in the lower right corner although filaments are normally absent through decay. Soft and spongy oncolite, Untersee of Lake Constance. Field of view is 4 1 0 p.m across.
Primary cavities within oncolites (fenestrae; Monty, 1 976), which mainly originate through the irregular growth of the oncolites are commonly lined by smooth algal layers and contain algal fragments and detrital grains. A dense microsttructure is produced by algae comparable with Schizothrix (Fig. 1 5) and consists of horizontal interwoven algal filaments, much smaller in size than those of Phormidium-Calothrix/Dichothrix. As the structure produced by Schizothrix is rather dense, algal filaments are rarely detectable in transmitted light and are best seen under incident light after etching. The Schizothrix layers occur within the oncolites as well as on the surface of the smooth and hard types. The latter occurrence indicates that Schizothrix was the last alga to colonize the oncolites. Initial algal growth on grains and mollusc shells is commonly by Schizothrix (Schmidle, 1 9 1 0) and it may also
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Fig. 14. SEM-view onto the surface of a soft and spongy oncolite. Dried algal filaments occur in tubes composed of micritic calcite. One filament still shows its top cell, the other is torn off and allows a view into its interior. Untersee of Lake Constance. Field of view is 40 p.m across.
bore into the carbonate substrate (Pia, 1 937; Golubic, Perkins & Lukas, 1 975; Schneider, 1 977; Kobluk & Risk, 1 977). The borings into shells (Fig. 1 6) or detrital carbonate fragments (Fig. 1 7) constitute micritic envelopes (Kobluk & Risk, 1 977). These borings are identical to Schizothrix filaments of the dense outer micritic laminae, in terms of size and branching. The intergrowth of Phormidium Calothrix/Dichothrix and Schizothrix in the Lake Constance oncolites is comparable to the intergrowth of Scytonema and Schizothrix reported by Monty ( 1 976). This intergrowth is shown by long thick algal filaments which have grown vertically through a thick felt of horizontally-woven and much thinner algal threads (Fig. 1 8). A fan-like microstructure is formed by Rivularia (Figs 1 9 and 20), comparable with R. haematites described by Monty ( 1 976) and Schneider ( 1 977). Rivularia is sparsely distributed throughout the oncolites and the associated fabric consists of large, elongate sparry calcite crystals (Fig. 2 1 ). Fritsch ( 1 950) and Golubic ( 1 973) have noted that Rivularia with its corallinaceous habit is usually found forming hard layers in the splash zone of lakes. In the Ermatingen Embayment, however, the oncolites are growing below mean water-level and this could account for the paucity of Rivularia colonies in the nodules (Kann 1 959; Schneider, personal communication).
98
A. Schafer and K. R. G. Stapf
Fig. 15. Dense structure probably made bySchizothrix, overlying a porous structure due to Phormidium. The
dense structure forms a smooth and hard outer layer to the oncolite and protects it from abrasion. Hard and smooth oncolite, Untersee of Lake Constance. Height of thin section is l -4 mm.
The organization of the three different algal structures of a Lake Constance oncolite is shown diagrammatically in Fig. 22. The nucleus is covered by an initial layer of dense micrite which originates from Schizothrix activity. A spongy microstructure follows due to Phormidium-Calothrix/Dichothrix and this makes up large parts ofthe oncolite, with local dense layers of Schizothrix. In hard and smooth oncolites the Schizothrix layers are more numerous and also constitute the outer surface. In the soft and spongy oncolites the Phormidium- Calothrix/Dichothrix activity dominates and produces a loose structure in the interior as well as on the surface. Rivularia colonies in the oncolites are less common, occurring singly or randomly throughout the entire oncolite. The microstructures produced by Phormidium-Calothrix/Dichothrix and Rivularia are vertically arranged while that of Schizothrix is horizontally oriented.
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Fig. 16. Section of a mussel valve from an oncolite nucleus totally bored, probably by Schizothrix. The valve
is overlain by micrite which makes up the oncolite (boundary shown by a dashed line). Soft and spongy oncolite, Untersee of Lake Constance. Height of thin section is 0·55 mm.
Oncolites are often turned over by wave activity (Schottle & Muller, 1 968) and the new upper surface is then colonized. This change in growth direction has obviously occurred with many oncolites and leads to the formation of so-called basal ring (Golubic, 1973; and own observations). The underside now remains in a reducing zone and there is no further algal growth. Diagenesis
The micritic carbonate of the oncolites has commonly been replaced by a coarser fabric. This sparitization has mostly happened with the carbonate encrusting the algal threads and may perhaps be initiated by them in some way. In the smooth and hard oncolites a coarser fabric bas also replaced some of the micritic matrix. Commonly, cavities within the algal fabric or the voids left through decomposition of the algal
1 00
A. Schafer and K. R. G. Stapf
Fig. 17. Detrital carbonate grain with a micritic envelope apparently produced by boring activities, probably of Schizothrix. Hard and smooth oncolite, Untersee of Lake Constance. Width of thin section is 0·55 mm.
filaments are filled by amorphous silica. In one case quartz with undulous extinction has formed. Sometimes even the microsparitic calcite encrusting algal filaments has been partly replaced by silica. The precipitation of silica seems to be caused by the decay of algal filaments in the interior of the oncolites. This can take place while algae are still flourishing on the oncolite's surface. Algal filament decay produces ammonia and provides an alkaline micro-environment in which silica is dissolved from the abundant siliceous diatom tests embedded in the algal fabric. As a result of this dissolution diatom tests can only be detected within the outer layer of the oncolites. Dissolved silica in the U ntersee (about 5 mg/ 1 ; Muller, 1 969) is not present in sufficient concentration to cause silicification. Sediment formation
Baumann ( 1 9 1 1 ) and later SchOttle & Mtiller ( 1 968) argued that the breakdown of oncolites by wave action, and by frost during winter exposure contributes much to the carbonate sedimentation in deep-water areas of the Untersee. It is thought, however, that the oncolites are too restricted in this distribution to play a significant role in large scale sediment formation. Erosion of sub-Recent lake marl along the littoral zones is a more important source of sediment (Schafer, 1 972, 1 973). Zonation of algal communities
Kann ( 1 959) described a zonation of algal communities from the alpine Traunsee in Austria where water depth was the main control but shoreline morphology, and fetch
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Fig. 18. SEM-view o f araldite-impregnated, HC !-etchings o f a hard and smooth oncolite. Straight bundles
of thicker (branched) algal filaments of probably Phormidium-Ca/othrix/ Dichothrix are interwoven with much thinner divergent filaments, probably of Schizothrix. Untersee of Lake Constance. Field of view is 90 JLm across.
determined the distribution of the individual algal species. In the Untersee, the algal carbonates form beach bars (Baumann, 1 9 1 1 , p. 44) or occur in shallow sublittoral areas only exposed at low-water (B aumann, 1 9 1 1 , pl. III & VIII; Pia, 1 933, p. 1 74). As shown by Kann ( 1 94 1 , p. 529) and Mattern ( 1 970) Calothrix, Dichothrix, Rivularia and Phormidium are mostly present in higher parts of the littoral zone whereas Schizothrix occurs at greater depths. Hard and smooth oncolites acquire their shape by movement in the breaker zone during autumn storms when the spongy surface layers are abraded (Kann, 1 94 1 ) until a deeper hard layer forms the surface. Mattern ( 1 970) thought that the movement of nodules severely hinders algal growth. In the Ermatingen Embayment Rivularia is sparsely distributed throughout the smooth and hard as well as the spongy and soft oncolites. The latter are restricted to deeper water most of the year which, according to Kann ( 1 959), p. 1 54, is too deep for Rivularia. As this alga does not occur in greater amounts in shallow water either, it is suggested that its paucity might be caused by the waste-water influx from the Konstanz sewage plant into the Rhine River. The latter is thought to produce exuberant blooms of Cladophora on the subaqueous Ermatingen sandfl.at (Kann, 1 959, p. 165).
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Fig. 19. Colony of Rivularia cf hae matites within a fabric built by mostly (?) Phormidium. Soft and spongy on co lite, Untersee of Lake Constance. Width of thin se ction is 3·5 mm.
Fig. 20. Incident-light view of Rivularia cf haematites showing its filament tubes. (Araldite-impregnated,
HC l -etched). Soft and spongy oncolite, Untersee of Lake Constance. Width of section is 1-4 mm.
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Fig. 21. Sequence of SEM-views of Rivularia cf haematites: corallinaceous habit of this blue-green algal colony is caused by sparry calcite precipitated within the mucus around the algal filaments. The latter have since decayed. Direction of growth is downwards in the photographs. Soft and spongy oncolite, Untersee of Lake Constance. Fields of view are (a) 4·1 mm, (b) 1·5 mm, (c) 375 p.m and (d) 1 50 p.m across.
The bundle-shaped algae Phormidium-Calothrix/Dichothrix occur in all the oncolites investigated and do not appear to be confined to a certain water depth. Schizothrix layers are intercalated with other algal colonies and only become numerous near the mean water-level, as the abundance of hard oncolites at the Untersee shores indicates. This concentration is due to erosion of friable oncolite surfaces and contrasts with the findings of Kann ( 1 959).
CONCLUSION The lacustrine environments of the Permian Saar-Nahe Basin and the modern Lake Constance are clearly very different. The Untersee of Lake Constance is of
1 04
A. Schafer and K. R. G. Stapf
Schizothrix (hard and smooth Ioyer) Rivularia (corallinaceous colony)
with snail and rock fragment
Fig. 22. Composite sketch of an oncolite showing both a soft and spongy as well as a smooth and hard surface, which originate from different algae producing carbonates. Erosion by water agitation affects the soft and spongy layer. Although on the basis of surface structure there are two types of oncolite, the internal fabric of both types is the same.
restricted size, whereas the Saar-N ahe B asin was a lake-river system on a much larger scale, with the lakes themselves varying in size throughout Rotliegend time. The water supply to Lake Constance fluctuates little compared with the water-level fluctuations which took place during the Permian in the Saar-Nahe Basin. Sedimentation in the rapidly-subsiding Saar-Nahe Basin was much more intensive than it is in Lake Constance today. The climate of the Lake Constance area is temperate and humid, whereas the Lower Rotliegend was humid to subtropical and the Upper Rotliegend subtropical. In spite of these differences both environments are lacustrine and have been conducive to the formation of algal carbonates. From both the ancient and modern occurrence of lacustrine oncolites described in this paper it can be concluded that: (a) the algal carbonates indicate a photic shallow water environment; (b) they are restricted to near-shore areas; (c) they occur in areas of low sedimentation; (d) the water agitation in these near-shore areas is generally low although periodically it may be intense; (e) sediment generation through break-down of algal carbonates is limited; (f) there is a wider range of stromatolite growth forms in the Permian case ( domical mounds, tabular mats and oncolites) whereas only oncolites are present in the modern example; (g) there is a lack of algal filaments in the Permian stromatolites whereas the modern ones show some filament preservation; and (h) sparitization destroyed the microstructure in the fossil stromatolites completely whereas this process is at an early stage in the modern oncolites. The investigation of modern freshwater algal carbonates and their microstructures presented here might contribute towards an understanding of the formation and internal organization of ancient ones such as those of the fluvio-lacustrine Permian Saar-Nahe B asin.
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ACKNOWLE DGMENT S
We thank Drs R. H. Osborne, M. H. Link and J. P. Smoot for critically reviewing the original draft of this paper. In addition we are gratefully indebted to M. E. Tucker for his help looking through a further version of non-Anglicisms.
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Origin of the carbonate sediments in the Wilkins Peak Member of the lacustrine Green River Formation (Eocene), Wyoming, U.S.A.
J O S E P H P . S M O OT* Department of Earth and Planetary Sciences, Johns Hopkins University, Baltimore, Maryland, U. S.A .
AB STR ACT The Eocene Green River Formation is one of the largest non-marine carbonate deposits in the world. The Wilkins Peak Member. which is made up predominantly of'primary' dolomite micrite, has been interpreted as a playa-lake complex. The following subenvironments are recognized: alluvial fans. fringing sandflat with broad, sheet-like basinward tongues, evaporative mudflat and ephemeral lake. The major mode of sedimentation was by sheetwash in a normally subaerially-exposed, evaporative environment. The carbonate sediments are mostly sand to silt-sized dolomite peloidal intraclasts which are found in traction-deposited bed forms. These intraclasts are thought to have formed by the disintegration of three basic kinds of syndepositional precipitates: ( 1) dense dolomite surface crusts formed on the evaporative mudflat; (2) mound-shape travertine tufa precipitated along spring-fed streams; and (3) caliche crusts and cements formed in the alluvial fans. A chemical mass-balance shows that all of the carbonate sediments in the Wilkins Peak could have been produced by this mechanism in the available time. The important consequences of this study are: ( 1) it provides a mechanism of lacustrine carbonate production without a standing lake; (2) it shows that hard, carbonate crusts, which are very abundant in some lake basins, can provide fine-grained sediment: and (3) a large accumulation of 'primary' dolomite can take place through the erosion of syndepositional crusts. *Present address: Department of Earth and Space Sciences, State University of New York, Stony Brook, New York, U.S.A.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
1 10
Joseph P. Smoot INTR O DU CTION
Understanding carbonate sediment production processes through geological time may help to resolve problems of the chemical and biological evolution of the earth. Although the origin of modern marine carbonate muds is still a matter of some uncertainty (Bathurst, 1 97 1 , pp. 276-292), most of the aragonite micrite sediment of the shallow carbonate platforms appears to have resulted from the breakdown of skeletal organisms, such as calcareous algae (Stockman, Ginsburg & Shinn, 1 967; Neumann & Land, 1975). Deepwater carbonate oozes of the oceanic abyssal plains consist primarily of the skeletal remains of planktonic organisms (Milliman, 1 974, pp. 229-236). Further, the importance of a skeletal source of carbonate mud in a hypersaline marine setting is illustrated in Shark Bay, Western Australia: where the waters are too saline for calcareous algae to exist there is no carbonate mud, despite the aragonite supersaturation of the waters, but where such algae thrive the bottom sediments include aragonite mud (Hagan & Logan, 1 974). Not all ancient marine carbonate mudstones can be of skeletal origin, however, because Precambrian carbonate deposits, older than the oldest known calcareous organism (Ginsburg, Rezak & Wray, 197 1, p. 55; Raup & Stanley, 1 97 1, pp. 3 1 4-3 1 5), are abundant. This suggests that a non-skeletal source for carbonate sediments must be called on for Precambrian times. The origin of dolomite in ancient carbonate deposits remains one of the maj or unresolved problems in sedimentology. While many dolomite deposits in the geological record are obviously the result of late diagenesis, some thick and extensive deposits seem to be facies-dependent and so are believed to be 'primary' (in the sense of being formed contemporaneously with deposition). In modern marine environments high Mg calcite and protodolomite (possible progenitors of dolomite) are found in minor amounts (e.g. Andros Island, Bahamas, Hardie, 1 977, p. 1 6 1 , or the Persian Gulf, Shinn, 1 973, p. 1 85). This has led some workers to suggest that ancient oceans had a quite different water composition from that existing today (e.g. Chilingar, 1 956). Unexpectedly, some light can be shed on these carbonate production problems by the lacustrine Eocene Green River Formation of Wyoming, U.S.A., which is a thick and extensive, non-skeletal carbonate deposit rich in dolomite. By providing details of the physical and chemical conditions which affected production of carbonate in this deposit, it is hoped that some insight may be gained into the origin of other non skeletal carbonate deposits, non-marine and marine, modern and ancient. T HE GREEN RIVER FORM A TION OF WYOMING
The Eocene Green River Formation occupies a series of downfaulted, closed, non marine basins in the western United States (Picard & High, 1 972, p. 1 39). The present study has concentrated on the Wilkins Peak Member (Table 1 ) which crops out in the Green River (or Bridger) Basin of Wyoming (Fig. 1 ). The Wilkins Peak deposit is a saline, non-skeletal carbonate deposit, approximately 300 m thick, which is composed predominantly of 'primary' dolomite micrite and has a large suite of well-preserved sedimentary features. These deposits are virtually undeformed and the surrounding mountains are still more or less like they were then. This allows a detailed reconstruction of the drainage patterns of the basin through the provenance of clastic sediments derived from the mountains, making it possible to postulate a reasonable model of the ground water chemistry during Wilkins Peak time.
Il l
Origin of carbonate sediments
1-80 N
CD � � {;\ \:::V
Outcrop belt of Wilkins Peak Member Location of
10
measured section -
km
-
-
� Mop location
_
�yomi�'1_
______
_____ _
Utah
Fig. l. Outcrop belt of Wilkins Peak Member in the Bridger Basin south of Rock Springs, Wyoming. Numbers show locations of measured sections (modified from Bradley, 1964).
The longstanding theory for carbonate production in the Wilkins Peak was precipitation from a stratified, perennial salt lake (Bradley, 1 964; Bradley & Eugster, 1969). Eugster & Hardie ( 1 975, p. 330), presented evidence indicating a playa ephemeral lake complex depositional setting and they postulated a diagenetic source at the basin margin for the carbonate which was periodically eroded and washed into the basin centre by storm floods.
1 12
Joseph P. Smoot Table l. Stratigraphy of Green River Formation
Member
Maximum thickness (m)
Laney Wilkins Peak Tipton
520 370 60
Age Middle Eocene Lower Eocene Lower Eocene
These two models, as well as others, were tested in the present study by analysing measured sections from the basin edge of the Wilkins Peak deposits near the Uinta Mountains (at the Wyoming-Utah border northward to near the ipferred hydrogra phic centre of the depositional basin (based on bedded trona and halite) near the town of Green River (Fig. 1 ) These measured sections were correlated using marker units similar to those of Culbertson ( 1 96 1 ) which include (I) siliciclastic braided stream and floodplain deposits forming nine wedges extending from the southeast corner of the basin northward (see Eugster & Hardie, 1 975, p. 325) and (2) numerous thin volcanic tuff beds (0·01-0·1 m thick). The measured sections provide control on the geographic and temporal distribution of rock facies, thus allowing detailed reconstruction of the physical conditions of carbonate sediment production in the Wilkins Peak and thus establishing the initial constraints on any production model. .
DEP O SITIONAL ENVIRONMENT OF THE WILKIN S PEAK SE DIMENT S
Figure 2 illustrates schematically the distribution of subenvironments (see Hardie et a!., 1 978) in Wilkins Peak time as determined from the interpretation of the facies in the measured sections. there are five major subenvironments: the alluvial fan, the sandflat, the dry mudflat, the perennial lake and the ephemeral lake (saline mudflat and salt pan). Only a very brief description will be given here, for more detail the reader is referred to Smoot ( 1 97 8, in preparation). The alluvial fan deposits are restricted to a narrow area (about 8 km wide) adjacent to the Uinta Mountain front. They are composed of poorly sorted boulder conglomerates and cross-bedded gravels close to the mountains and of broad lenses of quartz sandstones dominated by planar horizontal and low inclined lamination in more distal areas. These sediments are interpreted as having formed in sporadically flooded braid channels in an alluvial fan apex and fan toe. The provenance of the boulders and sands of the alluvial fan subenvironment reflect the limestone and orthoquartzite rocks of the Pennsylvanian Permian sequence in the Uinta Mountains (see Hansen, 1 965). These rocks provided no source for silt-and mud-sized sediments either by physical or chemical weathering.
1 13
Origin of carbonate sediments Perennial lake
Dry mudflat
or
Saline mudflat
salt pan
II
9
8
5
2
Fig. 2. Schematic drawing showing distribution of subenvironments in the Wilkins Peak Member. Numbers
show locations of measured sections.
Fig. 3. Sandflat sandstone. Gravelly sand with horizontal internal lamination (Section 6, Fig. 1 ) . Base of
specimen reflects an erosional surface while the top was covered by a thin, mudcracked dolomite mud drape. The sand is about 60% dolomite peloids, the rest being quartz. The delicate, thin white clasts are dense dolomite micrite. Scale bar is l em long.
The sandflat deposits are sheets of sandstone composed of mixtures ofquartz grains and dolomite peloids. These are in the form of thin beds with slightly scoured bases (usually <0·02 m deep) and horizontal to low-inclined internal lamination (Fig. 3) reflecting upper flow regimes (Simons, Richardson & Nordin, 1 965). Each thin bed is
1 14
Joseph P. Smoot
typically graded with a thin mudcracked dolomite mud drape indicating waning current deposition followed by subaerial exposure. These are interpreted as sheetflood deposits where the floods initiating on the alluvial fans spread out and decelerate on reaching the flat basin floor (Hardie et al., 1 97 8).
Figs 4-7. (4) Laminated dolomite mudstone from the dry mudflat subenvironment (Section 5, Fig. 1 ) . Sand
layers (dark) are lensoid and scour into underlying muds. Mudcracks cut across several laminae, but initiate at almost every bedding plane. Laminae are graded, particularly those in the upper part of the photograph. Scale bar is I em long. (5) Dolomite mudstone laminite from the perennial lake subenvironment (Section 9, Fig. !). Note lensoid character of coarse layers (light) and lack of any bedding disruption. Black dots at top of photograph are oxidized pyrite. Scale bar is I em long. (6) Saline mudflat massive dolomite mudstone ( Section 2, Fig. 1 ) . White crystals are calcite fillings ofshortite crystal molds. Scale bar is I em long. (7) Thin section of mudstone from dry mudflat. Dark grains are coarse silt-size dolomite peloids, rounded white grains are quartz silt. Dark area cutting across picture is a sheet crack filled in with fine-grained dolomite. Photograph is 1 · 5 mm wide.
The dry mudflat deposits are laminated dolomite mudstones with abundant desiccation features (Fig. 4) which are interpreted as having formed during subaerial exposure (Smoot, 1 977, pp. 97-105; Eugster & Hardie, 1975, p. 323). The laminae are typically graded with a lensoid, scouring peloidal fine sand base and a continuous peloidal silt cap. These laminae are also interpreted as having been deposited by sheet floods, but in areas farther away from the alluvial fan than the sandflat deposits. The dry mudflat deposits reflect an area where only fine-grained deposits could be transported and that was predominately subaerially exposed.
1 15
Origin of carbonate sediments
The perennial lake deposits are made up of oil-shales and finely-laminated dolomite mudstones. The lamination in both rock types is typically fiat and non-scouring, but occasional layers show a pinch-and-swell structure (Fig. 5) which suggests that these are made of at least silt-sized particles that could be reworked by gentle currents. These deposits are interpreted as having formed in a shallow but stable lake that occupied the centre of the basin part of the time. The ephemeral lake deposits are laminated to massive dolomite mudstones which reflect deposition in a standing lake (like the perennial lake) but also contain evidence for sporadic subaerial exposure such as mudcracks or evaporites. In the basin centre the ephemeral lake subenvironment is represented by saline mudflat deposits which are dolomite mudstones disrupted by numerous intrasediment salt crystals (Fig. 6) or salt pan deposits made up of alternating layers of massive trona or halite and dolomite mud. The saline mudflat deposits are interpreted as having formed by precipitation of intrasediment salts by evaporative concentration of pore water and the salt pan deposits are interpreted as the hydrographic low where salts precipitated from standing brines when the shallow perennial lakes dried up.
C ARBONATE PRO DUCTION MODEL S: HYPOTHE SE S AND PHY SICAL CON STRAINTS
Four models of carbonate production are possible for the Wilkins Peak: ( 1 ) direct precipitation from a standing lake, (2) intrasediment growth from pore waters, (3) detrital Palaeozoic rock fragments, and (4) erosion of syndepositional crustose precipitates. Each of these must first be judged against the physical constraints imposed by the sedimentary model just described. The constraints are: ( 1 ) the environment was predominantly one of subaerial exposure and sporadic sedimenta tion (over 70% of the carbonate sediments away from the basin centre are dry mudflat or sand fiat deposits), (2) conditions were arid (evaporation > inflow) but much water was coming in to account for the thick layered salt deposits and abundant intrasediment growths of soluble salts in the basin centre, (3 ) a perennial lake only occurred sporadically and its deposits are restricted primarily to the basin centre, (4) sediment transport was mostly by basinward sheetwash, (5) the source area for detrital sediments in the Uinta Mountains did not produce appreciable amounts of siliciclastic fine-grained sediments, (6) the hydrologic system forced the coarsest material to accumulate at the basin edge, while only the finest material was transported basinward, and (7) the carbonate sediments are predominantly silt to fine-sand sized dolomite micrite peloids (Fig. 7) which form traction-load bedding structures (see also Eugster & Hardie, 1 975, p. 3 30). Precipitation from a standing lake has long been the model for carbonate production in the Wilkins Peak Member (Bradley, 1 964; Bradley & Eugster, 1 969). This model obviously conflicts with the subaerial conditions inferred from the sedimentary model and also fails to explain the predominance of intraclast carbonate sediments as opposed to fine crystalline muds one would expect to precipitate out of standing lake water. In the in situ precipitation from pore water model one would expect massive growth of dolomite crystals to disrupt bedding, which is inconsistent with the abundance of well-preserved depositional structures in the Wilkins Peak sediments. Also it would not explain the consistent relationship of quartz grain size being similar to micrite clot size (Eugster & Hardie, 1 975, p. 3 24) The .
1 16
Joseph P. Smoot
Palaeozoic limestones and dolomites have definitely provided sediment to the alluvial fan deposits from boulder down to at least coarse sand size. However, that they also provided mud-size sediments is unlikely becanse the Palaeozoic carbonates are mostly sparry and they are composed of both calcite and dolomite, while the Wilkins Peak muddy sediments are overwhelmingly dolomite micrite peloids. The last model, erosion of syndepositional crustose precipitates, will concern the rest ofthis paper as it is believed to be the most compatible with the physical evidence.
SYN DEPO SITION AL C AR B ON ATE CRU ST S IN THE WILKIN S PE AK DEPO SIT
There are three types of syndepositional crust that appear to have provided significant amounts of sediment to the Wilkins Peak deposit: ( 1 ) dolomite surface crusts, (2) caliche crusts and cements, and (3) travertine tufas. Each of these will be described below in detail, with an explanation of its origin using modern analogs. Surface dolomite crust
The surface dolomite crusts appear as 0· 1 -2 mm thick, white layers of dense, pure dolomite micrite, which cap heavily desiccated dry mudflat or ephemeral lake mudstones (Fig. 8). The layers have a sharp top and gradational base, which as can be seen in Fig. 9, may be discordant with the underlying lamination, so that they cannot be mud drapes. The hardness of these layers is further demonstrated by the fact that mudcracks stop abruptly or extend horizontally as sheetcracks upon intersecting a crust (Fig. I 0). Also, the layers form very thin delicate fiat intraclasts (see Fig. 3), the size of the clasts being dependent on the thickness of the layer from which they are eroded; 0· 1 mm thick crusts readily provide fine sand and silt-sized intraclasts. The Wilkins Peak crusts are very similar to thin aragonite 'paper crusts' which are forming now on Andros Island, B ahamas, where capillary pore waters are concentrated by evaporation causing precipitation at the surface (Hardie & Ginsburg, 1 977, figs 52 and 53). The Wilkins Peak crusts are also very similar to thin crusts of 45 mol per cent Mg calcite that are scattered on the surface of the modern playa mud fiat in Salt Flat Graben, Texas. High Mg calcite and protodolomite crusts are reported from Shark Bay, Western Australia (Logan, 1 974, p. 225), the Persian Gulf (Shinn, 1 969), around Florida Bay (Atwood & Bubb, 1 970), and Andros Island, Bahamas (Shinn, Ginsburg & Lloyd, 1 965; Hardie, 1 977). The environment in which all these are found is a subaerially exposed (at least periodically evaporative) carbonate mud fiat with only sporadic introduction of sediment, the same conditions postulated for the Wilkins Peak dry mudflat deposits based only on the sedimentary structures. The crusts in the Wilkins Peak, then, are believed to have formed by the evaporative concentration in the vadose zone of pore water within the evaporative dry mudflat sediments causing the precipitation of high Mg calcite as an expanding cement at the sediment surface (see Hardie, 1 977, p. 1 73). Dolomite crusts were found in dry mud fiat mudstones in Section 2 to Section 1 1 (Fig. 1 ) The actual abundance ofthese crusts is difficult to estimate as only the thickest crusts ( > 1 mm) are visible in outcrop (and even then only on good outcrops). However, some samples of dry mudflat mudstones that were slabbed show crusts on .
Origin of carbonate sediments
1 17
almost every lamina (see Fig. 1 0), although other mudstones contain no in situ crusts. Dense, white dolomite clasts of sand size are very common in dolomitic sandstones and sandy dolomite mudstones. These are probably derived from erosion of the crusts.
Figs. 8-10. (8) Dolomite crust (light layer in centre of photograph) overlying a heavily desiccated ephemeral
lake mudstone (light wisps are very compacted mudcracks) (Section 7, Fig. l ). Overlying silt scours the mudstone just to the right of the crust. Scale bar is l em long. (9) Thin-section of dolomite crust shown in Fig. 8 . The crust (dark area) grades into and unconformably overlies the laminated peloidal silt underneath. The crust in turn is sharply overlain by a quartz-rich silt. Scale bar is 0·5 mm long. ( 1 0) Dry mudflat mudstone with numerous dolomite crusts (thin white layers) (Section 4, Fig. l ). Note flat clasts of crusts in sediments. Mudcracks (dark vertical disruptions) stop abruptly and form horizontal sheetcracks at the crust just right of centre, while cracks extend across the interface where the crust has been eroded on the left. Mudcracks stop abruptly at the very thin crust at top of the basal layer in this photograph. Scale bar is l em long.
Caliche crusts and cements
Caliches are found in the alluvial fan deposits (Sections 1 0 and 1 1 , Fig. 1 ) as crusts, cements and rinds around the Palaeozoic limestone and dolomite clasts. The caliches occur as both calcite and dolomite, the former as cements in the deposits close to the Uinta Mountain front (unmeasured sections south of Section 1 0, Fig. 1 ), the latter as crusts, rinds and cements in the fan toe deposits. The crusts and rinds are dense micrite with colour banding that has diffuse boundaries while the cements are usually simple micrite intergranular pore fillings. In thin section, the Wilkins Peak caliches have a clotted texture with irregular spar-lined voids, spar-coated quartz grains and disseminated quartz silt grains (Fig. 1 1 ), as are seen in Recent caliche crusts.
1 18
Joseph P. Smoot
Fig. 1 1 . Comparison of thin-sections of a caliche crust from an alluvial fan deposit in Wilkins Peak (a)
(Section 10, Fig. l) and a modern caliche crust from Texas (b). Note the clotted texture, the irregular sparlined voids (v), spar-coated quartz grains (q), and finely disseminated quartz silt (white specks). Wilkins Peak sample is 2 mm wide and the Texas caliche is 3 mm wide.
In modern alluvial fans caliches also commonly occur as crusts, cements and rinds on limestone boulders (Bull, 1 972, p. 65; Lattman, 1 973). The composition of modern caliches ranges from low-Mg calcite to protodolomite (Gevers, 1 930; Friedman, 1 965, p. 266; Goudie, 1 973, p. 21 ). The Wilkins Peak caliches are commonly found as recognizable cobble to granule-size fragments of crusts and cemented lumps of quartz sand in the alluvial fan toe deposits. The more delicate rinds and cements almost certainly provided some fine-grained micrite but such grains would be too small to preserve fabrics diagnostic of their source.
Fig. 12. Mound-shaped, stromatolitic travertine (Section 5, Fig. l) (centre of photo to left of knife) coating eroded surface of siliciclastic sand and in turn covered with siliciclastic silt (sediment in outer fringe of photograph). Note porous internally laminated character of the travertine crust. Knife is 35 em long.
Origin of carbonate sediments
1 19
·Travertine tufas
The travertine tufas occur as stromatolitic mounds on the eroded surfaces of siliciclastic and dolomitic mudstones and sandstones (Fig. 1 2) and as bumpy crusts on desiccated dry mudflat mudstones. The evidence for these carbonate features being travertines is based on their composition and fabric. The composition ofthe stromatolitic structures is independent of the surrounding sediments, for example, a pure dolomite mound will drape a pure quartz sandstone and in turn be covered by siliciclastic silt (Fig. 1 2). This indicates that the mounds are not sediment trapping algal stromatolites, but rather precipitated structures. The internal fabric of the mounds consists of alternating porous micrite layers, commonly with a crude palisade structure, and dense micrite layers with internal lamination (Fig. 13). In thin section the dense laminar layers are either sparry radial 'sinter' fabric (Irion & Muller, 1 968, figs 1 2 and 1 3 ) or very dense with vague vertical structures (the densely calcified zones of Monty, 1 976, fig. 7).
Fig. 13. Internal textures of travertine crusts. A travertine crust from the Wilkins Peak (a) is compared to a
modern travertine from a stream deposit near San Antonio, Texas. (b) The key characteristics are the dense laminar layers with vague internal colour banding alternating with fenestral layers with a tubular, clotted fabric. Wilkins Peak travertine with a vertical palisade fabric (c) is similar to modern stream tufas shown by Irion & Muller ( 1 968, fig. 5) and Monty ( 1 967, fig. 7). Scale bar on each photo is 0·5 em long.
The porous layers are made of micrite clots separated by vertical irregular voids (Fig. 14) which are like the tufa layers of modern travertines where carbonate precipitates around algal filaments or moss thalli (Irion & Muller, 1968; Monty, 1976, p. 207).
1 20
Joseph P. Smoot
Modern travertines of this variety are precipitating at spring orifices and along spring fed streams (Slack, 1 967; Irion & Mi.iller, 1 968) where ground waters degas C02 to equilibrate with pC02 in the air and the stream waters are evaporated while flowing along their course. Spring outlets are common in modern playa systems, particularly at the toes of fans, where they feed small streams or shallow ponds on the mud flats (Jones, 1 965; Hardie et a!. , 1 978). The scoured surfaces upon which the Wilkins Peak stromatolitic structures are precipitated are interpreted as small stream cuts and the crust-covered mudstones are believed to result from precipitation of travertine from shallow sheets of water flowing on the mud flat surface. The dolomitic composition of the structures is not inconsistent with the travertine origin as high Mg calcite and protodolomite travertines have been reported in the Recent (Barnes & O'Neil, 197 1 ; Stalder, 1 975), their composition apparently dependent on the MgH /Ca2 + ratio of the waters.
Figs. 14 and 15. ( 1 4) Thin section of porous tufa layers in travertine crust (Section 4, Fig. 1). Note clotted texture of the porous layers and the crude vertical orientation of voids. Dense layers are either dark calcified zones or light-coloured sinter layers. Photograph shows area 1 ·5 em high. ( 1 5) Thin-section showing close up of porous tufa layer in Fig. 14. Note the delicate micrite fabric and fine-sand to silt size clots. Compare this to the textures of the peloids in Fig. 7. Photograph shows area 2·5 m m high.
Coarse sand- and gravel-sized fragments of the dense laminar stromatolitic layers are commonly found in the Wilkins Peak sands, but fragments of the porous tufa layers are rare. An analogous situation exists in the modern deposits at Deep Springs Lake, California, where travertines precipitated along spring-fed stream channels on the Birch Creek alluvial fan are almost completely eroded with each major storm (every 3-4 years) (Slack, 1 967) but only a few dense fragments have been found near the fan toe (Jones, 1 963). Specimens of the B irch Creek travertine provided by Slack easily broke down with gentle rubbing to fine sand- and silt-sized micrite peloids. The porous layers of the Wilkins Peak travertines are believed to be an excellent source for the fine-grained dolomite micrite peloids that make up the bulk of the carbonate sediments (Fig. 1 5). The abundance of these structures is difficult to ascertain as they
Origin of carbonate sediments
121
have limited lateral extent (usually less than 5 0 m) and are frequently covered in weathered outcrops. For instance, in Section 4 (Fig. 1) seventeen travertine layers were found while in Section 5, twenty horizons were found. However, only eight layers were in equivalent horizons between the measured sections, which means there are at least twenty-nine horizons of travertine layers that were not completely removed by erosion there and many more horizons were found in unmeasured portions of the outcrops. Other evidences of travertine layers where no in situ mounds were found are sand layers with gravel fragments that are travertine. These layers are usually traceable into typical sand flat sandstones. Some of the gravelly sands are oolitic, with the ooids being poorly sorted, irregular shaped and often with only one or two layer thick cortices (Fig. 16). These features are very similar to those ofthe Pleistocene fluviatile ooids described by McGannon (1975) which he interpreted as having precipitated from a spring-fed river.
Fig. 16. Poorly sorted oolitic sand from sand flat sandstone (Section 4, Fig. I ). Note the irregular shapes of
the ooids and the grains with only one or two coatings (for instance the two large quartz grains in the upper left hand corner). Compare this with fig. I I in McGannon ( 1 975). Width of photograph represents about 4 mm.
M A S S B A L ANCE OF C A RB ON ATE P R O DU CTI ON
The dolomite crusts, caliches and travertines apparently provided some carbonate sediments in the Wilkins Peak, but did they provide enough to account for all of the carbonate? To test this out the following mass balance simulates the precipitation of travertine from spring waters using a computer programme devised by Hardie & Eugster ( 1 970). A ground water with low average values of solute concentrations from a limestone terrain in an arid or semi-arid region (White, Hem & Waring, 1963) was allowed to 'degas' to equilibrium with atmospheric pC0 2 then was evaporated to four times the initial concentration as would occur in modern springs as they surface and
1 22
Joseph P. Smoot
flow along open stream channels (Slack, 1 967). In this simulation the initial calcium in solution (0·01 g/1) was 90% removed as calcite while the water was still fresh enough for drinking (ionic strength < 0· 1 )! The simulated evaporative concentration and precipitation of calcite led to significant increase in alkali and CO/ - concentration, which is consistent with the sodium carbonate salts (mostly trona) which precipitated from the final brines in the basin centre during Wilkins Peak time. To produce 8-44 X 10'4 kg of carbonate sediment (Bradley & Eugster, 1 969) at the basin edges just by travertine precipitation in 1 X 1 06 years (estimated by Bradley, 1 964, but could be low by a factor of three, Eugster & Hardie, 1 975, p. 327), the yearly total spring water input would need to be 2·8 X 1 0'2 1/year. Table 2 shows a comparison of this value with measured spring inputs into two modern saline basins. Note that the differences in drainage area vary at the same scale as the spring water input, suggesting that the required Wilkins Peak spring inflow is not at all unreasonable. Table 2. Comparison of yearly spring inputs measured for Recent Deep Springs Lake and Great Salt Lake with the calculated value for the Wilkins Peak's Lake Gosiute.
Spring input (!/year) Area (km') Deep Springs Lake Great Salt Lake Lake Gosiute
·
5-4 X 1 09 ! ·O x 1 0 ' ' 2·8 X 1 0 1 2
< 10 350 9500
The computer program in this mass balance could not simulate the precipitation of Mg calcite, which is what is needed for the Wilkins Peak model. The difference in spring input, however, is not considered important enough to invalidate the mass balance for the following reasons: ( 1 ) Fuchtbauer & Hardie ( 1 976) demonstrated experimentally that the MgC0 3 content of a Mg calcite is directly related to the Mg/Ca ratio of the solution and that high Mg calcite (up to 60 mol per cent MgC03 !) will precipitate easily under the appropriate kinetic conditions (see discussion in Hardie, 1 977, p. 1 74), (2) modern high Mg calcite and protodolomite travertines have been reported, so the kinetic conditions in those circumstances must be favourable, (3) the Mg2+ /Ca2 + ratio of most ground waters will rise as calcite precipitates from them (Hardie & Eugster, 1 970) a phenomenon observed in modern alluvial fans as calcite cements precipitate from subsurface waters (Eugster & Hardie, 1 978). That this also occurred in the Wilkins Peak alluvial fans is indicated by the distribution of caliche cements, and (4) the mass balance is based only on travertine precipitation and the initial precipitation of calcite, while the actual conditions included the caliches and dolomite crusts and probably occurred until all of the Ca2+ and Mg2+ ions were removed.
C AR B ON ATE SE DIMENT P R O D UCTION : THE MO DEL
B ased on the environmental reconstruction of the basin, the observed carbonate textures and the hydrologic and chemical constraints imposed by the mass balance, the
Origin of carbonate sediments
1 23
following model is proposed for carbonate sediment production in the Wilkins Peak deposit. Ground waters draining the Palaeozoic limestone-dolomite-orthoquartzite terrain of the Uinta Mountains precipitated low Mg calcite cements within the sediments of the alluvial fan apices, thus increasing the Mg2 +/Ca2 + ratio ofthe waters. At the toes of the fans, high Mg calcite precipitated as caliche crusts, cements and rinds near and at the sediment-air interface. Where the ground waters surfaced as springs at the toes of the alluvial fans, high Mg calcite and protodolomite travertines (chemical stromatolites) precipitated as the waters degassed. More travertine precipitated along shallow cut stream channels as the spring outflow coalesced and drained onto the inactive mudflats. Some travertine also precipitated on the mudflat surface as dense laminar crusts, where the waters were reduced to thin flowing films on the surface. On the dry mudflat surface, very thin high Mg calcite (protodolomite?) crusts precipitated from vadose waters. All of the precipitation occurred during each of the many long periods between sedimentation events. Small storms initiating in the mountains caused flash floods on the fans which eroded caliche crusts and travertines at the toes and as the waters spread out into sheetfloods on the playa surface they continued to erode travertines and dolomite crusts on the surface. The hydrology of the basin ensured that the coarse material was left at the basin edge (quartz and Palaeozoic carbonate sand) while the finer sediment was washed into the basin centre. The fine sediments were almost exclusively fine-sand and silt fragments of travertine, caliche and dolomite crust since those grain sizes were not produced from the source rocks. It should be stressed that this absence of siliciclastic mud is the main reason why the Wilkins Peak Member is such a pure carbonate deposit. Each flood not only carried freshly eroded precipitates, but also intraclasts of desiccated, previously deposited carbonate sediment. Therefore, the accumulation in the basin centre is made up of a large fraction of carbonate sediments that have been through several stages of transport followed by subaerial exposure and desiccation then re-erosion. Floods capable of moving sand flat sands into the central portions of the basin eroded the crusts and carbonate muds (particularly the desiccation mudcrack polygons), thus the initially almost pure quartz sand was gradually diluted by carbonate intraclasts during transport into the centre. The coarser sand grains were left at the basin edge (which includes most of the quartz fraction) so the quartz component of the sands was also decreased by size sorting (the sorting may have been aided by the lower density of the micrite grains due to their porosity). The perennial lake laminites (including oil shales) are made of the small finer-than-silt fraction of the eroded crusts plus a minute amount of clay derived from the source area or brought in by wind. These settled out in the standing water and were reworked by bottom currents. Some calcite-rich oil shales may have been precipitated out of the water column. This model is favoured as the best explanation for carbonate sediment production in the Wilkins Peak Member for the following reasons: ( 1 ) the carbonate sediments are predominantly detrital intraclasts deposited in a subaerially exposed, evaporative environment; (2) the three crustose precipitates are the only obvious source of material similar to the intraclasts in texture and composition; (3) modern dolomite crusts, caliches and travertines all form in conditions similar to those postulated for the Wilkins Peak; (4) the various crusts all can easily break down into the necessary grain sizes, especially the tufa layers of the travertines; and (5) the mass balance suggests that adequate amounts of carbonate minerals can be produced by precipitating crusts at the basin edge to provide all of the carbonate sediments in the Wilkins Peak within a
1 24
Joseph P. Smoot
reasonable depositional time period. The possibility of direct precipitation from standing lake water, intrasediment growth or erosion of Palaeozoic rocks as sources for carbonate sediments have not been disproven. However, the evidence presented does not favour any of these as a major source in the Wilkins Peak, although they may be important in other conditions.
IMPLIC ATION S OF THI S STU D Y
This study has demonstrated the importance of knowing the physical conditions of a sedimentary basin before developing models for 'chemical problems' such as carbonate sediment production. As a mechanism of non-skeletal carbonate production the following implications can be made: ( 1 ) a major carbonate deposit can be made without a major body of standing water; (2) hard, precipitated crusts can provide fine-grained carbonate sediments; and (3) large 'primary' dolomite deposits can be made by the erosion of syndepositional high Mg calcite and protodolomite crusts and their deposition in a place where no other sediment is brought in. This last implication could include the Recent marine crusts, which only form a small fraction of the surface sediments, but could be deposited as the major sediment particles under the proper conditions. The implications for Precambrian carbonate deposits are centred on the aspect of precipitated stromatolites as sources of sediment. The Precambrian contains many examples of stromatolites which seem to be precipitated structures as opposed to sediment-trapping algal structures (see Bertrand-Sarfati, 1 976; Donaldson, 1 976; Hoffman, 1 976). These stromatolites are interpreted as having formed under marine conditions (although some are not obviously so) but this does not mean they cannot be providing sediment. The modern fresh-water algal marshes of the basically marine deposits on Andros Island, Bahamas and the Everglades, Florida are accumulating carbonate muds derived from in situ precipitated algal tufas (Monty, 1 967; Monty & Hardie, 1 976; Gleason & Spackman, 1 973) . This suggests that at least some of the Precambrian carbonate deposits could be entirely derived from precipitated stromatolites. Also some of the large Precambrian carbonate deposits could conceivably be non-marine! The presence of precipitated stromatolites should alert the worker to these possibilities.
A CKNOWLE DGMENT S
This paper is part of the author's doctoral thesis under the guidance of L. A. Hardie and H. P. Eugster. I would like to thank them for introducing me to the Green River Formation and teaching me the tools I needed to work in it. I am especially grateful to Dr Hardie for being a constant source of encouragement and insight. Dave Yanko provided invaluable assistance in measuring sections and collecting specimens for two field seasons and also helped in the preparation of the samples. I am very thankful to Mr James Montgomery and his family for letting Dave and me stay on their ranch
Origin of carbonate sediments
1 25
while working in Wyoming and for providing assistance in finding outcrops. Keith Slack of the U.S.G.S. kindly provided specimens of the Birch Creek travertines. I would like to thank the reviewers, Dr Hans Fuchtbauer and Dr C. C. Reeves, and the editors for their comments and criticisms. Field work was supported by NSF grant No. GA 3 1 076, GSA student grant No. 2 1 22-76 and the Shell Oil Research Fund at Johns Hopkins University. The I.A.S. graciously provided funds to allow me to attend this symposium.
RE FERENCE S ATWOOD, O.K. & BUBB, J.N. ( 1 970) Distribution of dolomite in a tidal flat environment, Sugarloaf Key, Florida. J. Geol. 78, 499-505. BARNES, I. & O'NEIL, J.R. ( 1 97 1 ) Calcium-magnesium carbonate solid-solutions from Holocene con glomerate cements and travertines in the Coast Range of California. Geochim. cosmochim. A cta, 35, 699-7 1 8 . BATHURST, R.G.C. ( 197 1 ) Carbonate Sediments and Their Diagenesis. Developments in Sedimentology, 12. Elsevier Publishing Co., Amsterdam. BERTRAND-SARFATI, J. ( 1 976) An attempt to classify Late Precambrian stromatolite microstructures. In: Stromatolites (Ed. by M. R. Walter), Developments in Sedimentology, 20, pp. 25 1-260. Elsevier, Amsterdam. B RADLEY, W.H. ( 1964) Geology of the Green River Formation and associated Eocene rocks in southwestern Wyoming and adjacent parts of Colorado and Utah. Prof Pap. U.S. geol. Surv. 496-A. BRADLEY, W.H. & EuGSTER, H.P. ( 1 969) Geochemistry and paleolimnology of the trona deposits and cssociated authigenic minerals of the Green River Formation of Wyoming. Prof Pap. U.S. geol. Surv. 496-B. BULL, W.B. ( 1972) Recognition of alluvial-fan deposits in the stratigraphic record. In: Recognition of A ncient Sedimentary Environments (Ed. by 1. K . Rigby and W. K. Hamblin), Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 16, 63-83. CHILINGAR, G . V. ( 1956) Relationship between Ca/Mg ratio and geologic age. Bull. Am. Ass Petrol. Geol 40, 2256-2266. CULBERTSON, W.C. ( 196 1 ) S tratigraphy of the Wilkins Peak Member of the Green River Formation, Firehole Basin quadrangle, Wyoming. Prof Pap. U.S. geol. Surv. 424-D, 170- 1 73 . DoNALDSON, J.A. ( 1976) Paleoecology of Conophyton and associated stromatolites i n the Precambrian Dismal Lakes and Rae Groups, Canada. In: Stromatolites (Ed. by M. R. Walter), Developments in Sedimentology, 20, pp. 523-534. Elsevier Publishing Co., Amsterdam. EuGST:2R, H.P. & HARDIE. L.A. ( 1 975) Sedimentation in an ancient playa-lake complex: the Wilkins Peak Member of the Green River Formation of Wyoming. Bull. geol. Soc. Am. 86, 3 1 9-334. EUGSTER, H.P. & HARDIE, L.A. ( 1 978) Saline Lakes. In: Physics and Chemistry ofLakes (Ed. by A. Lerman). Springer Verlag, New York. FRIEDMAN, G . M. ( !965) Occurrence and origin of Quaternary dolomite of Salt Flat, West Texas. J. sedim. Petrol. 36, 263-267. FOCHTBAUER, H . & HARDIE, L.A. ( 1 976) Experimentally-determined homogeneous distribution coeffici ents for precipitated magnesium-calcites: application to marine carbonate cements. A bs. Prog. geo/. Soc. A m. Meetings, 8, 877. GEVERS, T.W. ( 1 930) Terrester Dolomit in der Etoscha-Pfanne, Sudwest-Afrika. Z. Mineral. 6, 224-230. G INSBU RG , R.N., REZAK, R. & WRAY, J.L. ( 197 1 ) Geology of Calcareous A lgae (Notesfor a Short Course), Sedimenta I. Comparative Sedimentology Laboratory, University of Miami. GLEASON, P.J. & SPACKMAN, W. ( 1 973) The algal origin of a freshwater lime mud associated with peats in the southern Everglades. Abs. Prog. geol. Soc. Am. Meetings, 5, 398-399. GOUDIE, A.S. ( 1 973) Duricrusts in Tropical and Subtropical Landscapes. Oxford University Press, London. HAGAN, G.M. & LOGAN, B.W. ( 1 974) Development of carbonate banks and hypersaline basins, Shark Bay, Western Australia. Am. A ss. Petrol. Geol. Mem. 22, 61- 1 39.
1 26
Joseph P. Smoot
W.R. ( 1 965) Geology of the Flaming Gorge Area Utah-Colorado-Wyoming. Prof Pap. U.S. geol. Surv. 490. HARDIE, L.A. ( 1 977) Algal structures in cemented crusts and their environmental significance. In: Sedimentation on the Modern Tidal Flats of Northwest Andros Island, Bahamas (Ed. by L. A. Hardie), pp. 1 59-177. Johns Hopkins University Press, Baltimore. HARDIE, L.A. & EUGSTER, H.P. ( 1 970) The evolution of closed-basin brines. Spec. Pap. Miner. Soc. A m. 3, 273-290. HARDIE, L.A. & GINSBURG, R.N. ( 1 977) Layering, the origin and environmental significance of lamination and thin bedding. In: Sedimentation on the Modern Tidal Flats of Northwest Andros Island, Bahamas (Ed. by L. A. Hardie), pp. 50- 1 23. Johns Hopkins University Press, Baltimore. H ARDIE, L.A., SMOOT, J.P. & EuGSTER, H.P. ( 1 978) Saline lakes and their deposits: a sedimentological approach. In: Modern and A ncient Lake Sediments (Ed. by A. Matter and M. E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 7-4 1 . HOFFMAN, P . ( 1 976) Environmental diversity o f Middle Precambrian stromatolites. In: Stromatolites (Ed. by M. R. Walter), Developments in Sedimentology, 20, pp. 599-6 12. Elsevier, Amsterdam. IRION, G. & M OL LER , G. ( 1968) Mineralogy, petrology and chemical composition of some calcareous tufa from the Schwabische Alb, Germany. In: Recent Developments in Carbonate Sedimentology in Central Europe (Ed. by G. Muller & G. M. Friedman), pp. 157- 1 7 1 . Springer-Verlag, New York. JONES, B.F. ( 1 963) The hydrology and mineralogy of Deep Springs Lake Inyo County, California. Ph.D. Dissertation. Johns Hopkins University, B altimore. JoNES, B. F. ( 1 965) The hydrology and mineralogy of Deep Springs Lake, In yo County, California. Prof Pap. U.S. geol. Surv. 502-A, l -56. LATTMAN, L.H. ( 1 973) Calcium carbonate cementation of alluvial fans in southern Nevada. Bull. geol. Soc. A m. 84, 301 3-3028. LOGAN, B.W. ( 1 974) Inventory of diagenesis in Holocene-Recent carbonate sediments, Shark Bay, Western Australia. Am Ass. Petrol. Geol. Mem. 22, 1 95-249. MCGANNON, D.E. ( 1 975) Primary fluvial oolites. J. sedim. Petrol. 45, 7 1 9-727. MILLIMAN, J.D. ( 1974) Marine Carbonates. Springer Verlag, Berlin. MONTY, C.L.V. ( 1967) Distribution and structure of Recent stromatolitic algal mats, eastern Andros Island, Bahamas. Ann. Soc. geol. Be/g., Bull. 90, 55-100. MONTY, C.L.V. ( 1 976) The origin and development of cryptalgal fabrics. In: Stromatolites (Ed. by M. R. Walter), Developments in Sedimentology, 20, pp. 1 93-250. Elsevier, Amsterdam. MONTY, C.L.V. & HARDIE, L.A. ( 1976) The geologic significance of the freshwater blue-green algal calcareous marsh. In: Stromatolites (Ed. by M. R. Walter), Developments in Sedimentology, 20, pp. 447-478. Elsevier, Amsterdam. NEUMANN, A.C. & LAND, L.S. ( 1975) Lime mud deposition and calcareous algae in the Bight of Abaco, Bahamas: a budget. J. sedim. Petrol. 45, 763-786. PICARD, M . D . & HIGH, L.R. ( 1 972) Criteria for recognizing lacustrine rocks. In: Recognition of A ncient Sedimentary Environments (Ed. by J. K. Rigby & W. K. Hamblin). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 16, 1 08-145. RAUP, D.M. & STANLEY, S.M. ( 1 97 1 ) Principles of Paleontology. W. H . Freeman & Co., San Francisco. SHINN, E.A. ( 1 969) Submarine lithification of Holocene carbonate sediments in the Persian Gulf. Sedimentology, 12, 109-144. SHINN, E. A. ( 1 973) Carbonate coastal accretion in an area of longshore transport, NE Qatar, Persian Gulf. In: The Persian Gulf ( Ed. by B. H. Purser), pp. 1 79- 1 92. Springer Verlag, New York. SHINN, E.A., GINSBURG, R.N. & LLOYD, R.M. ( 1 965) Recent dolomite from Andros Island, Bahamas, In: Dolomitization and Limestone Diagenesis, A Symposium (Ed. by L. C. Pray & R. C. Murray). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 13, 1 1 2- 1 23 . SIMONS, D . B . , RICHARDSON, E.V. & NORDIN, C . F . ( 1965) Sedimentary structures generated b y flow in alluvial channels. In: Primary Sedimentary Structures and their Hydrodynamic Interpretation (Ed. by G . V. Middleton). Spec. Pubis Soc. econ. Pa/eon. Miner., Tulsa, 1 2 . 34-52. SLACK, K.V. ( 1 967) Physical and chemical description of Birch Creek a travertine depositing stream, Inyo County, California. Prof Pap. U.S. geol. Surv. 549-A. SMOOT, J.P. ( 1977) Sedimentology of a saline closed basin: the Wilkins Peak Member, Green River Formation (Eocene), Wyoming. Unpublished Ph.D. Thesis, Johns Hopkins University, B altimore. SMOOT, J.P. ( 1 978) Sedimentary subenvironments in the Wilkins Peak Member of the Green River Formation (Eocene), Wyoming. (In preparation.) HANSEN,
Origin of carbonate sediments
1 27
P.J. ( 1975) Cementation of Pliocene-Quaternary fluviatile clastic deposits in and along the Oman Mountains. Geologie Mijnb. 54. 148- 1 56. STOC KMAN, K . W . , G IN S B U RG, R.N. & S H I N N . E.A. ( 1 96 7 ) The production of lime mud by algae in South Florida. J. sedim. Petrol. 37, 633-648. WHITE, D. E., HEM, J . D . & WARING, G . A . ( 1963) Chemical composition of subsurface waters. Prof Pap. U.S. geo/. Surv. 440- L.
STALDER,
Spec. Publs int. Ass. Sediment. (1978) 2, 129-145
Late Neogene chemical sedimentation in the Black Sea*
K E N N E T H J . H S 0 and K E R R Y K E L T S Geological Institute, Swiss Federal Institute of Technology, Zurich, Switzerland
AB STR ACT
Deep-sea drilling penetrated a largely lacustrine sequence more than 1 km thick in the Black Sea. The oldest sediments are Late Miocene black shales, deposited in a brackish water body, which was a part of the Paratethys lac-mer. The middle (and the bulk) of the penetrated section, ranging from Late Miocene to early Pleistocene in age, was laid down during a time of periodic chemical sedimentation, when calcite, magnesian calcite, aragonite, dolomite, and siderite were precipitated. The youngest sediments are middle to late Quaternary and are largely terrigenous. The Black Sea was a lake during much of the late Neogene; only rarely was the Black Sea, as it is now, marine brackish, when sea water from the Mediterranean could enter. Calcite chalks were deposited in a deep freshwater lake environmentally similar to the Holocene Lake Zurich. Aragonite and magnesian calcite were laid down in the Black Sea at times when it was brackish-marine or hypersaline. Dolomite muds were precipitated during a period when the Black Sea changed from a brackish lac-mer into a shallow salt lake. Stromatolitic dolomite was formed diagenetically in a sabkha-like environment. Siderite occurred when deep-weathering of low-lying coastal plains resulted in a high influx of dissolved iron to the Black Sea.
*Contribution No. 1 06 of the Laboratory of Experimental Geology, Swiss Federal Institute of Technology, Zurich, Switzerland.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
130
Kenneth J. Hsii and Kerry Kelts INTR O DUCTION
The Black Sea is now a giant brackish lagoon of the Mediterranean Sea. However, the Black Sea basin hosted a lake for much of the time during the last 6 or 8 million years. The lower Quaternary and the Plio-Miocene sequences there include considerable chemical and biogenic sediments. Calcitic lacustrine chalk is the main chemical sediment; sideritic, dolomitic, and aragonitic deposits are also present. Diatoms comprise the most common biogenic component. Ostracodes, molluscs, and other calcitic fossils are present as minor constituents. This paper is an attempt to interpret the chemical sediments of the Black Sea and their changing depositional environments, with reference to Holocene analogues, in particular Lake Zurich. We would like to mention at the outset that this work resulted from a post-cruise study of cores obtained by the DSDP Leg 42B cruise to the Black Sea. Our interpretations rely heavily on geological data to be published in a forthcoming cruise report (Ross, Neprochnov et a!., 1 978); we refer particularly often to the work by Peter Stoffers, Alfred Traverse, and Hans Schrader. We are grateful to the JOIDES organization, and to the Deep Sea Drilling Project for having given us the opportunity to study the Black Sea cores. We would also like to thank the many scientists who served as the shipboard staff and who engaged in post-cruise studies; discussions with them have contributed greatly to our understanding of the Neogene geology of the Black Sea. We are indebted to the help of our colleagues, especially Judy McKenzie who furnished data on stable isotopes and Helmut Franz, who made the excellent SEM photographs. Peter Stoffers, John Milliman, and the Editors of this Volume read the first draft of the manuscript; their comments led to a substantial improvement.
Geology of the Black Sea
The Black Sea basin probably owed its origin to a marginal basin, south of the Eurasian continent, and behind a Cretaceous island arc, which extended from Bulgaria, through Anatolia, to the Caucasus (Hsii, Nachev & Vuchev, 1 977). The sedimentary sequence of the Black Sea is probably more than 10 km thick (Neprochnov, Neprochnova & Mirlin, 1 974). A deep-sea drilling cruise to the Black Sea bored six holes at three sites, with a maximum penetration of 1 073·5 m at Site 380 (Ross, Neprochnov et a!., 1 975). A generalized stratigraphy of the section penetrated at DSDP Site 380 is shown by F ig. l . The sediments are believed by us to range from Late Miocene to Quaternary in age, on the basis of interpreting palaeomagnetic stratigraphy, climatic variation, sedimentation rate, and other data (Hsii, 1 978). It should be noted that this interpretation is tentative and has not been accepted by a consensus of the shipboard staff (see Ross, 1 978). The Black Sea was a part of the ancient Tethys and a site of marine sedimentation until the Middle Miocene when its connections to the Mediterranean were severed (Fig. 2a). The restricted sea of eastern Europe extended from Vienna to regions beyond the Aral Sea and was called the Paratethys. It became brackish with time and its salinity at the start of Late Miocene was reduced to 1 5-20%0 (Koj umdgieva, 1 976) when black shales were deposited. With the disintegration of the Paratethys the Black Sea became a giant lake (Fig. 2b). Periodic chemical sedimentation took place under changing climate. The youngest sediments of the Black Sea are upper Quaternary and mainly terrigenous, laid down in a freshwater or brackish environment.
Late Neogene chemical sedimentation in the Black Sea
Salinity
Freshwater
Climate
Time
13 1
Lithologic Units
Fig. I. Stratigraphy, climate and water salinity of the deep Black Sea. (Modified from Figures in Hsii, 1978; Traverse, 1978; Schrader, 1978.)
Chemical sediments of the Black Sea
The late Neogene chemical sediments of the Black Sea are intercalated in a dominantly marly sequence, which is more than 600 m thick at Site 3 80 on the western side of the abyssal plain (Ross, Neprochnov et a!., 1 978). The sequence has been divided into the following units on the basis of the different carbonate interbeds (Fig. 1); they are in descending order: (7) Upper Siderite Unit; (6) Upper Chalk; (5) Lower Siderite Unit; (4) Lower Chalk; (3) Aragonite Unit; (2) Gravels and Dolomites; ( 1 ) Laminated Carbonates, including dolomite.
132
Kenneth J. Hsii and Kerry Kelts
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Fig. 2. Palaeogeography of the Tethys and Para tethys region. (a) Middle to Late Miocene. A, Atlantic; M, Mediterranean; P, Paratethys. (b) Late Miocene (Messinian). Vertical lines are Mediterranean, horizontal are Paratethys remnant lakes.
These sediments were deposited in a large lake of considerable water depth. The only exceptions are the Late Miocene gravels and dolomites, which are deposits of shallow subaqueous or subaerial origin (Ross, Neprochnov et a/. , 1 978). The intercalation of these sediments in a deep-water sequence is believed to signify an episode of desiccation when there was a hydrographic deficit for the Black Sea, and much of the Paratethys water was drained into the desiccated Mediterranean (Hsu, 1 978; see also Fig. 2b). The late Neogene was a time of gradual climatic deterioration. The cooling trend led to the extinction of thermophyllic floras during the Late Miocene. With the onset of glaciation in northern Europe and in the Alps and Caucasus, forest vegetations were replaced by steppe floras. Traverse ( 1 978) portrayed the climatic fluctuations with the help of steppe index, which is the percentage of steppe pollen in a total pollen assemblage (see Fig. 1). Chemical sedimentation started in Late Miocene when the climate was still warm, continued throughout the Pliocene, reached a zenith of chalk
Late Neogene chemical sedimentation in the Black Sea
133
deposition during the Alpha Glacial Stage in the early Pleistocene, and was replaced by terrigenous sedimentation at the beginning of the Beta Glacial Stage* . The correlation of lithology with palaeoclimate suggests that calcite was on the whole deposited during colder and siderite during warmer tim.::s . Diatom floras are common in the sediments of Black Sea. They indicate that the lake water was mostly fresh or brackish; only rarely did the salinity approach that of seawater (Schrader, 1978; see also Fig. 1). C H ALK S
Calcite is present as a chemical sediment in practically all units, but the precipitation of calcite exclusive of all other carbonate minerals took place only during the times when the Upper and Lower Chalks were laid down. Nearly pure calcitic sediments occur either as paper-thin laminations in varves (Fig. 3), or as homogeneous layers (Fig. 4).
Figs 3 and 4. (3) Cycles of Laminated Chalks alternating with grey marl. Lower Chalk Unit, Black Sea.
(Sample 380A-5 5 - l / 1 08- 1 25 em.) (4) Cycles of homogeneous chalk alternating with grey marl. Upper Chalk Unit, Black Sea, showing the sharp basal contact and a burrowed transition zone (Samples 379A-58-4/35-50 em corresponds to cycles b and c referred to in Table 1). Scales in centimetres. *Traverse proposed a terminology specifically for the Black Sea because of the difficulty of correlating D SD P data with established European glacial stages.
1 34
Kenneth J. Hsii and Kerry Kelts
Laminated chalk
Laminated chalks are characterized by dark laminae rich in organic matter (up to 2%) and light laminae rich in calcite. Several samples were examined under the scanning electron microscope (Fig. 5). The calcitic grains are nearly uniform in size, ranging from 5 to 1 5 Jlm. The blocky subhedral crystals show well-defined crystal edges. Twinning is present, but uncommon. Surfaces of crystals show signs of selective etching or fluting through dissolution; skeletal surfaces with etch pits resulted from advanced dissolution. Some small, relic, rod-shaped crystals 0·2-0· 5 Jlm in size are present (Fig. 6). They are calcareous, and may have been derived from the breakdown of larger crystals. Diatoms are present, and many tests show advanced degrees of dissolution.
Fig. 5. Lacustrine Chalk, Black Sea. (a) Dissolution features on euhedral to subhedral calcite polyhedra from a cyclically deposited Seekreide of the Upper Chalk. Note the very selective surface etching of the larger grains. Dissolution proceeds along preferred lattice directions and some faces are more resistant than others. X-ray diffraction indicates that some calcite in this almost pure chalk may contain up to 4% Mg in solid solution. (b) same, detail of grain surface. (c) same, some crystals are sculptured, other hollowed by dissolution. (Sample 380A- 1 2-4/7 1 -7 2 em.)
Laminated chalks are varve-like, similar to the varved chalks of Lake Zurich (Fig. 7). White laminae are essentially pure or very low Mg calcite, whereas the dark laminae consist of detrital clay-sized particles, diatom frustules and organic material. Calcite grain-size, form and dissolution features also match those from Lake Zurich (Kelts & Hsu, 1 978). The similarity suggests that individual laminae-couplets represent an annual cycle. Calcite precipitated during the summer and detrital laminae were deposited during the winter. Cyclic deposition of chalk
Laminated chalks are most common in the Lower Chalk Unit. They are interbedded with grey marls (Fig. 3). The alternation indicates periodic environmental changes. Not all chalks are varves. In fact many of the chalks from the Upper Chalk Unit form thin homogeneous layers a few centimetres thick and these layers are interbedded with marls (Fig. 4). Examined under SEM, the chalk is found to consist of the same blocky polyhedra of subhedral calcite crystals as the laminated chalks. A qualitative estimate, however, indicates more advanced dissolution effects in the non-laminated beds.
Late Neogene chemical sedimentation in the Black Sea
135
Fig. 6 . Lacustrine Chalk, Black Sea. A n example o fdissolution of fine-grained calcite grains i n a greyish-tan homogeneous chalk, Upper Chalk Unit. (a) a rhombic surface with knobby dissolution textures. (b) selective in situ dissolution leaves a mere skeletal relic of a calcite grain. (Sample 3 80A- 1 7-6/49-5 1 em.)
Fig. 7. Recent non-glacial varves, Lake Zurich. Dark layers comprise diatom frustules, organic sludge and detrital minerals; light layers mainly non-magnesian calcite crystals, 2-5 lim large in diverse stages of dissolution. Thick layer at the base is a lutite bed deposited from a low-density turbidity current in 1 9 1 8. Scale in centimetres.
136
Kenneth
J.
Hsii and Kerry Kelts
The cyclic sediment commonly ranges from 2 to 8 em and is typically 4-5 em thick. The base of the cycle is invariably a clastic sediment, which may be a silt or mud. This sediment lies with a sharp contact on the chalk of the previous cycle. The silt, if present, grades upward into a dark grey mud or marl, which contains some calcite. A zone of burrows is commonly present between the mud (or marl) and the chalk, the burrows having a light-coloured infilling from the overlying chalk. (Fig. 4). The burrows belong to the ichnogenus Chondrites, which is known mainly from marine sediments (Ekdale, personal communication). However, the annelids that produced the Black Sea burrows probably belonged to the genus Nereis, which is present in the Caspian Sea today (Zenkevich, 1 957). Nereis can compete successfully where there is insufficient circulation, as it has the remarkable ability to withstand short periods of absence of oxygen. The highest lithology of a cycle is the light-grey chalk. The chalk was probably also laminated before it was reworked by burrowing. The burrowing processes were apparently interrupted by the turbidity-current deposition of the overlying terrigenous clastics; the upper contact of the chalk is sharp and is not affected by bioturbation. So the cycle repeats itself.
Table I . Isotopic composition of calcite in cyclically deposited mud and chalk, Black Sea. Analyses by J. McKenzie, ETH, Zurich
Cores
"0
PDB
"C PDB
Cycle a 379A/5 8-4-27-28, chalk 379A/5 8-4-29-30, burrowed chalk 379A/5 8-4-30-3 l , mud 379A/5 8-4-32·5-33·5, silt
- 5· 1 2 - 5 ·57 - 5 ·86 - 6·22
+ 2·22 + 1 · 64 + 0·64 - 0·58
Cycle b 379A/5 8-4-3 8-39, chalk 379A/5 8-4-39-5-40·5, burrowed chalk 379A/5 8-4-4 l-42, mud
- 4·94 - 5 ·34 - 5 ·73
+ 1 ·60 + 1 ·37 + 0·00
Cycle c 379A/5 8-4-48-49, chalk 379A/58-4-50-5 l , mud
- 5 ·20 - 5 ·53
+2·19 + 0·57
Cycle d 3 80A/l 4-3-4-5, chalk 380A/ l 4-3-5·5-6·5, mud
- 5 -46 - 5 ·75
+ 1 ·39 + 1 ·39
Cycle e 380A/ l 4-3-7 ·5-8·5, chalk 380A/ 14-3-9·5- 1 0·5, chalk 380A/14-3- 1 2·5- 1 3 ·5, mud
- 5 ·52 - 5 ·44 -5·19
+ 1 · 72 +2·18 + 1 -49
Cycle f 380A/ 1 4-3- 1 4·5- 1 5 ·5 , mud
- 5 ·35
+ 1 · 96
The isotopic composition of calcite in the cyclical deposits has been determined and the results are shown in Table 1. The changes of the o180 and oUC values are systematic; the calcite in the chalk almost invariably has a less negative o180 and more positive oUC value than the underlying mud: the change in o180 is more than 2 per
Late Neogene chemical sedimentation in the Black Sea
137
mille and was apparently caused by fractionation a s a consequence o f increasing evaporation. The increase in o 1 3C is up to 3°,00 and the trend is similar to that found in cyclically deposited Cretaceous pelagic sediments of the southern Alps, where the change was related to increasing stagnation (H. Weissert, personal communication). Exclusion of clastics
A prerequisite for chemical sedimentation is the exclusion of terrigenous clastics. For example, Lake Zurich has been a site of chalk sedimentation since the turn of the century, because the waters emptying into the lake contain hardly any suspended particles. The detritus from the main tributary, the Linth, has been deposited in the lakes upstream (Walensee, Obersee) and the sediments from the minor side-tributaries have been trapped behind small-reservoir dams which were built for the purpose of flood-control prior to the turn of the century. Considering the material balance of the Black Sea, we note that the inflowing rivers are now supplying annually 1 50 X 1 06 tons of detritus (Shimkus & Trimonis, 1 97 4). If the influx is evenly spread out over the 0-42 x 106 km of the area, the rate of terrigenous sedimentation should be 1 5 em per 1 000 years. However, the rate of terrigenous accumulation on the abyssal plain is only 6 em per 1 000 years (t.y.). The discrepancy between the calculated and actual rates indicates that more than two thirds of the detritus has been trapped on the shelf in shallow waters (Ross & Degens, 1 974). The present slow rate of terrigenous influx permits the deposition of biogenic calcitic sediments (Muller & Stoffers, 1 974; Stoffers & Muller, 1 978). The rate of terrigenous sedimentation during the last glacial stage was as high as 90 cm/t.y.; the detrital supply then must have been at least six times greater than the present and the bulk of that was dumped onto the abyssal plain. Any pelagic chemical or biogenic components produced were obscured by the detrital influx. One was thus inclined at first to attribute a critical role to climatic factors; chemical sediments should only be forming during tern perate to warm preglacial or interglacial times, when erosion was moderate and when much of the terrigenous influx was deposited on the shelf. It was, therefore, a great surprise when the palynological results (Traverse, 1 978) proved that the main period of the chalk deposition took place during a glacial stage! A regional synthesis has revealed that the palaeogeography of eastern Europe during the early Pleistocene was significantly different from that of the present (Hsu, 1 978). The Bosphorus was then not yet open, and the Danube was emptying into a lake in Romania east of the Carpathian Range. Although the Danube water may have been emptied into the Black Sea, the Danube detritus, which constitutes more than half of the terrigenous influx to the Black Sea, was then trapped in this peri-Carpathian lake. The Dnestre detritus might also have been directed there, and that of the Dneper and Kuban may have been trapped by the Sea of Azov. Deprived of those sources, the Black Sea received annually only about a quarter as much terrigenous detritus as it does now, so that the rate of terrigenous sedimentation during an early Quaternary glacial stage should have been about 22 cm/t.y. The annual load of dissolved Ca2 + of rivers draining into the Black Sea is 14 X 1 06 tons (Shimkus & Trimonis, 1 954) and may have been similar in quantity then. If this load was combined with dissolved carbonate and deposited in the Black Sea, the rate of chalk deposition should have been about 9 cm/t.y. Thus, the total rate for clastic and chalk sedimentation should
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have been 3 1 cm/t.y. This rate is about the same as that calculated on the basis of the thickness ofthe early Pleistocene sediments (Hsii, 1 97 8). Also the ratio (9:22) of calcitic to terrigenous sedimentation coincides more or less with our visual estimates of the proportion (about 1 :2) of carbonate to clastic sediments. This analysis suggests that the deposition of the Upper Chalk Unit during the Alpha stage of glaciation took place because much of the detritus carried down by rivers was trapped in peripheral basins. The Lower Chalk was deposited during a warm preglacial epoch at 5·4 cm/t.y. This low rate indicates not only a trapping of clastics in peripheral basins, but also that the rivers carried much smaller terrigenous and dissolved loads than they do now. Depositional environments of chalks
The Lower and the Upper Chalks were deposited in considerably different environments. The Lower Chalk was deposited during relatively warm Pliocene time; the Upper Chalk was deposited during a glacial stage. The Lower Chalk has a rich planktonic diatom-flora and the Melosira-Stephanodiscus assemblage is typical of an oligohaline environment (3-5o/00)(Schrader, 1 97 8). Diatoms are very abundant in some layers, constituting more than half of the bulk volume. The Upper Chalk contains an unusual planktonic flora of freshwater dinoflagellates (Traverse, 1 97 8). Diatoms are commonly absent; they are either not produced, or they are not preserved from dissolution (see Fig. 8). The one exception is an almost monospecific assemblage ofthe freshwater diatom Melosira undulata at one level, indicative of a water salinity less than 3o/00 (Schrader, 1 97 8). The Lower Chalk contains no benthonic fossils. The Upper Chalk contains a rich benthonic ostracod fauna; the Candona-Loxoconcha assemblage indicates a deep freshwater lake environment like that of the Great Lakes of North America (Olteanu, 1 97 8). The chalks of the lower unit are commonly laminated, and the structure indicates completely stagnant bottom conditions, devoid of life and
Fig. 8. Dark lamina in a varve, Lower Chalk, Black Sea. (a) Dark greenish-grey part of a lamina couplet in
the Lower Chalk shows mostly layered silicates and some scattered diatom frustule remnants with some quartz and low Mg calcite. (b) as above. In some rare instances there is evidence of overgrowth in diatom pores rather than dissolution. (Sample 380A-49-5/ 1 1 9- 1 20 em.)
Late Neogene chemical sedimentation in the Black Sea
139
current activity. The chalks of the upper unit contain both laminated and structureless varieties; the latter has been burrowed by annelids in bottom environments not completely devoid of oxygen. The Upper Chalk also contains abundant silts and mud layers, which were probably deposited by low-density turbidity currents. The Black Sea lacustrine environments of chalk deposition had much in common with the environments of Lake Zurich. The lake has been eutrophic since the turn of the century, and no benthonic life exists on the deep lake floor. Despite the general stagnation of the deepest bottom, partial annual overturn of the lake water is still taking place. The water movement at the beginning of the spring brings cold and bicarbonate-rich deep water to the surface. C alcite precipitation takes place in the late spring and throughout the summer in response to seasonal temperature and pH changes (Kelts & Hsu, 1 978). A maj or diatom bloom occurs early in the spring; diatom tests sink to the bottom, where they form the base of an annual varve. Coarser calcite crystals are also formed during the early to late spring. The summer harvest of calcite tends to be fine-grained. This chalk is overlain by a dark lamina of organic-rich clay deposited during the late autumn and winter. The lack of bioturbation permits the preservation of annual deposits as carbonate varves (Fig. 7). The present-day Lake Zurich environment is thus an excellent analogue of Lower Chalk deposition, except the Lake Zurich water is probably even less saline. The stagnation of the deepest bottom of Lake Zurich is caused by lack of bottom circulation. Studies of the water movement in Swiss lakes indicated that bottom current activities are commonly related to the emplacement of density underflows (Lambert & Luthi, 1 977). Cores from Lake Zurich show no major turbidite sedimentation after 1 900. Flood-control constructions since that date have apparently prevented the continuation of stream-floods on land as turbidity underflows in the lake. The absence of turbidity-current activity may in fact be responsible for the oxygen-deficiency of the lake. Stagnation of the Black Sea at the time of the Lower Chalk deposition might be likewise explained, because this unit contains practically no turbidites; a density-stratification of the slightly brackish water may have contributed further to the stagnation. The present chalk-deposition of Lake Zurich is not a unique event during the Holocene. The previous period was the Preboreal to Boreal time, 1 0,300-7,500 years ago. This older unit consists mainly of light grey chalk (65-85% calcite) with intercalations of silt laminae and was deposited at a time when there was little detrital input to Lake Zurich (Kelts, 1 97 8). Those chalk beds are mostly bioturbated and homogeneous. Like the Boreal chalk, the Upper Chalk of the Black Sea is also characterized by the predominance of a non-laminated variety; its lack oflamination can also be attributed to bioturbation. The Boreal chalk of Lake Zurich and the Upper Chalk of the Black Sea are furthermore similar because diatoms are rarely preserved or altogether absent in these sediments. One maj or difference is the prevailing climate. The Upper Chalk of the Black Sea contains mainly steppe pollen, indicative of deposition during a glacial stage. The Boreal sediments contain abundant tree pollen, and were deposited during a post-glacial climatic optimum. The data confirm our interpretation that the exclusion of clastics, not climate, is the critical factor for chalk deposition. This condition was achieved for the Upper Chalk deposition in the Black Sea, when the climate was cold and arid, and for the Boreal chalk of Lake Zurich when the climate was warm and when erosion was minimized by forest-growth. Another factor is the
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Kenneth J. Hsii and Kerry Kelts
physical limnology. A freshwater lake in the Black Sea basin may have had a thermal structure similar to that of Lake Zurich during the Boreal time. Glaciers were far away, but the lake water was cold in the winter and warm in the summer. Surfacing of cold waters rich in dissolved carbonate during the annual spring-overturn and the subsequent warming provided the condition for supersaturation and precipitation of calcium carbonate. Cause of cyclic deposition
Cyclical deposits of the Upper Chalk in the range of centimetres in thickness cannot be annual varves. The cycles signify climatic oscillations of a few hundred years in period. However, the influence of climate is complicated. The Black Sea cycles were initiated when chalk sedimentation was interrupted by the deposition of terrigenous silts, carried to the abyssal plain by turbidity currents. The intrusion of the underflows, like their counterpart in Swiss lakes, has apparently promoted bottom circulation. With the exclusion of turbidites the bottom became more stagnant, while chalk deposition began. The periodical increase of the terrigenous influx was probably related to periodic retreat ofglaciers during the Alpha stage of glaciation. Pollen studies of cyclically deposited sediments by Traverse ( 1 978) indicated that the steppe pollen are dominant in the chalk whereas the forest pollen are dominant in the terrigenous layers of the cyclic units. The steppe index (S.I.) reached as high as 83 in a chalk (sample 3 80A/ l 2/4/4 1 -42 em), but went down to as low as 1 in the terrigenous sediment 2 em below. Other S.l. values include variations from 73 to 19, 17 to 22, 48 to 36, 5 5 to 9, 83 to 35, 93 to 1 8, 67 to 67, 85 to 1 0, 63 to 60 in the chalk and mud layers, respectively; all except one show that the chalk was deposited during colder and mud during warmer times of short-termed climatic cycles. Such drastic changes of climate took place in cycles of several hundred years; the periodicity is comparable to the Holocene climatic variations as evidenced by the advances and retreat of mountain glaciers in the Swiss Alps, where four 'miniglacial' and 'mini interglacial' stages during the last 2,000 years have been recorded (Schneebeli & Rothlisberger, 1976).
M AGNE SI AN C A LCI TE AN D AR AGONI TE
Magnesian calcite is present only in appreciable quantity in the Aragonite Unit. Aragonite is present in this unit, in the Laminated Carbonates unit, and near the top of the Upper Siderite unit. Aragonite Unit
The Aragonite Unit was deposited during the warm Early Pliocene time. Aragonite and magnesian calcite occur as thin layers of laminated sediments in a mudstone sequence. The laminae are less than 1 mm in thickness and have a varved structure. The Aragonite Unit contains brackish-marine floral assemblages. The nannofossil assemblage is monospecific, and consists of a Braarudosphaera species, which is tolerant of a less than normal marine salinity down to 22°/00 (Bukry, 1 974). The diatom assemblage A ctinocyclus ehrenbergli - Hermesinum adriaticum is mixo-euhaline, 300 40 /00 salinity (Schrader, 1 97 8). The absence of a foraminiferal and a diversified
Late Neogene chemical sedimentation in the Black Sea
141
Figs 9- 1 1 . (9) Aragonite, Laminated Carbonate Unit, Black Sea. Aragonite prisms i n a marl. X-ray diffraction indicates quartz and aragonite are the dominant minerals in this sample with minor amounts of low Mg calcite and ankeritic dolomite. Aragonite is particularly concentrated in clusters or radial sprays of subhedral prisms in the SEM. Note a tapering off of some prisms at one end. (Samples 3 80A-64-4/93-95 em.)
( 10) Dolomite in Gravel Unit, Black Sea. The euhedral dolomite rhombs occur together with low Mg calcite in thin white laminae from green marl stones in a mud-pebble horizon. The knobby grains are dissolution surfaces of a Mg calcite (5-7% Mg). Non-stoichiometric dolomite constitutes over 65% of the sample, with calcite (5%) and minor quartz (7%). Dolomite crystals appear to have grown diagenetically as replacement and interstitial filling. Several grains interpenetrate and many are twinned. (Sample 380A-58-l /6 1-63 em.) ( l l a) Siderite from Upper Siderite Unit, Black Sea. The sample is from a reddish brown hard layer, and shows smooth, rounded, anhedral siderite crystals growing on a substratum. X-ray diffraction indicates that the sample contains only 1 0% manganosiderite, in a calcitic sediment, with quartz and some dolomite. (Sample 380A- l -3/28-3 l em.) ( l i b) Siderite from Lower Siderite Unit, Black Sea. The sample is from a laminated reddish brown hard layer in pale grey sediments. Manganosiderite, with up to 1 5 % Mn in solid solution, constitutes over 80% of the bulk. It occurs as irregular 'wheat grains' suggestive of a primary rather than replacement origin. Fine grained matrix contains mainly quartz and iron oxides. (Sample 380A-37-3/57-58 em.)
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nannofossil assemblage indicates that the Black Sea then, like now, was not exactly marine. There was probably a density layer then, as now, with a stagnant bottom where varve-like laminations could be preserved. The bulk of the carbonates in the Aragonite Unit are aragonite and magnesian calcite. The aragonite occurs as stubby, imperfect prisms, commonly clustered into sprays with a tapered shank (Fig. 9). The magnesian calcite contains commonly 6% Mg, but more than 1 2% Mg in some instances. That aragonite and Mg calcite should occur in place of calcite indicates an environment where abundant Mg2 + ions inhibit the crystallization of the stable phase calcite (e.g. Berner, 1 975). Mudstone was deposited in an environment different from that of laminated carbonates. Diatoms are common in the mudstone, but rare or absent in the carbonate. Apparently the water was alternating from CaC03 saturation and Si02 undersatura tion to CaC03 undersaturation and Si02 saturation. The cause ofthe changes in water chemistry was ultimately related to climatic variations; an excess ofCa2 + was imported during temperate times, but an excess of dissolved Si02 was imported because of deep weathering under tropical or subtropical conditions. Laminated Carbonate Unit
Aragonite, but not magnesian calcite, is present in the Laminated Carbonate Unit, which was deposited at times of fluctuating lake-levels and changing salinity within the Black Sea basin. The alternation of calcite, dolomite, and aragonite sedimentation also indicates changes in water-chemistry. Calcite was precipitated when the lake was freshwater or oligohaline, and aragonite precipitated when the Mg2 + ion concentration was high. DOLOMITE
Dolomite in the Gravel Unit is stromatolitic, similar to that formed in modern sabkha environments. SEM investigations (e.g. Fig. 1 0) show euhedral, smooth, rhombs 4- 10 p,m in size, some twinned, which appear to have been formed in pore spaces of intertidal sediments. Dolomite in the Laminated Carbonate Unit was formed in a lake of fluctuating water-levels. lts origin may be similar to the Holocene dolomite of Lake Balaton (described by Muller, Irion & Forstner, 1 972; Muller & Wagner, 1 978). SI DERITE
Siderite occurs as hard layers or as soft muds. Some hard layers are apparently concretionary, with encrusting anhedral crystals growing on a substratum. Others are apparently cemented mud, consisting of wheat-grain-like crystals in a finer-grained mat:ix. Depositional environments of siderite
The steppe index indicates warm climate at the time of siderite deposition (Fig. 1 ). With a warming trend after a cold episode during the Pliocene, Lower Chalk deposition was replaced by Lower Siderite. Similarly the change from the Alpha-
Late Neogene chemical sedimentation in the Black Sea
143
Glacial to the A(Anna)-Interglacial Stage coincides with the change from the Upper Chalk to the Upper Siderite sedimentation. Diatoms are common in both units of sideritic deposits. The Lower Siderite has an oligohaline assemblage, and the Upper Siderite has an assemblage indicative of oligohaline to mesohaline conditions. The diatomaceous sediments are commonly grey and laminated. The light grey diatom-rich laminae contain up to 80% diatoms whereas the clay-rich contain only 10 or 1 5 % diatoms. The existence of a rich diatomaceous flora suggests a eutrophic water-mass, and the laminated structure indicates anoxic bottom waters. Origin of siderites
Siderite occurrences are rare in Holocene environments. A small amount of siderite is formed diagenetically in the brackish-water sediments of Chesapeake Bay (Bricker & Troup, 1 97 5), and layered siderite occurs in Late Quaternary deposits of Lake Kivu in equatorial Africa (Stoffers, 1 975). The conditions for the Black Sea siderite deposition were comparable to that of Lake Kivu. The climate of a middle latitude region like the Black Sea during an interglacial stage was similar to the climate of the tropical Lake Kivu during a glacial stage; it was temperate rather than tropical. The eutrophic watermass of the Black Sea then was evidently low in Eh, as Lake Kivu was 1 2,000 years ago. That an iron carbonate should form instead of a calcium carbonate requires that the Ca/Fe ratio of the water be less than 20/ I (Berner, 1 97 1 ). Sluggish streams of the southeastern United States where relief is low and rainfall is high, have a high content of organic matter, iron and silica, and a low content of other dissolved ions such as Ca2+ and MgH (Beck, 1 972). The present-day stream from the coastal plain of Georgia, for example, has a Ca/Fe of about I (Garrels, Perry & McKenzie, 1 973). The rivers draining into the Black Sea today have a Ca/Fe ratio far higher than 20 (Shimkus & Trimonis, 1 974). However, their water-chemistry may have been different during the times of siderite deposition. We can assume that the low-lying coastal plains surrounding the Black Sea were then heavily forested, the prevailing climate being similar to that of the southeastern United States. A deep chemical weathering led to a breakdown of silicates so as to yield abundant iron and silica to river waters. The formation of iron carbonate in place of iron sulphide in a reducing environment suggests that the sulphate ion concentration of the Black Sea was low. Since sea water is rich in dissolved sulphate, the salinization of the Black Sea during the time of siderite deposition was probably not caused by an invasion from the Mediterranean, but resulted from evaporitic excesses. A marine invasion terminated siderite formation, and led to the deposition of pyritic and aragonitic muds on top of the Upper Siderite. Sideritic concretions are a diagenetic phenomenon, formed at or near the sediment-water interface like the siderite of Chesapeake Bay. On the other hand, wheat-like siderite grains in the soft sideritic muds are probably subaqueous precipitates.
SUMM ARY
Chemical sedimentation in the Black Sea started in Late Miocene when the water body was changing from a brackish sea into a lake. During the first stage of isolation,
1 44
Kenneth J. Hsii and Kerry Kelts
the lake was being progressively desiccated. Oscillations of water level and variations of water chemistry resulted in the deposition of a wide assortment of sediments. Calcite was precipitated when the water was fresh; dolomite, magnesian calcite and aragonite were deposited when the water was brackish to saline. The sediments are commonly laminated and indicate stagnant bottom conditions. Eventually the edge of the Black Sea abyssal plain was exposed and the Laminated Carbonates were overlain by gravels, sands and stromatolitic dolomites. At the beginning of the Pliocene, the Black Sea was again submerged; aragonite and magnesian calcite were deposited. The marine connection was, however, soon severed, and the lake was rapidly desalinified. The Lower Chalk Unit includes both calcitic and diatomaceous sediments. The water mass was probably eutrophic and slightly brackish, and the bottom stagnant. Chalks are commonly laminated like the carbonate varves of Lake Zurich. During a warming trend of the late Pliocene, the lake became brackish. The influx of iron-rich waters led to the formation of siderite. With the onset of continental glaciation in northern and central Europe and a changed climate, the Black Sea was again desalinified. Cyclically deposited chalks were laid down then, when the bottom was not completely anoxic. After a change from a glacial stage to an interglacial, chalk sedimentation was replaced for a second time by deposition of siderite. The latter was terminated by a marine invasion during the middle Quaternary; aragonite was then precipitated. The interval of periodic chemical sedimentation was ended soon afterward, when the D anube brought in enough clastics to herald an epoch of terrigenous sedimentation. However, the Black Sea remained a great lake until the Bosphorus Strait came into existence and marine waters spilled over. During the last glaciation, when the sea level dropped below the sill of Bosphorus, the Black Sea basin again became a freshwater lake. Finally, as the Holocene sea level rose, Mediterranean waters re-entered via the Bosphorus. The Black Sea became a brackish lagoon; the density-layering of the watermass led to the present bottom stagnation and anoxic condition.
REFERENCE S BECK, K.D. ( 1 972) Sediment-water interactions of some Georgia rivers and estuaries. Office of Water Resources Project B-033GA, Rept. Georgia Inst. Tech. BERNER, R.A. ( 1971) Principles of Chemical Sedimentology. McGraw Hill, New York. BERNER, R.A. ( 1 975) The role of magnesium in the crystal growth of calcite and aragonite from sea water. Geochim. cosmochim. A cta, 39, 489-504. BRICKER, O.P. & TROUP, B.N. ( 1 975) Sediment-water exchange in Chesapeake Bay. In: Estuarine Research (Ed. by L. E. Cronin), pp. 3-27. Academic Press, New York. BUKRY, D. ( 1 974) Coccoliths as paleosalinity indicators. In: The Black Sea, Geology, Chemistry and Biology (Ed. by E. T. Degens and D. A. Ross). Am. Ass. Petrol. Geo/. Mem. 20, 353-363. GARRELS, R.M., PERRY, JR, E.A. & McKENZIE, F.T. ( 1973) Genesis of Precambrian iron formations and the development of atmospheric oxygen. Econ. Geol. 68, 1 1 73- 1 179. HsO, K.J. ( 1 978) Stratigraphy of the lacustrine sedimentation in the Black Sea. In: Initial Reports of the Deep Sea D rilling Project, Vol. XLIIB (D. A. Ross and Y. Neprochnov et a/.). U.S. Government Printing Office, Washington D.C. (In press.) HsO, K.J., NACHEV, I.K. & VucHEV, V.T. ( 1 977) Geologic evolution of Bulgaria in light of plate-tectonics, Tectonophysics, 40, 245-256. KELTS, K . ( 1 978) Geological and sedimentary evolution of Lakes Ziirich and Zug, Switzerland. Dissertation, ETH, Ziirich, NR 6 1 46. KELTS, K. & HsO, K.J. ( 1 978) Calcium carbonate sedimentation in freshwater lakes and the formation on non-glacial varves in Lake Ziirich. In: Lakes, Physics, Chemistry, Geology (Ed. by A. Lerman). Springer-Verlag, New York. (In press.)
Late Neogene chemical sedimentation in the Black Sea
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E. ( 1 976) Paleoecologie des communautes des mollusques du Miocene en Bulgarie du Nord Ouest. IV Communautes des Mollusques du Bessarabien et du Chersonien (Sarmatien moyen et superieur). Geologica Balk. 6, 37-56. LAMBERT, A. & L OTH !, S. ( 1977) Lake circulation induced by density currents: an experimental approach. Sedimentology, 24, 735-7 4 1 . MOLLER, G . & STOF FERS , P. ( 1 974) Mineralogy and Petrology o f Black Sea Basin sediments. I n : The Black Sea, Geology, Chemistry and Biology (Ed. by E. T. Degens and D. A. Ross). Am. A ss. Petrol. Ceo/. Mem. 20, 200-248. MOLLER, G., IRION, G. & FORSTNER, U. ( 1972) Formation and diagenesis of inorganic Ca-Mg carbonates in the lacustrine environment. Natunvissenschaften, 59, 1 5 8 - 1 64. MOLLER, G . & WAGNER, F. ( 1 978) Holocene carbonate evolution in Lake Balaton (Hungary): a response to climate and impact of man. In: Modern and A ncient Lake Sediments (Ed. by A. Matter and M . E. Tucker). Spec. Pubis int. Ass. Sediment. 2, 57-8 1 . NEPROCHNOV, Yu. P., NEPROCHNOVA, A.F. & MtRLIN, YE. G . ( 1 974) Deep structure o f Black Sea Basin. In: Black Sea, Geology, Chemistry and Biology (Ed. by E. T. De gens and D . A. Ross). Am. Ass. Petrol. Ceo/. Mem. 20, 35-49. OLTEANU, R. ( 1978) Ostracoda from DSDP Leg 42B . In: Initial Reports ofthe Deep Sea Drilling Project, Vol. XLIIB (D. A. Ross and Y . Neprochnov et al.). U.S. Government Printing Office, Washington D.C. (In press.) Ross, D . A. & DEGENS, E.T. ( 1 974) Recent sediments of the Black Sea. In: The Black Sea, Geology, Chemistry and Biology (Ed. by E. T. Degens and D. A. Ross), Am Ass. Petrol. Ceo/. Mem. 20, 1 8 3 - 1 99. Ross, D.A. ( 1 978) Black Sea Stratigraphy. In: Initial Reports of the Deep Sea Drilling Project, Vol. XLIIB (D. A. Ross, Yu Neprochnov et a/.). U.S. Government Printing Office, Washington D.C. (In press.) Ross, D .A., NEPROCHNOV, Yu. et a/. ( 1 975) G LOMAR CHALLENGER drills the Black Sea. Geotimes, 20/10, 1 8-2 1 . Ross, D . A., NEPROCHNOV, Y u . et a/. ( 1978) Initial Reports ofthe Deep Sea Drilling Project, Vol. XLIIB. U.S. Government Printing Office, Washington D.C. (In press.) SCHNEEBELL, W. & ROTHLISBERGER, F. ( 1976) 8000 Jahre Walliser Gletschergeschichte. Die A /pen, Zeitschrift Schweiz. A lpenclub, 52, 3/4, 5- 1 52. ScHRADER, H.J. ( 1978) The Diatom Units and the Paleogeography of the Black Sea in the Late Cenozoic (DSDP Leg 42B). In: Initial Reports of the Deep Sea Drilling Project, Vol. XLIIB (D. A. Ross, Yu Neprochnov et a!.). U.S. Government Printing Office, Washington D.C. (In press.) SHIMKUS, K.M. & TRIMONI S, E.S. ( 1 974) Modern sedimentation in Black Sea. In: The Black Sea, Geology, Chemistry and Biology (Ed. by E. T. Degens and D. A. Ross). Am. Ass. Petrol. Ceo/. Mem. 20, 249-278. STOFFERS, P . ( 1 975) Sedimentpetrographische, geochemische und paldoklimatische Untersuchungen an ostafrikanischen Riftseen. Habilitationsschrift UniversiUit Heidelberg. STOFFERS, P. & M O L LER, G. ( 1978) Mineralogy and Lithofacies of Black Sea Sediments, Leg 42B, Deep Sea Drilling P roj ect. In: Initial Reports of the Deep Sea Drilling Project Vol. XLIIB (D. A. Ross. Yu Neprochnov et a!. ). U.S. Government Printing Office, Washington D.C. (In press.) TRAVERSE, A. ( 1 978) Palynological Analysis of DSDP Leg 42B ( 1975) Cores from the Black Sea. In: Initial Reports of the Deep Sea Drilling Project, Vol. XLIIB (D. A. Ross, Yu Neprochnov et a/. ). U.S. Government Printing Office, Washington D.C. (In press.) ZENKEVICH, L.A. ( 1 957) Caspian and Aral Seas. In: Treatise on Marine Ecology and Paleoecology (Ed. by J. W. Hedgpeth). Mem. Ceo!. Soc. A m. 67, 89 1-9 1 6. KOJUMDGIEVA,
Spec. Pubis int. Ass. Sediment. ( 1 978)
2,
147- 168
Turbidites and varves in Lake Brienz (Switzerland): deposition of clastic detritus by density currents
M I C H A E L S T U R M * and A L B E R T M A T T E R Geologisches Institur der Universitdt Bern, Sahlistrasse 6, CH-3012 Bern, Switzerland
AB STR ACT Lake Brienz is a 14 km long and 26 1 m deep oligotrophic valley lake which lies in the front ranges of the Swiss Alps. Sedimentation is entirely clastic and is dominated by two rivers which enter the lake at opposite ends. The sediment load is transported and deposited in the lake by overflows, interflows and underflows (low- and high-density turbidity currents) depending on the density difference between river and lake water. Whereas high-density turbidity currents, which deposit up to 1 50 em-thick graded sand layers. occur only once or twice per century after catastrophic flooding. low-density turbidity currents occur annu ally during periods of high discharge and deposit centimetre-thick faintly graded sand layers. Fine-grained sediment supplied by overflows and interflows rains down continuously during summer thermal stratification to form the dark-grey summer half-couplet of a varve; at turnover in the autumn the remaining sediment trapped at the thermocline settles out and forms the light-grey winter layer. Turbidites grade distally into thin dark-grey layers indistinguishable from the dark-grey summer half-couplet. Turbidites on the basin plain can be correlated with varves on the slopes. Therefore, in Lake Brienz the formation of varves and turbidites is genetically related and depends on the existence of over- and interflows. turbidity currents and seasonal thermal stratification.
INTRO DU CTION
While working on turbidite sequences of ancient Flysch basins in the Alps we became interested in turbidity currents and their deposits that occur today in lakes. First observed in lakes by Forel ( 1 8 85) and Heim, Moser & B iirkli-Ziegler ( 1 8 88), turbidity currents are now recognized as an important depositional process. In * Present address: Swiss Federal Institute of Technology, EA WAG, CH-8600 Diibendorf/Zii rich, Switzerland
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
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Michael Sturm and A lbert Matter
comparison with ocean basins, lakes can be regarded as closed, natural sedimentation laboratories whose parameters are comparatively easy to measure and monitor, and hence to a certain extent lakes can be used for modelling processes in the oceans. We believed that by applying modern coring and seismic techniques and sedimentological methods to lake sediments much information could be obtained about sedimentary regimes in lakes, especially on the origin, nature, abundance and distribution of turbidity currents, which would also be applicable to marine sedimentation. Furthermore, we wanted to study the origin of clastic varves in lakes and find out whether they are related to turbidity currents as postulated by Kuenen ( 1 95 1 ) and B anerj ee ( 1 968) or to overflows as suggested by Antevs ( 1 9 5 1 ).
SWITZ[RLANO SAMPLE STATIONS •
box core ( BK-
•
piston core 0 lOcm ( BP-
, 1n
o
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Fig. I. Index map of Lake Brienz showing bottom topography, sample stations and position of cross sections. The two main tributaries, the Aare and Liitschine Rivers, are at either end of the longitudinal basin. Note steep slopes and narrow shore terraces of the basin. Solid contour lines indicate 20 m intervals; dashed lines indicate 10 m intervals.
In order to investigate these problems Lake Brienz was chosen for this study, firstly because it is a deep longitudinal trough with a sediment source at each end and therefore has much in common with ancient flysch basins of the Alps (Hesse, 1 974), and secondly because it has a long hydrologic record, including a known input of sediment load through time and a well studied current pattern and physical limnology. Lake B rienz is one of several longitudinal valley lakes within the front ranges of the Swiss Alps. It is situated about 70 km southeast of Berne, 5 64 m above sea level (Fig. 1 ) The background data summarized in Table 1 show that 18% of the drainage basin is covered by glaciers and lies at a mean altitude of 1950 m. .
Turbidites and varves, Lake Brienz
1 49
Table 1. Physical parameters of Lake Brienz. Data from Nydegger ( 1 967) and from the Swiss Fed. Bureau of Water Management.
Maximum length Maximum width Maximum depth Mean depth Surface area Volume Altitude Drainage area Mean altitude of drainage area G laciated part of drainage area Ratio drainage area : lake area Aare water discharge Liitschine water discharge Total water discharge Mean water retention time Aare suspended load Liitschine suspended load Total suspended load
14 km 2·5 km 26 1 m 174 m 30 km' 5·2 km3 566 m above sea level 1 1 40 km' 1 950 m above sea level 1 8·2% 38 : l 33 m3 sec- ' (mean 1 954-1975) 1 9 m3 sec- ' (mean 1924- 1975) 61 m3 sec - ' (mean 1 935- 1975) 2·8 years 1 23·7 x 103 tons year - ' (mean 1965- 1 975) 1 24· 1 X 103 tons year - ' (mean 1 965-1975) approx. 285 x 1 03 tons year - ' (mean 1 965- 1 975)
Most of the catchment is drained by two major rivers, the Aare and the Uitschine, which carry enormous amounts of glacial and snow-melt waters and sediment into the lake, especially in spring and early summer (Table 1 ). Based on our experience in the neighbouring Lake Thun (Sturm & Matter, 1 972c) we expected that because the Aare drains largely crystalline terrane whereas the Liitschine drains sedimentary and crystalline terrane the rivers would carry detritus of distinctly different mineralogic composition, and also that the steep-gradient streams debouching from lateral slopes into the lake would influence sedimentation in only a minor way and could be neglected. This was confirmed during a study of the surficial sediments that revealed that each sample containing coarse silt to sand sizes could be attributed to either the Aare or Liitschine source and that sediment load of lateral tbrrents had very little influence on the sedimentation within the basin (Sturm, 1 976). Lake Brienz is a holomictic, oligotrophic lake (Nydegger, 1957, EAWAG, 1 967), and the lake sediments are therefore predominantly allochthonous in origin (Sturm, 1976). Thermal stratification develops from March to November with an average thickness of the thermocline of about 25 m. Current measurements made at several depths within the uppermost 60 m reveal a counter-clockwise rotation of the water mass (Nydegger, 1967, 1 976). This circulation pattern is attributed to the continuous driving force of the inflowing rivers which are deflected to the right as a result of the earth's rotation. Flow velocities measured at 1 9 and 25 m depth reach 5 ·6 em sec - 1 and 5·0 em sec - 1 (Nydegger, 1 967, 1 976). Nydegger also reported the existence of a turbid layer within the thermocline. He thought that it was built-up through accumulation of suspended matter carried into the lake by the rivers. The basin morphology of Lake Brienz is rather simple; a flat central basin plain is flanked by steep (30-40.) lateral slopes with a delta at each end of the lake. The shore terrace is generally very narrow or missing entirely where the shores are formed by cliffs. A slightly broader shore terrace occurs only near the village oflseltwald (Fig. 1 ).
1 50
Michael Sturm and A lbert Matter
The lateral slopes as well as the central basin plain show a smooth bottom morphology whereas in the delta areas distinct morphological features such as channels and levees are present. The channels on the Aare delta foreslope for example are up to 200 m wide and up to 35 m deep (Sturm, 1 975; Sturm & Matter, 1 972b). The lake trough has been eroded deeply into Mesozoic bedrock by fluvial and glacial action. The maximum depth of the basement is about 860 m below lake level i.e. approx. 300 m below sea level. During Pleistocene and Holocene times about 600 m of clastic sediments accumulated (Matter et al., 1973). We expected that turbid meltwater that enters the lake each spring would provide material for turbidity currents in Lake Brienz. By collecting a large number of cores from the entire basin we wanted to correlate individual turbidites across the basin, use the mineralogy to identify which of the two rivers was the source and then determine the path of each turbidity current and the amount of erosion, if any, at the base of each turbidite. If successful, we would thus be able to reconstruct the pattern of basin filling and arrive at a model for clastic sedimentation in such lakes.
METH O D S
In order to determine the sedimentary pattern and the sediment dispersal in the lake 80 cores and 140 surface samples were collected during several field campaigns between 1973 and 1 976. The cores were taken with a 3-6 m Reineck piston corer with l 0 em diameter (Reineck, 1967) and a modified 8- 1 0 m Kullenberg piston corer with a diameter of 6 em. The surface sediment cores were taken with a Reineck box corer (Reineck, 1 963) and a 1 m gravity corer. The locations of the coring sites were determined from sextant readings on distinct shore features. Water depths were measured by echo sounder. The cores were taken to the laboratory, cut lengthwise with an electro-osmotic guillotine (Sturm & Matter, 1 972a), photographed and sampled. Samples and cores were sealed in plastic bags or polyethylene foils and stored in a core library at about Ire. The analytical programme included grain-size analyses by standard sieve and sedimentation balance procedures, heavy mineral counts, gasometric analysis of the carbonate content, and clay mineral analysis of the clay size fraction ( < 4,um) by X-ray diffraction. The sequence of lithologies in each core and the core-to-core correlations were worked out using core photographs. During 1 973 and 1974 transmissometer and temperature profiles were measured at fixed stations in the Aare delta to identify turbid layers in the water column. At the same time 1 litre water samples were taken at various depths at these stations to determine the amount of suspended matter.
RE S U LT S Description of sediments
Inspection of a large number of cores revealed that the sediment lithology could be described in terms of four different sediment types, (a) delta sand and mud, (b) laminated mud, (c) homogeneous mud and (d) laminated mud with interbedded graded sand/silt.
Turbidites and varves, Lake Brienz
151
Fig. 2. Box cores from delta area of the Aare River: (a) core B K-42, 240 m water depth, from bottom of main channel, showing massive, poorly graded sand; (b) core BK-42, 240 m water depth, from levee of main channel showing alternations of thin sand and silt beds. Note lenticular shape of some of the layers. Scale in centimetres.
Delta sand and mud is found on the foreslopes of the Aare and Liitschine deltas. Coarse-grained well sorted massive sand beds occur in channels as 1 5-30 em thick layers with sharp lower and upper contacts (Fig. 2a). Occasionally pebbles up to 2 em in size were found at the base of these layers, and very rarely a faint lamination was seen. The delta muds consist of lenticular bedded sand/silt alternations with subordinate clay layers. The sandy and silty layers are commonly less than 1 ·5 em thick and frequently swell and pinch out over a distance of a few centimetres. Ripple cross bedding, usually characteristic of lenticular bedding (Reineck & Singh, 1975, p. 1 00) has not yet been recognized, but grading was observed in silt layers (Fig. 2b). Lenticular bedded sand/silt is found in the levee and interchannel areas as well as in the lowest foreslopes of the deltas. This kind of bedding was also reported from lake sediments in front of small developing deltas by Coleman ( 1 966, cit. in Reineck & Singh, 1975, p. 1 02). Because of the irregular topography in the channelled portion of the delta, there are abrupt lateral lithologic changes from a channel to a levee or interchannel area. The irregular nature of the lenticular bedding and the abrupt changes from channel to levee deposits preclude core-to-core correlation in the delta areas. Laminated muds consist ofregularly alternating dark and light grey laminae 1 -3 mm thick which form varve-like couplets (Fig. 3 a). The maximum thickness of a couplet never exceeded 10 mm. Microlaminae of about 1 mm in thickness were sometimes observed. Contacts between individual lamina includingthe contact between dark and light coloured laminae of a couplet are sharp and nongradational. From visual inspection the dark lamina always appeared to be slightly coarser grained and less
152
Michael Sturm and A lbert Matter
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Fig. 3. Piston cores from Lake B rienz: (a) core B L-6, 22 m water depth, from shore terrace near Iseltwald.
Regular varves are devloped on top and bottom of core divided by a thick homogeneous mud section; (b) core BL-8, 26 1 m water depth, from centre of the basin plain. Irregular varve bedding is occasionally interrupted by graded beds of Aare (TA) and Liitschine (TL) turbidites. Youngest sediment on top of turbidite TA -1 is more regularly bedded than older sediments. Note frequent gas expansion holes. Scales in centimetres.
1 53
Turbidites and varves, Lake Brienz
clayey than the light lamina. This was confirmed by grain-size analysis of thirteen couplets from core B L-6 (Fig. 3a and Fig. 1 2). Although the mean grain sizes of both kinds of laminae lie in the silt range, the dark laminae are generally about 0·5 coarser grained. Moreover the light laminae contain very little carbonate ( < I 0%) whereas the dark laminae contain 8-36% carbonate. Therefore, if these couplets are varves, the dark basal lamina represents the summer and the light lamina the winter layer, which is j ust the opposite of classical glacial varves. Laminated muds accumulate on the lateral slope and on the basin plain (see below). Homogeneous light grey muds were present in several cores, for example in the middle section of core B L-6 (Fig. 3 a). Homogeneous muds occasionally show a very faint lamination. No macroscopic evidence of bioturbation was observed, however. They are virtually carbonate free ( < I %) and contain no sand-sized admixtures. Their mean grain size lies within the fine silt grade. The observed thickness ranges from 70- 1 80 em. Homogeneous muds were never observed in cores from water depths greater than 40 m, nor in the uppermost sediments in shallower water. It is thought that they are not forming in the lake today. Their distribution in the sediments is restricted to a small area on the upper slope and shore terrace near Iseltwald (Fig. 4). Laminated muds interbedded with graded sand or silt beds were found in cores from the basin plain. The graded beds are either of dark grey or light grey colour. They vary not only in colour but also in thickness, carbonate content and heavy mineral composition. The light grey graded beds are generally much thicker and contain less carbonate ( < 1 0%) than the dark grey graded beds ( > 1 5 % carbonate). The heavy mineral suite of the dark grey beds is dominated by garnet, apatite and hornblende whereas the light beds contain a hornblende-epidote suite. These graded beds are the deposits of turbidity currents generated by the Aare River (light grey) and the Li.itschine River (dark grey).
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surface but have been mapped from occurrence in cores. Contour lines indicate 50 m intervals.
The distribution of these four sediment types was plotted on a bathymetric map (Fig. 4). It is obvious that each sediment type or facies is closely related to a morphologic unit. Delta sand and mud cover the delta areas at each end of the lake. Laminated mud is found on the lateral slopes. Homogeneous mud is restricted to the shelf area near Iseltwald. Layers of graded sand/silt intercalated with laminated muds are found on the fiat basin plain at depths between 250 and 26 1 m.
1 54
Michael Sturm and A lbert Matter
Characteristics and correlation of turbidites
Turbidites found in cores from the basin plain are of particular interest. One of the more prominent Aare turbidites (TA-3), nearly 70 em thick in core 1 7 from mid-basin, was analysed in detail in order to show the structural and textural variations within an individual turbidite (Fig. 5). The base of this turbidite is a scoured surface, overlain by coarse sand, grading upwards into fine sand, then silt and finally silty clay at the top. Wood fragments and plant leaves occur at the base. Simple graded bedding is the only sedimentary structure present in this and in the rest of the large number of turbidites that we have studied. Other structures such as parallel lamination or ripple bedding are not developed. A distinct very light coloured 2 em thick clay layer overlies with sharp contact the graded sequence (Fig. 5). Within this single turbidite the carbonate content increases from 2·5 to 8 · 5% at the top. Changes in carbonate content and grain size are more pronounced in the upper half to two-thirds of the turbidite. Clay is virtually absent in the lower half of the bed. The overlying clay layer contains very little carbonate. Sorting is poor throughout the bed, but slightly less so in the clay deficient lower part.
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after Folk & Ward ( 1 957).
155
Turbidites and varves, Lake Brienz
Liitschine and Aare turbidites can b e identified i n cores o n the basis o f different grey hues, carbonate content and heavy mineral composition. We succeeded in correlating eleven Liitschine turbidites (TL- 1 to 1 1 ) and six turbidites that had been deposited by Aare turbidity currents (TA- 1 to 6) (Figs 6-7). The correlation is based on the sequential pattern of dark and light turbidites, their thicknesses and on varve counts and correlations in the laminated muds between the turbidites. Varve correlation also provided estimates of the amount of erosion caused by a turbidity current (see erosional contacts in Figs 6 and 7). Two particular turbidites (TA- l and T A-3) were closely sampled and grain-size analyses were carried out to check the textural variations and the core-to-core correlation.
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Fig. 6. (a) Cross-section A-A' from the eastern part of basin about 5 km offthe mouth of River Aare (Fig. 1 ). Here Aare turbidites of varying thickness clearly dominate the uniformly thin Liitschine turbidites. Note erosional base of A are turbidites in the southern and central parts of profile. Upper panel shows mean grain size and sorting of basal sediment in Aare turbidite TA- 1 . Middle panel shows bottom profile ofbasin plain and location of cores. Note that slopes which extend to lake surface at 564 m above sea level are not shown. (b) Cross-section B-B' from the western part of basin about 4 km off the mouth of River Liitschine. Aare turbidites are thinner than in section A-A' and rarely exceed 10 em. Liitschine turbidites reach 35 em in thickness.
Individual turbidites vary in thickness from 2 to 1 5 0 em (Figs 6 and 7). Thick turbidites are much less common than thin ones. Lateral variation in thickness is
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Fig. 7. Longitudinal section I-I' from the basin plain of Lake B rienz (Fig. 1). Aare turbidites thin out significantly towards the western end of the basin plain. Turbidites of Liitschine interfinger with thick Aare turbidites and are partly eroded in the eastern half of the basin plain. Upper panels are as in Fig. 6a.
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Turbidites and varves, Lake Brienz
157
greatest i n thick turbidites such a s TA- 1 whereas the thin turbidites are more uniform in thickness. Because the basin plain sediments overlying TA- 1 contain no thick turbidites approximately even amounts of sediments (50-60 em) have accumulated over the entire basin plain subsequent to the deposition ofT A- 1 (which probably took place in 1 896, see below). Aare turbidites attain maximum thickness and frequently have eroded underlying sediments in the eastern part of the basin plain (Fig. 6a). Liitschine turbidites are thickest at the other end of the basin but no erosion was observed. Whereas Aare turbidites can be traced up to the foot of the Liitschine delta, Liitschine turbidites are often absent in cores from the Aare delta region due to erosion by subsequent Aare turbidites (Fig. 7). For example in core 1 7 TA- l eroded at least 40 em of previously deposited silt and clay including Liitschine turbidites TL- 1 , 2 and 3 (Fig. 6a). Even at a distance of 9 km from the mouth of the Aare River this turbidity current still was able to erode at least 25 em of mud including the 3-4 em silt layer of TL- 1 (Fig. 7, core 3 2). As would be expected mean grain size of the sediment deposited by an individual turbidity current decreases with increasing distance from the source, for example TA- 1 (Fig. 7). Sorting, however, remains fairly constant. Correlation of varve-like couplets
The correlation of turbidites was only possible within the central basin plain where cores of sufficient length could be obtained. On the steep lateral slopes box corer and gravity corer were used and yielded cores 1 0-30 em long. These cores were used in an attempt to correlate individual varve couplets from slope to slope across the profundal basin plain (Fig. 8). Key couplets showing distinct features such as microlamination, sandy admixtures, and a more pronounced dark colour (Fig. 8, 0- 1 5 , Q- 1 0, T- 1 ) aided in correlation of the other more uniform varves. The regular couplets on the slopes never exceed 1 0 mm in thickness and the dark basal silt lamina of each couplet is generally thinner than the light-coloured clayey lamina above it. Towards the basin plain couplet thickness becomes more irregular and varies between 5 and 20 mm. This is also true for the couplets of the laminated intervals ofthe basin plain facies. On the basin plain the dark lamina is thicker than the light lamina, and would thus correspond to proximal deposition according to Ashley ( 1 972). Our observations indicate that the darker sand/silt layer is up to twenty times thinner on the slopes than in a proximal core of the basin plain whereas the light clayey layer remains nearly constant. Core BP-56 from the basin plain near the Aare delta shows an Aare turbidite that consists of a 57 mm thick graded sand/silt and a 4 mm thick clay layer (Fig. 8, BP-56, 01 5) . In cores from the basin plain (Fig. 8, BK-69, 70, 84, 85) this sand/silt layer decreases to an ungraded 5-7 mm silt lamina with increasing distance from the Aare mouth, whereas the clayey lamina above is of constant thickness (about 3 mm) in all cores. In cores from the northern and southern slope (Fig. 8, B K -48, 68, 77, 83) the silt layer is even thinner, 1 -3 mm, but the light layer is constantly 3 mm thick. Apparently a distinct turbidite may grade laterally into a thin lamina indistinguish able from a dark lamina of a varve. A similar observation was reported by Ludlam ( 1 969, 1 974) in Fayetteville Green Lake (New York). The most striking point is that a turbidite on the basin plain can be correlated with a dark lamina ofwhat we believe are true varves on the upper lateral slope, an area certainly beyond the reach of turbidity currents.
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Fig. 8 . Correlated surface samples from slopes and basin plain. Varves on the slopes are more regular and show less frequent microlamination within the dark basal layer than
do varves on the basin plain. Note change of key-bed 0- 1 5 from a graded bed in the eastern basin plain to an ungraded basal lamina of varve on the slopes. Gas expansion holes are responsible for cracks within some of the cores. Index map gives location of correlated cores. Scale in centimetres.
Turbidites and varves, Lake Brienz
159
DI S C U S SION Turbiditic sedimentation
Sedimentation of the uppermost Aare turbidite TA- 1 is typical of the larger turbidites in Lake Brienz. Fig. 9 shows the areal distribution of TA- 1 and a longitudinal cross-section. Unfortunately no cores were obtained in the proximal part of the turbidite because of the difficulty in coring the deltaic sands. By analogy with modern turbidity currents observed in Lake Brienz TA- 1 probably originated as sediment-laden river water flowed downslope following a delta channel to the basin plain. Where the turbidity current overflowed the delta channel, lenticular sands and silt layers were deposited. The channel becomes less deep and loses its identity at about 240 m water depth. From this point the main body of the turbidity current spread out and eroded sediments across a large part of the fiat lake floor (Figs 6a and 7). The head of the current lost its erosional capacity about 9 · 5 km from the Aare inflow. This limit is indicated by a dashed line in the western part of the basin on Fig. 9a.
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Near its source the current deposited a coarse-grained and poorly sorted thin sand layer of about 22 em along the axis of the flow (BP-25, Fig. 9a), which thins out laterally to a 10 em thick layer in core BP-23 (Fig. 9a). The much deeper erosion (higher velocity) and greater deposition in the centre of the basin as shown on Fig. 6a indicates that at cross-section A-A' the main mass of the flow was still more or less confined and thinned and slowed laterally towards the base of the slopes. The 1 40 em thick graded bed which had rapidly accumulated in the basin centre only 1 km downcurrent of B P-25 suggests that the turbidity current may have undergone an hydraulic j ump (Komar, 197 1 ) from an erosional supercritical flow to a subcritical depositional flow. By about 6 km from the Aare inflow the turbidity current had spread a uniform sediment layer over the entire basin plain as evidenced by the constant thickness observed in cores 7- 1 2 and B P-40 (Fig. 9a). Beyond core 32 the thickness of TA- 1 decreases rapidly to 4 em in core 25, and it finally remains as a dark thin lamina at the foot of the Liitschine delta.
160
Michael Sturm and A lbert Matter
Assuming a bulk sediment density of 1 ·7 g. cm - 3 (Jackli, 1958) this particular turbidity current deposited a total mass of about 8 x 1 06 tons (dry weight). This corresponds to about thirty times the mean amount of detritus delivered annually to the lake (Table 1). According to Forstner, Muller & Reineck ( 1 968) an exceptional flood of the Rhine River carried three times the average annual suspended load into Lake Constance within a few hours. During the past 30 years the yearly load carried by the Aare River was up to ten times the 30 year-average. It is concluded therefore that in Lake Brienz TA- 1 was deposited by a turbidity current of much larger size than any other that had occurred during the past decades and may have been the result of a catastrophic flood in the drainage basin. The last such event was a catastrophic mudflow which moved out from a lateral valley onto the Aare delta and into the lake in 1 896 (von Steiger, 1 896). According to varve counts TA- l occurred about 70 to 90 years ago which would fit the date of the mudflow. TA- l , however, has a typical Aare River heavy mineral suite that is different from that of the river draining the mudflow source area. Two alternative explanations are then: ( 1 ) the turbidity current depositing TA- l originated when part of the upper delta slope consisting of Aare-derived material became unstable due to overloading by the mudflow from the lateral valley and slid off, merging eventually with the much smaller turbidity current generated directly by the mudflow; or (2) a similar catastrophic event occurred in the main valley of the Aare causing the exceptionally large turbidity current in the lake which picked up additional material by eroding the delta foreslope. Approximate varve counts of the laminated intervals separating large turbidites indicate that large turbidity currents occurred on the average once or twice per century. This is compatible with historical records reporting catastrophic events such as large floods or mudflows. Such reports, however, are often too vague and incomplete to be precisely correlated with turbidites in the lake. Whereas this first type or 'catastrophic' turbidity current .is able to carry large amounts of sand -sized detritus far out onto the basin plain and to erode the lake floor, a second kind of turbidity current is caused by river floods (annual spring floods etc.). This type deposits a much thinner graded sand/silt layer that decreases much faster in thickness with increasing distance from its source and becomes soon indistinguishable from the basal silty layer of varves. If several floods produce turbidity currents of this kind during a single year, a microlaminated silt layer results. Although no direct current measurements are available for Lake Brienz, we can crudely estimate the velocities of turbidity currents from the mean grain size at the base of a graded bed. Using the sediment threshold curve of Miller & Komar ( 1 977) for quartz grains in pure water at 2o·c, a sloping lake floor of 1 ·, and a maximum grain size of 1-2
Turbidites and varves, Lake Brienz
161
The fact that a complete Bouma sequence has never been observed in turbidites of Lake Brienz (Sturm, 1 975) suggests rapid deceleration of the depositing turbidity currents in a manner shown by the experimental studies of B anerjee ( 1 977). Hence catastrophic and flood-generated turbidity currents are shortlived events. Possibly steady-state long-lived (days to weeks) turbidity currents fed by the inflowing river do exist. Such currents, however, will have little effect on sedimentation, because most of the detritus which accumulates through weathering in the catchment is normally carried into the lake during river floods. The large number of turbidites, some with an erosional base, clearly show that the inflowing water had a density greater than lake water at any level and flowed along the depositional gradient as an underflow to the deepest part ofthe basin (Fig. 1 0). The flood-induced underflows have densities of 1 ·0 1 g cm - 3 (Table 2), that is, they are low density turbidity currents. The catastrophic turbidity currents on the other hand, may have been of much higher densities as judged from the grain size measured at the base of large turbidites.
Delta sedimentation
There is abundant evidence of erosion in the Aare delta over the past hundred years, almost exclusively because of man's influence. Between 1 866 and 1 875 a new subaerial channel was excavated to channellize the Aare River near its mouth. Instead of flowing into the lake at the northern part of the alluvial plain, it was routed laterally to the south. Within the past hundred years the Aare has built a new delta with channels that have cut into parts of the old delta. Seismic investigations indicate that interchannel areas are underlain by older channels which were abandoned and filled suggesting that the Aare shifted many times before man influenced the natural processes. Comparison of two bathymetric maps prepared by the Swiss Federal Bureau of Water Management on the basis of intensive surveys of the Aare delta in 1 898 and 1932 indicates that during this 34 year period 5 m of sediment accumulated in some areas while in others 3-4 m had been eroded (EAW, 1 939). These changes are almost certainly related to the rerouting of the river course and migration of the delta channels which funnel coarse-grained sediment to the lower delta foreslope and the basin plain. An additional bathymetric survey undertaken in 19 54 revealed that some depositional areas of the 1 932 survey had become erosional areas and vice versa. Other important man-made alterations include the construction of several dams and reservoirs after 1932 which trap a large amount of the Aare detritus and control water discharge to some extent.
'Pelagic' sedimentation
Density currents in lakes and reservoirs have been known for a long time, but early authors (e.g. Bell, 1 942; Kuenen, 195 1 ; Bates, 1953; Gould, 1 960) assumed a more or less unstratified water body in their depositional models. Most temperate lakes, however, show a distinct thermal stratification during the summer. The existence of a thermocline and therefore a pycnocline is of utmost importance in sedimentation as schematically shown by Ashley ( 1 972, fig. 3).
162
Michael Sturm and A lbert Matter
shore terrace
basin slope
basin plain
delta area
inrerflows (underc rrents)
����;·· ·· ·· ···f'\\;:;��;in�.��������2P?.', �11Peiogic11 sedimentation ·
. . I
.
:l
Fig. 10. Distribution mechanisms and resulting sediment types proposed for clastic sedimentation in oligotrophic lakes with annual thermal stratification. Note that hypothetical shore terrace is situated higher than depth of thermocline. Width of basin and sediment thickness are not to scale.
Of the detritus delivered by the Aare and Liitschine Rivers to Lake Brienz the coarsest bedload is dropped near the river mouths on the upper delta. The suspended load may be transported farther into the basin by a turbidity current. During April to October, however, when the lake is thermally stratified, if the density of the infl.owing sediment-laden river water is less than that of the lake water, it will flow at the lake's surface as an overflow (Fig. 1 0). Most frequently, however, the river water is denser than that of the epilimnion, and less dense than that of the hypolimnion and it will be inj ected at the epilimnion/hypolimnion boundary as an interjlow (Fig. 1 0). Large amounts of suspended matter are thus injected into the thermocline (Fig. 1 1 b). This sediment is moved by counter-clockwise currents around the lake as shown by studies of surficial sediments (Sturm, 1 976). Coarse silt-sized particles settle out from the thermocline during this transport and accumulate as a dark silty layer.
Table 2. Suspended sediment concentration and temperature in the Aare River and Lake B rienz water on 28 May 74.
Sample water-depth m
Aare River* o•• 0·5** 10·0** 26·0** 1 1 0·0**
Temperature ·c
5·4 13·8 13·7 1 1 ·0 8-4 5·6
S uspended matter mg 1 - '
2 1 400 10·7 4·3 40·0 4·8 54·7
Density H,O § g cm - 3
0·9999842 0·9992987 0·999 3 1 23 0·9996328 0·9998509 0·9999795
Density suspension §§ g cm - 3 1 ·0 1 28435 0·999305 1 0·9993 149 0·9996568 0·9998538 1 ·00001 23
!:J.p Density suspension - density H20
128593 64 26 240 29 328
*Sample was taken at 7:40 a.m. I km upstream from the mouth of river. •• Samples were taken at 1 0:00 to 10:40 a.m. 2 km off the mouth of river Aare. § Calculated for pure water at measured temperature. §§ Calculated using an average particle density of 2·5 g em _ ,
X X X X X X
10-6 1 0 -6 10-6 10-6 1 0- 6 1 0 -6
Turbidites and varves, Lake Brienz
1 63
Most of the finest-grained particles remain in suspension at the thermocline. This occurs primarily because the vertical density gradient is more dependent on temperature than on an increase in density due to suspended particles. This is supported by the data in Table 2 showing the maximum concentration of suspended matter in the Aare river water during a large flood. The profile (Fig. 1 1 ) in the lake 2 km from the river mouth was measured a few hours after maximum flood stage when the main body of the turbidity current had already passed the station. Highest concentrations of suspended matter were found in the thermocline at 1 0 m water depth and near the bottom. The turbid layer· at the bottom thus represents the tail of the turbidity current. The density increase due to suspended matter in the thermocline is less than the density increase due to temperature between 1 0 and 26 m. Matthews ( 1 956) noted that a concentration of suspended particles of about 50 mg/ 1 results in only insignificant changes in bulk density. Furthermore the viscosity increase and turbulence in the thermocline will counteract settling. The thermocline therefore acts as a sediment trap (Nydegger, 1 967). Apparently, only after turnover do the fine grained particles of the suspension begin to settle out as a 'suspension blanket' forming the light-coloured winter layer of a varve and the light-coloured layer capping a turbidite. Constant thickness of light-coloured layers in proximal and distal areas results from the counter-clockwise currents which distribute suspended particles in the
(b)
(a) Temperature 0(
Electric resistivity 103
.n.
10
10
20
20
30
30
40
40
50
50 20-ll-1973 60
60
70
70
Fig. I I . Temperature and light transmission profiles in Lake Brienz measured 2 km off the mouth of the Aare
River. Turbid layers (b) coincide with levels of maximum temperature gradient (a). The complex turbidity distribution in the water profile on 24 May 73 becomes smoother during summer and autumn.
1 64
Michael Sturm and A lbert Matter
epilimnion uniformly over the lake with Aare River detritus deposited mainly along the north shore and Liitschine River detritus along the south shore (Sturm, 1 976). Fig. 1 1 shows that with the beginning of turnover in November the suspension layer and the thermocline had descended from about 1 5 m in September to 45 m in November. The non-gradational basal contact of the light lamina is further evidence of a hiatus in deposition between dark and light-coloured laminae. This mechanism implies that the dark/light couplets are indeed annual layers and clastic varves (0strem, 1 975; Gustavson, 1 975). On the slopes beyond the reach of turbidity currents only 'pelagic' sedimentation takes place and leads to uniform varved sequences. Dark laminae of the slopes represent the summer fallout i.e. the 'summer' layer of a varve. Dark laminae on the lowest slope and the basin plain, however, are of multiple origin. They may be the distal part of a larger turbidite, normally single or, when microlaminated, multiple flood-induced turbidites, and 'summer' fallout from the thermocline and possible deeper interfiows such as the one shown at 80 m in Fig. 1 1 b. The multiple origin
8
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Fig. 12. Mean grain size and sorting in turbidite T A-3, varves and suspended matter in water. Sediments are
from core 8 taken in the centre of the basin (see Fig. 7) ; water samples were collected on 28 May 1974--2 km from the mouth of the A are River and filtered through 0-45 fLm Millipore filters (see Table 2). In the turbidite mean grain size grades upwards from medium sand to finest silt. Mean grain sizes in the top of the turbidite, varves and suspended matter are approximately the same. All samples are poorly to very poorly sorted, • and ... indicate mean grain- size (M,) and sorting ( a1 ) of samples of turbidite and dark varve laminae; oand 6 indicate M, and a1 of samples of light laminae and of suspended material.
Turbidites and varves, Lake Brienz
1 65
explains the greater thickness of dark laminae in these areas compared to those from the slopes. This may be checked by a comparison ofgrain sizes of a large turbidite (TA3), varves and suspended material of water samples (Fig. 1 2). In all cases sorting is poor. Mean grain sizes of TA-3 are distinctly coarser than mean grain sizes of varves and suspended material, but similar grain sizes can be observed within the top layer of the turbidite, in light-coloured laminae of varves and constituting the suspended material. Regardless of their origin, dark layers in the basin and on the slope generally form between spring and autumn, and both are capped by a light-coloured 'winter' layer. Thus, although deposited during the course of only a few hours, turbidites on the basin plain correspond with half-couplets on the slope. If pelagic sedimentation is continuous over the year, an unlaminated sediment would accumulate such as the homogeneous mud found only in cores on the upper slope and shore terrace off Iseltwald. This happens when the thermocline does not develop or if it is at water depths greater than the depositional surface. This was also reported by Ashley ( 1 975) who noticed that no varves were formed on morphologic highs in glacial Lake Hitchcock. No such homogeneous muds are accumulating today in Lake Brienz. The occurrence of homogeneous muds, however, in cores from the upper slope indicates that according to varve counts about 600 years ago the level of Lake Brienz and hence its thermocline was about 10- 1 5 m lower than today. S UMM A R Y
Sedimentation in Lake Brienz is dominated by two rivers, the Aare and the Liitschine, which enter at opposite ends of the long, narrow and deep basin. The sediments are clastic. The rivers drain areas of distinctly different lithologies, and thus the source of the sediments can be clearly delimited. Four major types of sediment can be distinguished, massive and lenticular sands and silts near deltas, laminated muds and homogeneous muds on slopes and laminated muds interbedded with graded sands and silts on the basin plain. The sediments are related to two different depositional processes, (a) turbidity currents (underflows) along the lake bottom and (b) interflows and surface currents (overflows) in and on the lake water. Massive turbidites up to 150 em thick apparently are the product of high-density turbidity currents that occur only once or twice per century as a result of catastrophic flooding or landslides. These turbidites represent only the graded A-interval of the Bouma sequence, suggesting rapid deceleration of turbidity currents. The currents cut channels into deltas and erode basin plain sediments up to 9 km from the river mouth. Such turbidity currents cover most of the basin plain with coarse sand fining distally into graded silt layers and finally into thin ungraded silt/clay laminae. One such turbidite can be clearly mapped over 1 1 km2 and contains a volume of about 5 X 106 m3 sediment. Turbidites from both sediment sources interfinger. A second type of turbidite is represented by thin dark grey layers apparently deposited by low-density turbidity currents which occur one or more times each year as a result of spring meltwaters and heavy rainfalls. These turbidites consist of slightly graded coarse-to-fine sands in the delta region but thin rapidly into ungraded silt laminae on the basin plain. Neither type of turbidite was found on the lateral basin slopes or morphological highs.
166
Michael Sturm and A lbert Matter
Laminated muds on the lateral slopes and on the flat basin plain (here interbedded with the thick turbidites) apparently result from an annual sedimentation cycle related to thermal stratification in the lake water and seasonal inputs of sediment-laden river water. If the suspended load and water temperature are such that river water is less dense than cold bottom lake water, instead of flowing down along the bottom as an underflow, the river water will form a turbid layer at or above the thermocline (interflow or overflow). Counter-clockwise currents will spread the entrapped suspended sediment over the entire basin. Coarser particles fall out continuously during the summer months and form the dark grey basal laminae of varve couplets. Light-coloured layers represent fine-grained particles that become trapped by the density gradient at the thermocline . during the summer and settle out after the thermocline is destroyed in the autumn. On the basin plain, the occurrence of one or more low-density turbidity currents during the spring and summer results in additional thin, dark grey layers (microlaminations) which are indistinguishable from the basin-wide 'summer' layer. Dark layers are therefore thicker on the basin plain than on the lateral slopes which are beyond the effects of turbidity currents. The light-coloured 'winter' layer is of uniform thickness over the entire basin. Lake Brienz sediments show that the mechanisms ofturbidite deposition and clastic varve formation are integrally related. A half-couplet on the lateral slope can grade into a thick graded turbidite on the basin floor and both have a common sediment source but are the products of an overflow or interflow on the one hand and an underflow, i.e. a turbidity current, on the other hand. Thus, neither the models of de Geer ( 1 9 1 0), Kuenen ( 1 95 1 ) and Laj tai ( 1 967), and others, which invoke only turbidity currents as the varve-forming mechanism, nor the model of Antevs ( 1 95 1 ) which requires only overflows, provides a complete explanation for deposition of laminated clastic sediments in lakes. Our study suggests that both over- and interflows as well as turbidity currents are required, and that the existence of seasonal thermal stratification is of critical importance.
A CKN OWLE DGMENT S
We wish to thank especially R. F. Wright for criticizing earlier drafts of the manuscript and suggesting important changes as well as for correcting the English of the final draft. Further thanks are due to K. Kelts and P. Shannon for useful comments in reviewing this paper and to A. Lerman for stimulating discussions. This study was supported by the Swiss National Science Foundation, grant No. 2.427.75.
REFERENCE S ANTEVS, E.( 195 1 ) Glacial clays in Steep Rock Lake, Ontario, Canada. Bull. geo/. Soc. A m. 62, 1 223- 1 262. ASHLEY, G . M . ( 1 972) Rhythmic sedimentation i n glacial Lake Hitchcock, Massachusetts-Connecticut. Contr. Geo/. Dept Univ. Mass. 10. ASHLEY, G.M. ( 1 975) Rhythmic sedimentation i n glacial Lake Hitchcock, Massachusetts-Connecticut. In: Glaciofluvial and Glaciolacustrine Sedimentation (Ed. by A. V. Jopling and B . C . McDonald). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 23, 304-320.
Turbidites and varves, Lake Brienz
1 67
BANERJEE, I. ( 1 968) A study of glacial varves as turbidites. A bstr. Meet. A m. geol. Soc. Wash. D. C. Feb. 1 968, 335-336. BANERJEE, I. ( 1 977) Experimental study on the effect of deceleration on the vertical sequence of sedimentary structures in silty sediments. J. sedim. Petrol. 47, 77 1-783. BATES, CH. C. ( 1 953) Rational theory of delta formation. Bull. Am. A ss. Petrol. Geol. 37, 2 1 1 9-2 1 62. BELL, H.S. ( 1 942) Density currents as agents for transporting sediments. J. Geol. 50, 5 12-547. EAW (Eidg. Amt. f. Wasserwirtschaft) ( 1 939) Deltaaufnahmen des Eidgenossischen Amtes fiir Wasserwirtschaft. Mitt. 34. EA WAG ( 1 967) Einfluss der veriinderten Zuflussverhaltnisse auf den Thuner- und Brienzersee. Report Nr. 3435, EAW AG-ETH, Ziirich. FoREL, F.A. ( 1 885) Les ravins sous-lacustre des fleuves glaciaires. C. r. hebd. Seanc. A cad. Sci., Paris, 101, 725-728. FOLK, R.L. & WARD, W.C. ( 1957) B razos River bar: a study in the significance of grain size parameters. J. sedim. Petrol. 27, 3-26. FORSTNER, U., MOLLER, G. & REINECK, H.-E. ( 1 968) Sedimente und Sedimentgefiige des Rheindeltas im Bodensee. Neues. lb. Miner. A bh. 109, 33-62. DE GEER, G. ( 1 9 10) A geochronology of the last 1 2,000 years. Int. Geol. Congr. XI. Sess. Stockholm, 241-253. GoU L D , H.R. ( 1 960) Turbidity Currents. In: Comprehensive Survey of Sedimentation in Lake Mead (Ed. by W. 0. Smith et al.). Prof Pap. U.S. geol. Surv. 295, 201-207. GusTAVSON, T.C. ( 1975) Bathymetry and sediment distribution in proglacial Malaspina Lake, Alaska. J. sedim. Petrol. 45, 738-744. HElM, ALB., MOSER, R. & B ORKLI-ZIEGLER, A. ( 1 888) Die Catastrophe von Zug. 5. Juli 1887. Veri. Hofer & Burger, Zurich. HESSE, R. ( 1 974) Long-distance continuity of turbidites: possible evidence for an early Cretaceous trench abyssal plain in the East Alps. Bull. geol. Soc. A m. 85, 859-870. Ji\cKLI, H. ( 1 958) Der rezente Abtrag der A!pen im Spiegel der Vorlandsedimentation. Eclog. geol. Helv. 51, 354-365. KOMAR, P.O. ( 1 97 1 ) Hydraulic jumps in turbidity currents. Bull. geol. Soc. A m. 82, 1477-1 488. K u EN EN, P.H. ( 195 1) Mechanics of varve formation and the action of turbidity currents. Geol. Fiiren. Fiirhandl. 73, 69-84. LAJTAI, E.Z. ( 1 967) The origin of some varves in Toronto, Canada. Can. J. Earth Sci. 4, 633-639. L AM B ERT, A.M. ( 1 978) Eintrag, Transport und Ablagerung von Feststoffen im Walensee. Eclog. geo/. Helv. 71. (In press). LAMBERT, A.M., K E LTS, K . R. & MARSHALL, N.F. ( 1 976) Measurements of density underflows from Walensee, Switzerland. Sedimentology, 23, 87- 1 05. LUDLAM, S.D. ( 1 969) Fayetteville Green Lake, New York. III. The laminated sediments. Limn of. Oceanogr. 14, 848-857. LUDLAM, S.D. ( 1 974) Fayetteville Green Lake, New York. VI. The role of turbidity currents in lake sedimentation. Limnol. Oceanogr. 19, 656-664. MATTHEWS, W.H. ( 1 956) Physical limnology and sedimentation in a glacial lake. Bull. geo/. Soc. A m. 67, 537-552. MATTER, A., 0ESSOLIN, D., STURM, M. & S O SST RU N K , A. E. ( 1973) Reflexionsseismische Untersuchung des Brienzersees. Eclog. geol. Helv. 66, 7 1-82. MILLER, M.C. & KOMAR, P.O. ( 1 977) The development of sediment threshold curves for unusual environments (Mars) and for inadequately studied materials (foram sands). Sedimentology, 24, 709-72 1 . NORMARK, W.R. & DICKSON, F . H . ( 1 976) Man-made turbidity currents i n Lake Superior. Sedimentology, 23, 8 1 5-832. NYDEGGER, P. ( 1 957) Vergleichende limnologische Untersuchungen an sieben Schweizerseen. Beitr. Geol. Schweiz-Hydrol. 9, l-80. NYDEGGER, P. ( 1 967) Untersuchungen iiber Feinstofftransport in Fliissen und Seen, iiber Entstehung von Triibungshorizonten und zuflussbedingten Stromungen im Brienzersee und in einigen Vergleichsseen. Beitr. Geol. Schweiz-Hydrol. 16, l-92. NYDEGGER, P. ( 1976) Stromungen in Seen. Untersuchungen in situ und an nachgebildeten Modellseen. Beitr. Geol. Schweiz, kl. Mitt. 66, 1 4 1 - 1 77.
1 68
Michael Sturm and A lbert Matter
0STREM, G. ( 1 975) Sediment transport in glacial meltwater streams. In: Glacioflu vial and Glaciolacustrine Sedimentation (Ed. by A. V. Jopling & B. C. McDonald). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 23, 1 0 l- 1 22. REINECK, H.-E. ( 1 963) Der Kastengreifer. Natur Mus. Frankfurt, 93, 1 02- 108. REINECK, H.-E. ( 1 967) Ein Kolbenlot mit Plastik-Rohren. Senckenberg. leth. 48, 285-289. REINECK, H.-E. & SINGH, I. B. ( 1 957) Depositional Sedimentary Environments. Springer-Verlag, New York. voN STEIGER, H. ( 1 896) Der Ausbruch des Lammbaches vom 3 1 Mai 1 896. Mitt. Naturf Ges. Bern, 265-275. STURM, M. ( 1 975) Depositional and erosional sedimentary features in a turbidity current controlled basin (Lake Brienz). !Xth Int. Congr. Sedim., Nice, 5, 385-390. STURM, M. ( 1 976) Die OberfHichensedimente des Brienzersees. Eclog. geol. Helv. 69, l l l - 1 23 . STURM, M. & MATTER, A. ( l972a) The electro-osmotic guillotine, a new device for core cutting. J. sedim. Petrol. 42, 987-989. STURM, M. & MATTER, A. ( l972b) Geologisch-sedimentologische Untersuchungen im Thuner- und Brienzersee. Jb. Thuner- und Brienzersee, 52-72. STURM, M. & MATTER, A. ( l 972c) Sedimente und Sedimentationsvorgiinge im Thunersee. £clog. geo/. Helv. 65, 563-590.
Spec. Pubis int. Ass. Sediment. ( 1 97 8) 2, 169- 1 87
Lacustrine facies in the Pliocene Ridge B asin Group: Ridge B asin, California
M A R T I N H . L I N K and R 0 B E R T H . 0 S B 0 R N E Department of Geology, Los Angeles Harbor College, Wilmington, California 90744, and Department of Geological Sciences, University of Southern California, Los Angeles, California 90007, U.S.A.
A B STR ACT
The Ridge Basin is a wedge-shaped trough 1 5 by 40 km which contains over 9000 m of lacustrine sedimentary rocks. Lacustrine sedimentation in the Pliocene Ridge Basin occurred in this elongated trough formed during active strike-slip displacement along the San Gabriel fault. The lacustrine and other terrestrial deposits reflect syntectonic deposition related to steep faults with components of slip along the eastern and northeast and strike-slip to oblique slip components along the western margin of the Ridge Basin. Asymmetrical development of facies resulted from this differential tectonism along the margins of this basin. The Violin Breccia, which crops out along the western margin, consists of breccia, conglomerate and sandstone deposits that chiefly represent narrow talus and small alluvial fans that accumulated along the San Gabriel fault scarp. The Peace Valley 'beds' which occur in the central part of the basin, contain nearshore to offshore dark analcimic and ferroan dolomitic mudrock, organic-rich shale, and localized turbidite sandstone. This unit has a lacustrine origin as indicated by molluscs, ostracods, stromatolites, plants, and insect and vertebrate remains. The Ridge Route 'formation' crops out along the eastern margin and consists of sandstone and conglomerate which accounts for the greatest volume of sedimentary rock exposed in the Ridge Basin. It interfingers with both the Peace Valley 'beds' and Violin Breccia to the west. The sediments comprising this large alluvial fan-fluvial, marginal lacustrine and offshore turbidite complex were derived from the north and east. The lower part of the Ridge Basin Group is transitional from underlying marine strata. Ridge Basin lake evolved from an externally-drained, relatively deep lacustrine and/or marine system with thick sequences of turbidite and slump-folded strata to an internally drained, rather shallow, closed-lake system with thick sections of dolomitic mudrock. Lacustrine sedimentation in the Ridge Basin ended with the termination of movement along the San Gabriel fault. At that time the basin was filled with fluvial and fanglomerate deposits assigned to the Hungry Valley Formation.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
1 70
Martin H. Link and Robert H. Osborne INTRO DUCTION
The Ridge Basin, California (Fig. 1) is a relatively small wedge-shaped basin (Eaton, 1 939) or trough ( 1 5 X 40 km) which contains an extremely thick ( > 1 2,000 m) composite marine, lacustrine and fluvial stratigraphic section (Fig. 2). The Ridge Basin was tectonically active from Mohnian (late Miocene) to Blancan (early Pleistocene), which is an interval of about 10 million years (Crowell, 1 976, personal communication). The siliciclastic deposits have been deformed, uplifted and eroded forming spectacular and nearly continuous outcrops across the width of the basin (Fig. 3); therefore, it provides an excellent three-dimensional view of the structure and sedimentology of a recently dissected lacustrine basin (Fig. 4). Inasmuch as the Ridge Basin occurs at the j unction of the San Gabriel and San Andreas fault systems, it has been the subject of intensive tectonic investigations (Crowell, 1 974a,b) and is of economic interest because of oil shows, potential source beds, and its proximity to oil provinces. The authors are engaged in a comprehensive study of the sedimentology of the Ridge Basin principally to integrate its sedimentary and tectonic evolution. Data from stratigraphic sections, electric logs from boreholes, petrography, and mapping are integrated in this paper. The purpose of this paper is to concisely describe and discuss the major lacustrine facies in the Ridge B asin.
Mojave Desert
0
20
40
60
80
Ki lometers
Fig. I. Index map of the Ridge Basin, major faults, and location of diagrammatic geologic cross-section A-B.
G EOLOGI C AL S ETTING
Sedimentation in the Ridge Basin was the result of infilling of an elongated trough formed during active strike-slip (transcurrent) movement along the San Andreas transform fault margin (Crowell, 1 973, 1 975). The Ridge B asin was adjacent to the San Gabriel fault system which was the major Pliocene strand of the San Andreas fault. The basin formed at a bend along this fault where one side was stretched, depressed, and filled with a thick sequence of sediments by gradual overlapping of
171
Lacustrine facies, Pliocene Ridge Basin, California BASIN
R I DGE
T H I C K NE S S
lUT[RI
'UT tS,OOO
4500
10,000
3000
4000
CL EARWATER FAUL T ZONE
2000
sooo
0000
0
2 2
HORIZONTAL
WILU 4 IIIL.OMETt:ltS
SCALE
D •
EX PLA NATION SANDSTONE SHALE - MUDSTONE
� CONGLOMERATE
� GNEISS
ri::'.J
GRANITIC ROCKS
� BRECCIA
Crowell
1975
Fig. 2. Diagrammatic geologic cross-section along line A-B (Fig. I) showing major stratigraphic and
structural relationships in the Ridge Basin (after Crowell, 1 975).
older strata by younger as the depocentre moved northward through time (Fig. 4). Although restricted source areas to the west shed sediment along the San Gabriel fault (Violin Breccia), the maj or sediment input came from the east and northeast (Ridge Route and Hungry Valley Formations). An idealized cross-section of the Ridge Basin (Fig. 2) shows a half-graben bounded by Mesozoic and older basement rock on both sides and separated on the west by the San Gabriel fault and on the east by the Clearwater, Liebre, and associated fault systems. The San Gabriel fault was active from Miocene into Pleistocene time and shows over 60 km of right slip displacement (Crowell, 1 975). The Clearwater and Liebre fault systems are dominantly high-angle reverse faults that were active during sedimentation in the Ridge Basin (Fig. 4). In Pleistocene time the modern trace of the San Andreas fault originated in this area as strike-slip displacement splayed eastward from the San Gabriel to the modern San Andreas fault. This composite zone may be considered as the San Andreas transform fault margin. A total of 240 km of right slip on the San Andreas fault system since early Pliocene for the Ridge Basin is suggested by Ehlert & Ehlig ( 1 977). The lowest deposits in the basin are the Miocene nonmarine Mint Canyon Formation (Fig. 2). It is overlain by the late Miocene marine Castaic Formation which is about 2200 m thick and is transitional upward into the Ridge Basin Group. The Violin Breccia is laterally and vertically transitional into the Castaic Formation and the Ridge Basin Group. It is 1 1 ,000 m thick but extends along strike for a maximum distance of only 1 500 m (Fig. 3). The Ridge Basin Group is about 9000 m thick and consists of four units, which are the Violin Breccia, Peace Valley 'beds', Ridge Route 'formation', and the Hungry Valley Formation. The Peace Valley 'beds' and Ridge Route 'formation' do not yet have formational status. Based on faunal and
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Fig. 3. Generalized geologic map of Ridge Basin, California (modified after Crowell, 1 954, 1 975, and unpublished map).
Lacustrine facies, Pliocene Ridge Basin, California
1 73
Fig. 4. Isometric sketch of the Ridge Basin, Southern California (diagrammatic and not to scale). Strata of the Ridge Basin Group not labelled. Stratigraphic section at A is overlapped at B. Symbols: BF, basin floor; gn, gneiss; gr, granitic rocks; Eo, mainly Eocene; Mv, Miocene volcanic rocks; Msm, Miocene Santa Margarita Formation; Mm, Miocene Modelo Formation. From Crowell, 1 974a, p. 298.
stratigraphic evidence, the lowermost 600 m of the Ridge Basin Group have been interpreted as marine, whereas the remainder (8400 m) has been interpreted as dominantly lacustrine (David, 1 945; Axelrod, 1 950; Stock, in Crowell, 1 950; Crowell, 1 975). This is among the world's thickest and most spectacular lacustrine sequences. Asymmetrical development of rock units and facies resulted from differential tectonism in this lacustrine basin (Figs 2 and 3). The Violin Breccia, which crops out along the western margin, consists of a sequence of breccia, conglomerate, dark-grey sandstone and grey to red mudstone over 1 1,000 m thick. The Violin Breccia is thickest near the San Gabriel fault and only extends laterally about 1 500 m into the basin. The Peace Valley 'beds' which crop out in the central part of the basin, contains dark-grey analcimic and ferroan dolomitic mudrock, organic-rich shale, and localized sandstone. Its composite thickness is about 6000 m and it interfingers with the Violin Breccia to the west and the Ridge Route 'formation' to the east. The Ridge Route 'formation', which crops out along the eastern margin, is composed predominantly of white to brown conglomeratic sandstone, and some conglomerate and breccia. It is about 8000 m thick and several tongues laterally extend across the width of the basin where they locally interfinger with the Violin Breccia (Fig. 2). The Hungry Valley Formation is about 1 1 00 m thick and is composed of white to red sandstone and conglomerate. It represents the final clastic infilling of the Ridge Basin and onlaps the
1 74
Martin H. Link and Robert H. Osborne
San Gabriel fault (Fig. 2). These strata overstep unconformably on Mesozoic crystalline basement rock or sedimentary strata of Palaeocene to Eocene age (San Francisquito Formation) (Fig. 4).
B IOT A
Much of the Ridge Basin Group is lacustrine, as documented by David ( 1 945), Axelrod ( 1 950), Crowell ( 1 950), and Jennings ( 1 953). Molluscs, ostracods, stromato lites, plants, a few insects and vertebrate remains are the most abundant faunal and floral representatives known from the Ridge Basin Group. Bivalves, including Anodonta A. cygnea Linne, a few gastropods, probably the genus Lymnaea, and smooth-shelled ostracods have been reported and are considered to be of freshwater origin (Crowell, 1 950). Stromatolites are common and contain both coccoid and filamentous algae (Link, Osborne & Awramik, 1 978). Abundant floral remains include members of woodland, chaparral and border-forest plants which suggest a semi-arid climate (Axelrod, 1 950). Several dragonflies (Shepard, 1 962; Squires, 1 978), stickleback fis h - Punctiduis (David, 1 945), and pupfish - Empetrichthys and Fundulus (Shepard, 1 962) also have been reported from the Ridge Basin. Other vertebrate remains associated with the lacustrine sediments include fragments of Pliocene horses, camel, mastodon, antelope, rhinoceros, sloth, and turtles (see Crowell, 1 950, p. 1 638; Miller & Downs, 1 974, and Whistler & Downs, personal communication, 1 977). L A C U STRINE F ACIE S
Sediments of the Ridge Basin Group were deposited in lacustrine environments . adjacent to areas of high relief which at times contributed large quantities of clastic sediment. Three major non-marine sedimentary palaeoenvironments are represented in the Ridge Basin Group: alluvial fan-fluvial, marginal lacustrine, and offshore lacustrine facies. The lithologies, sedimentary features and biota that characterize these facies are summarized in Table 1 and are discussed below. Alluvial fan-fluvial facies
Coarse-grained fluvial sediments were deposited as a series of coalesced alluvial fan complexes along the eastern and western margins of the Ridge Basin (Fig. 2). Thick accumulations of interstratified conglomerate, sandstone and breccia, characteristic of the alluvial fan complexes, interfinger with finer-grained lacustrine mudstone, shale and sandstone toward the centre of the basin. The eastern and western alluvial fans differ markedly with respect to size, composition, texture and distribution, and will be discussed separately. Eastern alluvial fan-fluvial facies
Along the eastern margin, the alluvial fan-fluvial facies (Ridge Route 'formation') is very coarse-grained (gravel- to cobble-sized clasts are common), contains well rounded to angular clasts, is moderately to poorly sorted, lacks well-defined
1 75
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stratification, and is devoid of fossils (Fig. 5B, C and D). Well rounded, friable basement clasts (Fig. 6A) are a common constituent. Finer-grained (sand- and gravel sized clasts predominate), well-stratified (Fig. 6B and C) and moderately to well-sorted deposits suggest deposition near the lacustrine shoreline. These deposits include tabular beds of cross-stratified and horizontally-laminated sandstone (Fig. 6D), erosive channels filled with conglomerate (Fig. 6E) and cross-stratified sandstone, and contain load and flame structures (Fig. 6F), convolute laminae, ripple marks, mudcracks and a few vertebrate remains. At times, perhaps during seasonal flooding, coarse-grained sediment was transported by braided streams on the alluvial fan complexes into the lake. These streams greatly modified the marginal lacustrine environments. Such sediment influxes are indicated by channelling and by basal conglomerate or sandstone containing intraclasts of stromatolites, mudstone, siltstone and sandstone. The eastern alluvial fan-fluvial facies consists of light-coloured granitic debris and minor amounts of volcanic and sandstone clasts derived from adj acent basement terrain and from Palaeocene to Eocene strata (San Francisquito Formation) along the eastern side of the basin (Crowell, 1 954). These clasts may have been derived even farther east from the Mojave Desert region (Ehlert & Ehlig, 1 977). Local fanglomerate
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is found around basement highs and along major steep faults on this side of the basin (see Fig. 2). This facies contains the greatest volume of sedimentary rock in the Ridge Basin and a great abundance and diversity of sedimentary structures. The well rounded nature of basement clasts suggests reworking or moderate transport on large alluvial fan complexes. Western alluvial fan-fluvial facies
Along the western margin, the alluvial fan-fluvial facies (Violin Breccia) consists of narrow but continuous exposures of dark-coloured breccia, conglomerate and diamictite (Fig. 7 A). The clasts in the Violi� Breccia consist of gneiss, schist, and Mesozoic 'granitic' basement. Near the San Gabriel fault (Fig. 7C, section 1), the Violin Breccia lacks well-defined stratification (Fig. 7B-A), is poorly sorted (Fig. 7B B), sheared and fractured, and devoid of fossils. It displays few sedimentary structures and the clasts are extremely angular and large (up to 2 m in diameter). This formation is extremely thick ( 1 1 ,000 m) and extends as a band along the San Gabriel fault for about 30 km with a width of only 1 500 m (Fig. 4). The relatively small fans of the Violin Breccia dramatically interfinger with fine-grained lacustrine sediment a short distance (500-1 000 m) into the basin. Here the distal parts of the Violin Breccia are
Lacustrine facies, Pliocene Ridge Basin, California
177
Fig. 6 . Photographs of t h e eastern alluvial fan-fluvial and marginal lacustrine facies. ( A ) Basal
conglomerate with subordinate breccia; (B) parallel-laminated sandstone; (C) low-angle tabular and festoon cross-bedding in conglomeratic sandstone; (D) tabular cross-bedded conglomerate cutting into mudstone (E) U-shaped channel conglomerate cutting into sandstone; (F) load structures in lower part of a sandstone; (G) ooids and oncolites; (H) stromatolites (S) in an oncolitic sandy mudstone overlain by a conglomeratic sandstone; (I) domal stromatolite with concentric layers.
developed as fine-grained strata, which are dark grey to red, mudcracked, cross bedded, parallel-laminated, burrowed (Fig. 7B-C), and contain occasional stromato litic horizons. The Violin Breccia formed as talus, debris flow, and landslide deposits along the San Gabriel fault scarp. With continued strike-slip to oblique-slip displacement along this fault, sediment was deposited continuously within a narrow belt of great stratigraphic thickness, but confined lateral extent. The western side was near the axis of the Ridge Basin trough and shows the greatest composite thickness of sedimentary rock which probably was related to local subsidence along this dominantly strike-slip margin. Marginal lacustrine facies
The marginal lacustrine facies in the Ridge Basin are interrelated and complex. Sediments representing shoreline, nearshore mudflat and sandflat, bar complexes, and fluvial-deltaic palaeoenvironments have been identified. These facies occur through out the section on both sides of the basin. Examples of each facies are discussed below.
178
Martin H. Link and Robert H. Osborne
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fan marginal lacustrine facies (Violin Breccia). Photographs include: (A) massive to poorly-bedded Violin Breccia; (B) diamictite; and (C) fine grained facies of the Violin Breccia showing low-angle cross-bedding in the upper part and burrowing in the lower part.
Shoreline facies
Terrigenous and carbonate sediments were deposited and reworked along the margins of the Ridge B asin lake. The shoreline facies is especially common and well defined along the eastern side where broad alluvial fans from the east extended into the lake (Fig. 5C and D). Well-sorted, horizontally-laminated and low-angle large-scale cross-stratified sandstone, interbedded with and transitional to lacustrine mudstone, characterize this facies. The following features are common in the shoreline facies: small-scale cross stratification, climbing ripples, ripple marks, flaser bedding, desiccation cracks, ooids and pisoliths (Fig. 6G), and lag conglomerate. Associated fossils include molluscs (some in growth position), ostracods, vertical burrows, oncolites, rootlets and other plant remains. Wave agitation and local reworking is indicated by cross-stratified and horizontally-laminated sandstone and intraclast-lag conglomerate. The shoreline probably consisted of beaches with localized mudflats and lagoons. A narrow sandy shoreline is suggested because of the high local relief in the Ridge B asin and the general coarseness of the detrital sediment. The widespread lateral and vertical extent of this facies is probably related to fluctuations in lake level.
1 79
Lacustrine facies, Pliocene Ridge Basin, California Nearshore lacustrine facies
Carbonate and terrigenous sediments deposited in nearshore environments display the greatest lithologic diversity of any lacustrine facies (Fig. 5C and D). This facies commonly includes blue-grey mudstone, fine- to medium-grained sandstone, ooids, pisoliths, oncolites, stromatolites (Fig. 6H and 1), and intraclasts. Strata are horizontally-laminated, cross-stratified and massively bedded, from poorly- to well sorted, and commonly burrowed. Conglomerate- and sand-filled channels and desiccation cracks are present. Stromatolites are most common in this facies (Link et a/., 1978) and are commonly associated with molluscan, ostracod and plant remains. Ooids, pisoliths, oncolites, and ostracod-bearing strata are interbedded with the stromatolites and are transitional into adj acent facies. The nearshore facies is transitional between the shoreline and offshore lacustrine facies, and is characterized by at least three depositional factors. ( 1 ) Coarse-grained sediment influx by fluvial processes from the alluvial fan and shoreline facies is indicated by channelling, textural changes, and the incorporation of stromatolites, ooids, pisoliths and oncolites into dominantly terrigenous sediment. (2) Wave-agitated lacustrine sedimentation in the nearshore facies is indicated by cross-stratified, well sorted strata, oscillation ripple marks, the formation of ooids, pisoliths and oncolites, and the presence of intraclasts, which in part may reflect storm activity or fluctuations
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1 80
Martin H. Link and Robert H. Osborne
in lake level. (3) Subaerial exposure of at least the shallower parts of this facies is indicated by normal mudcracks, desiccation cracks within stromatolites, and the occurrence of intraclasts. Repetition of these three factors in response to climatic and tectonic conditions produced the shoreline and nearshore sedimentary sequences illustrated in Fig. 5D. B ar-complex facies
Thin- to thick-bedded isolated packages of well-sorted, cross-bedded, white sandstone occur within thick molluscan-rich mudstone notably along the western margin (Fig. 8A and B). These sandstone bodies are commonly convex upward with low angle cross bedding dipping both on and offshore and contain mudstone rip-up clasts, pebble lags, molluscan debris, and burrows (Fig. 8B). Individual bar complexes have erosional to gradational lower contacts with low-angle accretionary (Fig. 9E), tabular and festoon cross bedding which grades upward into pebble and mudstone chip lags with small-scale cross stratification (climbing ripples) and shell debris (Fig.
Fig. 9. Photographs of the western marginal lacustrine facies; (A) interfingering stratal packages of
composite sandstone with mudrock; (B) close-up of composite cross-bedded sandstone interfingering with mudrock (C) asymmetrical ripple marks on top of a bed; (D) mudcracks; (E) low-angle (accretionary) cross bedding in sandstone; (F) small-scale cross stratification (xb) and shell debris (sh); (G) convex-upward sandstone interpreted as a bar complex.
Lacustrine facies, Pliocene Ridge Basin, California
181
9F). The upper contact is commonly convex upward (Fig. 9G) with low-angle cross bedding dipping both onshore and offshore, and gradational upward into blue-grey mudstone. These sandstone bodies are up to 3 m thick and at least from 1 0 to 3 0 m long. It is difficult to tell what type of bar complex these represent, but their geometry and stratigraphic position within mollusc-bearing mudstone suggest these are most likely lacustrine bars deposited parallel to shoreline. Fluvial-deltaic facies
Several well-developed fan deltas are present in the Ridge B asin Group (Fig. 8A and C). Individual depositional lobes are from 30 to 90 m thick and consist of sandstone interbedded with mudstone (Fig. 9A and B). The sandstone is white to blue grey, fine- to coarse-grained, cross-bedded, ripple marked (Fig. 9C), burrowed, and commonly amalgamated. Slump folds, pull-apart beds, load structures, mudcracks (Fig. 9D) and rootlets are present. The mudstone is blue-grey to black, ripple marked, mudcracked, burrowed, and contains much organic material and many rootlets. A typical vertical sequence (Fig. 8C) displays a few thin sandstone beds overlain by a composite thick cross-bedded channel sandstone with accretionary foresets (?). This sequence is, in turn, overlain by massive and gradational sandstone and mudstone sequences which are commonly mudcracked. These cycles may be interrupted or repeated several times. These fluvial-deltaic complexes clearly entered shallow water and built into the lake, prograding over initial bottomset and foreset strata. Locally thick fluvial sequences cap these complexes and with delta shifting or abandonment, thick deposits of interdistributary mud and sand accumulated. Slumping, contempo raneous faulting and the generation of turbidites occurred off the front and flanks of these prograding fan deltas. Offshore lacustrine facies
The offshore lacustrine environment consists of two end members: one is an off shore turbidite facies and the other is predominantly a mudrock-carbonate facies. Each of these lithologies accumulated at different times in or near the centre of the basin. Offshore turbidite facies
An offshore turbidite facies is characterized by thin- to thick-bedded sandstone interbedded with organic-rich mudstone and shale (Fig. l OA and B). Strata are commonly graded; contain mudstone rip-up clasts, dish structure (Nilsen et a/., 1977), and sole marks; and are slump folded. Individual thin turbidite sandstone units are intercalated with the mudrock-carbonate facies and are graded, contain Bouma Ta, Tab, Tbc, and Tcde intervals, sole marks, mudstone rip-up clasts, and may be highly burrowed or ripple marked. Thick-bedded turbidite sandstone units are best exposed in the lowermost sandstone package of the Ridge Route 'formation' (Figs 2 and 3). This package contains the marine-lacustrine transition in the Ridge Basin and probably represents prodelta fan deposits accumulating in relatively deep-water lacustrine or brackish marine environments (Figs 5, l OA and B). This sandstone package is over 1 300 m thick and thickest in the central part of the basin where it interfingers and changes facies to coarser-grained marginal lacustrine units along the flanks of the basin. In this lens-shaped package palaeocurrent directions suggest
1 82
Martin H. Link and Robert H. Osborne
sediment transport to the southeast along the axis of the basin. Individual sandstone units are thick-bedded, commonly amalgamated, and contain dish structures (Fig. l i D), mudstone rip-up clasts, and Bouma Ta and Tab intervals (Fig. l i E). They are interbedded with massive to laminated mudstone that commonly is slump-folded (Fig. l l A and C) or pulled apart (Fig. l i B) (brecciated). Molluscs, pebbly mudstone, and small-scale cross stratification occur locally in the mudstone sequences. Numerous thickening- and thinning-upward cycles (Mutti, 1 974) occur in the lower part of this sandstone package and these are interpreted as outer and middle fan deposits (Fig. lOA and B). A laterally continuous conglomeratic sandstone body overlies this turbidite sequence (Fig. l l F). This conglomeratic sandstone, which is commonly amalgamated (Fig. l l G) contains channels, cross bedding, molluscs, and ostracods. This overlying sequence is interpreted as a deltaic and/or inner-fan distributary channel complex, which initially supplied sediment to the underlying turbidite sequence and later prograded over it.
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Mudrock-carbonate facies
A mudrock-carbonate facies is characterized by massive to horizontally-laminated mudstone (Figs 1 2A, B and 1 3A), shale and siltstone, which contain ferroan dolomite, other iron carbonates, analcime, pyrite, j arosite, and gypsum nodules (Irvine, 1 977). This facies was deposited lakeward of the coarser-grained marginal facies. Sedimentary structures include parallel, varve-like laminations (Fig. 1 3B), concre tions (Fig. 1 3C), soft sediment deformation (Fig. 1 3 D), small-scale internal faulting
Lacustrine facies, Pliocene Ridge Basin, California
1 83
Fig. I I . Photographs of the offshore turbidite facies: (A) Large-scale slump-folded strata showing growth
faults with laterally continuous overlying and underlying strata; (B) brecciated bed bounded by laterally continuous strata; (C) small-scale slump-folded and disrupted strata; (D) dish structures with pillar columns (water escape features); (E) graded sandstones (Bouma a intervals) which contain rip-up clasts separated by darker mudstone; (F) thick-bedded laterally continuous sandstone separated by darker mudstone; and (G) cross-bedded and amalgamated conglomeratic sandstone interbedded with thinner mudstone.
(Fig. l 3 E), desiccation features and locally graded and brecciated beds. Biological constituents include plant remains, fish, insects, peloids, shell debris, and minor burrowing (Fig. l 3 E). Total carbon (as C%) ranges from 1 ·5 to 3 · 8 and bulk ferrous iron content ranges from 1 ·5 to 5·0% FeO for a few selected samples (Irvine, 1977). The analcime and ferroan dolomite both appear to be early diagenetic with the analcime formed from the reaction of montmorillonite with saline, alkaline waters and the ferroan dolomite formed by diagenesis of a high Mg calcite precursor phase in a reducing environment (Irvine, 1977). The association of black, highly-organic mudstone and shale, various carbonates, analcime, pyrite, and the paucity of in situ fossils and burrowing may suggest reducing conditions at least within the sediment. Reducing conditions may have been enhanced by chemical and/or thermal stratification within the water column. Minor burrowing also may indicate anaerobic conditions along the sediment-water interface. Periodic subaerial exposure of these sediments or subaqueous syneresis cracks are suggested by desiccation features and brecciated beds.
1 84
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(modified after Irvine, 1 977).
CONCLU SION S
Ridge Basin lake (Fig. 14) formed in a narrow, small, tectonically active basin in which a great quantity of coarse-grained terrigenous sediment accumulated in alluvial fan-fluvial complexes which flanked the lake. Stromatolitic algae, fresh-water molluscs, ostracods and fish comprised the maj or biota, which lived in marginal lacustrine environments that were subj ect to periods of major terrigenous influx, subaerial exposure and wave agitation. In the offshore lacustrine environments, pro deltaic turbidites and mud-rich carbonate rocks accumulated. These offshore, non turbidite sequences contain dark, organic-rich dolomitic and analcimic mudstone and shale with pyrite. The physical and chemical nature of the Ridge B asin lake changed considerably through its 10 million year history. The lacustrine phase started as an infilling of a lake or semi-restricted marine embayment as represented by the thick turbidite-deltaic sequence of the lowermost Ridge Route 'formation'. At that time the lake or embayment was relatively deep (Fig. 1 4) with some external drainage into the adj acent
Lacustrine facies, Pliocene Ridge Basin, California
1 85
Fig. 13. Photographs of the offshore lacustrine mudrock-carbonate facies;(A) Carbonate mudrock
composed of ferroan dolomite and analcime at the base of Pyramid Dam; (B) light-coloured dolomite and analcime interbedded with dark-coloured, organic-rich shale; (C) nodular concretions of gypsum (?); (D) small-scale deformed beds; and (E) mottled (burrowed) and internally faulted carbonate mudrocks.
West
� A LLUVIli.L FAN OlE POSITS [§] LACUSTRINE tr.IUO D TUA8LOITES
Eosf
� SLUNP FOLDED STRATA � SLOPE OR DELTA FRONT CHANNELS � SHORELINE SAND
Fig. 14. Diagram depicting inferred 'deep-water' lacustrine and/or marine palaeoenvironments in the Ridge
Basin, California.
1 86
Martin H. Link and Robert H. Osborne East
West
LAKE
-�
- C STR ��'b�l oNCOLITES � DOLOMITE (FE RROAN) I++ +- [ LA
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INE
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L
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Fig. IS. Diagram depicting inferred 'shallow-water' lacustrine palaeoenvironments in the Ridge Basin,
California.
Ventura marine basin. These conditions are suggested by palaeocurrent directions to the southeast, the extremely thick turbidite section (== 1000 m) and the large-scale (�30 m thick) slump folded intervals within the turbidite section. With continued strike-slip displacement along the San Gabriel fault, external drainage from the Ridge Basin lake was blocked to the south and thus it became a closed lacustrine system. A relatively shallow, internally-drained lake formed as suggested by thick carbonate-mudrock sequences, organic-rich shale, evidence of desiccation and the dominance of 'shallow water' lacustrine environments that interfinger with marginal alluvial fan-fluvial deposits (Fig. 1 5). The final major event in the history of the Ridge Basin lake was the termination of strike-slip displacement on the San Gabriel fault coupled with infilling of the lake with fluvial and fanglomerate deposits of the Hungry Valley Formation.
A CKNOWLE DGMENT S
We wish to thank J. C. Crowell for his interest, advice, encouragement and use of unpublished information concerning the geology of the Ridge Basin. Mark Newton assisted with part of the field work on which this paper is based. J. C. Crowell, J. D. Cooper, and C. E. Turner-Peterson kindly reviewed this manuscript. The illustrations were drafted by Janet Dodds and the manuscript was typed by Anne Snell. A Research and Publication Grant from the University of Southern California partially offset the cost of illustrations. Contribution No. 369, Department of Geological Sciences, University of Southern California.
Lacustrine facies, Pliocene Ridge Basin, California
1 87
REFERENCE S
D.l. ( 1950) The Piru Gorge flora of Southern California. Carnegie Inst. Washington, Pub/. 590, 1 59-2 14. CROWELL, J.C. ( 1 950) Geology of Hungry Valley area, Southern California, Bull. A m. Ass. Petrol. Geol. 34, 1 623- 1 646. CROWELL, J.C. ( 1 952) Probable large lateral displacement on the San Gabriel fault, Southern California. Bull. A m. Ass. Petrol. Geo/. 36, 2026-2035. CROWELL, J .C. ( 1 954) Geologic Map of the Ridge Basin area, California. California Div. Mines Bull. 170, Mapsheet 7. CROWELL, J . C . ( 1 973) Ridge Basin, Southern California. Soc. econ. Paleont. Miner., Pacific Sec. , Guidebook, Trip 3, l-7. CROWELL, J.C. ( l974a) Sedimentation along the San Andreas fault, California. In: Modern and Ancient Geosynclinal Sedimentation (Ed. by R. H. Dott, Jr and R. H. Shaver). Spec. Pubis Soc. econ. Pa/eont. Miner., Tulsa, 19, 292-303. CROWELL, J . C . ( l 974b) Origin of late Cenozoic basins in Southern California. In: Tectonics and Sedimentation (Ed. by W. R. Dickinson). Spec. Pubis Soc. econ. Paleonl. Miner., Tulsa, 22, 1 90-204. CROWELL, J . C . ( 1 975) The San Gabriel fault and Ridge Basin, Southern California. In: San A ndreas Fault in Southern California (Ed. by J. C. Crowell). Spec. Rep. California Div. Mines Geol. 1 18, 208-233. DAVID, H.B. ( 1 945) A Neogene stickleback from the Ridge formation ofCaalifornia. J. Paleont. 19, 3 1 5-3 1 8. EATON, J.E. ( 1 939) Ridge Basin, California. Bull. Am. A ss. Petrol. Geo/. 23, 5 1 7-558. E H LERT, K.W. & EHLIG, P.L. ( 1 977) The 'Polka-dot' granite and the rate of displacement of the San Andreas fault in Southern California. A bs. Prog. geol. Soc. Am. Meetings, 9, 4 1 5-4 1 6 . IRVINE, P.H. ( ! 977) The Posey Canyon Shale - A Pliocene lacustrine deposit of the Ridge Basin, Southern California. M.S. Thesis, University of California, Berkeley. JENNINGS, C.W. ( 1 953) Geology of the southern part of the Quail Quadrangle, California, Spec. Rep. California Div. Mines Geol. 30, l - 1 7 . LINK, M . H . , OsBORNE, R . H . & AWRAMIK, S . M . ( 1 978) Lacustrine stromatolites and associated sediments of the Pliocene Ridge Route Formation, Ridge Basin, California. J. sedim. Petrol. 48, 143- 1 5 8 . MILLER, W.E. & D ow Ns, T. ( 1 974) A Hemphillian local fauna containing a new genus ofantilocapvid from Southern California. Contrib. in Science, Los A ngeles County Nat. Hist. Museum. No. 258. M U TTI, E. ( 1 974) Examples of ancient deep-sea fan deposits from Circum-Mediterranean Geosynclines. In: Tectonics and Sedimentation (Ed. by W. R. Dickinson). Spec. Pubis Soc. econ. Paleont. Miner. , Tulsa, 22, 92- 1 05. NILSEN, T.H., BARTOW, J.A., STUMP, E. & L IN K , M.H. ( 1 977) New occurrences of dish structure in the stratigraphic record. J. sedim. Petrol. 47, 1 299- 1 304. SELLEY, R.C. ( 1 970) Ancient Sedimentary Environments. Cornell University Press, Ithaca. S HE PARD, J . B . , J R ( 1 962) San Gabriel fault zone. Bull. Am. Ass. Petrol. Geol. 46, 1 938- 1 94 ! . SQUI RES, R.L. ( 1 978) Medial Pliocene dragonfly nymphs Ridge Basin, Transverse Range, California. Abs. Prog. geol. Soc. Am. Meetings, Cordileran Sec. WI LLIAMSON, C.R. & PICARD, M.D. ( 1 974) Petrology of carbonate rocks of the Green River Formation (Eocene). J. sedim. Petrol. 44, 738-759.
AXELROD,
Spec. Pubis int. Ass. Sediment. ( 1 978) 2, 1 89-203
Lacustrine sedimentation in an evaporitic environment: the Ludian (Palaeogene) of the Mormoiron basin, southeastern France
GEORGES TRUC Departement des Sciences de Ia Terre, 15-43 Bd du 1 I Novembre 1 918, Universite Lyon 1, 69621 Villeurbanne, France et Lab. associe au C. N. R. S. No. 1 I 'Paleontologie stratigraphique et Paleoecologie'
A B STR A CT
The Palaeogene basin of Mormoiron is one of many sedimentary basins and grabens developed along the Rh6ne-Sa6ne axis of southeastern France. During the Palaeogene the Rhodanian trough, between the Massif Central and the Alps, consisted of several subsiding grabens separated by stable platforms. From west to east small grabens were located at Ales, Nimes and at Manosque, bordering the Durance fault, and basins at Mormoiron and Apt Forcalquier. The graben structures in the Ludian were formed as a result of an east-west extensional tectonic phase following the Lutetian and Bartonian compression of the 'pyreneo-proven<>ale' phase. The structural regime produced a palaeogeography of isolated lake basins in which evaporites were formed. In the basin of Mormoiron, a thin sequence of evaporites developed on the platform areas, and a much thicker sequence was deposited in the more rapidly subsiding basin centre. In the Lower Ludian, terrigenous clastics were deposited around the margins of the basin, with finer material reaching the basin centre where organic rich clay accumulated. In the eastern part of the basin, near the Monts-de- Vaucluse, limestones with brackish-water molluscs (Potamides, Tympanotonus and Me/anopsis), ostracoda and foraminifera were deposited. Elements of a marine fauna indicate adaptation of certain species derived from a nearby marine area. Evaporites also formed in the basin centre during the Lower Ludian so that lenticular gypsum crystals are scattered in the organic-rich clay. During the Upper Ludian, a decrease in supply of terrigenous material led to desiccation and the Mormoiron lake basin became part of a much larger evaporitic area. Dolomite, sepiolite and magnesian smectites formed in the basin centre together with gypsum. At the basin margins limestones became more dolomitic. A fauna persisted in this evaporitic situation through ground-water springs from the Monts-de-Vaucluse. Periodically,
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
1 90
G.
True
desiccation of the basin produced mud-cracked horizons. locally with bird and mammal footprints. At the end of this arid period, an extensive development of algal mats covered the whole area of the basin. This was followed by a renewed phase of terrigenous clastic sedimentation indicating a return to a more humid climate. Diagenetic phenomena in the sulphate deposits include anhydrite after gypsum and secondary gypsum (alabastrine-type) after anhydrite.
Re sume
L e bassin Paleogene de Mormoiron fait partie d'un ensemble de bassins et de grabens repartis le long de !'axe Rh6ne-Sa6ne, dans le sud-est de Ia France. Une tectonique en extension est-ouest ayant pris naissance au Ludien apres les phases compressives nord-sud du Lutetien et du Bartonien, conduit il ia formation de grabens separes les uns des autres par des plates-formes plus stables ou se developpent quelques bassins it subsidence moderee. Au Ludien, ce schema morphotectonique favorise !'existence d'un endoreisme important et l'isolement des bassins ou s'accumulent des depots evaporitiques. Au Ludien inferieur, le bassin de Mormoiron est envahi par une formation detritique d'origine fluviatile. Le materiel grassier est localise sur les marges du bassin tandis que seules les particules les plus fines atteignent le centre du bassin, milieu euxinique ou se forment les premiers cristaux de gypse. Sur les marges orientales du bassin, au pied des Monts-de- Vaucluse. apparaissent des calcaires it faune saumatre avec mollusques (Potamides, Tlmpanotonus et Melanopsis) ostracodes et foraminiferes adaptes it un rriilieu oligo- it mesohalin et provenant d'un reservoir marin peu eloigne. Pendant le Ludien superieur. Ia decroissance des apports detritiques, correlative d'une aridification de plus en plus affirmee du milieu. conduit it une sedimentation ou alternent dolomite, sepiolite, smectites magnesiennes et gypses tan dis que sur les marges Ia sedimentation carbonatee devient de plus en plus dolomitique. La me me faune persiste dans le secteur oriental grace aux emergences provenant de Ia nappe aquifere des Monts-de Vaucluse. A Ia fin de cette periode chaude et sou vent a ride, Ia plus grande partie du bassin est recouverte par un vaste manteau stromatolitique immediatement surmonte par des apports detritiques qui marquent le retour it des conditions climatiques plus humides. Les phenomenes diagenetiques les plus caracteristiques se manifestent au niveau du gypse et plus particulierement sur les marges du cirque evaporatoire (transformation/remplacement gypse-anhydrite at gypse secondaire de facies albatre).
INTRO D UCTION
The Palaeogene basin of Mormoiron is part of a complex of sedimentary continental basins in southeastern France, developed in the region of the Sa6ne and Rhone valleys. From the Valence graben situated between the Massif Central and the western foothills of the Alps, the Rhodanian trough enlarges to the south and is replaced by several subsidiary grabens (Ales, Nimes, Manosque) separated by stable platforms with very localized zones of subsidence (Fig. I ) such as the Mormoiron and Apt basins. In this region, all the grabens and basins contain Ludian evaporitic deposits and/or bituminous limestones. Structural framework
Before the Ludian, the region had been submitted to north-south compressive stresses from the Middle Cretaceous until the Bartonian. Some anticlines formed during the Cretaceous are still recognizable, such as the Comtadin anticline (Masse & Philip, 1 976). During the Lutetian and Bartonian the north-south 'pyn!neo-proven �ale' orogenic phase produced a strong anticlinal/synclinal fold pattern over the whole
191
Lacustrine sedimentation
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MORMOIRON
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sedimentary rocks others besides Ludian eruptive and metamorphic rocks
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halite
gypsum
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Fig. I . Location map and structural framework showing the position and the palaeogeography o f the most
important Ludian basins in the Rhone valley, France.
area, with the intensity of folding decreasing from south to north (Triat & True, 1 97 4). From the end ofthe Cretaceous and especially during the Palaeocene the area between
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8
springs of the southeastern margins of the basin of Mormoiron
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Mts-de-Vaucluse water table (during the Ludian)
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stromatoUtic dolomites ("Calcaires des Pitis et de Ia Nesque") gypsum ("Gypses de Mormoiron'')
dolostones (dolomitic belt with scarse gypsum) ("Dolomies bl:lnches de Blauvac")
�
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limestones with calc. algae and oncolites ("Calcaires .i Algues des Bastides") limestones with lignitic levels ("Lignites et calcaires de Mtthamis") green sandy marls and days ("Argiles vertes de Mormoiron")
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black sh les of the dee�r part of the . eva pontiC area northern conglomerates and marly sands ("Conglomtrats de Crillon-le-Brave") cretaceous limestones (kustified during 1he Paleocene)
Fig. 2. Evolution of depositional environments and of sediments during the Ludian in the basin of Mormoiron. I, early evaporitic phase (Lower Ludian); 2, main evaporitic
phase (Upper Ludian); 3, development of basin-wide stromatolite unit (Uppermost Ludian).
;:: '"'
Lacustrine sedimentation
1 93
Valence and Mt Luberon suffered intense weathering that resulted in karstification of the Barremian (Urgonian) limestones (Fig. 2). Calcretes and silcretes also developed at this time. Terrigenous material derived from the Valence-Mt Luberon area was transported into a lacustrine basin extending from Mt Luberon to Marseille. The beginning of the Ludian period was marked by an episode of east-west tension following the Lutetian and Bartonian compression. This rifting affected the 'pyreneo-proven�ale' folds and resulted in northeast-southwest oriented structures, such as at Ales, Ni'mes, Manosque and Valence, in which great amounts of halite, anhydrite and bituminous limestones were deposited. More stable platforms in the neighbourhood of Apt and Mormoiron separated these grabens and were areas oflittle subsidence. Shallow-water carbonate sedimentation took place on these platform areas and sulphates were deposited in local areas of enhanced subsidence (Fig. 1). This structural framework produced a palaeogeography of grabens and basins of interior drainage which were isolated from each other and separated from contemporaneous marine areas by an important highland region (Gubler et al. , 1 975). Evaporites and their structural context
The Palaeogene of the Rhone valley (Ludian to Chattian) is comprised of several thick sequences of persistent clastic horizons which pass up into evaporites (Triat & True, 1 974). The alternation of clastics and evaporites, and the gradual passage of the former into the latter, are taken to indicate a climatic control on sedimentation of more humid followed by more arid phases. Within this tectonic framework, however, deposits of marls and limestones occur containing no trace of evaporites but with some freshwater faunas which existed during more humid periods. Penecontemporaneous tectonism, chiefly in the form of fault movements (Fig. I ), controlled the fluviatile clastic contribution to the basins during some periods of the Palaeogene. These sediments, commonly of stream-flood type, were concentrated at basin margins with only silts and clays spreading out towards the basin centres. Such a pattern is similar to other evaporitic basins and grabens (such as the Dead Sea) where marginal areas are dominated by a narrow strip of coarse detrital material which passes basinward into laminated mudstones (Gubler et a!., 1 975). Evaporite mineralogy is also controlled by the morphology and structure of the sedimentary basin. Halite is always localized in rapidly subsiding grabens while on the adj acent platforms there is no halite at all but many occurrences of gypsum. The preferential concentration of halite in basin centres is attributed to solution by undersaturated waters of any halite precipitated on platform areas and the progressive migration of these waters to basin centres. Evaporative concentration during arid periods then produces substantial halite deposits. Busson ( 1 968, 1 974) and others have involved a similar mechanism for the preferential concentration of halite in localized areas of subsidence.
T HE EV A PORITE B A SIN OF MORMOIRON
During the Ludian, the basin of Mormoiron behaved as a platform, with only moderate and even subsidence so that the formation of evaporites was not as substantial as in graben structures.
194
G.
True
Evaporites and their diagenesis
(a) Lower Ludian: initial phase of gypsum precipitation At the beginning of the Ludian the compressive stresses of the B artonian 'pyreneo-proven<;ale' orogenic phase gave way to a tensional tectonic regime as mentioned above. The northern margins received a strong detrital influx (Fig. 2. 1 ) of which only the finest materials spread over the entire basin. During this time, small lenticular gypsum crystals developed within laminated bituminous shales (Fig. 3a). This type of gypsum seems to have formed within the mud, a few centimetres beneath the sediment-water interface. These crystals were entirely replaced by anhydrite, without any deformation of the original gypsum crystal shape. Displacive enterolithic veins of 'primary' anhydrite also developed (Fig. 3b). Most of this anhydrite after
Laminated bituminous shales with sulphates of the first evaporitic phase. (a) Small lenticular gypsum crystals, which formed a few centimetres beneath the sediment-water interface, are entirely replaced by anhydrite. (b) Lenticular gypsum crystals pseudomorphed by anhydrite are displaced by an enterolithic layer of 'primary' anhydrite. Core samples from Quarry at Mazan (Vaucluse).
Fig. 3.
gypsum has been preserved through rapid burial, following this replacement. As a result of these observations, it is suggested that the original crystallization of gypsum was followed by its dissolution (not dehydration) and then precipitation of anhydrite
Lacustrine sedimentation
1 95
took place in the ghosts of the gypsum crystals and in the form of enterolithic veins. lt is thought likely, therefore, that the replacement resulted from the introduction of undersaturated waters to the sediment and from the later crystallization of anhydrite, following a period of evaporation. (b) Upper Ludian: main phase of gypsum precipitation The supply of detrital material (Fig. 2. 1 ) had become insignificant by this time so that the basin of Mormoiron was largely an evaporitic area. Also during this time tectonic movements promoted active subsidence (Fig. 2.2). The gypsum facies of the Upper Ludian is particularly homogeneous. In the centre of the basin, where terrigenous detritus was scarce, the scouring effects of water are reflected in the nature and orientations of the gypsum crystals. Thin beds full oflens-shaped gypsum crystals alternate with dolomicrite (a in Fig. 4). The gypsum crystals grew just beneath the sediment-water interface during very early diagenesis and include a 'grass-like' form
4. Typical facies of the gypsum in the basin centre. Thin beds rich in lens-shape gypsum crystals (a), alternating with dolomicritic layers containing 'grass-like' gypsum. Growth of the latter occurred during early diagenesis and the crystals grew through overlying carbonate laminae and gypsum (b). Quarry at Mazan (Vaucluse).
Fig.
Growth of some gypsum crystals above the sediment-water interface formed small sedimentary traps in which finely laminated sediments were deposited. Local erosion of the gypsum crystals (a) produced gypsarenites. Quarry at Mazan (Vaucluse).
Fig. 5.
G.
1 96
True
Fig. 6. Top
of a large gypsum crystal corroded and overlain by a thin layer of gypsarenite. Quarry at Mazan (Vaucluse). Crossed nicols.
7. Stromatolitic fabric, consisting of alternations of dark, bituminous, dolomitic mud, with fine-grained gypsum reworked and trapped by algal filaments, occurring above gypsum crystals of early diagenetic origin. Quarry at Mazan (Vaucluse). Crossed nicols.
Fig.
Fig.
8. Algal filaments scattered within gypsum crystals. Quarry at Mazan (Vaucluse). Crossed nicols.
Lacustrine sedimentation
1 97
of gypsum and twinned crystals of swallow-tail type. Frequently, the crystals have grown through overlying non-lithified carbonate laminae (b in Fig. 4) to protrude above the sediment surface and provide sheltered sedimentary traps (Fig. 5). In such hollows, thin laminated deposits accumulated. Rapid changes in the sedimentary regime led to the reworking of the mud containing thin gypsum crystals. This resulted in gypsarenite beds and layers, and corrosion of the largest gypsum crystals (Fig. 6). These data suggest that the growth of gypsum within mud is related to the saturation of the interstitial waters during desiccation of the basin. The gypsarenites and corroded gypsum crystals imply the existence of undersaturated surface water entering the basin during a more humid phase. Stromatolites are not uncommon in this type of environment and in this Upper Ludian situation, algal filaments are present within bituminous dolomitic mudstones with fine-grained gypsum (Fig. 7). Algal filaments are also scattered within some large gypsum crystals themselves (Fig. 8). For the most part there have been few later diagenetic modifications to the early diagenetic fabrics described above. On the margins of the evaporite basin, green laminated marls are interbedded with gypsum. The green laminated marls, commonly bioturbated, are related to more humid (wetter) periods while the gypsum beds represent periods of aridity and evaporation. Farther landwards, dolomite constitutes a belt around the evaporites and shows a similar type of alternation with green marls. This lateral change is also accompanied by an important decrease in thickness of the deposits (Fig. 2). Some of the most ·interesting observations concern the gypsum facies in this marginal part of the basin. Some beds are composed of an alabastrine facies characterized by a nebulous appearance (similar to alabastrine gypsum of Holliday, 1 970) and containing displaced carbonate laminae. Thin sections show nodules (within a dark dolomitic clay) of gypsum containing relics of lath anhydrite (fibroradiate fabric). Rare lenticular crystals filled by secondary gypsum are present in the dolomitic clay (Fig. 9), as well as 'primary' lath anhydrite, also preserved as secondary gypsum. Commonly, the original sedimentary lamination is not entirely destroyed and ghosts of crystals with poorly defined outlines (lens-shaped or swallow-tail) are visible here and there. Thin sections from levels where such crystals are still abundant show relics of lath anhydrite in a gypsum matrix (Fig. 1 0). In this particular example, the crystal terminations are well preserved but in many cases the outlines are deformed and the crystals are infilled by a secondary gypsum composed of irregular small crystals with diffuse extinction. Satin spar veins are abundant, filling fractures and vugs in the gypsum beds and desiccation cracks in adjacent green marls. The satin spar is thought to result from the volume increase which characterizes the transformation of anhydrite to gypsum as elucidated by Shearman et al. ( 1 972). These diagenetic replacements also affect sediments in the areas where gypsum passes laterally into dolomitic beds (Fig. 2.2) and where the calcitic/dolomitic mud contains rare gypsum crystals, stromatolitic layers and evidence of current activity. After dissolution and/or dehydration of gypsum, displacive anhydrite infilled ghosts of gypsum crystals and vugs in stromatolitic layers. Carbonate particles were also corroded at this time. Secondary rehydration of anhydrite promoted the development of a thin homogeneous alabastrine facies containing corroded dolomitic laminae and horizontal satin spar veins (Fig. 1 1 ).
198
G. True
9. Nodules of nebulous secondary gypsum after anhydrite, showing relics of lath anhydrite. Lens shaped gypsum crystals and lath anhydrite, without a gypsum precursor, developed in the mud and are infilled by secondary gypsum. Quarry at Mazan (Vaucluse). Crossed nicols.
Fig.
Fig. 10. Gypsum ghosts filled by secondary gypsum after anhydrite. Relics oflath anhydrite are present and well preserved terminations of gypsum crystals. Quarry at Mazan (Vaucluse). Crossed nicols.
II. Replacement of calcitic/dolomitic layers (commonly stromatolitic) by secondary gypsum after anhydrite. Satin spar veins are present between two convoluted sheets of stromatolitic fabric. Mormoiron (Vaucluse).
Fig.
Lacustrine sedimentation
1 99
Before continuing, it should be stated that these diagenetic events characterize almost solely the margins of the evaporitic area, while in the basin centre the gypsum facies is primary. This is taken to indicate that in this case burial is not responsible for the transformation of gypsum to anhydrite. In fact, the depth of burial of the basin centre gypsum has not exceeded 200 m. The diagenetic events can be summarized as follows: (a) crystallization of gypsum as lenticular, twinned (swallow-tail) and 'grass-like ' crystals, precipitated through saturation of the water-body in the basin centre or of ground water within the sediments; (b) increasing aridity leading to desiccation and a fall in level of the water table; (c) gypsum was sub jected to the high temperatures that occur near the surface and changed into anhydrite, while further anhydrite ('primary') was precipitated directly in the mud without a gypsum precursor (examples described by Kinsman, 1 969; Bush, 1 973); and (d) a more humid climatic phase during which anhydrite was buried and converted (during burial or the ensuing uplift) into secondary gypsum and satin spar gypsum was formed (as discussed by Shearman et al. , 1 972). Fauna and flora of the evaporite sequences
Over the whole evaporitic area, well exposed bedding surfaces of both the gypsum and green marls show mud-cracks and are covered with footprints of birds and the trails of mammals. The marls and gypsum beds are also extensively burrowed. This can explain the abundance of bird footprints since the birds (probably waders) would have been feeding on the infaunal annelids and insect grubs responsible for the bioturbation. The presence of an infauna might indicate that the formation of gypsum was not a result of excessive salinities. Trails of mammals, chiefly of the herbivore Palaeotherium, are generally straight suggesting that they were obliged to cross the playa of Mormoiron by the most direct route to find some suitable vegetation. These points, together with pollen analyses, show that the Mormoiron basin was surrounded by forests and that the gypsum did not form under typical arid, desert conditions. Analysis of stable isotopes (see below) corroborate this hypothesis. It is envisaged that the gypsum formed under oligo-mesohaline conditions during warm and/or arid periods. Such a situation can be compared with the well documented Purbeck evaporites of southern England (West, 1 975) which contain evidence for a warm but not necessarily arid climate. With regards to the ultimate source of the salts constituting the Palaeogene evaporites in southeastern France, it has to be remembered that the environment is strictly continental with no marine connection. Analysis of stable isotopes of carbon, oxygen and sulphur (Fontes, personal communication; Fontes, Triat & True, in preparation) shows that the leaching of sedimentary formations (in particular Triassic evaporites) surrounding the evaporitic basins is the main source of the Palaeogene evaporites. End of the evaporitic cycle
At the end of the Ludian, the tensional tectonic phase waned and a climatic change took place. Rippled silts and sands spread over the entire surface of the basin, including the evaporitic areas. Above the clastics, a final evaporitic phase is represented by an extensive stromatolite horizon together with convoluted limestone which is commonly dolomitized. A small amount of gypsum was developed in the
200
G. True
formerly subsiding basin centre (Fig. 2.3). In some places, especially around the gypsum zone, limestones pass into dolomite containing thin lenses of gypsum and stromatolites. Subaerial exposure ensued and the undersaturated or fresh waters led to the dissolution of gypsum and the brecciation of dolomite. This process is similar to that described for the Broken Beds of the Purbeck in southern England (West, 1975), although for the latter, a later diagenetic event was invoked. Southeastern limestone deposits and the dolomitic belt
Throughout the Ludian, a local area in the southeast of the Mormoiron B asin was the site of pelloid-rich limestone deposition, with intercalations of lignite in the lower part of the sequence. The limestones are composed of numerous molluscs especially gastropods of the genera Potamides and Tympanotonus. These gastropod genera are adapted to environments of changeable salinities but are more common in mesohaline waters. Ostracods (studied by G. Carbonnel) and some genera of foraminifera are also present. As with the evaporites, burrowing organisms were abundant and have resulted in extensive bioturbation of the micritic limestones. The flora contains calcareous algae (related to Cayeuxia) and locally blue-green algae in the form of oncolites (Fig. 1 2). Charophytes are common, especially in the lignitic beds. Elements of a marine fauna are found in these micritic limestones and represent species which were derived from a nearby marine area, and which have adapted to this continental environment. Chemically the latter would have been somewhat similar to the environment from which they were derived. The analysis of the stable isotopes of carbon and oxygen from these limestones (Fontes, personal communication) shows that all the carbonate was derived from the leaching of the hinterland surrounding the
12. Oncolitic fabric showing well preserved filaments, in the limestones of the southeastern part of the basin. Les B astides, near Methamis (Vaucluse).
Fig.
Lacustrine sedimentation
20 1
basin, especially the karstified Monts-de-Vaucluse (Fig. 2). Dolomites which constitute a belt around the evaporitic area seem to be related to a replacement of calcite or aragonite. Fragments of organisms (molluscs, ostracoda, foraminifera) are strongly dolomitized, whereas sedimentary feature s (burrows, currents features etc.) are always well preserved. The relative abundance of magnesium arises from the fact that gypsum precipitation tends to raise the Mg/Ca ratio which promotes dolomitization of calcite (or aragonite). Some of this magnesium causes aggradation of smectites by the process of 'cation pumping' (Trauth, 1 974). Mudstones
Clays accompanying the formation of gypsum are p artly authigenic and partly aggraded or replaced. The most common clay minerals are authigenic magnesian smectites and sepiolites (Trauth, 1 974). In the northern area (Fig. 2) the clays are chiefly detrital, derived from the hinterland of marine Jurassic and Cretaceous sediments, with dominant illite and chlorite, or from Tertiary palaeosoils with montmorillonite and attapulgite. It is important to note the fundamental difference between clays of halite basins and those of platforms like Mormoiron. In halite basins, such as the Valence and Manosque grabens (Fig. l) clays are totally composed of well crystallized chlorite and illite diagenetically aggraded from an initial detrital origin (Sittler & Triat, 1 975). On the platforms, however, magnesian smectites and sepiolite dominate (Triat & Trauth, 1 972). These observations can be explained in terms of the structural framework. Strong subsidence and the proximity of active faults produced an influx of very fine-grained silt to the grabens, accompanied by a rapid burial. The platforms, on the other hand, distant from such zones, received detrital sediments more infrequently so that authigenesis of clays through the action of magnesium enriched waters was more likely. Basin-margin clastics
Coarse sediments characterize the northern part of the basin ('Conglomerats de Crillon-le-Brave') (Fig. 2). They resulted from the tectonic activity in the Lower Ludian and the existence ofsignificant relief in the hinterland. Conglomerates contain boulders up to 0·5 m 3, derived from the B arremian limestones which constitute the Mont-Ventoux (Fig. 1 ). The features of the sediments, with little sorting or regularity suggest stream-flood deposition in a wadi-like system. Detrital clays are the main component of associated marls which locally show a mottling which is typical of hydromorphic soils in humid areas.
L A K E S ED I M E N T A T I O N H I S T O RY
The first event, following the initiation of graben structures, consisted of widespread deposits of detrital material coming from the northwest. In the centre of the basin an increasing structural confinement promoted the formation ofbituminous marls (black shales with hydrocarbons) which interfinger with lenticular gypsum horizons. In the southeastern part of the basin (Fig. 2.C) limestones alternating with lignitic marls constitute a quiet deposit in an often marshy environment sheltering an abundant fauna of ostracods and molluscs. Ground water spring existed (Fig. 2.B), connected
202
G. True
with the karstified 'Mts-de-Vaucluse' water-table (Fig . 2.A). Increasing subsidence and climatic changes, towards an arid environment led to the formation of a great thickness of gypsum while laterally a dolomitic belt developed around the evaporitic area (Fig. 2.2). Limestones continued to be deposited in the quiet southeastern area. The decreasing rate of subsidence coupled with near-complete filling of the basin, controlled the appearance of a stromatolitic flat over the entire basin. A detrital influx at this time reflected climatic and structural changes that terminated the period of evaporite formation.
CONCLUSIONS
Great thicknesses of evaporites can occur in continental environments when a structural control leads to the development of an interior drainage basin. Such continental basins do not necessarily require very arid or desert conditions for the formation of evaporites. Marine-derived faunas can adapt to such conditions, particularly in the vicinity of an evaporitic area. The precipitation of gypsum does not require high evaporation rates when there is a source of evaporites in the surrounding hinterland which can be leached out. The Ludian basin ofMormoiron shows a pattern where a structural confinement due to active subsidence, together with a warm climate were the two main factors controlling the formation of evaporite. The movement of the water in the basin, and especially the movement of ground water, controlled the early diagenesis of the evaporites deposited.
A C K N O W L ED GM E N T S
I would like to thank the 'Centre National de la Recherche scientifique ' (Paris) for providing support for this work within the structure of the 'Laboratoire associe' No. 11, 'Stratigraphic Palaeontology and Palaeoecology' managed by Dr L. David to whom goes my profound gratitude. I wish to thank also the 'Pl iitrieres de France' and especially the manager of the plaster factory of Mazan, M . P. Tiffou, who has always given a friendly reception and facilitated this work. Thanks are due also to Dr K. Stapf (Mainz), Dr I. West (Southa mpton) and Dr A. Matter (Bern) who have read the manuscript and made helpful suggestions, and also to Dr M. E. Tucker (Newcastle upon Tyne) for the extensive rewriting and checking of this contribution.
REFERENCES BusH, P. (1973) Some aspects of the diagenetic history of the Sabkha in Abu Dhabi, Persian Gulf. In: Persian Gu lf (Ed. by B. Purser), pp. 395-407. Springer-Verlag, New York. BussoN, G. (1968) La sedimentation des evaporites; comparaison des donnees sahariennes :i. quelques theories, hypotheses et observations classiques ou nouvelles. Mem. Mus. nat. His/. nat. , Paris, 19, 128167. BussoN, G. (1974) Sur les evaporites marines: sites actuels ou recents de depots d'evaporites et leur transposition dans les series du passe. Revue Geogr. phys. Geol. dyn. 16, 189-207. G u B LER, Y. et a/. ( 1975) Dynamique des depots dans un bassin sedimentaire continental; exemple d'un bassin paleogene de Haute-Provence. 9eme Congr. Int. Sedim. , Nice, 5, 1 99-203.
Lacustrine sedimentation
203
D.W. ( 1 970) The petrology of secondary gypsum rocks: a review. J. sedim. Petrol. 40, 734-744. D.J. ( 1969) Modes of formation, sedimentary associations, and diagnostic features of shallow water and supra-tidal evaporites. Bull. Am. A ss. Petrol. Geol. 53, 830-840. MASSE, J.P. & PHILIP, J. ( 1976) Paleogeographie et tectonique du Cretace moyen en Provence: revision du concept d'Isthme durancien. Revue Geogr. phys. Geol. dyn. 18, 49-66. SHEARMAN, D.J., MOSSOP, G., DUNSMORE, H. & MARTIN, M. ( 1 972) Origin of gypsum veins by hydraulic fractures. Trans. Ins/. Min. Metal/. 81, 149- 1 5 5 . SITTLER, C. & TRIAT, J.M. ( 1 975) Les argiles du fosse d e Manosque; synthese des donnees e t interpretations. 9eme Congr. Int. Sedim., N ice, fieldtrip guide-book, A2, 44-45. TRAUTH, N. ( 1974) A rgi/es evaporitiquesdans Ia sedimentation carbonatee continentale tertiaire. Unpublished thesis, University of Strasbourg. TRIAT, J. M . & TRAUTH, N. ( 1 972) Evolution des mineraux argileux dans les sediments paleogenes du bassin de Mormoiron. Bull. Soc. Jr. Miner. Cristal/ogr. 95, 482-494. TRIAT, J .M. & TRue, G. ( 1974) Evaporites pateogenes du do maine rhodanien. Revue Geogr. phys. Geol. dyn. 16, 235-262. WEST, I. ( 1975) Evaporites and associated sediments of the basal Purbeck Formation (Upper Jurassic) of Dorset. Proc. geo/. A ss. 86, 205-225.
HOLLIDAY, KINSMAN,
Spec. Publs int. Ass. Sediment. ( 1 978) 2, 205 -224
Triassic lacustrine sediments from South Wales: shore-zone clastics, evaporites and carbonates
MAURICE E. TUCKER Department of Geology, University of Newcastle upon Tyne, Newcastle, NEI 7R U, U.K.
ABSTRACT Lacustrine shore-zone clastics, evaporites and carbonates pass laterally and vertically into the offshore lacustrine Keuper Marl (Norian, Triassic) in Glamorgan, South Wales. The shore-zone clastics are well-sorted breccias of beach origin which show a rapid lateral passage, into sandstones and siltstones containing wave-formed ripples, graded beds (of storm origin) and slump folds. Where the Triassic lake margin was in contact with Carboniferous Limestone, shore platforms and wave notches were cut into the bedrock during stillstands of lake level. Screes accumulated upon the terraces and were locally reworked by wave activity. Calcretes developed within the colluvium during the extended periods of lake shoreline retreat. Where the lake margin was distant from contemporaneous Palaeozoic exposures, fine-grained lake shore-flat sediments are interbedded with stream and sheet flood deposits. The evaporites are continental sabkha-type gypsum-anhydrite deposits (commonly replaced), exhibiting nodular and enterolithic textures. Replaced sulphate crusts with polygonal cracks are also present. The lacustrine carbonates consist chiefly of calcarenites with wave ripples and fenestrae, and cryptalgal limestones.
A minor facies is
rhythmically-laminated limestone comprised of calcisiltite laminae (suspension current deposits) in calcilutite.
Desiccation and tepee horizons and simple calcretes are
intercalated. The carbonates are regarded as shallow sublittoral to littoral in origin, with deposition taking place upon a near-horizontal shore-zone mud-flat. Interpretations of the lacustrine sediments suggest that there were significant fluctuations in lake level; terraces and coarse shore-zone clastics formed at high lake level stands and calcretes, evaporites and tepee horizons developed during retreats of the lake shoreline. The Keuper Marl was the offshore sediment of this Triassic lake, deposited subaqueously, but periodically exposed during times of regression when the sulphates developed within the sediments.
A consideration of the factors controlling
sedimentation suggests that both climate and tectonics were important. The sequence of lacustrine facies suggests that there were changes in chemistry of the lake waters.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
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Tucker
INT R O D U C T I O N
In Britain, the Triassic rocks below the Rhaetic are almost entirely continental 'red bed' deposits, formerly divided into two parts, the Bunter and Keuper. The latter is dominated by the Keuper Marl, a homogeneous red dolomitic mudstone, locally with evaporites. The depositional environment of the Keuper Marl has been much discussed (eg. Lomas, 1 907; Bosworth, 1 9 1 3 ; Wills, 1 970, 1 976; Warrington, 1 970, 1 974a; Tucker, 1 977). Views have ranged from a hypersaline lake, inland sea or epeiric sea, in which deposition was for the most part subaqueous to a megaplaya, desert plain or supratidal flat which was chiefly subaerial, save for local brine pools and salt lakes (comparisons made with the Rann of Cutch, Lake Eyre, or Colorado Delta). This paper is concerned with Triassic sediments in Glamorgan, South Wales, which formed at the edge of the Keuper Marl 'basin' of deposition and which, it is shown, are marginal lacustrine in origin. Although these deposits are not extensively developed laterally or vertically, they contain a great variety of lithologies, textures and structures. There is abundant evidence, including wave-cut platforms and beach sediments (described in this paper), to show that these sediments accumulated at the margin of a substantial water-body. The Keuper Marl is interpreted as the offshore sublittoral deposit of this water-body, and thus of subaqueous origin. The chief reasons for a lacustrine interpretation are the rapid lateral facies changes of littoral to sublittoral, the rapid vertical facies changes of subaqueous to subaerial, certain sedimentological features, a lack of evidence for tides and the absence of any faunal indications of a truly marine environment.
S T R A T I G R A P HY A N D F A C I E S
In South Wales, the Triassic below the Rhaetic consists of two contrasting facies, the Keuper Marl as mentioned above and its lateral equivalent, referred to locally as the Marginal Triassic (formerly the Littoral Triassic or Dolomitic Conglomerate, Strahan & Cantrill, 1 9 12; I vimey-Coo k, 197 4). The Triassic sediments rest with marked unconformity upon folded Carboniferous and Devonian strata, with the Marginal Triassic resting upon and directly against the Palaeozoic rocks, passing laterally and vertically into the Keuper Marl (shown diagrammatically in Fig. 1). The Marginal Triassic is very heterogeneous and, apart from those of lacustrine origin described in this paper, contains sediments of alluvial fan, stream and sheet flood, colluvial, playa and pedogenic origin (Tucker, 1 977). The Keuper Marl, which is to be re-named the Mercia Mudstone Group (Warrington, l 974b), reaches a maximum thickness of 1 20 m in the Cardiff district (Strahan & Cantrill, 1 9 1 2). It is likely that the Glamorgan Keuper deposits belong to the Norian stage, Division 5 of Audley-Charles, 1 970a (Tucker, 1 977). The lacustrine facies of the Marginal Triassic are well exposed on the coast of south Glamorgan, between Barry and St. Mary's Well Bay and on Sully Island, and can be located inland at Dinas Powis (Fig. 2). The Keuper Marl itself is seen on the coast between Lavernock Point and Penarth, and at Barry Island. Inland exposures are generally poor or non-existent. Lacustrine sediments within the Marginal Triassic of South Wales are very variable in lithology and include limestones and dolomites, breccias through to siltstones and marls, and evaporites. Three main lithofacies can be
207
Triassic lacustrine sediments
M ARGINAL
TRIASSIC
alluvial fan and plain
lacustrine shorepzone clastics,
KEUPER
sediments.
carbonates and evaporites.
offshore lacustrine.
M ARL
1. Schematic diagram showing the relationship between the Marginal Triassic and Keuper Marl in South Wales.
Fig.
recognised, (i) lacustrine shore-zone clastics, (ii) lacustrine evaporites and (iii) lacustrine carbonates. In the region of Sully, these three lithofacies are developed in sequence (Figs 3 and 4) and the carbonates then pass up into the Keuper Marl. These lithofacies are described and interpreted in the following sections, and then discussed in terms of Keuper lake sedimentation and its controls.
A
, '-�,'
0
Fig. 2. A
2km
Keuper MRr/
, '
'
\
5km
& B: Sketch maps showing location of area and localities referred to in text. The Triassic considered in this paper (southwest of Cardiff) was deposited on the northern margin of the B ridgewater Bay-Somerset Basin (B.B.S.B. in Fig. 2A), located to the south of the South Wales Coalfield (S. W.C.) and Mendip Hills (M.), which formed upland areas. The locations of the Cheshire Basin (C.B.) and Worcester Graben (W.G.), sites of thick Triassic sediments including halite, are also shown. C: Triassic palaeogeography of the area around Cardiff (for details see Tucker, 1977) showing context of lacustrine sediments described in this paper.
Maurice E. Tucker
208
SUL L Y ISL AND
Keuper Mart rippled and cryptalgal
lacustrine
limestones
carbonates
with tepees and c alcretes ---channelled base-
lacustrine evaporites (replaced)
---E��g�S��2J
nodular dolomite replaced sabkha -sulphate
shore- zone
clastic s
beach to offshore lake sediments Car b. Limestone
Fig.
3. Section at Sully Island, Glamorgan.
carbonates
replaced evaporites
shore- zone clastics Carboniferous Limestone 4. Marginal Triassic sediments at Sully Island, Glamorgan; succession shown on right. Length of hammer, 0·4 m.
Fig.
L A C U S T R I N E S H O R E -Z O N E C L A S T I C S
C lastic lacustrine sediments resting on Carboniferous Limestone at Sully Island
Triassic lacustrine sediments
209
(Grid Ref: ST 169669) form a wedge of sediment showing very rapid lateral facies changes. At B arry Island [Treharne's Point ( 1 1 0662) and Nell's Point ( 1 1 9663 ) ] Triassic clastic sediments are associated with shore-zone erosion and solution features (shore platforms and wave-notches) cut into Carboniferous Limestone. Marginal lake clastics of shore-fiat type, with replaced evaporites, are interbedded with stream flood conglomerates, sheet flood sandstones and soil horizons near Hayes Point ( 1 40672), a locality between Sully and Barry Islands (Fig. 2). Shore-zone clastic wedge, Sully Island
A wedge-shaped unit of clastic sediment rests upon a planar surface of Carboniferous Limestone, dipping about Y to the Triassic sediment above (Figs 3 and 4). At the thick end of the wedge, 5 m of sediment is capped by a dolomitized fenestral limestone (0·3 m thick). The latter passes laterally towards the thin end of the wedge into a hematitic dolomite which eventually rests upon the Carboniferous Limestone surface itself. There is a very rapid increase in grain size passing eastwards from the
Figs 5-8 all from the lake-margin clastic wedge (Figs 3 and 4), Sully Island. Well sorted breccias interpreted as a lacustrine beach deposit from the thin (coarse) end of the clastic wedge.
Fig. 5.
Small ridge developed on surface of sorted breccia, interpreted as a lacustrine beach berm. Length of hammer head, 0·2 m.
Fig. 6.
7. Dolomitic fine sandstone and siltstone from thick (fine) end of clastic wedge, interpreted as shallow sublittoral in origin. Wisps of laminae of coarser sediment are probably of storm origin. The well-sorted breccias of Fig. 5 pass laterally into the sediments of Fig. 7 in only 50 m. Length of hammer, 0-4 m. Fig.
Fig.
8. Recumbent slump fold formed in laminated dolomitic siltstone of shallow sublittoral origin.
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Maurice E. Tucker
thick to the thin end of the wedge. At the thin end of the wedge the sediments consist of breccias and conglomerates interbedded with pebbly sandstones (Fig. 5). Many of these rocks are calclithites (Folk, 1 97 4) locally dolomitized. Although some coarse beds are poorly sorted, a characteristic feature is the presence of single beds of well sorted breccia, from 0·08 to 0· 1 5 m thick (Fig. 5). Pebbles in these beds are angular or subangular, generally 1 0-30 mm in diameter. These well-sorted breccias do not have sharp scoured bases, but gradational upper and lower surfaces and can be traced over several square metres. The breccias have a grain-supported texture, but possess a matrix which is probably infiltrated. Lenses of well-sorted breccia are present with a poorly-defined, low-angle planar cross-stratification and local pebble imbrication. Two ridges of breccia and pebbly sandstone seen on one bedding surface (Fig. 6) may represent the structure responsible for the lenses. The ridges, straight-crested and parallel to each other, are asymmetric and between 0·3 and 0·4 mm wide. In the direction of wedge thickening, the well-sorted breccias and associated pebbly sandstones pass rapidly (over 1 0-20 m) into red pebbly sandstones, sandstones and siltstones. These finer calcilithites contain thin, often impersistent graded b eds, less than 50 mm in thickness, with pebble concentrations at the base. Lenses and more continuous beds of cross-lamination occur (single sets, height 1 0-20 mm) as well as thinly bedded and laminated units. Symmetrical, straight-crested ripples with rounded and locally sharp-crested profiles are seen on bedding surfaces. They have wavelengths of 5 0-100 mm and heights of 5- 1 0 mm. Continuing in the direction of wedge thickening, these dominantly sand-sized sediments pass laterally into dolomitic fine sandstones, siltstones and marls (Fig. 7). Many of these finer sediments are thinly bedded or laminated. More massive, structureless sediments contain graded layers and thin wisps of coarser sediment. An important feature of the finer grained, thin-bedded sediments is the presence of various bedding irregularities and contortions, and in particular, recumbent folds, from 0·05 to 0·3 m in height (Fig. 8). Symmetrical anticlinal folds also occur, located at specific horizons and containing disrupted laminae within them. Interpretations
Perhaps the most conspicuous feature of these sediments is the very rapid lateral passage, in only 50 m, from coarse breccias (Fig. 5) to fine siltstones (Fig. 7). There are few environments where such a transition occurs but one which it characterizes is the lake shore-zone. The narrowness of gravel lake shore-zones and the rapid passage into finer grained sublittoral sediments have been noted by Morrison ( 1964), Picard & High ( l 972a) and Donovan ( 1 975). These features basically result from the reduced wave activity and lack of tidal effects in lakes, particularly when compared with open marine coasts. The well-sorted breccias at the thin end of the wedge (Fig. 5) have all the characteristics of beach gravels developed along lake shores, where they form through wave swash-backwash effects, coupled with small fluctuations in lake level. The well sorted but angular nature of the breccias is also a reflection of the reduced wave activity of lake shore-zones. Morrison ( 1 964) has pointed out the importance of fetch in determining the texture of beach gravels of the Pleistocene Lake Lahontan; sorted gravels developed where the fetch was long, but in sheltered areas, gravels were not sorted or rounded. Similar textures, it should be noted, can develop in protected marine and quasi-marine situations such as along the shorelines of Scottish sea lochs
Triassic lacustrine sediments
21 1
and estuaries. The ridges of breccia (Fig. 6), giving rise to poorly defined cross stratification, are similar to the berms which develop on gravel lake shores during periods of more persistent wave activity ( eg. Dulhunty, 1 975). These lake beach ridges are on a much smaller scale than their marine counterparts. The pebbly sandstones, sandstones and siltstones into which the breccias pass laterally formed principally in the very shallow sublittoral zone, probably down to depths of several metres, although they would have been affected by breaker and swash-zone processes during falls in lake level. The lenticular, single sets of cross lamination probably formed through the development of wave-current ripples which are common in shallow sublittoral zones (Picard & High, 1 972b). The symmetrical ripples are wave-formed oscillation ripples, with the small wavelength indicating very shallow depths. The finest sediment, which constitutes the thickest part of the wedge, would have accumulated in slightly deeper water, generally below lake wave -base (shallower than in the marine environment). The sediment itself would have been transported to these depths in suspension and been derived principally from rivers draining into the lake but also from the turbulent shore-zone. Graded units of coarse material probably represent storm deposition, of sediment derived from shallower water (cf. McLeroy & Anderson, 1 966). The recumbent folds (Fig. 8) are interpreted as slump structures due to downslope mass-movement of sediment. Recumbent slump folds have been described from several other lacustrine sequences, eg. Smith ( 1 959), Belt ( 1 968), Sanders ( 1 968) Donovan ( 1 975) and Link & Osborne ( 1 978). Other folds and disturbed strata probably result from sediment dewatering. Earthquake shock is frequently invoked as a trigger mechanism for dewatering and slump structures so that contemporaneous fault movements may be indicated. The relatively high gradient of the surface upon which the sediments were deposited may also have been a causal factor. The thin limestone which caps the clastic wedge suggests that relief at the lake margin had been removed and a shore-zone mud-flat established. Emergence is suggested by the presence of fenestrae. The lateral hematitic replacement of the limestone towards and beyond the thin end of the wedge is interpreted as a pedogenic feature with the hematite horizon representing an iron-rich soil or ferricrete. Taken as a whole, the clastic wedge represents a fairly constant stand of the lake, with beach gravels passing offshore into sandstones and siltstones, and then (although not seen at this locality) into the Keuper Marl. Sedimentation eventually buried the lake shore-zone slope to establish a near-horizontal depositional surface upon which the fenestral limestone accumulated. A substantial contraction of the lake followed, during which an iron-rich soil horizon developed. Shore-zone sediments and geomorphic features, Barry Island
The lacustrine sediments at Barry Island are developed around inliers of Carboniferous Limestone and cut into the latter are several horizontal to sub horizontal steps or terraces (Figs 3, 9 and 10). The terraces are overlain by Triassic sediments (Fig. 1 0) so that any confusion with Pleistocene wave-cut platforms is precluded. The terraces are up to 1 5 m wide and have a slope which may reach 5 ' . Cliffs between successive terraces are vertical o r subvertical and in some cases are defined by prominent joint planes within the Carboniferous Limestone. The vertical distance between successive terraces varies from 0·5 to 5 m. Although the terraces have
212
Maurice E. Tucker
commonly been affected by Recent weathering, some appear to have had serrated surfaces as a result of differential Triassic erosion of the steeply inclined Carboniferous Limestone strata. At the foot of one former cliff, at the back of a prominent platform, a 0·4 m high notch is preserved, containing a little Triassic sediment (Fig. l l ) Less well preserved notches are seen elsewhere. .
Fig. 9. Triassic shore platforms cut into Carboniferous Limestone. Three wave-cut platforms (A, B and C) are shown, with B and C dipping slightly to left (NW). Scree sediments (D) rest at the back of platform B against the former cliff-line, and are overlain by beach sediments (E) deposited on platform C. The scree sediments were also planated during formation of platform C. Barry Island, west side of Treharne's Point.
The sediments associated with the terraces are of two types: (a) breccias consisting of poorly sorted, angular clasts of Carboniferous Limestone up to boulder size; and (b) better-sorted breccias and conglomerates, with clasts up to cobble size, and finer sediments. The first type occurs at the back of the terraces, banked up against the former cliff-lines (Fig. 1 0). A thick, large scale asymtotic cross-stratification with dip up to 20· is commonly developed and may be cut by prominent trough-shaped surfaces. Dripstone cements occur on the underside of some clasts and infiltrated silt may partly infill cavities. Elongate nodules of fine-grained calcite have commonly developed within the red matrix of the breccias, replacing and displacing the original sediment. This unsorted angular breccia is clearly of colluvial origin (cf. Caine, 1 967), and represents scree which accumulated at the foot of the cliffs, after formation of the terrace upon which it rests. The large-scale cross-stratification would have been controlled by angle of repose and surface movement of clasts. Periodic failure of the scree slopes resulted in small landslips and debris flows which produced the trough shaped surfaces. The calcite nodules are interpreted as pedogenic in origin, being
Triassic lacustrine sediments
213
Fig. 10. Detail ofTriassic shoreline: Triassic wave-cut platform (A) and former cliff(B) cut i n Carboniferous Limestone, with a second platform (C), 2 m higher. Ill-sorted, coarse breccias (scree sediment D), which accumulated at the back of first platform (A) were truncated by the later planation (C) and overlain by thin bedded beach sediments (E) deposited on platform (C). Low angle cross-bedded calclithites above (F) may be toe of a later scree slope. B arry Island, west side of Treharne's Point.
II. Triassic wave-notch cut into Carboniferous Limestone at the back of a shore platform. A higher platform is developed 0·5 m above the notch and is overlain by sorted breccia (beach sediment). Barry Island west side of Treharne's Point. Length of hammer, 0-4 m.
Fig.
214
Maurice E. Tucker
identical to the calcretes described from many present-day semi-arid areas (eg. Reeves, 1970; Goudie, 1 973), the Triassic e lsewhere in Britain (eg. Stee l, 1 975) and the O ld Red Sandstone (eg. Allen, 1 974). Sediments of the second type, sorted breccias and finer sediments, rest upon the terraces and are generally p lanar bedded (Fig. 1 0). Locally, cross-stratified lenses are present and shallow channel structures up to 2 m across and 0.5 m deep. Wave ripp les and desiccation cracks are developed within finer-grained sediments. In one instance, the sediments at the foot of a Triassic scree consist ofvery well sorted but angular clasts of Carboniferous Limestone, similar in texture to the sorted breccias (beach gravels) described above from Sully Island. Pebbly sandstones and siltstones (calclithites) with wave-formed ripples and desiccation cracks are interbedded. The well sorted breccias and finer sediments resting upon the terraces are c learly beach gravels derived from cliff retreat and cutting of the terraces. Channels within these shore-zone sediments were probably cut during s light ly lower lake- leve l stands (cf. Donovan & Archer, 1 975). Well-sorted breccias at the foot of scree s lopes resulted from wave-reworking in the lake shore-zone and this may have induced the failure of the scree s lopes. In several instances the colluvium at the back of one terrace is overlain by bedded and sorted breccias and finer shore-zone sediments deposited on the terrace above (Fig. 1 0). This re lationship shows that many of the terraces were cut in ascending order, and that earlier Triassic sediments were a lso cut by the terraces. The lateral passage of the coarse scree and beach sediments into the Keuper Marl can be seen at both Nell's and Treharne's Points. At the first locality the transition is very rapid (over 1 0 m). The Keuper Marl passes into a buff-coloured, do lomitic coarse mudstone, which rests on co lluvium and Carboniferous Limestone. Small channels, infilled with pebb les of Carboniferous Limestone occur within the Keuper Marl, together with packets of grey wave-ripp led and laminated calcisiltite. At the second locality, dolomitized fenestral limestones and fine calclithites with channel structures are the transitional sediments. Interpretations
Terraces or shore p latforms and notches are a characteristic feature of many present day shorelines, both of seas and lakes (eg. Twidale, 1976). Shore p latforms, now abandoned, were a common feature of the P leistocene lakes of the U. S.A. (eg. Lakes Lahontan and Bonneville, Morrison, 1 964). The formation of terraces has frequently been attributed to wave a,ction a lone (eg. Edwards, 195 1 ), hence the term wave-cut p latforms. Bradley ( 1 958) and Dietz ( 1 963) have both emphasized the importance of such surf-zone erosion in the production ofwave-cut p latforms. The importance is a lso recognised of water-level weathering, promoting solution of the bedrock, and of organic effects, particularly a lgal, a lso leading to so lution (eg. Wentworth, 1938; Hills, 1 949; Twidale, 1 976). The cliff-foot notches, too, may have formed through a combination of wave erosion and solution. The terraces and their sediments can be interpreted to give an indication of the sequence and time-scale of events a long the lake margin. From the data of Bowman ( 197 1 ), terraces developed during pauses in contraction of the Dead Sea wou ld each have take some 3000 years to form (28 terraces developed in 80,000 years). Most of those terraces, however, are cut into a lluvial fan or earlier lacustrine sediments and not into solid bedrock. Thus longer periods of relative lake-leve l stillstand, of several to
Triassic lacustrine sediments
215
tens o f thousands o fyears, were probably involved in the cutting o feach o fthe Triassic terraces. During the stillstands wave abrasion and possibly solution effects led to cutting of cliff-foot notches, retreat of the cliff-line and development of the shore platforms. Wave attrition of liberated clasts produced well-sorted beach gravels and locally finer sediments with ripples. The stillstands were followed by falls in lake level, 'regressions ', when the terraces were abandoned and talus slopes accumulated upon the terraces at the cliff-foot. The build up of colluvium was mainly due to the lack of wave activity to rework and redistribute the screes as lake shore-zone sediments. Scree material was probably liberated through temperature changes, sheeting and exfoliation. The effects of these processes have been described from a nearby locality (Tucker, 1 974). Substantial periods of regression are indicated by the presence of calcretes within the screes. Calcretes of this type, simple rod-like nodules (glaebules) fall into the young to mature category of Reeves ( 1 970) and would probably have required up to 7000 years in which to form (Leeder, 1 975). Later lake-level rise or 'transgression' and then another stillstand resulted in a new and higher terrace being cut. The geomorphic features and the sequences of sediments show that there were significant fluctuations in lake level. Although such fluctuations would have had a more marked effect in the marginal lake area than within the lake centre, there should be some expression of the regressions in offshore sediments when they became emergent. The Keuper Marl, which is the offshore sediment of this lake, possesses evaporite horizons. It is shown in a later section that these evaporites are the type which develop within sediment through subaerial exposure. Shore-flat sediments with non-lacustrine interbeds, Hayes Point
Sandstones, siltstones and marl at Hayes Point contain many symmetrical, wave formed ripples, horizontal laminations and desiccation cracks. Disruptions and contortions of laminae also occur, as well as nodular carbonates of probable pedogenic and replaced evaporite origin. Interbedded are discrete packets of sorted conglomer ates, with channelled bases, ascribed to stream flooding and thin, but persistent graded sandstones deposited from sheet floods. (Tucker, 1 977, Tucker & Burchette, 1 977). Interpretation
The sedimentary structures of the finer-grained rocks indicate that they were deposited upon a lacustrine shore-flat, periodically covered by shallow water. During emergence at times of lake-shoreline retreat, evaporites (see below) and soil horizons formed and sporadic stream and sheet floods deposited coarse sediment. The situation would have been similar to present-day playas, which marginally interdigitate with fluviatile sediments.
LACUSTRINE EVAPORITES
Evaporites developed with the British Triassic are sulphates (gypsym-anhydrite) and halite. The halite chiefly occurs in the lower part of the Keuper Marl in areas of thicker sediment accumulation, such as the Cheshire Basin and Worcester Graben (Fig. 2A) (Warrington, 1 974a). Halite is not present within the Keuper Marl of South
216
Maurice E. Tucker
Wales, although it could well be present beneath the Bristol Channel as a continuation of the Somerset salt-field (Warrington, l 974a). Gypsum-anhydrite is developed as a lateral equivalent of the halite in the lower part of the Keuper Marl, and in the upper part occurs at several horizons of generally similar stratigraphic position in different parts of Britain (Warrington, 1 974a). In South Wales, sulphate is present as large but composite nodules of alabaster with associated fibrous gypsum, forming several horizons (up to 0·5 m thick ) in the upper 25 m of Keuper Marl at Penarth (Grid Ref. ST 1 907 1 5) . Within the Marginal Triassic, gypsum-anhydrite was present, but has been largely replaced by dolomite, quartz and calcite. A unit of replaced evaporites is developed above the clastic lake shore-zone sediments described in an earlier section from the Sully area (Fig. 4). The unit is up to 1 2 m thick and consists of wave-rippled, laminated and structureless dolomitic sanjy-silty marl in the lower part (2 m) which passes up into red nodular dolomite and dolomitic marl (Fig. 12). The latter contains scattered quartz nodules 50- 1 50 mm in diameter, which possess relic inclusions of anhydrite. This, together with other fabrics in the quartz nodules and dolomites, shows that the whole nodular dolomite and marl horizon is a replacement of anhydrite (Tucker, in preparation). The packed nature of the dolomite and quartz nodules (Fig. 1 2) is directly comparable with the chicken-wire texture of modern anhydrite as reported from the sabkhas of the Trucial Coast (Shearman, 1 966; Butler, 1 970). Irregular but continous layers of dolomite also occur and are analogous to the enterolithic texture of anhydrite. Sabkha-type sulphates form during early diagenesis by precipitation within subaerial sediments just below the surface. The context of the replaced evaporites from South Wales (occurring above and below undoubted lacustrine deposits), coupled with the lack of any definitive marine fauna, argues for a continental sabkha situation rather than a marine coastal situation. The latter was suggested by Stevenson ( 1 970) for slightly younger sediments (the Grey Marls) in Somerset. The sequence in the replaced evaporite unit of laminated and rippled
Fig. 12. Nodular dolomite from the central part of the cliff at Sully Island, formed by replacement of anhydrite. Length of scale, 0·2 m.
Triassic lacustrine sediments
217
sediment (shallow sublittoral), overlain b y nodular and enterolithic sulphate (now largely dolomite), is the sequence to be expected through sab kha progradation (Shearman, 1 966). A gradual fall in lake level, such as through persistent evaporation, would also favour development of sabkha-type evaporites. Evaporites are developed within the sediments around many present-day arid-zone lakes and playas as a result of flooding, groundwater evaporation and capillary rise (eg. Jones, 1 965; Hunt & Washburn, 1 966; Kinsman, 1969; Hardie, Smoot & Eugster, 1 97 8). Apart from forming within sediments, evaporites in desert lakes occur in two other ways, as surface crusts and capillary efflorescences (Jones, 1 965; Hardie et a!. , 1 978). There is evidence within the South Wales Marginal Triassic for the first of these two other modes of evaporite occurrence. At Hayes Point, near Barry (Grid Ref. ST 140672), a large-scale polygonal pattern (polygons 1 ·0- 1 ·5 m across) is developed on the surfaces of several nodular dolomite horizons. Evidence for the former presence of evaporites (gypsum-anhydrite) occurs within the dolomite. Large-scale polygonal cracking is a characteristic feature of the evaporite crusts developed on dried-up salt lakes (eg. Death Valley, California, Hunt & Washburn, 1 966; Lake Eyre, Australia, Madigan, 1 930). The polygonally cracked dolomite horizons are interpreted as replacements of such evaporite crusts (Tucker, in preparation). For the halite present within the British Triassic, precipitation from large, relatively deep, stratified salt lakes has been proposed by Evans ( 1 970) and others, while Arthurton ( 1 973) and Warrington ( 1 974a) advocated deposition from very shallow, periodically desiccated, brine pools or salinas. An aeolian origin has also been suggested (Holland, 1 9 12). The features of the South Wales Marginal Triassic, with its evidence for periodic and probably extensive lake shoreline regressions, would favour the second, salina model of halite formation. The identification of sabkha-type sulphate would also suggest this since many present-day salt lakes have sulphate precipitates within marginal sediments and halite (plus other salts) within the shallow, occasionally dry, lake basin centre (eg. Madigan, 1 930; Hunt & Washburn, 1 966).
LACUSTRINE CARBONATES
In the coastal area around Sully, limestones and dolomites occur above the replaced evaporite horizon (Figs. 3 and 4) (Tucker, 1975). The carbonates are chiefly calcitic in the lower part ( l -5 m thick) and dolomitic in the upper (3-5 m). At Dinas Powis (Grid Ref. ST 1 547 1 4) 2 m of similar limestone, but precise stratigraphic position unknown, rest upon a conglomerate of fluviatile origin. Three principal lacustrine facies can be recognized within the carbonates: (i) calcarenites, commonly rippled and fenestral intrapelsparites (Fig. 13); (ii) cryptalgal carbonates (Fig. 1 4) and (iii) rhythmically laminated limestones (Fig. 1 5 ). Thin pedogenic horizons (calcretes) also occur. These lacustrine carbonates are briefly described, details will be presented elsewhere. Calcarenites are composed of sand- and silt-sized clasts of Carboniferous Limestone, peloids, and carbonate silt ancf mud. The peloids are probably aggregates similar to those described from the lacustrine Green River Formation by Williamson & Picard ( 1974). Fenestrae or birdseye structures are commonly developed (Fig. 1 3) and may be equant or laminar types. Symmetrical, straight-crested wave-formed ripples are the characteristic sedimentary structure, although parallel lamination is also developed. Cryptalgal limestones are mainly of the planar laminated and columnar varieties, and
218
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may be composed of sediment or possess numerous fenestrae. Pelsparite laminae and layers of fibrous calcite which formed as crusts, are also present. Brecciation of laminites is common. Columnar stromatolites are generally on a small-scale with diameters between 1 ·5 and 20 mm (Fig. 14) . Linked domes are less frequent, but may be larger. Concentrically laminated pisolites are associated and some are oncolites. Rhythmically-laminated carbonates are a minor but distinctive lithofacies consisting
Triassic lacustrine sediments
219
of calcilutite with thin graded units o f calcisiltite, 0·0 1 -5 mm in thickness (Fig. 1 5). Coarser laminae have scoured bases and are noteworthy for containing disarticulated ostracod valves. Desiccation cracks have locally disrupted the laminae. One further feature of the lacustrine limestones is the presence of pseudoanticlines or tepee structures (Fig. 1 6), similar to those described by Assereto & Kendall ( 1 977). They occur at particular horizons and are spaced at 1 -3 m intervals. The tepees are developed in laminated and fenestral carbonates. Interpretations
The features of these sediments indicate shallow subaqueous deposition, inter rupted by periodic exposure, of a lacustrine shore-zone carbonate mud-fiat distant from terrigenous influence. Algal horizons are not uncommon in lacustrine limestones (eg. papers in Walter, 1 976) and in this case the columnar and domed varieties appear to have formed in a more permanently subaqueous situation compared with the planar type. Rhythmically laminated sediments are a characteristic feature of lake sequences (eg. Anderson & Kirkland, 1 960; B e lt, 1 968; Donovan, 1 975), since, compared with marine areas, seasonal variations in sediment supp ly and organic productivity are often more marked, and deposition, even in shallow regions, is less agitated. Some calcisiltite laminae are the result of low density suspension currents, which may have been seasonally controlled. The tepee structures formed through contraction and expansion of surface sediments through repeated desiccation and cement precipita tion, during extended periods of subaerial exposure. During more prolonged periods of lake shoreline retreat, the calcretes developed within the lacustrine limestones.
DISCUSSION Donovan ( 1 975) introduced the terms concident and non-coincident for lake margins, referring to the relationship of the lake margin to the basin margin. These terms can usefu lly be applied to the Triassic lake sediments of this paper and serve to illustrate the contro l that the position of the lake shore line has with regard to lithologies and facies variations (Fig. 1 7). At points of coincidence, lake sediments abut against surrounding basement and are generally coarse grained. In areas of non coincidence lake sediments are finer grained (since the sediment source has been buried or is distant) and may be interbedded with fluviatile sediments and soils. With the Welsh Triassic lake, the shore-zone gravels (as at B arry and Sully Islands) and reworked co lluvium (as at Barry Island) developed when the lake margin was coincident with the basin margin and lake waters lapped up against the surrounding Palaeozoic rocks. It was during these periods of coincidence that the terraces and wave notches were cut. Very rapid lateral facies changes from coarse shore-zone clastics out into the much finer-grained offshore deposits (the Keuper Marl) characterize the lake at this stage. In areas of non-coincidence of the Welsh Triassic lake, the paucity of coarse c lastics a llowed siltstones, marls and carbonates to form. A wide near horizontal mud-fiat would have exi�ted at this time and been subject to rapid changes from a subaqueous to subaerial state (and vice versa) as a result of s light changes in lake water level. Rapid vertical facies changes from lacustrine to subaerial characterize the lake at this stage and are exemp lified in the fluviatile horizons interbedded with
220
Maurice E. Tucker
lacustrine shore-fiat sediments at Hayes Point, and in the evaporite, calcrete and tepee horizons intercalated in the upper part of the succession at Sully Island (Fig. 3). Areas of non-coincidence may be due to the original topographic configuration (probably the case with Hayes Point), to burial of the marginal basement rocks (as is the case with the upper part of the Sully Island succession) or to lake shoreline retreat. ( a } H1gh lake-level stage Upland of Corb. Lst
Non- coinc1dent lake margin - i nterbedded lake and fluviatile sediments :
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:: ;jjh=JQ7
Fine dolo m i tic s i l t and cloy (Keuper Marl}
Scree
Local halite
Sabkha loco l ly w1th
Nodular anhydrite
17. Model of sedimentation of the Triassic lacustrine sediments in South Wales. High lake-level stage: shore platforms and beach breccias developed at coincident lake margins, and fluviatile sediments interbedded with lake sediments at non-coincident lake margins. River-transported silt and clay was deposited in the shallow sublittoral to constitute the Keuper Marl. Low lake-level stage: shoreline abandoned, calcretes developed in marginal sediments, sulphate deposited within subaerially exposed sediment as nodules and on surface as crusts, and halite precipitated from shallow brine pools in basin centre.
Fig.
As described above, the Marginal Triassic of the Sully area consists of a clastic wedge, overlain by an evaporite unit (replaced gypsum-anhydrite) and then a carbonate unit, before passing up into the Keuper Marl (Figs 3 and 4). Although there are strong lithological differences between the units, they all indicate similar fi uctuations of the lake shoreline, beginning with a transgressive event and followed at some stage by a regression, during which subaerial conditions prevailed and soils and evaporites formed. Similar sequences of events were deduced for the Barry Island Marginal Triassic where terraces were cut during a stillstand following a lake-level rise, and then calcretes developed during the lake-level fall. The question arises as to the underlying control of the lake-level fluctuations. There are two possibilities, tectonics and climate. The importance of tectonic movements as a control on lake base level is difficult to assess in this Triassic case. The Triassic (and Jurassic) of northwest Europe was generally a period of extension, related to the opening of the Atlantic Ocean (Hallam, 197 1 ) . Triassic sedimentation in B ritain was located within fault-bounded basins, such as the Cheshire B asin and Worcester Graben, and upon adj acent horsts, with much
Triassic lacustrine sediments
22 1
thicker sequences accumulating within the former (Audley-Charles, 1 970b). The South Wales Triassic was deposited on the northe rn side of the Somerset-Bridgewater Bay Basin, a fault-bounded basin to the south of the South Wales Coalfield and Mendips (Fig. 2A). It is possible that lake base-level was affected by contemporaneous movements upon the bounding faults. Fault movements are likely to have been normal and given rise to regressive phases at the basin margin. Warrington ( 1 970) has suggested from micro-palaeontological evidence that there were transgressions into the Triassic depositional basins of marine water derived from southern Europe by way of the North Sea or western approaches. Such transgressions would presumably have been tectonically controlled, a reflection of some change in continent-ocean configuration of the Tethyan (or other) region. If such influxes of seawater into the Triassic basins took place they would have produced transgressive events at basin margins. There are many cited examples of the effects of climatic changes on Pleistocene and Recent lakes (eg. Morrison, 1 964; Bowman, 1 97 1 ) and climatic variations have been invoked by Van Houten ( 1 962), and others, for ancient lacustrine sequences. Wills ( 1 970, 1 976) recognised cycles of several orders within the Keuper Marl of the Midlands, based on evaporite content, which he ascribed to climatic changes by comparison with Lake Eyre, S. Australia. If such variations occurred during the Triassic, then periods of more intense rainfall could have resulted in high lake-level stands during which terraces and notches were cut at coincident margins, and beach gravels derived from cliff erosion. Periods ofdrought and aridity could have produced low lake-level stands during which calcretes formed within marginal lake sediments and evaporites precipitated in finer sediments around the contracting lake. That there were variations in the Triassic climate can be deduced from the presence of ferricrete near the base of the Sully section and calcretes towards the top; the former indicating more humid conditions than the latter (Goudie, 1 973). One of the main effects ofthe climate on lakes is in its control of the water chemistry, through determining rainfall and evaporation rates. Several previous interpretations of the Keuper Marl have invoked hypersaline depositional waters (eg. Warrington, 1 970; 1 974a) and evidence for this is seen in the evaporite deposits. The clay mineral suite which characterizes the Keuper Marl (illite, corrensite, swelling chlorite and montmorillonite, with subsidiary palygorskite and sepiolite) is typical of alkaline, hypersaline, magnesium-rich waters, such as occur in arid-zone lakes and areas of inland drainage (Dumbleton & West, 1 966; Millot, 1970). In terms of water chemistry, the replaced gypsum-anhydrite unit of the Sully area shows that sediment porewaters were hypersaline (chlorinity in excess of 1 45o/00, Butler, 1 970) and so suggests that the lake water itself was to some extent saline. The lacustrine carbonates which occur above the replaced evaporite horizon show no evidence for the development of evaporites; this is in spite of the frequent subaerial exposure which occurred and as a result of which one would have expected gypsum and perhaps halite to have formed. This absence of evaporites is thought to be significant and is taken to indicate that the lake waters were substantially fresher during deposition of the carbonates, compared with the evaporite unit below. This postulated difference in lake water chemistry could well be climatically controlled: an arid period of little rainfall leading to strong evaporation, the development of a more saline lake and precipitation of evaporites from hypersaline sediment pore waters; followed by a more humid period of high rainfall when lake waters would have been relatively fresh and too dilute for evaporite
222
Maurice E. Tucker
formation. An additional factor could be proximity to rivers draining into the lake, especially in view of the carbonate-rich terrain (Carboniferous Limestone) which formed nearby upland areas in Triassic times. There is in fact abundant evidence for fluviatile activity elsewhere in the Marginal Triassic of South Wales (Tucker, 1 977). It is envisaged that the fine-grained offshore sediments of the lake (the Keuper Marl) were transported there by rivers, rather than being wind-borne, as has been suggested (eg. Lomas, 1 906; Sherlock & Hollingworth, 1 93 8 ; Taylor, Price & Trotter, 1 966; in part Wills, 1976). Aeolian reworking of subaerially exposed lake sediments, in particular the Keuper Marl, during periods of lake shoreline retreat, cannot be discounted, and indeed is likely to have taken place. CONCLUSION
On the basis of sedimentology, Triassic deposits from South Wales are attributed to a lake environment. Two diagnostic features are the rapid lateral facies changes of the lake shore-zone sediments and the rapid interbedding of lacustrine and subaerial facies. The former are functions of the reduced wave and current activity which distinguishes a lake from the open sea while the latter arises from the marked fluctuations in water-level, which are also characteristic oflakes. Climatic changes can be inferred from the sedimentology and could certainly account for the fluctuations in lake level. Contemporaneous tectonic movements controlled the locus of Triassic sedimentation, but it is difficult to assess any part they may have played in controlling facies sequences. As far as previous suggestions for the Keuper Marl depositional environment are concerned subaerial (megaplaya/supratidal flat) versus subaqueous (hypersaline lake/inland sea), the evidence presented in this paper indicates that both situations applied at different times (Fig. 1 7). Periodically the depositional area was subaqueous, with lake waters extending to the margins of the basin where wave-cut platforms and beaches developed in coincident situations (Fig. 1 7A). The Keuper Marl was deposited at these times, in the offshore environment, with the sediment transported into the basin by rivers. The presence of sabkha sulphates and calcretes show, however, that there were significant periods of subaerial exposure during lake shoreline retreat (Fig. 1 7B). At these times the depositional environment would have taken on more the appearance of a megaplaya with the undoubted possibility of aeolian reworking of the Keuper Marl. The sedimentological evidence also suggests that there may have been significant changes in lake water chemistry, so that it was not permanently or ubiquitously hypersaline. A C K N O W L ED G M E N T S
I am grateful to Pat Shannon and Geoffrey Warrington for comments on the manuscript, to Jennifer Brown for typing this and other papers in this volume, to Christine Cochrane for artwork and to Vivienne Tucker for field support. REFERENCES ALLEN, J.R.L. ( 1 974) Studies in fluviatile sedimentation: implications of pedogenic carbonate units, Lower Old Red Sandstone, Anglo-Welsh outcrop. Geo/. J. 9, ! 8 ! -208. ANDERSON, R.V. & KIRKLAND, D.W. ( 1 960) Origin, varves and cycles of Jurassic Todilto formation, New Mexico. Bull. Am. Ass. Petrol. Geol. 44, 37-52.
Triassic lacustrine sediments
223
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Permo-Triassic lacustrine deposits in the Eastern Karoo Basin, Natal, South Africa
D . E . V A N D I J K , D . K . H O B D A Y * and A . J . T A N K A R D Department of Zoology and Department of Geology, University of Natal, Pietermaritzburg, South Africa 3200, South African Museum, Cape Town, South Africa 8000
ABSTRACT Beaufort strata of the eastern Karoo Basin contain a classic lacustrine assemblage of freshwater fish, amphibians, reptiles, therapsids, crustaceans, insects and plants. The basin was subject to sediment influx from the south and east, with a subordinate contribution from the north. Maximum water depth was about 150 m. Fine-grained sediments settled over wide areas of the lake floor and were reworked by organisms. A predominantly massive lacustrine siltstone facies contains occasional complete skeletons of reptiles. A complex, shifting pattern of fluvio-deltaic progradation produced a highly irregular shoreline characterized by broad, open embayments, with smaller isolated ponds on the coastal plain. Lake shore encroachment generated upward-coarsening sequences with wave ripples, coprolites and fish trails in the silty lower part and traction-current structures in the sandstone above. Coarse-grained sediments were concentrated as mouth bars, distributary channel fills and overbank splays. Thin, upward-coarsening bay-fill sequences contain a great variety of fossils. Three bay-fill subtypes are recognized. One shows a gradual upward increase in grain size into fine-grained silty sandstones and is attributed to overbank flooding. The second contains erosively-based, lenticular crevasse-splay sandstones near the top, while the third was sediment-starved and has a high organic content. The lakes were eventually infilled by northward-prograding fluvial systems, and fluvial sedimentation finally ceased with the onset of arid conditions. Arid conditions also characterized the later Stormberg environment of dunefields, playa lakes and ephemeral streams. Typically the playa lake deposits show shallow water and emergence structures, together with intercalated evaporites. These lakes were a favourable habitat for dinosaurs, and ten different types of vertebrate footprints are present, including quadrupedal and bipedal walking tracks and bipedal hopping prints. Karoo sedimentation was terminated by widespread Jurassic volcanism. * Present address: Bureau of Economic Geology, University of Texas at Austin, Austin, Texas 787 12.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
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D. E. Van Dijk, D.
K.
Hobday and A. J. Tankard
I N T R OD U C T I O N
Climatic amelioration following the extensive late Palaeozoic Dwyka glaciation led to the development of the Ecca 'sea' in the intracratonic Karoo Basin (Fig. 1). Deglaciation was accompanied by a short-lived marine incursion from the west (McLachlan & Anderson, 1 973). This marine influence extended briefly across much of the Karoo Basin, as evidenced by microfossils from the Dwyka-Ecca transition at Pietermaritzburg (Fig. 2) near the eastern basin margin (1. McLachlan, personal communication). The Ecca basin probably comprised a single large body of water, deep in the south, with a broad, shallow shelf flanked by deltas and coal swamps in the north (Ryan, 1 967). Despite local micropalaeontological evidence of marine conditions in the Ecca (Hart, 1 964), an enigmatic nautiloid cephalopod (Rilett, 1 963) and occasional glauconite, the prevailing opinion is that the waters were for the most part only slightly brackish (McLachlan, 1 973), representing a transition into the freshwater conditions that prevailed during Beaufort deposition.
1. The Karoo Basin with study area enclosed. Palaeocurrents for the eastern part of the basin in part after Theron ( 1973).
Fig.
A diachronous relationship between the Ecca and Beaufort Groups has been suggested by Ryan ( 1 967) and Theron ( 1 973) as resulting from northward fluvial progradation fed by a tectonically-active southerly source. Generally lower energy conditions characterized the northern and eastern basin margins, but minor tectonism briefly exposed gneissic basement, resulting in local arkosic deposition (Theron, 1 973). Encroachment of the basin shores was probably accompanied by segmentation of the water body. In Natal there was a continuously changing pattern of lakes and embayments along the shores of the shrinking Beaufort basin. Renewed Triassic tectonism of the Gondwanide orogeny in the south initiated the Molteno fluvial wedge (Turner, 1 970) which is the basal stratigraphic unit of the Stormberg Group. A return to lower energy conditions is evidenced in the Elliot redbeds (Botha, 1 967), with extensive low-gradient alluvial floodplains building northward, and a generally warm climate. The final phases of Karoo sedimentation were influenced by a progressively more arid climate. The Clarens Formation (Cave
227
Permo- Triassic lacustrine deposits
Lavas
J
Sediments
Fig.
Storm berg Group
10 0 •••
20
40
Km
2. Generalized locality and stratigraphic map of study area.
Sandstone) was deposited in an environment of aeolian dunes and playa lakes
(Beukes, 1 970). Finally, outpourings ofthe Drakensberg Volcanics were accompanied
by widespread dolerite intrusion and dislocation of strata. This paper focuses on the Beaufort and Stormberg lake deposits exposed in the dissected country below the Drakensberg Escarpment in Natal (Fig. 2). Sedimento logical and palaeontological information is used to develop facies models which characterize known variations. Sediment supply was complex, and included a relatively fine-grained contribution from the distant southern tectonic terrain, together with influx from the nearer but less elevated sources to the east and north. As a result, several different, and sometimes opposing, directions of progradation are apparent (Fig. 1). N evertheless, relatively similar conditions prevailed under each system until the pattern was disrupted by the northward progradation of a major fluvial wedge.
228
D. E. Van Dijk, D.
K.
Hobday and A. J. Tankard
LACUSTRI N E SHELF FAC I E S
This facies includes grey, green and red siltstones and mudstones in units up to 40 m thick. Units thinner than 2·5 m are assigned to the 'bay' facies discussed below. The lacustrine siltstones vary from finely laminated to massive, the latter being more characteristic. Randomly oriented micas suggest that bioturbation has been largely responsible for destruction of original bedding. Laminated siltstones occur mainly in the greyish lowermost Beaufort and closely resemble the underlying Ecca. Intense bioturbation is not a widespread phenomenon in the Ecca, although individual traces are frequently well preserved (Hobday & Tavener-Smith, 1 975). This is ascribed to higher sedimentation rates in the Ecca, with the importance of biological processes increasing upward through the Lower Beaufort. In addition to millimetre-scale lamination, some of which is graded, there is also fine colour banding, with alternation of dark and light hues in texturally-uniform siltstone. These respective alternations may have resulted from differences in settling velocity of suspended fines, and periodic changes in water chemistry. In a restricted basin, it is possible that both were seasonally controlled. Larger-scale alternations of maroon and green-grey siltstone are particularly characteristic of parts of the Middle Beaufort. Occasionally the colours are admixed in a mottled or sublenticular arrangement. The colours may represent different oxidation states of iron. The green siltstones occasionally contain finely disseminated organic matter and rare pyrite, not present in subjacent red siltstones. The influence of a reducing microenvironment is also suggested by the occurrence of green colouration enveloping fossils in red siltstone. Skeletons of vertebrates, mainly therapsids (mammal-like reptiles), are frequently found in this facies, and include complete and perfectly articulated skeletons of the small therapsid Thrinaxodon. Large coprolites, fish scales and freshwater conchostraca are less common. Trace fossils include paired arthropod tracks and burrows and trails of sediment-ingesting organisms, but are seldom recognizable because of intense reworking. Deposition of this lacustrine facies was largely accomplished by suspension settling in an area of negligible traction current activity below effective wave base. Despite the presence of graded lamination there is no evidence of density flows. The sedimentation rate was low, with a varying proportion of organic contribution. The substrate varied from oxygenated to reducing.
FLUVIO-LA C U S T R I N E O FF L A P FAC I E S
The lacustrine shelf facies is generally overlain by an upward-coarsening succession with wave and current ripples (Fig. 3). Vertical gradation from suspension deposits to coarser traction and wave-reworked sediments takes place over a 2- 1 0 m interval. Wavy bedding oflow amplitude marks the base of the transition, and is associated with silty sandstone showing abundant Scolicia, arthropod tracks and horizontal spreiten structures. Occasional sinuous traces are ascribed to fish swimming in contact with the substrate, the more posterior fins leaving traces of greater amplitude. Impressions of plant leaves and stems are common, and tend to be concentrated on hummocky, bioturbated surfaces, along with fish scales, large bone fragments and coprolites.
229
Permo- Triassic lacustrine deposits liTHOLOGY AND STliUCTlJRES
m
16
scour
and fill. trough cross-
bedding and cross-lamination with
14
Fossil wood fragments
Fine to medium-grained sandstone
with
BIOLOGIGAL FEATURES
westarly azimuths Fine-grained sandstone with uni-
INTERPRETATION
Proximal mouth bar, locally SCOind by small distributary channels
.,. Rare vertical burrows
directional cross-lamination
Current dominated
12 Silty sandstone with combined wave10
Horizontal spraiten structures Linear wave-ripples. striking N-S
8
Wave dominated
current ripples
Sandy siltstone with oscillation ripples
l
Fish swimming trails Horizontal tracks and grazing
Distal ofllap deposits
trails Sandy si�stona. wavy and flasar structures
Vertebrate skulls and
4
miscellaneous bones; fish Homogeneous si�stone.
2
Fish scales Oriented plant stems
massive
sandy.
Finely laminated siltstone
scales. coprolites
Highly bioturbated
Lacustrine shelf
Homogeneous mudstone
Fig_ 3_ Section at Winston Hill, near Estcourt, showing regressive oftlap deposits. Wave-ripple trends and fore-set azimuths are indicated.
Ripples become more abundant upwards and show a progressive increase in amplitude. They are predominantly linear to slightly sinuous, with either sharp or rounded crests, and a symmetrical form. Low in the succession the ripples are internally symmetrical, and are ascribed to oscillatory wave motion. Ripples higher in the succession show fore-set laminae, and thus evidence of migration, probably under the impetus of translatory waves. Ripple spacing of between 8 and 16 em is attributed to short period, shallow water waves (Harms, 1 969). The crests commonly display a consistent north-south trend (varying between NNW-SSE and NNE-SSW, Fig. 3). which probably reflects the direction of the prevailing wave system. At some localities subordinate modes with varying degrees of obliquity may have been produced by local winds, shoreline configuration, or refraction.
230
D. E. Van Dijk, D.
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Hobday and A. J. Tankard
Upward gradation from sandy siltstone to silty sandstone is accompanied by a change from wave to current-produced structures, preserved mainly as microtrough cross lamination. Maximum thickness of the major sandstone units is between 5 and 20 m, and occurs above a scoured base. As the sandstones thin laterally the basal disconformity gives way to a transitional relationship. The thick medial sandstones commonly contain lenses ofmudclast conglomerate or siltstone in their lower parts, with the remainder consisting of fine-grained sandstone with cross-bedding, cross-lamination, and climbing ripples. Mud drapes occur on fore sets and along set boundaries. Medium to large-scale trough cross bedding is most common towards the top of the sandstone bodies, where it occupies scours and channels. Graded fore-sets and preserved brink points are observed in some of the thicker sets (30-80 em). Fore-set inclinations are unidirectional within a particular outcrop, but display regional variation from southwest to north. The maximum observed lateral extent of the sandstone bodies is 20 km. The extremities are uniformly ripple cross-laminated with siltstone partings and interfinger with bay or shelf-type sediments. The upper parts of the sandstone unit contain plant rootlets and vertical burrows, overlain by carbonaceous siltstone with vertebrate remains, desiccation cracks and plant stems in growth position. Pedogenic horizons are occasionally observed. At one locality near Oliviershoek there are concentrations of complete Lystrosaurus and other therapsid s keletons. The offiap sequences are interpreted as resulting from progradation of small delta systems. The lower, wave-reworked deposits would be analogous to the upper prodelta of a modern subdelta (Coleman & Gagliano, 1 965), and the upward-coarsening, current-dominated upper part to the mouth bar. Sedimentation varied between intermittent and rapid, as indicated by clay drapes on fore-sets, preserved brin k points and climbing ripples (Harms et al., 1 975). The medially scoured base and local lag concentrations are typical of the central part of a modern mouth bar influenced by high energy flow projected from the channel mouth (Horne & Ferm, 1 977). As the system prograded the upper surface was traversed by multiple smaller distributary channels. These were flanked by low levees bearing plants and subject to periodic desiccation. Concentration of vertebrate remains at this level may have been due to the stranding of floating carcasses along the river margins during floods. Alternatively, animals like Lystrosaurus may have foraged in swampy environments in the immediate vicinity (Colbert, 1 972).
B AY F A C I E S
Bay sequences are characteristically upward-coarsening and result from the encroachment of the bay margins by overbank flooding and the development of crevasse splays (Coleman & Gagliano, 1 964; Elliott, 1 974). Similar sequences in the Beaufort, ranging in thickness between 2 and 15 m, contain laminated to massive, slumped siltstone (Fig. 4) overlain by interlaminated siltstone and mudstone, attributed to current-induced differential settling. Included lenticular laminations formed as starved silt ripples on a muddy bottom. These structures are common in interdistributary bay environments where overbank flooding has introduced fine sediment (Coleman & Gagliano, 1 965; Horne & Ferm, 1 977).
·
23 1
Permo- Triassic lacustrine deposits m
LITHOLOGY ANO STIIUCTIJRES Foreset azimuths
5
BIOLOGICAL FEATURES
,.
Oesia:ation cracks
current ripples. lenticular
wavy
Overbank
sheetflow
deposits
and Haser bedding interference ripples
Strike of linear ripples 3
lavae
Sihstone and fine-grained
sandstone. vvave ripples. climbing
4
INTERPRETATION Emergent
-
Alternating sihstonelsandstona,
Plant
rootlets.
Fossil fish
Rhythmically laminated siltstone. lenticular and
wavy
bed-
ding
Subaqueous
lavae
Minor crevasse splays
slumps and load structures Siderite
rare
vertical burrows
Restricted conditions
Leaves Grazing trails. insect wings leaves Conchostraca,
Carbonaceous sihstone, slumped
Bay centre
fish scales leaves
Fig. 4. Section at Wagondrift, near Estcourt, showing upward-coarsening bayfill. Wave and current ripple directions are indicated.
Fossils in the lower parts include insects, freshwater conchostraca and plant debris. The interlaminated siltstones and mudstones contain Glossopteris fossils (Fig. 5) which retain carbon, indicating a reducing bottom environment. 20 em higher pyrite infilling of carbonaceous Glossopteris and sphenopsids (horse-tail rushes) has facilitated preservation of cellular detail. There are also minute pyrite-filled burrows. Above this level there is evidence of increased current activity, and the Glossopteris leaves are preserved as iron oxide impressions. The overlying sideritic layer contains a large freshwater fish A therstonia. Sideritic layers with overlying rock compacted around them are common constituents of bayfill sequences (Horne & Ferm, 1 977), and indicate mildly reducing, slightly alkaline conditions. Sandstone lenses in the succeeding layer were probably introduced by minor overbank splays. Flow rolls indicate subsequent downslope gravity displacement. Above this, and extending to the top of the sequence (Fig. 4) is siltstone and very fine-grained sandstone containing abundant wave ripples, lenticular, wavy and fiaser bedding, and climbing ripples including all three varieties of Walker ( 1 963). Bedding surfaces show interference ripples of various kinds, including ladderbacks. There are abundant desiccation cracks throughout, and convolute lamination occurs at several levels. These conformably-based silty sandstones were probably a product of overbank flooding and the concentration of the coarsest sediment fraction along levees flanking an embayment. Infilling of a bay by levee encroachment generates an upward coarsening sequence characterized by textural alternations (Elliott, 1974) such as are observed in the Beaufort. Convolute lamination associated with climbing ripples is also a common feature of subaqueous levees (Coleman & Gagliano, 1 965). Lenticular, wavy and flaser bedding, desiccation cracks, wave and ladderback ripples, and polymodal foreset azimuths are common attributes of a tidal fiat environment (Reineck, 1 967; Klein, 1 970). In the Beaufort, because this assemblage of structures is confined to the bayfill facies, we suggest that they were generated by winds blowing over shallow water from different directions, and by the refracted and highly
232
D. E. Van Dijk, D.
K.
Hobday and A. J. Tankard
5. Glossopteris from level l ·5 m above base ofWagondrift section (Fig. 4), preserved mainly as carbon, but with pyritization near the midrib and iron oxide at number of spots. Scale I em'.
Fig.
degraded swell from the open lake. Otherwise, there is no evidence for tides in the Karoo Basin. Some bayfill sequences include erosively-based sandstones overlying siltstones and mudstones (Fig. 6). The sandstones are thickest close to the subaerial levee deposits, beyond which they thin and fan out distally into the bays. They are interpreted as crevasse splays, as distinct from sheet-flood deposits, and resulted from local breaching of the levee and incursion of sediment-laden floodwater into the bay (Arndorfer, 1973; Elliott, 1 974).
[] Sandstone
Q Alternating rn Siltatone � Mudstone Fig.
VV Burrows sandstone and siltstone
Hl
Plant rootlets
(§)
Concretions
x x x
Siderite
.,111 Plant
fossils
'" � 0
40
6. Erosively-based crevasse sandstone cutting through rooted levee and bay deposits.
The sandstone in Fig. 6 is cross-bedded and cross-laminated and cuts obliquely through levee and bay deposits with burrows, rootlet beds and plant remains. A notch in the upper surface of the sandstone probably represents one of the radiating channels
Permo- Triassic lacustrine deposits
233
which commonly develop over a crevasse splay (Saxena, 1 976). According to Saxena crevasse sedimentation is generally an episodic and short-term process. Following abandonment of the crevasse channel and plugging of the levee the distal splay deposits are subject to wave and current reworking. Figure 7 illustrates several thin, reworked crevasse sandstones.
7. Thin reworked distal crevasse splays in bay sediments overlying mouth bar sandstone (at level of scale).
Fig.
Some crevasse sandstones indicate multiple episodes of sedimentation, with overlapping sandstone splays separated by siltstone drapes. Occasionally a hiatus between two sandstone splays is demonstrated by the presence of plant rooting. The crevasse splay sandstones are commonly overlain by highly carbonaceous siltstones and mudstones. This probably records the gradual reversion to marsh and bay conditions. Longer term occupation of a crevasse channel is indicated by sequences which resemble the fluvio-lacustrine offiap facies, but are considerably smaller, and overlie relatively shallow bay sediments. These resemble the sub-deltas described by Coleman & Gagliano ( 1 964). A further variation of the bay-fill facies is represented by an exposure of coal overlain by carbonaceous mudstone preserving spores, cuticles and complete specimens of the delicate moss Buthelezia (Fig 8f) which probably grew on an emergent levee nearby. Associated with these are numerous nymphal exuviae (shed skins) of Plecoptera (stone flies), the nymphs of which are aquatic. Other fossil insects include the adult plecopteran Euxenoperla, the paraplecopteran Mioloptera, a beetle and numerous homopteran sap-sucking insects (Fig. 8b ), including Beaufortiscus (Fig. 8d) and the scale insect A leuronympha (Fig. 8c). The insect remains consist mainly of wings, but whole insects besides the nymph exuviae do occur, suggesting proximity to source. The insects occur with a Glossopteris flora, including leaves of two or more species of Glossopteris, the male fructification Eretmonia and the female fructifications Lidgettonia andPlumsteadia. Some of the leaves show evidence of having been eaten (Fig. 8i). There is an upward increase in frequency and size ofsphenopsids from small side branches with leaves to large stems and rhizoids with roots attached. The rhizoids indicate uprooting of the plants, which were buoyant. Conchostraca (Fig. 8e) and fish scales are also present.
234
D. E. Van Dijk, D.
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Fossils of bay facies: (a) B eetle X 5; (b) Sap-sucking Homopteran X 5 - note p, thin proboscis, h head, t thorax, a abdomen and w wing; (c) A leuronympha (Scale insect) x 5; (d) Beaufortiscus (homopteran) wing X 5; (e) Conchostracan x 20; (f) Buthelezia (moss) X 2!; (g) Sphenophyl/um whorl (sphenopsid) X I ; (h) Equisetites rhizome with roots X I; (i) Glossopteris leaf with chewed margins x I - note penetration by root or burrow; (j) Conducting tissue of Glossopteris infilled by pyrite ( X 1 000).
Fig. 8.
Permo- Triassic lacustrine deposits
235
The overlying siltsones display small wave ripples, arthropod tracks, feeding traces and pyritized rootlets. Also present are pyritized Glossopteris leaves and sphenopsid stems. These sediments probably originated in a small, restricted pond, surrounded by marsh, and receiving only the finest overbank sediments. The substrate was highly reducing, with plant decay generating anaerobic conditions and favouring the formation of pyrite.
C HA N N E L F A C I E S
Erosively-based, upward-fining channel sandstones which display a pebbly base and a progressive vertical decrease in the thickness of cross-bed sets commonly overlie or cut through bay sediments. These sequences are interpreted as resulting from the avulsion of major distributary channels of the lower delta plain, a process described by Elliott ( 1 974). The channels switch their courses to an interdistributary depression because of the gradient advantage. Coarser-grained arkosic fluvial deposits form an insignificant proportion of the westward- and southward-prograding floodplain sediments. Northward influx is recorded in a 50 m thick braided channel complex in which the number of superimposed channel sequences varies between 1 6 and 28.
P L AY A L A K E F A C I E S
Above the Beaufort is a fluvially-dominated interval (Molteno and Elliott Formations) followed by the predominantly aeolian Clarens Formation (Beukes, 1 970) . An exposure of the Clarens Formation at Giants Castle (Fig. 2) consists of waterlain-deposits, which are bounded laterally by aeolian sandstones. A fluvial channel sequence at the base of the succession (Fig. 9) displays the typical upward-fining pattern with root-penetrated silty sandstone at the top. Above this is a highly bioturbated silty mudstone, suggesting the development of a shallow pond. The overlying sandstones and siltstones display rill marks, etch marks and runzel marks (Fig. 10), indicating fluctuating water levels. Subaerial exposure is evidenced by abundant desiccation cracks and thin gypsum layers. Graded beds with a variety of sole marks including flutes, grooves and prods are similar to sheet flood deposits recorded by Tucker & Burchette ( 1 977). The overlyin,g trough cross-bedded sandstones with abundant scour and fill structures suggest that the water body was surrounded by ephemeral streams. Within the limited outcrop area the patterns of etch marks, rills and linear desiccation cracks define an arcuate lake margin. The lake dimensions are not known, but adj acent exposures suggest that the maximum diameter was less than 2 km. The lateral transition into aeolian sandstones suggests that the lakes may have formed where watercourses were dammed by desert dunes. Trace fossils include gastropod trails, arthropod tracks, rare vertical burrows and a variety of vertebrate footprints. There are ten different types of vertebrate prints, including quadrupedal and bipedal walking tracks (Fig. 1 1 ) and bipedal hopping types. The latter were probably produced by dinosaurs which ranged from sparrow to ostrich sized (see Ellenberger, 1 970, pp. 3 50-354, figs. 1 1 4, 1 1 5, 1 1 8 and 1 1 9 for diagrams of four species based on photographs from this site). The shallow margin of
236
D. E. Van Dijk, D.
m
14
12
K.
Hobday and A. J. Tankard
LITHOLOGY ANO STRUCTURES
BIOLOGICAL FEATURES
Trough cross-bedded sandstones at variable grain size
INTERPRETATION
Ephemeral stream
Scour and fill
deposits (?)
Rare burrows
10
8
Graded coarse-grained Sheet flow
sandstone-si�stone Quadrupedal tracks Rute and tool marta
6
Playa lake. subject
Desiccation cracks
Bipedal dinosaur
Gypsum
tracks Shallow pond
Siltstone and fine-grained
sandstone, with 4
2
etch marta
rill marta,
and
runml
marta
with
Snail trails and
fluctuating water
artll ropod tracks
level
Silty mudstone
Highly bioturbated
Climbing ripples
Roollets
Trough cross-bedding
to drying out
l.rM Ruvial channel
Giants Castle section showing vertical gradation from fluvial channel into playa lake and ephemeral stream deposits.
Fig. 9.
an ephemeral lake was probably attractive to dinosaurs in an arid environment. Perfect preservation of the footprints is probably accounted for by the low degree of invertebrate activity in a high-stress environment, and by the cohesiveness of the sediment in which they formed. More intense biological reworking may explain the rarity of tetrapod tracks in the Beaufort, despite the wealth of amphibian, reptilian and therapsid fossils. In contrast, numerous tetrapod trackways are recorded elsewhere in the Clarens Formation, mainly in Lesotho (Ellenberger, 1 970).
D I S C U S S I O N A ND C O N C L U S I O N S
As has been pointed out by Picard & High ( 1 972) the deposits oflarge lakes resemble those of epicontinental seas, particularly in terms of their physical properties. Thus the Beaufort shelf and offlap sequences are similar in many respects to the deposits of shallow marine shelves and prograding shoal-water deltas (Gould, 1 970). Bay facies in the Beaufort are analogous to parts of the lower delta plain of the modern Mississippi (Coleman & Gagliano, 1964; Saxena, 1976). However, the palaeontological attributes of the Beaufort, particularly the fish, conchostraca, insects and abundant plant remains, indicate a non-marine basin.
237
Permo- Triassic lacustrine deposits
Fig.
10. Rills, etch and runzel marks in playa lake deposits, Giants Castle. Scale
I
em'.
11. Natural cast of footprints of a bipedal dinosaur. Note also the scour marks. About one quarter natural size.
Fig.
23 8
D. E. Van Dijk, D.
K.
Hobday and A. J. Tankard
Because of the Stormberg cover it has not been possible to trace the outline of the Beaufort waterbodies, or even to estimate their dimensions. Fluvio-deltaic palaeocur rent patterns (Fig. 1) suggest progressive reduction in lake area due to shoreline encroachment from the south, east and, to a lesst:r extent, from the north. The basin was probably elongated NNE-SSW along the line of maximum subsidence (Theron, 1 973). The outline was highly irregular, with broad, open embayments and isolated ponds on the low-lying coastal plain. There were relatively few distributary trunk channels, and a large proportion of sedimentation occurred by overbank processes. Maximum water depths in the main basin, as estimated from the thickness of shelf and offlap deposits (Klein, 1 974), was about 1 50 m, but the average was considerably less. Variations in lake level were probably brought about by changes in the main loci of sediment influx, compaction, tectonic subsidence, and the interplay between rainfall, freshwater inflow and evaporation. Shallow transgressions were accomplished without any evidence of shoreline erosion or reworking of sand bodies, probably because the limited fetch and water depth meant that wave energy was low. Regional palaeocurrent patterns persisted throughout the Beaufort (Theron, 1 973) despite changes in the relative contributions of the different source areas. Each episode of local accretion and lake filling is represented by an upward-coarsening offiap sequence which resembles the ideal lacustrine sequence (Visher, 1 965). However, there were several independent regressive systems, each with its own association of bay and channel deposits, complexly related to one another. Major regressive sequences are repeated as many as five times vertically, and some of the major sandstone units overlap or coalesce laterally. This pattern of lenticular sandstones of limited areal extent suggests that the overall geometry was determined by sedimentological, as opposed to tectonic mechanisms (Walker & Harms, 1 9 7 1 ) or independent changes in lake level. The lakes were eventually filled by major influx from the south. Fluvial sedimentation prevailed until the onset of desert conditions, when Beaufort sediments were reworked by westerly winds into transverse and barchan dunes (Beukes, 1970). According to Beukes, playa lake deposits are common in the Stormberg, and fish, crustacean and plant remains are recorded locally. Light-limbed dinosaurs indicate an adaptation to the arid climate (Haughton, 1924). Karoo sedimentation was terminated by widespread Jurassic volcanism. ·
REFERENCES ARNDORFER, D.J. ( 1 973) Discharge patterns in two crevasses in the Mississippi River delta. Mar. Geo/. 15, 269-287. BEUKES, N.J. ( 1 970) Stratigraphy and sedimentology of the Cave Sandstone Stage, Karroo System. Second Gondwana Symposium, South Africa, 1970, Proceedings and Papers. pp. 3 2 1 -3 4 1 . C.S.l.R., Pretoria. BoTHA, B.J.V. ( I 967) The provenance and depositional environment of the Red Beds Stage of the Karroo System. Gondwana Stratigraphy, pp. 763-774. U.N.E.S.C.O. Paris. COLBERT, E.H. ( 1 972) Antarctic Triassic tetrapods. In: A ntarctic Geology and Geophysics (Ed. by R. J. A die), pp. 393-40 1 . Universitetsforlaget, Bergen. COLEMAN, J . M . & GAGLIANO, S.M. ( 1 964) Cyclic sedimentation in the Mississippi River deltaic plain. Trans. Gulf-Cst A ss. geol. Sacs, 14, 67-80. CoLEMAN, J.M. & GAGLIANO, S.M. ( 1 965) Sedimentary structures: Mississippi River deltaic plain. In: Primary Sedimentary Structures and their Hydrodynamic Interpretation (Ed. by G. V. Middleton). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 12, 1 1 3- 1 48.
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ELLENBERGER, P. ( 1 970) Les niveaux paleontologiques de premiere apparition des Mammiferes primordiaux en Afrique du Sud et leur ichnologie. Second Gondwana Symposium, South Africa, 1 970, Proceedings and Papers. pp. 343-370. C . S.l.R. Pretoria. ELLIOTT, T. ( 1 974) Interdistributary bay sequences and their genesis. Sedimentology, 21, 6 1 1-622. GouLD, H.R. ( 1 970) The Mississippi Delta complex, In: Deltaic Sedimentation Modern and A ncient (Ed. by 1. P. Morgan). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 15, 3-30. HARMS, J.C. ( 1 969) Hydraulic significance of some sand ripples. Bull. geol. Soc. A m. 80, 363-396. HARMS, J.C., SOUTHARD, J.B., SPEARING, D.R. & W.UKER. R.G. ( 1 975) Depositional Environments as Interpreted From Primary Sedimentary Structures and Stratification Sequences. Soc. econ. Paleont. Miner., Tulsa, Short Course No. 2. HART, G.F. ( 1 964) Where was the Lower Karroo sea' Scientific S. Afr. I, 289-290. HAUGHTON, S.H. ( 1 924) Fauna and stratigraphy of the Stormberg Series. Ann. S. Afr. Mus. 12, 323-495. HOBDAY, O.K. & TAVENER-SMITH, R. ( 1 975) Trace fossils in the Ecca of northern Natal and their palaeoenvironmental significance. Pa/eont. Afr. 18, 47-52. HORNE, J.C. & FERM, J.C. ( 1 977) Carboniferous Depositional Environments in the Pocahontas Basin, Eastern Kentucky and Southern West Virginia. Dept of Geology. University of South Carolina. KLEIN, G. deY. ( 1 970) Tidal origin of a Precambrian quartzite - the Lower Fine-grained Quartzite oflslay, Scotland. J. sedim. Petrol. 40, 973-985. KLEIN, G. deY. ( 1 974) Estimating water depths from analysis of barrier island and deltaic sedimentary sequences. Geology, 2, 409-4 1 2 . MCLACHLAN, l.R. ( 1 973) Problematic microfossils from the Lower Karroo Beds in South Africa. Palaeont. Afr. 15, 1-2 1 . McLACHLAN, I.R. & ANDERSON, A . ( 1 973) A review of the evidence for marine conditions i n Southern Africa during Dwyka times. Pa/aeonr. Afr. 15, 37-64. PICARD, M.D. & HIGH, L.R. ( 1 972) Criteria for recognizing lacustrine rocks. In: Recognition of Ancient Sedimentary Environments (Ed. by J. K. Rigby and W. K. Hamblin). Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 16, 108- 145. REINECK, H.-E. ( 1 967) Layered sediments oftidal fiats, beaches and shelf ofthe North Sea. In: Estuaries (Ed. by G . H. Lauff). Am Assoc. Adv. Sci. Pub/. 83, 1 9 1 -206. RILETT, M.H.P. ( 1 963) A fossil cephalopod from the Middle Ecca Beds in the Klip River Coalfield near Dundee, Natal. Trans. Roy. Soc. S. Afr. 37, 73-74. RYAN, P.J. ( 1967) Stratigraphic and Paleocurrent Analysis of the Ecca Series and lowermost Beaufort Beds in the Karroo Basin of South Africa. Unpublished Ph.D. Thesis, University of Witwatersrand. SAXENA, R.S. ( 1 976) Sand Bodies and Sedimentary Environments of the Modern Mississippi Delta - An Excellent Model for Exploration in Deltaic Sandstone Reservoirs. Egypt General Petroleum Corp. Exploration Seminar, Cairo, Egypt, 1 5 - 1 7 November, 1976. THERON, J.C. ( 1 973) Sedimentological evidence for the extension of the African continent southwards during the Later Permian-Early Triassic times. Third Gondwana Symposium, Australia, 1 973, Proceedings and Papers, 6 1 -7 1 . TUCKER, M . E. & BURCHETTE, T.P. ( 1 977) Triassic dinosaur footprints from South Wales. Pa/aeogeogr. Pa/aeoc/im. Palaeoeco/. 22, 195-208. TURNER, B.R. ( 1 970) Facies analysis of the Molteno sedimentary cycle. Second Gondwana Symposium, South Africa, 1970, Proceedings and Papers, pp. 3 1 3-3 1 9 . C.S. l . R. Pretoria. YISHER, G.S. ( 1 965) Use of vertical profile in environmental reconstruction. Bull. Am. Ass. Petrol. Geo/. 49, 4 1 -46. WALKER, R.G. ( 1 963) Distinctive types of ripple-drift cross-lamination. Sedimentology, 2, 1 73- 1 88. WALKER, R.G. & HARMS, J.C. ( 1 97 1) The 'Catskill Delta': a prograding muddy shoreline in central Pennsylvania. J. Geol. 79, 3 8 1-399.
Spec. Publs int. Ass. Sediment. ( 1 978) 2, 24 1 -257
Subaqueous clastic fissure eruptions and other examples of sedimentary transposition in the lacustrine Horton Bluff Formation (Mississippian), Nova Scotia, Canada
R E I N H A R D H E S S E and H A R O L D G . R E A D I N G Department of Geological Sciences, McGill University, Montreal and Department of Geology and Mineralogy, Oxford University
ABSTRACT In the lower Mississippian lacustrine Horton Bluff Formation of Nova Scotia remarkable examples of sedimentary transposition structures occur. Transposition structures are secondary sedimentary structures that result from intrastratal viscous or hydroplastic flow. These include intermediate-sized circular and elliptical collapse structures associated with sediment extrusion from adj acent fissure vents and a swarm of more than I 00 parallel clastic dykes, many of which show evidence for extrusion at the sediment surface. The dykes are strongly folded suggesting very early diagenetic initiation of intrastratal sediment movement. The Horton Bluff Formation is made up of dark-grey to black and greenish-grey shale alternations grouped into sequences. The greenish-grey shale divisions contain fossil soil horizons (mainly in their upper parts). Signs of temporary emergence (mud-cracks, isolated tree stumps) are also found in the dark-grey shale division. Lithology, fossil content and association with fluvial deposits below and above suggest a lacustrine origin. The formation of the transposition structures is attributed to contemporaneous Mississippian earthquake activity in the neighbourhood of the Mississippian rift of Nova Scotia. Similar structures found elsewhere may therefore indicate palaeoseismicity and may be of help in evaluating the diagenetic history of sedimentary basins. For an extended German abstract of this paper see Hesse ( 1976).
INTRODUCTION
In the sea-cliffs on the south shore of the Minas Basin, Bay of Fundy, Nova Scotia, excellent outcrops of the lacustrine Horton Bluff Formation of the Mississippian Horton Group occur. Exposure is nearly three-dimensional with vertical sections in the cliffs and foreshore where bedding planes are accessible at low tide.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
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In the past the Horton Bluff Formation has received attention primarily from palaeontologists who were attracted by the findings of fish remains, amphibian bones and amphibian tracks, the latter being among the oldest known tetrapod footprints (Carroll et al., 1972). From a sedimentological point of view the Horton Bluff Formation is of interest in two respects. First, it represents an example of lacustrine sediments, which form a characteristic element of Late Palaeozoic-Early Mesozoic continental deposits in the Maritime Provinces of Canada. The lacustrine origin of these rocks has only been recognized by modern workers (Belt, 1 968a; Carroll et al., 1 972). They were considered largely fluviatile in origin by older investigators (Bell, 1 92 1 , 1 929; Crosby, 1 962). Second, the formation contains unusual sedimentary transposition structures in the cliff section at Horton Bluff on the west bank ofthe Avon River near Avonport Station of the Canadian Pacific Railway. Transposition structures are secondary sedimentary structures that result from intrastratal viscous or hydroplastic flow (Elliott, 1 965). The most impressive structures are large circular or elliptical collapse structures apparently caused by subsurface soft-sediment movement. Almost as remarkable is a swarm of parallel clastic dykes which apparently led to subaqueous sediment eruption at the lake bottom. Although these structures are well exposed, they are not mentioned in any of the older geological reports.
W ol f v i l l e S s Tr i a s s i c
j:::;:i;I . ..
.
.
.
.
Pennsylvanian
.
W i n d sor Group
C h evarie
Fm.
H or t o n B l u f f
Fm.
H a l i fa x S l a t e Or d o v i c i a n
1. Geological map of Wolfville area (after Bell 1929, Crosby 1 962; G.S.C. Map 1 1 28A). Nos I and 2 indicate the locations of measured sections. Inset: possible maximum extent of Hortonian lakes (dot pattern) in the Maritimes and relationship to the Fundy Basin (rift-valley) and Platform tectonic framework (after Belt 1968a).
Fig.
Subaqueous clastic fissure eruptions
243
S T R A T I G R A P HY A ND T E C T O N I C S E T T I N G
In the Wolfville map area of Nova Scotia (Crosby, 1 962) the Horton Bluff Formation represents the middle part of the Tournaisian (Lower Mississippian) Horton Group. It is conformably overlain by the Cheverie Formation (Bell, 1 929) and rests unconformably on older Palaeozoic rocks (Ordovician Halifax Slate, Fig. 1). The basal part of the Horton Group is missing at this locality. Age determinations are by plant fossils, invertebrates (Bell, 1 960), and miospores (Playford, 1 963). The area is located immediately south of the Carboniferous rift valley of Nova Scotia (Fig. l) on the edge of the stable Meguma Platform. On this Carboniferous platform a relatively thin and incomplete stratigraphic sequence accumulated, contrasting with the continuous succession in the rift valley which is five to ten times as thick (Belt, 1 968b). Hence the lower Horton Group fanglomerates; fluvial sediments and andesitic and rhyolitic volcanics known from the rift are absent from the Wolfville area. Three lithostratigraphic members of the Horton Bluff Formation may be distinguished (Bell, 1 929).
2. 'Main dyke bed' exposed in a slightly eastward plunging anticline. The dykes appear as northwestward sandstone projections (arrows) within the underlying shale where the 'dyke bed' has been denudated.
Fig.
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Reinhard Hesse and Harold G. Reading
The lower member consists oflight grey to reddish feldspathic quartz conglomerate, coarse- to medium-grained feldspathic quartz-sandstone (or quartz-arenite) and arkose, commonly cross-bedded, and interlayered with minor amounts of thin bedded, cross-laminated or parallel-laminated siltstone and shale. Colours of the latter are generally grey to dark grey but rarely red. The thickness varies from about 200 m in the south to over 600 m in the north (Bell, 1 929). The middle and upper members are dominantly fine-grained consisting of fine to very fine grey to dark grey sandstone and siltstone alternating with dark grey and greenish-grey to pale yellowish silty shale. Pyrite concretions and nodular dolomite form irregular bands, pyrite being more abundant in the middle, dolomite more abundant in the upper member. Dolomite also occurs as thin lutitic beds often showing polygonal crack patterns. The green shale horizons are most abundant in the upper member which also contains some brown to reddish shales (not at Horton Bluff) and intercalations of grey, medium to coarse arkose. These may form sheets of over 1 m thickness but usually are less than 30 em thick. In the Cheverie Formation these arkosic sandstones predominate; the silty shale becomes subordinate. At Horton Bluff the formation forms part of a broad east-west trending syncline. Within this syncline are minor, gently dipping, folds (Fig. 2).
dolomite concre t i o n s
Sm
HORTON
BLUFF
FM.
4
vreenish silty shale
main dike b e d
dark g re y shale
0
rippled s a n d stone
silty sandstone
�§ -
2
3
4
5
6 a,b
7
8
9
10
11
3. Vertical succession of eleven lacustrine black-green shale sequences (No. I oldest; No. I I youngest) measured in Horton Bluff Formation on the northern flank of the syncline at Horton Bluff (see Fig. I for location).
Fig.
-
-
Subaqueous clastic fissure eruptions
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L I T H O F A C I E S A ND L A C U S T R I N E E N V I R O N M E N T OF DEPOSITION Description
The section near Avonport Station (section 1 , Fig. 1) starts within the top portion of the middle lithologic member but the bulk of its stratigraphic thickness is in the lower half of the upper member (Fig. 3). Section 2 was measured on the southern fiank ofthe syncline where the succession is symmetrically repeated, including lower parts of the middle member. The Horton Bluff Formation consists of a succession of black and green shale sequences. Each sequence starts with dark grey to black shale which passes up into greenish-grey shale. Sandstones are intercalated in both the dark grey to black shales and the greenish-grey shales. The relative proportion of dark-grey and black shale to greenish-grey shale varies stratigraphically. The dark grey shales dominate the sequences in the upper half of the middle member while the greenish-grey shales become relatively more important in the upper member. Thirteen sequences were measured in the upper member at Horton Bluff and individual bands could be correlated between the two sections across the syncline for a distance of about 1 km.
4. Moulds of fossil tree stumps. Upper part of middle member, Horton Bluff Formation, section 2, southern flank of syncline at Horton Bluff.
Fig.
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Reinhard Hesse and Harold G. Reading
The dark grey shales contain abundant pyrite and bear bivalve, arthropod and fish fragments. Trace fossils made by limulid arthropods are also common. Infrequent mud-cracked horizons and one upright tree stump have been seen. The greenish-grey shales include bedded dolostone and algal stromatolites. They also contain evidence for palaeosols such as roots and in situ trees and dolomitized carbonate concretions resembling calcrete horizons. In the middle member of section 2 (south of the lighthouse) there are many moulds and casts of tree stumps varying in size from a few to 30 or 40 em in diameter (Fig. 4). In one bed, many trunks several decimetres or metres long were found more or less aligned (075-090°) presumably destroyed by a Mississippian windstorm. Bell ( 1929) counted fifty-six palaeosol horizons in the upper member in cliff exposures near Cambridge east of the Avon River mouth, but we now know that many of his root horizons are bioturbated layers (Belt, personal communication). Mud-cracks are abundant in the upper member. Fine-grained sandstones are intercalated throughout the succession though there is a tendency in sequences 7 - 1 1 for a concentration of sandstones to occur near the top of the dark grey shale and the base of the greenish-grey shale. The sandstone beds are normally less than 30 em thick and show ripple cross-lamination, lenticular or flaser bedding. In section 2, there is one coarse arkosic sandstone, 90 em thick, that has small scale trough cross-bedding at the top and low-angle cross-bedding at the base.
Interpretation
The dark grey shales were formed in a standing body of fresh water, their dark colour and abundance of pyrite reflecting reducing, anaerobic conditions at or close to the sediment/water interface. The waters were sufficiently shallow for wave action to ripple sand occasionally. Very rare emergence is indicated by mud-cracks and one tree stump. The greenish-grey shales were formed in a marginal environment, partly subaqueous but where emergence was frequent. Subaqueous deposition in a moderately oxygenated environment is indicated by the shales themselves, the bedded dolostone and the algal stromatolites. Emergence is shown by the palaeosol horizons. Sandstones are more frequent within some of the subaqueous greenish-grey shales and suggest wave reworking in the marginal environment. Although most sandstones appear to have been reworked, some of the sandstone layers, such as the 90 em thick cross-bedded layer mentioned above, were probably brought into a normally quiet environment as sheet sands during flood surges into the lake. Other layers, on the other hand, as we shall show, were ejected at the surface from below and their position in the sequence does not necessarily indicate a sand-carrying environment at the surface. The presence of non-marine fauna and absence of marine fossils (Carroll et a!., 1972) strongly support a lacustrine interpretation for the basin. The only known marine Carboniferous deposits of Nova Scotia occur in the stratigraphically younger Visean Windsor Group. Although a lagoonal environment cannot be completely ruled out, it is significant that Belt ( 1968a) showed that the deposits of the middle and upper members of the Horton Bluff Formation are part of a more extensive development of similar lacustrine facies in southern New Brunswick and Nova Scotia (Fig. 1). There were two scales of fluctuating lake levels, those shown by the sequences as a whole and those shown by the more abundant emergent horizons. Both probably
Subaqueous clastic fissure eruptions
247
Two dykes (No. 1 5 b and 1 5 c) of the 'main dyke bed' developing from the same vent which bifurcates about 1 5 em above the hammer. The vent does not cross dark, soft shaly layer (at hammer handle) which presumably is the source horizon for the dyke sediment. Original height of dyke in uncompacted state may have been 2·5 m.
Fig. 5.
Dyke 68 of 'main dyke bed'. lts upper horizontal portion (above hammer handle) may have once been in a more or less vertical position, but was rotated into a horizontal position during shale compaction.
Fig. 6.
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Reinhard Hesse and Harold G. Reading
reflect a combination of climatic fluctuations, effecting periodic variations in runoff, and isostatic subsidence of the platform due to its marginal position close to the rapidly subsiding sediment-accumulating rift valley. Fluctuations of ancient lake levels from the Triassic East Berlin Formation of the Newark Group, Connecticut (Hubert, Reed & Carey, 1 976) differ slightly from those of the Horton Bluff by the more symmetrical nature of the black to grey shale sequences they produced.
S ED I M E N T A R Y T R A N S P O S I T I O N S T R U C T U R E S
1. Dykes associated with subaqueous clastic fissure eruptions
Clastic dykes are common throughout the lacustrine members of the Horton Bluff Formation. Their distinction from sand-filled mud-cracks is not always easy (see below), particularly in the case of small dykes. Unquestionable dykes of intermediate size occur in the uppermost sequence of the middle member associated with a single sandstone bed, the 'main dyke bed' (Fig. 3), to which the top of the dykes are attached.
7. Sandstone horizon immediately above hammer shows oblique stratification dipping in opposite directions away from assumed central vent, which, however, was not exposed at the stage of cliff recession when the photo was taken. An alternative explanation that the structure may represent the upper portions of two dykes which were rotated in opposite directions during compaction is less likely in this case because both structures can be seen to terminate abruptly downwards against the dark-grey shale. The central part of the fissure eruption which presumably had originally a dome-like shape was levelled off during subsequent redistribution of the extruded sediment by wave or current action. Lenticularity of sand layers above the 'main dyke bed' appears to be caused by the root system of a large tree. Fig.
249
Subaqueous clastic fissure eruptions
More than a hundred dykes were counted in the cliff and on the tidal flat (Fig. 2) where they fill a swarm of parallel linear fissures striking uniformly at 1 50- 1 65 The dykes are wedge-shaped, measuring up to 50 em or more in width at the top, as little as 1 em at the base, and up to 75 em in height in the compacted state. The dykes usually show extreme ptygmatic folding (Figs 5 and 6), which proves that they were formed before the muddy host sediment underwent appreciable compaction. Thus the original height of the dykes may have been three times or more their present height, taking into consideration some flattening of the dykes after they had been transposed into a zig zag pattern during compaction of the surrounding muds. The fine to silty sand of the dykes contains small plant fragments. These dykes are due to upward movement of sand leading to extrusion at the lake floor. That sand movement was in fact upward and not downward can be seen in cases where evidence of surface extrusion and overspill is preserved such as on Figs 7 and 8. Oblique stratification, dipping in opposite directions away from the assumed central vent, is interpreted as a fissure eruption structure (Figs 7, 8 and 1 2). It can easily be confused with cross bedding, especially where exposure is imperfect. Polished slabs of the wedge-shaped tops of some dykes also provide evidence for extrusion rather than for downward fissure-filling (Figs 9 and 1 0). •.
Fig.
8. Schematic diagram illustrating inferred extrusion mechanism for the structure in Fig. 7.
The top parts of some dykes display either horizontal lamination or oblique lamination. Some of these probably developed as parallel vertical laminations during laminar extrusion of the dyke sediment and were later rotated into a horizontal position when the dyke was strongly folded during compaction of the surrounding shale (Fig. 1 1 ). Some of the oblique laminations represent incompletely rotated vertical laminations (Fig. 1 1 , left side) and the distinction between these and those due to surface extrusion and outpouring of the sediment is not always clear. The oblique laminations of some surface extrusion units contain small cross-laminae which formed during the outpouring of the material from the fissure (Fig. 1 2). Where such oblique laminations are overlain by horizontal laminae, the latter clearly formed after the extrusion as seen by s::nall irregularities in the surface of the obliquely bedded slab which were filled in by the horizontally laminated sediment (Fig. 1 2). Where the vent
250
Reinhard Hesse and Harold G. Reading (9)
(10)
Fig. 9. Top end of dyke showing polyphase extrusion ofsoft sediment. The lighter-coloured innerwedge and
its folded downward thread represent a somewhat later phase because they transect some laminae in the otherwise homogeneous dark-grey silty fine sand of the earlier extrusion. Downward continuation of dyke broken off at left side of specimen (as indicated by inset; arrows: direction of sediment intrusion.). Top of ' middle member in Section 2.
10. Top of dyke showing floating fragment of laminated sediment. It probably represents material that was extruded from the same dyke at an early stage and later sank into the homogeneous muddy sand of the main dyke fill in a manner similar to pseudonodules. Downward continuation of dyke broken off at lower left side of specimen. Same locality as Fig. 9.
Fig.
has been rotated into a horizontal position, it may be impossible to draw the boundary against the horizontally laminated sediment that was deposited later, as for example in Fig. 1 1 . On the other hand, sediment extrusion at the soft muddy lake bottom has led to loading and differential sinking of the extruded sediment (Fig. 13) adding further complications to the interpretation of these structures. In all observed cases, however, the degree of loading was moderate suggesting that the strength of the muddy surface sediment was generally sufficient to keep the extruded sand essentially at the surface. No pseudo-nodules were observed. The original fine sandy to silty layer from which the dykes originated seems to have disappeared. It was probably located at a depth of about 2·5 m below the surface of the soft, little-compacted sediment. A recession in the weathering profile due to a soft shale layer now located about 80-90 em beneath the dyke bed may mark the former position of the vanished sand-silt layer (Fig. 5). No dyke was seen crossing this layer. The layer is undisturbed except for some downbending seen under the more massive dykes which is caused by differential compaction (Fig. 5). In the neighbourhood ofthe 'main dyke bed' only one example of a small dyke was seen beneath this soft shaly layer, but it did not cross it. 2. Distinction between mud-cracks and dykes
In the Horton Bluff Formation both mud-cracks filled from above and dykes filled from below occur in the section and may be closely associated. In Fig. 1 4 the coarse sand in the fissures can be related to a sand layer on top and, therefore, had to come in from above, filling mud-cracks. In the same specimen, a broad fissure is seen filled with laminated silty shale clearly also coming from above. In most cases, however, the
Subaqueous clastic fissure eruptions
25 1
I I. Dyke 23 of'main dyke bed' isoclinally folded during compaction of surrounding shale. The parallel laminated dyke which originally was more or less vertical shows four or five deflection points in the upper part around which it was bent into isoclinal, horizontal folds. Oblique stratification left of hammer is probably upper termination of a less intensely folded satellite dyke rather than being a fissure eruption structure, because it appears to be connected to the small dyke at the left. The horizontally laminated sediment above represents material redistributed by surface processes.
Fig.
5 4 3 2 1
0
12. Oblique stratification similar to Figs 7 and 8, probably due to fissure eruption from vent immediately to the left of specimen. Presence of very small ripple cross-laminae in the light-coloured oblique sand laminae indicates outpouring of sediment from a cone- or ridge-shaped elevation at the site of the vent that was levelled off before deposition of the overlying horizontal laminae of dark-grey sandy siltstone and lighter fine sandstone which fill in depressions in the irregular surface of the obliquely laminated material. U nder the stereomicroscope the first thin lamina above the obliquely laminated material displays viscous flow structures and was probably injected at the boundary.
Fig.
source of the coarse sand in the fissure seems to have been at the base, indicating filling from below. Obviously, every case needs specific consideration. Dykes may certainly display polygonal patterns as modern examples show (Reimnitz & Marshall, 1 965), and we see no simple criteria for distinguishing the two.
252
Reinhard Hesse and Harold G. Reading
30
10
Fig. 13. Loading and differential sinking of extruded sediment into soft, muddy lake bottom sediment (sketch of dyke 6a).
Fig. 14. Polished slab containing folded mud-cracks filled with coarse sand identical to sand-layer above.
Subaqueous clastic fissure eruptions
253
3. Mechanism of dyke formation and fissure eruption
The trigger mechanism envisaged for the dyke formation and the associated fissure eruptions is the surface shock wave of an earthquake. Fissuring and subsequent infilling with water and sand from below as a secondary effect of earthquakes is well documented in the literature (e.g. Oldham & Mallet, 1 872; Dutton, 1 889; Reimnitz & Marshall, 1965). Generally it is assumed that some lateral movement of the surface deposits is caused by the earthquake that gave rise to the opening of fissures. Such movement could be due to slope instability along river banks (Oldham & Mallet, 1 872) or tidal channels (Reimnitz & Marshall, 1 965). We have no evidence that the fissures are related to an original slope of the lake bottom. The dykes show no persistent dip direction. Some dip toward the west, others toward the east. Sediment movement may have been caused by some excess pore pressure which developed in the sand layer when dewatering (and/or degassing) was iilhibited by the overlying mud during the start of compaction. The abundance of plant material in the silty sand layer indicates that decomposition of organic matter is a likely source for the production of carbon dioxide and probably methane during burial. Triggering of the upward movement of this overpressured silty sand was then effected by the seismic shock, which may have caused liquefaction of the sediment and at the same time, while the surface wave travelled across the lake bottom, opening of the fissures for a short moment. Ejection probably continued for a while after the seismic shock. Sediment extrusion took place in a laminar fashion where laminations parallel to the dyke walls are still preserved, but may have been more violent in the case of homogenized dyke material. Eye witnesses of the Alaska earthquake of 1 964 reported spouts of mud and water on the tidal flats being ejected into the air as high as 1 0- 1 3 m (Reimnitz & Marshall, 1 965). Other structures similar to Seilacher's ( 1 969) microseismites have been observed by us in the Horton Bluff section. As more than l OO dykes ofthe 'main dyke bed' occur at the same stratigraphic level, movement must have been triggered more or less simultaneously. Moreover, the regular horizontal spacing of the linear dykes of between 0·75 and 5 m suggests that the pressure release from the buried sand layer proceeded systematically across the lake bottom, probably in a direction perpendicu lar to the orientation of the dykes, which may have been the direction of propagation of the seismic wave, i.e. from WSW to ENE or vice versa. The upward widening of the dykes probably reflects the upward decrease of the degree of compaction of the host sediment at the time of ejection. It appears that part of the 'main dyke bed' is due to redistribution of the surficially extruded sediment by weak currents. Seismic activity can be related to Mississippian rifting in Nova Scotia.
4. Circular and elliptical collapse structures
Circular and elliptical structures, measuring between 3 and 1 5 m in diameter (Figs 1 5- 1 7), occur on a somewhat larger scale than the clastic dykes. For description of the structures see figure caption. We interpret them as collapse structures associated with sediment extrusion of liquefied (and possibly gas-rich) sediment, which escaped to the surface along the fissure vent, so creating a bowl-shaped collapse structure. The extruded material refilled the bowl by flowing back into it as indicated by foresets in the sandy silt filling the elongate ear-shaped depression (Fig. 1 7). These foresets dip toward the centre of the structure. Existence of the bowl as a surface depression is
254
Reinhard Hesse and Harold G. Reading
shown by broken ends of originally horizontal beds being bent down sharply towards the centre at the edge of the structure (Fig. 1 6).
Fig. 15. Circular structure, 3 m diameter: what appears t o b e a convex upward ring accompanied b y a n outer ring-shaped depression is the draping of the broken edge of a bed sharply bent downward towards the centre of the bowl-shaped structure (compare Figs 16 and 1 7). Hammer for scale.
16. Elliptical structure, 1 0 m diameter with marginal vent (right hand side of photo). The 'ear' of the structure in the foreground represents a linear extension of the collapse structure paralleling the vent through which the sediment was extruded. The extruded sediment partly flowed back into the adj acent depression as indicated by the oblique sand laminae (in the 'ear') which dip toward the structure.
Fig.
255
Subaqueous clastic fissure eruptions
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17. Diagramatic sketch of elliptical structure (Fig. 1 6) in plan view and cross section. Black arrows indicate flow direction of sediment extruded from the vent. Dotted bands: sand layers, ( l ) deposited before the disruptive event, bent upward near the vent; (2) deposited before the disruptive event, bent downward at the opposite margin of the structure, compare Fig. 1 6; (3) sediment that flowed back from the vent into the adjacent depression formed during the disruptive event; (4) deposited during and after the disruptive event and draping the structure; open arrows indicate transport direction by currents or waves which is about opposite to the flow direction of the material that emanated from the vent.
Fig.
Subsequent filling of the structure not only took place from the side of the vent, but was aided by wave- or current-transported material from the opposite direction as indicated by the orientation of ripple crests and foreset laminae (Fig. 1 7). The circular structure (Fig. 1 5) was completely refilled. What appears to be a convex ring is a draping by later sediment of the edge of the bed which stands more or less vertically at the margin of the structure. Genetically these structures appear to be closely related to the earthquake-triggered clastic dykes somewhat higher in the sequence, and possibly also record an earthquake. The formation of similar structures during recent earthquakes is well
256
Reinhard Hesse and Harold G. Reading
documented (Dutton, 1 889; Reimnitz & Marshall, 1 965) although deformation structures have been attributed to less catastrophic events such as the migration of dune bed forms on tidal flats (e.g. Parker, 1 973).
S I G N I F I C A N C E O F S ED I M E N T A RY T R A N S P O S I T I O N STRUCTU R E S
A prerequisite for the formation of these structures is inhibited dewatering and degassing of the source sediment. Thus clastic dykes or diapirs may form from deeply buried layers of uncemented and probably overpressured sand, silt and clay whose cover rocks may be well compacted and cemented at the time of transposition. If, as we believe, the clastic dykes were triggered by earthquake shocks then they are not only indicators of palaeoseismicity but also can be used to detect the direction of propagation of the seismic wave. This would be perpendicular to the trend of the dyke swarm. The conditions of sedimentation in many lakes are ideal for the development of transposition structures: alternations of mud and sand; a high organic, especially plant fragment, content; and the entrapment of gas and water under pressure. Modern examples from Californian reservoirs were described by Sims ( 1 975). They are, however, certainly not restricted to this environment and are common in many other environments (e.g. Johnson, 1 977). Unless outcrop conditions are very good, it is extremely difficult to distinguish between sedimentary dykes intruded from below and cracks filled by sand from above. It is also easy to confuse cross-bedding and cross-lamination resulting from surface currents with that due to sand extrusion. Many upward-widening dykes look like minor channels (e.g. Fig. 1 1 ) Collapse structures may be mistaken for erosional scoops or basal slide phenomena. In environmental reconstruction it is also important to remember that sand may have come by extrusion from below and does not necessarily reflect emplacement by currents from an exposed source. .
A C K N O W L ED G M E N T S
Sedimentologic study of the Horton BluffFormation was initiated in the summer of 1972 during a field trip to Nova Scotia led by Paul Schenk of Dalhousie University, Halifax, while one of us (H.G.R.) held a Nuffield Visiting Lectureship to Canada. It was continued that season by H.G.R. and during the following two field seasons by R.H. Financial support for this study came from National Research Council of Canada grant A7368 to R.H. We are grateful to E. S. Belt, Amherst, J. D. Sims, Menlo Park and an anonymous reviewer for a critical review of the paper and to R. L. Carroll, Montreal for comments.
REFERENCES BELL, W.A. ( 1 92 1 ) The Mississippian formations of the Horton-Windsor district, Nova Scotia. Am. J. Sci. (ser. 5) I, 1 53-173.
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BELL, W.A. ( 1 929) Horton-Windsor district, Nova Scotia. Mem. geol. Surv. Brch. Can. 155, 1-268. BELL, W. A. ( 1 960) Mississippian Horton Group of type Windsor-Horton district, Nova Scotia. Mem. geol. Surv. Brch. Can. 314, 1 - 1 1 2. BELT, E$. ( 1 968a) Carboniferous continental sedimentation, Atlantic Provinces, Canada. In: Symposium on Continental Sedimentation, Northeastern North America (Ed. by G. de V. Klein). Spec. Pap. geol. Soc. Am. 106, 1 26- 176. B ELT, E.S. ( 1968b) Post-Acadian rifts and related facies, eastern Canada. In: Studies ofAppalachian Geology, Northern and Maritime (Ed. by E-An Zen et a/.), pp. 95-1 1 3 . Wiley & Sons, New York. CARROLL, R.L., BELT, E.S., DINELEY, D.L., BAIRD, D. & MCGREGOR, D.C. ( 1 972) Vertebrate paleontology of eastern Canada. Guidebook, Exc. A 59, 24th int. Geo/. Congr. Montreal. CROSBY, D.G. ( 1 962) Wolfville map-area, Nova Scotia (2 1 H I ). Geol. Surv. Canada Mem. 325, 1-89. DUTTON, C. E. ( 1 889) The Charleston earthquake of August 3 1 , 1 886. U.S. Geol. Surv. A nn. Rept. 9, 209-528. ELLIOT, R.E. ( 1 965) A classification of subaqueous sedimentary structures based on rheological and kinematical parameters. Sedimentology, 5, 1 93-209. HESSE, R. ( 1 976) Einige ungewohnliche sekundiire Sedimentstrukturen im lakustrinen Unterkarbon Neuschottlands (Kanada) und ihre Deutung als Erdbebenanzeiger. Ec/ogae geol. Helv. 69, 196-20 1 . HUBERT, J.F., REED, A.A. & CAREY, P.J. ( 1 976) Paleogeography of the East Berlin Formation, Newark G roup, Connecticut Valley. Am. J. Sci. 276, 1 1 83- 1 207. JOHNSON, H.D. ( 1 977) Sedimentation and water escape structures in some Late Precambrian shallow marine sandstones from Finnmark, North Norway. Sedimentology, 24, 389-4 1 1 . OLDHAM, E . & MALLET, R . ( 1 872) Notice o n some of the secondary effects of the earthquake of lOth January, 1 869, in Cachar. Q. Jl geo/. Soc. Lond. 28, 255-270. PARKER, W .R. ( 1 973) Folding in intertidal sediments on the west Lancashire coast, England. Sedimentology, 20, 6 1 5-623. PLAYFORD, G . ( 1 963) Miospores from the Mississippian Horton G roup, eastern Canada. Geo/. Surv. Canada Bull. 107, 1-47. REIMNITZ, E. & MARSHALL, N.F. ( 1 965) Effects of the Alaska earthquake and tsunami on Recent deltaic sediments. J. Geophys. Res. 70, 2363-2376. SEILACHER, A. ( 1 969) Fault-graded beds interpreted as seismites. Sedimentology, 13, 1 5 5 - 1 59. SIMS, J.D. ( 1 975) Determining earthquake recurrence intervals from deformational structures in young lacustrine sediments. Tectonoph. 29, 1 4 1 - 1 52.
Spec. Publs int. Ass. Sediment. ( 1 978)
2,
259-278
A Proterozoic lacustrine interlude from the Zambian Copperbelt
HARRY C L EMMEY Department of Earth Sciences, The UniversiTy of Leeds, Leeds LS2 9JT, U. K.
AB STRACT O n the Zambian Copperbelt the Ore body Member of the Kit we Formation is represented by some
30 m of dolomitic and siliciclastic rocks. These meta-sediments illustrate the
problems of differentiating between pre-Phanerozoic sediments of lacustrine. lagoonal or epeiric sea origin. Facies analysis of the sequence indicates a four stage history. First. a marine transgression took place over a dominantly fluvial landscape. The basal sediments of this transgression contain a trace fossil fauna indicative of an advanced metazoan benthos. Secondly. isolation of the sea led to the establishment of lacustrine conditions. Incipient sabkha diagenesis occurred in the lake edge sediments and is indicated by anhydrite crystal lath mush. gypsum crystal mats and nebulitic anhydrite. Thirdly. a return to marine conditions occurred and finally a major regression. documented by a coarsening upward sequence. indicates the advance of alluvial sediments into the area. Siliciclastic sabkhas were developed during this final period. Strontium values lend some support to the facies interpretation. the highest values occurring in the marine stages. Five major transgressive periods impart a recognizable cyclicity to the sediments. The stratigraphic applications of this cyclicity are of potential economic importance for the major orebodies of the Copperbelt are located within the unit.
INTRODUCTION
Only l% of all described lacustrine deposits are of Precambrian age, although the existence oflarge supercontinents throughout much of the Proterozoic, as suggested by palaeomagnetic work, indicates a substantial lacustrine contribution to Precambrian sedimentation. The reasons for the dearth of recognized lake sediments are complex and include such factors as extensive metamorphism and the poor preservation
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
260
Harry C/emmey
potential of ancient rocks. Indeed the problems inherent in the recognition of any ancient lacustrine deposit (Picard & High, 1 972) are heightened in pre-Phanerozoic rocks by the p aucity of p alaeontological evidence and by a few hypothetical and non uniformitarian concepts regarding the chemistry of water bodies and the strength of tidal forces. If Cloud ( 1 968) and other workers are correct in postulating that tidal forces were exceptionally strong during the early periods of earth history, then significant tides may have affected large lakes to such an extent that facies analysis of sediments produced in such an environment might suggest a marine tidally dominated deposit. Conversely, if normal tidal forces prevailed, then the distal reaches of epeiric seas would be tideless on account of friction causing decay of the tidal waves. The marine facies produced would then resemble those produced in lakes. Further problems are presented by abandoned arms of epeiric seas. Seas of this type may have been common-place in the Precambrian as epeiric transgressions penetrated into the continents along intercratonic, tectonically active weaknesses only to be isolated by epeirogenic uplift. Abandonment would not be sudden and total and occasional marine connections would be maintained for some time. At what point does a lagoon formed in this way and fed with marine waters by exceptional wind tides [as in the case
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Fig.
26 1
Proterozoic lacustrine interlude
of the B aj a Lagoons described by Phleger ( 1 969) and Kinsman ( 1 969)] become recognizable as a lake? In spite of these problems, attempts should be made to differentiate between the marine and lacustrine contributions to Precambrian sedimentation, otherwise a major contribution to Precambrian lithostratigraphy, tectonics and crustal evolution will be lost. This paper represents one such attempt on a group of ore-bearing meta-sediments from the Zambian Copperbelt (Fig. 1 ) . Comparisons are made with a Recent ephemeral lake in Zambia.
Location
The Zambian Copperbelt formed within a trough oflate Proterozoic sedimentation between two major shield areas in the centre of Gondwana (Fig. 1 ) The late Proterozoic apparent polar wander track (Piper, Briden & Lomax, 1 973) indicates that throughout the period of sedimentation discussed in the paper, this central area lay at low latitudes (0-30.). .
Stratigraphy and general geology
The sediments described belong to the Roan Group of the Katangan (Cahen, 1974). The three formations recognized in the Roan by C1emmey ( 1 976a) are shown in Table 1. G ROUP
K u
N D E
l u
N G
u
FO R M AT I O N
A G E m y. 620
U pp e r 6 75
M id d l e
l uf i l ian Orog e n y
800
lower G r a nd � g l o m e r at e
950
M wa s h i a
R 0
B a n c r of t Do l o m i t e K itwe
A N
BAS E M E N T
•
. . . . . . ....c:
Mindola Clastics COMPLEX
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Table I. Stratigraphy oft he Precambrian rocks in the Zambian Copperbelt district. The Ore body Member, discussed in this paper, is the lowest member of the Kitwe Formation (position shown by an arrow at 1050 My ).
262
Harry Clemmey
The lowermost Mindola Clastics Formation was deposited over the landscape of an ancient landmass, herein named the Kafue Landmass. Alluvial wet and dry fan, aeolian and lacustrine facies are recognized and the period ended with the deposition of a coarsening and maturing upward cycle (the Kafue Arenites Member) which is recognizable all round the flanks of the landmass and thought to be chronostratigra phic. The overlying Kitwe Formation is dominated by fine-grained siliciclastics, dolomites and anhydrite. Several coarse clastic fluvial incursions are recognized but the main problem concerns the marine or lacustrine nature of the dolomite-evaporite-
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( b) 2. Desiccation cracks from the Orebody Member. (a) Median-ridge cracks, Unit 3; (b) diagram illustrating the formation of median-ridged cracks, not to scale.
Fig.
Proterozoic lacustrine interlude
263
silt rocks. The problems of facies analysis on these rocks are here illustrated by the basal Orebody Member of the Kitwe Formation. The Orebody Member is of great geological importance for it contains the oldest known advanced animal traces (Clemmey, 1 976a and b) and the oldest known unreplaced evaporites, approximately 1 000 My old.
F A C I E S A N A LY S I S O F T H E O R E B ODY M E M B E R Structures penetrating bedding planes
Desiccation cracks
Several types of desiccation cracks are found and include deep polygonal crack patterns infilled by overlying sediment and similar cracks in which the polygons are colonized by algal mat dolomite resulting in a cuspate edge and a convex upwards cross section. These polygons resemble examples described from Ordovician tidal fiats by Davis ( 1 975, p. 1 64). Both types of crack may form in lacustrine or marine environments and indicate periodic and relatively lengthy subaerial exposure. A further type of crack is unusual (Fig. 2a) and occurs in both polygonal networks and incomplete patterns only 2-5 mm deep and having a median ridge formed from the underlying laminae (Fig. 2b). Formation is envisaged as follows. First, sedimentation in a lacustrine or lagoonal environment results in a thin bottom lamina in which syneresis causes linear cracks to develop. Subsequently a pore water pressure increase in the underlying sediment forces their cohesive skin, as a sharp anticline, into the developing crack. Perhaps in 'median-ridged desiccation cracks' we have a new lacustrine lagoonal indicator? Syneresis cracks
Randomly oriented incomplete cracks, often having a characteristic bird's foot shape and formed during subaqueous dewatering (possibly aided by salinity changes;
Fig. 3. Syneresis cracks from the Orebody Member; characteristic bird's foot shape, Unit 2, scale bar in millimetres. The crack has the dolomite/anhydrite fill removed by intense tropical weathering.
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Harry Clemmey
Donovan & Foster, 1 972) are termed syneresis cracks. Such structures are well developed in the Orebody Member (Fig. 3). Donovan & Foster indicate the ephemeral nature of the cracks which, if unfilled, close up. The cracks of the Orebody Member often have an infilling of dolomite and anhydrite, unlike the sediment-filled examples described by Donovan & Foster, indicating quiet bottom conditions and long breaks in sedimentation. Whilst they are not unique to lake sediments, syneresis cracks are particularly abundant in ancient lacustrine formations (Devonian-Orcadian Basin, Triassic Lockatong Formation, Eocene-Green River Formation). Synsedimentary faults
Synsedimentary faults are preserved in three dimensions on some bedding planes (Fig. 4) and display exquisite preservation of the scarp thus providing additional evidence for calm depositional conditions.
Fig.
4. Synsedimentary faults. Two faults from bedding plane in Unit 3, scale in millimetres.
Textures and features formed on bedding surfaces
Gas bubble bursts
A further indicator of calm conditions is the preservation in relief of burst gas bubbles, which are differentiated from the burst tops of syneresis cracks by the lack of crack penetration below the burst and the greater number of radial cracks (Fig. 5). Rainfall texturing
Many clues to the depositional environment are obtained from rainfall textures on bedding plane surfaces. Pits with distinctive outlines characterize recently exposed but cohesive sediment (Fig. 6a). This situation pertains in lacustrine and lagoonal environments where exposure is preceded by a period of subaqueous dewatering. Preservation of the pits depends on the shower being of short duration, and on their burial before any subsequent rainfall. Some examples have been found in the Orebody
Proterozoic lacustrine interlude
265
Member (Fig. 6b). If the surface should be subject to further rainfall an embossed surface results (Fig. 6c) and again examples are found in the Orebody Member (Fig. 6d).
Fig. 5. Gas bubble burst, Unit 2. The great number of radial cracks in the top and the lack of penetration distinguish these structures from syneresis cracks, scale bar in millimetres.
Fig. 6. Rainfall texturing: (a) Recent example of 'one shower' pitting; (b) 'one shower' pitting preserved in the Ore body Member; (c) Recent example of texturing produced by prolonged rainfall or multiple showers ( 'embossing'); and (d) example of rainfall embossing from the Orebody Member, Unit 3 .
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Harry Clemmey
Palaeostrand-line positions
Lacustrine strand-line positions may have a fairly high preservation potential. As mentioned earlier, rain pitting affects only wet sediment, thus if a body of water is retreating slowly successive strand line positions may be recorded by parallel zones of differing rain pit intensity. The strands may be enhanced by small wave-cut notches, by the accumulation of debris and the folding of the littoral sediment. Such features are found in the Orebody Member (Fig. 7a) where the strandline is marked by rolled algal bun debris. Away from the strands the algal buns, thought by Clemmey ( 1 976a) to be analogous to the slimy gelatinous stromatolites found in Persian Gulf lagoons, are circular in plan and flattened in the plane ofbedding indicating in situ preservation. Wind-formed features
A variety of distinctive textures owe their origin, directly or indirectly, to wind blowing over exposed surfaces. Micro-ripples may form in non-cohesive sediment while micro-folds may develop in a more cohesive substrate. These are the runzelmarken of Reineck & Singh ( 1 973). Preservation potential is good and ancient examples are known from the Silurian (Hunter, 1 969) and the Precambrian (Klein, 1 975). Examples are found in the Orebody Member (Fig. 7c). Also figured by Reineck & Singh are adhesion warts of wind blown origin. Preservation potential is low but well developed examples are found in the Orebody Member (Fig. 7b). Exposure, strong desiccation (to provide some lithification) and quiet burial are all indicated. Synsedimentary fold belts
In the strand-line area of some Recent lakes, small-scale fold belts have been observed (Fig. 7d) when a falling lake-level results in exposure of the cohesive bottom layer of lake sediment to the energy of the littoral zone. It is thought that incipient folds are first formed by wind drag over the exposed bottom sediment. Then the advancing wind driven tide gains access beneath the top lamina 'floating' it on a cushion ofwater and carrying it forward. The pre-formed micro-folds are accentuated and a fold belt forms parallel to the strand. Several phases of deformation may be present (Fig. 7f) and the ultimate feature is a series of folded clay flakes. The preservation of synsedimentary fold belts in the Orebody Member (Fig. 7e) indicates a calm, possibly lacustrine environment. Summary
The features described above indicate that during Orebody Member sedimentation subaerial exposure was common, the climate was probably hot and semi-arid, the water body was at times quasi-static and possibly subject to salinity changes and a very calm depositional environment dominated. The occurrence of the features described, in close association, is suggestive of a lacustrine environment.
Sedimentary facies
There are two dominant facies assemblages developed in the Orebody Member.
Proterozoic lacustrine interlude
267
Fig. 7. Wind formed features. (a). Palaeostrand-line from Unit 3. Note the differential rainpit intensity, especially strong in the synsedimentary fold troughs (A), a phenomenon commonly found in Recent lake sediments as the moisture is retained longer in fold and ripple troughs and hence the sediment is receptive to rainfall texturing for a longer period. Note also the typical strand-line folding and trash (B), the latter consisting of rolled gelatinous stromatolite buns. The isolate flake (C) is a fragment of Recent rain-pitted lacustrine sediment for comparison. (b) Adhesion warts from Unit 3. Note the elongation from upper left to lower right (a palaeowind azimuth?). (c) Fossil runzelmarken from the Orebody Member. Synsedimentary folding. (d) Fold belt marking the strand-line area of a Recent lake. (e) A synsedimentary fold belt from the Ore body Member. Often these belts are found superimposed with differing orientations in individual belts, thus indicating their synsedimentary nature (f) Recent example of folded layer indicating polyphase folding.
268
Harry Clemmey
Dolomite-increasing-upward cycles
This term is given to cycles within siliciclastic silt/muds in which dolomite laminae increase in importance upwards as the rock passes into dolomite. In the ideal cycle there is a development of dolomite/anhydrite rock or pure anhydrite at the top (sabkha on Fig. 8). This latter feature, together with the evidence of desiccation indicates that the cycles reflect progressive shallowing. The dolomite layers are thought to be of algal origin (Clemmey, l976a) and analogous to the algal 'peat' of similar cycles described by Davies ( 1 970) from Shark Bay.
8. Dolomite-increasing-upward cycle from Mindola Mine. Two small cycles are in evidence: dolomite increases upward from the subaqueous silts of Unit 4 to form an incipient sabkha (indicated by dolomite/anhydrite, layer A); a return to low intertidal conditions is indicated by the succeeding dolomite-silt layers and finally the return to sabkha conditions by the dolomite bed with anhydrite nodules (arrowed). The cycle is truncated by a typical Type D bed 'storm' layer (B). Note the synsedimentary faults developed around the anhydrite nodules.
Fig.
Silt/fine sand facies assemblage
Many problems are inherent in the interpretation of fine-grained metamorphosed siliciclastic rocks. Clean fine-grained silts take on a cherty aspect and recrystallization
Proterozoic lacustrine interlude
269
often destroys any structures that may have been present. Subtle grain-size variations, however, can be preserved when the tectonic grade is low, and cross-lamination, even on the micro-scale, can be recognized when accentuated by heavy minerals. Five types of stratification are recognized in the Orebody Member. Type A laminates are characterized by normal grading on a scale of 2- 1 0 mm and their formation may be ascribed to micro-density flows. Such flows are common in lacustrine environments (e.g. the Walensee; Lambert, Kelts & Marshall, 1 976). Type B laminates lack the grading of type A, are characterized by even lamination and may have been formed by bed load transport, by migration of miniature bed forms or by fallout from suspension. Type C laminates consist of alternating mud and silt. The fine-grained portion is often wispy laminated on a scale of several layers to 1 mm and the degree of inheritance between anticlines and synclines is low. This type of lamination could form by selective algal binding in siliciclastic sediment and may be the only evidence for the existence of algal mats in such rocks. Gunatilaka ( 1 975) and Schwartz, Einsele & Herm ( 1 975) described non-carbonate producing algae selectively binding the fine grained sediment fraction in a siliciclastic environment. Type D beds are 2- 1 0 em thick, apparently structureless beds composed of silt or fine sand. The beds often occur at the base of transgressive sequences and were clearly erosive of the underlying algal mat material. Type D beds are also found within and at the top of dolomite-increasing-upward cycles and some bands are mappable for great distances. One such unit at Mindola is only 2-4 em thick and may be traced for several kilometres. The location of the beds within such facies strongly suggests an origin as storm-generated layers. Their location within algal mat prograding units (Fig. 8) is identical to the occurrence of Recent storm layers in the algal mat prograding sequence at Shark B ay cited by Davies ( 1 970, fig. 1 0) . Type E beds consist o f flaser bedded, rippled, slumped, channelled, wavy and lenticular bedded assemblages (Fig. 1 1 ). This association is indicative of deposition in a tidal and/or wave regime (Klein, 1 975; Reineck & Singh, 1 973) especially in view of the lag time indicated by clay fall out on ripple foresets (Allen, 1974). Whilst type A, B and C laminates and type D beds have an almost ubiquitous distribution within the siliciclastic units, type E beds are confined in development to the top and bottom of the Orebody Member. Depositional units of the Orebody Member
The Orebody Member sediments may be divided into seven distinct units which succeed each other in vertical order. Figure 9 shows the units and their nomenclature as used on the various mines of the Rokana Group. Each unit has a distinct lithological and facies content and Clemmey ( 1 976a) has shown them to be recognizable over most of the Copperbelt and up to 400 km distant in Zaire (Clemmey, 1 975). This lateral continuity of thin (2-4 m) depositional units over large areas suggests either a lake or an epeiric sea with weak tidal currents as the depositional environment. Depositional Unit
1
(Fig. l Oa)
Sedimentation of the Kitwe Formation commenced with a succession of dolomitic laminae of possible algal origin and fine grained siliciclastic sediments. A quiet transgression over a mature fan topography is indicated and reworking of the
Harry Clemmey
270
: -._:_....-- . - �. -.-......:..- · . · _-....:__:- · - -- � - - -� -- �- - - .
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No. 1
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UNIT4
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UN I T2
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S c h i st o s e Footwa l l
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UN I T l
Cong l o m e rate
9 . Log of the Ore body Member showing the parochial nomenclature a s used a t Rokana Mines and the Depositional Units recognized here.
Fig.
Proterozoic lacustrine interlude
27 1
underlying Mindola Clastics Formation is confined to minor winnowing of the matrix in the conglomerate top. Tiny sand fiasers, ripple bedding and argillite wisps are the only indicators of current activity. On bedding plane surfaces type A desiccation cracks, gas bubble bursts and dolomite crystal pseudomorphs are found. Internally the individual mats are broken and wavy bedding and small channels are present. Evidence of infauna is provided by backfilled burrows (Type A of Clernrney, 1 976b) and by intense swirl bioturbation. Two possible early sponge-like organisms are associated with this unit. Unit 1 is interpreted as a low intertidal facies. The lack of any high energy indicators is attributable to the supposed epeiric nature of the sea, and to the protected nature of the area (Clernrney, 1 976a). A marine interpretation is preferred due to the occurrence of the faunal elements and the marked change in sediment type from the Mindola Clastics. It is conceded, however, that none of these factors is unequivocal and we may be dealing with a lacustrine transgression. Depositional Unit 2
Unit 2 often overlies Unit 1 with an erosional contact. The unit is characterized by a succession of massive (type D beds) passing to laminated silts (types A and B laminae) in sub-units 5-20 ern thick. In some parts the laminated silts and argillites are dominant. No cross-lamination of any scale has been noted and individual laminae appear to be continuous. Further indications of a low depositional energy are provided by the bedding plane features, notably the preservation in relief of gas bubble bursts, synsedimentary fault scarps, and the lack of dispersal of collapsed, pellet-formed, burrow walls at the sediment/water interface (Clernrney, 1 976a). The burrows also provide evidence for very slow sedimentation rates as only rare millimetre-scale adjustments were required to keep the burrows at the sediment surface. Towards the top of the unit syneresis cracks begin to develop, possibly indicating a change in salinity of the water mass and type B burrows become smaller and more dendritic. Perhaps this latter fact reflects a fossil community under stress as is so common in lacustrine environments (Picard & High, 1 972). No evidence for desiccation has been observed on the bedding planes. The above evidence is interpreted to imply continuing transgression over the tidal fiats of Unit 1 , thus rendering the area subaqueous. The low energy depositional environment, incoming of syneresis cracks and the stunting of the fauna (implied by the change in burrow habit) could indicate isolation of the sea and the establishment of lacustrine conditions. Depositional Unit 3 (Fig. l Ob)
Interbedded silts and 1 -3 ern dolomites characterize this unit. The dolomites do not occur as increasing-upward cycles and they may be attributable to physico-chemical precipitation. Desiccation cracks of the median-ridged type, rain textures and strand marks on bedding planes, together with gypsum crystal mats, anhydrite lath mush, runzel rnarken and synsedimentary fold belts all indicate extensive subaerial exposure. Syneresis cracks and the existence of fossilized gelatinous stromatolites indicate fairly stable subaqueous environments which, as indicated by the preservation of synsedimentary fault scarps and the relief of gas bubble bursts, must have been calm.
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10. Representative logs of the depositional units referred to in the text. (a), Unit I; (b), Unit 3; (c), Unit 5; (d), Unit 6; (e), Unit 7; numbers at the base of the logs refer to specific levels and locations at Mindola Mine. The log guide is applicable to all other logs in this paper.
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Proterozoic lacustrine interlude
273
Two environments could provide the alternating but fairly lengthy periods ofwater level fluctuation, a seasonal lagoon or shallow lake edge. Depositional Unit 4
Apart from a lower interstitial carbonate content, Unit 4 is identical to Unit 2 in lithology. Ripple-drift lamination, however, is evident and indicates a marginal increase in depositional energy over Unit 2. The concentration of syneresis cracks relative to Unit 3 shows an initial decrease but then increases markedly toward the top. Gas-bubble bursts reach their maximum development in Unit 4 and are often spectacular. The lack of evidence for subaerial exposure, the syneresis cracking and gas-bubble bursts all indicate a subaqueous environment. On the lower contact of Unit 4 are rare isolated channels, up to 60 em deep and 2 m wide, cut into Unit 3 . Whilst the transgression of Unit 4 rocks may have been mildly erosive, it is surprising to find channels of this magnitude especially in the absence of any lag conglomerate in the channels or elsewhere at the base of Unit 4. A tidal origin for the channels is rej ected and it is suggested they may have been eroded by streams draining into a lake. Observations on a recent lake in a semi-arid area indicate that in years when the rainfall remains abnormally low during the rainy season and the lake level fails to achieve that of the previous year then the lake edge sediments are incised by the fluvial flood waters entering the lake. The resulting channels are isolate and steep sided. 5
(Fig. l Oc) This unit consists of one or a series of small dolomite-increasing-upward cycles (Figs 8 and l Oc). It is interpreted as a shoreline prograding facies and is the Unit in which sabkhas are best developed. Unit 5 is remarkable for its lateral continuity, a fact recorded in the local name for the unit of No. 1 Marker (Fig. 9). The whole unit rarely exceeds 0·5 m in thickness and the individual laminae, even on the millimetre scale show remarkable lateral continuity. The marker silt layer at the top of the unit (type D storm layer) can be traced for many kilometers. Though not diagnostic, all the features described are compatible with a lacustrine interpretation. Depositional Unit
6 (Fig. lOd) By contrast with the preceding units the sediments of Unit 6 are highly variable in a lateral and vertical sense. In some areas the upper shoreline facies of Unit 5 are repeated whilst in others the change is marked by the incoming of stromatolite bioherms which exhibit a vertical succession of growth forms from flat mats through open growth forms to tight algal buns; this is interpreted as a progressive shallowing of the water. Where stromatolites or shoreline facies are not developed lithologies vary rapidly. The dominant change from the preceding units is always an increase in the depositional energy, reflected by dolomite and sandy dolomite flas�r bedding, dolomite-flake conglomerates and erosion channels cut into the bioherms. The incoming of the bioherms and the changes in the sedimentation pattern (lack of lateral continuity and increased depositional energy) may indicate a fundamental change in the water body.
Depositional Unit
Depositional Unit 7 (Figs 1 Oe and 1 1 )
Initially Unit 7 is marked by a change from the brown and khaki sediments of Unit 6
274
Harry Clemmey
to grey and black silts and fine sands. The trend of increasing energy of environment continues (Fig. 1 1 ) and a marked coarsening upwards culminates in the incoming of arkosic sands and gravels. In the lower parts of the unit bands of anhydrite nodules are found. At the top of the unit, beds typical of the alluvial fan facies of the Mindola Clastics occur.
Fig. 1 1 . Exposure of lowermost Unit 7 in Mindola North Pit. Note the wavy bedding, channels, and low angle cross-lamination.
The interpretation of Unit 7 suggests a continuation of the higher depositional energies of Unit 6, periodic uplift of the Kafue landmass (as recorded by horizons of siliciclastic sabkha) and finally the advance into the area of rej uvenated alluvial fans, thus bringing an end to Orebody Member sedimentation. A close analogue for Unit 7 times is provided by the siliciclastic sabkhas of the Sinai coast (Gavish, 1 974), where the Holocene sabkha is backed by high land with attendent alluvial fans.
DISCUSSION Vertical facies relations
The succession described above indicates deposition dominated by alternating exposed and subaqueous conditions, with the sediments defined as subaerial exhibiting evidence for high temperatures and evaporation rates. The pattern of alternating subaerial/subaqueous conditions is superimposed on a model of an initial marine transgression, abandonment with the establishment of lacustrine conditions then a re-connection with marine waters. Finally a major regression resulted in the re establishment of continental conditions (see summary, Fig. 12).
275
Proterozoic lacustrine interlude UNI T
7
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--
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12. Diagram summarizing the facies interpretation of the depositional units.
A cyclicity is detected within the Orebody Members (Fig. 13) and is thought to reflect the gross subaqueous or subaerial character of the depositional units. Cycle 1 does not have a subaqueous phase at the base as it represents the inital transgression. All other cycles have a subaqueous phase at the base (Units 2, 4, lower 6 and 7), and a subaerial phase at the top, (Units 3, 5, Upper 6 and 7). The subaerial phase of cycle 3 is often represented by sabkha development and that of cycle 5 by the return to continental conditions. This cyclicity is of great value in chronostratigraphic correlation.
Strontium values in the Orebody Member
Whole rock Sr values have been determined for the Orebody Member (Fig. 1 4). The low values probably reflect the release of Sr during diagenetic and metamorphic recrystallization. Two distinct groupings of Sr are evident. The first of these shows a progressive reduction in values between Units 1 -5 almost ignoring the gross carbonate content of the various units. Significantly Reeves ( 1 968) and Clauer ( 1 976) indicated that the Sr content of freshwater is generally ten times lower than sea water. The Sr content of Units 1 -5 is therefore consistent with their interpretation as recording the transition from marine to freshwater environments. The marked increase in Sr values after Unit 5 times coincides with the postulated return to marine conditions and the low values of Unit 7 would be reflecting the increasing freshwater contribution as the alluvial fan progrades into the area.
ECONOMIC SIGNIFICANCE
Some 78% of the Copperbelt metal reserves are located within this Member. The stratigraphic correlations on which this figure is based were arrived at using the cycles of sedimentation described above, the lithological characters of the depositional units and a recognition of the Kafue Arenites at the base. This work also destroys the myth of the Copperbelt Ore Shale, for shale of sapropelic character is rare in economic orebodies and the sedimentation of the Orebody Member was complex. Furthermore the rapid vertical variation in facies and lithology within individual deposits has a limited reflection in the Ore grades; the only real lithologic control is an antipathy between the mineralization and sediments of chemical and biogenic origin. The
276
Harry Clemmey
Uni t
y
Cycle 5
7
-
-
- - - - - -
------
Cycle 4 6
Cycle 3
5 4 - - - - - - - -
------
3
Cycle 2 2
- - -
-
-
- - - ------
Cycle I
13. Depositional cycles in the Orebody Member. Each cycle indicates a gross deepening to shallowing succession. The top of cycle 3 sometimes coincides with the top of Unit 5 in areas that were distal with respect to the palaeoshoreline.
Fig.
277
Proterozoic lacustrine interlude
D E POS I T I O NA L U N I T S 2
1
3
1
4
1
5
1
6
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7
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Fig. 14. Sr values for the Orebody Member. Note the two distinct populations. Maximum and minimum values are shown for the better sampled units and solid squares indicate isolated samples.
genesis of the ore deposits is, however, complex, involving both detrital and diagenetic processes (Clemmey, 1 976a).
CONCLUSIONS
After a lengthy review Picard & High ( 1 972) came to the conclusion that there are no unequivocal indicators of lacustrine conditions. The problems are shown to be compounded in the Precambrian thus a similar conclusion must be reached here. Conversely no unequivocal evidence of marine sedimentation has been described and the Orebody Member may be wholly lagoonal or wholly lacustrine. The evidence afforded by the location of the Orebody Member between two unequivocal fluvial units and the close association of the varying rainfall textures, median-ridged desiccation cracks, synsedimentary fold belts and preserved strand-lines probably indicates on balance, a lacustrine contribution. Several lacustrine indicators have also been presented, from a study of a modern lake, for further consideration and application.
ACKNOWLEDGMENTS
My thanks are due to the following people and institutions. Much of the work was undertaken whilst in receipt of an Anglo-American Corporation scholarship and written whilst at the University of Leeds as an NERC research fellow. Mike Leeder
278
Harry Clemmey
provided stimulating discussion during tenure of the scholarship, Andy Gardiner, Rog Anderton and Paul Bridges critically read ihe manuscript and Maurice Tucker gave a very able presentation of the paper in the author's absence.
REFERENCES ALLEN, J.R.L. ( 1 974) Reaction, relaxation and lag in natural sedimentary systems: general principles, examples and lessons. Earth Sci. Rev. 10, 263-342. CAHEN, L. ( 1 974) Geological background to the copper-bearing strata of Southern Shaba (Zaire). In: Gisements Stratiformes et Provinces Cupriferes (Ed. by P. B artholome), pp. 57-77. Societe Geologique de Belgique, Liege. CLAUER, N. ( 1 976) 87Sr/86Sr composition of evaporitic carbonates and sulphates from Miocene sediment cores in the Mediterranean Sea (D.S.D.P. Leg 1 3). Sedimentology, 23, 1 33- 140. CLEMMEY, H. ( 1 975) Correlation of the copper-bearing sequences of Zambia and Zaire. (Abs) 1 9th A nn. Rep. Res. lnst. African Geol. Univ. Leeds. CLEMMEY, H. ( 1 976a) Aspects of Stratigraphy, Sedimentology and Ore Genesis on The Zambian Copperbelt with Special Reference to Rokana Mines. Unpublished Ph.D. Thesis, University of Leeds. CLEMMEY, H. ( 1 976b) World's oldest animal traces. Nature, 261, 576-578. CLOUD, P.E. ( 1 968) Atmospheric and hydrospheric evolution on the primitive earth. Science, 160, 729-735. DAVIES, G .R. ( 1 970) Algal-laminated sediments, Gladstone Embayment, Shark Bay, Western Australia. In: Carbonate Sedimentation and Environments, Shark Bay, Western A ustralia (Ed. by B. W. Logan et a!.). Mem. Am. Ass. Petrol. Geol. 13, 169-205. DAVIS, R.A. ( 1 975) Ordovician intertidal deposits, Upper Mississippi Valley. In: Tidal Deposits (Ed. by R. N . Ginsburg), pp. 299-306. Springer-Verlag, Berlin. DoNOVAN, R.N. & FosTER, R.J. ( 1 972) Subaqueous shrinkage cracks from the Caithness Flagstones Series (Middle Devonian) of northeast Scotland. J. sedim. Petrol. 42, 309-3 17. G AVISH, E. ( 1974) Geochemistry and mineralogy of a Recent sabkha along the coast of Sinai, Gulfof Suez. Sedimentology, 21, 397-4 14. GuNATILAKA, A. ( 1 975) Some aspects of the biology and sedimentology of laminated algal mats from Mannar Lagoon, N orth West Ceylon. Sedim. Geol. 14, 275-300. HuNTER, R.E. ( 1 969) Aeolian microridges on modern beaches and a possible ancient example. J. sedim. Petrol. 39, 1573- 1 578. KINSMAN, D.J.J. ( 1 969) Modes of formation, sedimentary associations and diagnostic features of shallow water and supra-tidal evaporites. Bull. Am. A ss. Petrol. Geol. 53, 830-840. KLEIN, G.V. ( 1 975) Paleotidal range sequences, Middle Member, Wood Canyon Formation ( Late Precambrian), Eastern California and West Nevada. In: Tidal Deposits (Ed. by R. N . Ginsburg), pp. 17 1- 177. Springer-Verlag, New York. LAMBERT, A.M., KELTS, K.R. & MARSHALL, N . F . ( 1 976) Measurements of density underflows from the Walensee, Switzerland. Sedimentology, 23, 87- 105. PHLEGER, F.B. ( 1 969) A modern evaporite deposit in Mexico. Bull. Am. Ass. Petrol. Geol. 53, 824-829. PICARD, M.D. & HIGH, L.R. ( 1 972) Criteria for recognising lacustrine rocks. In: Recognition Of A ncient Sedimentary Environments (Ed. by J. K. Rigby and W. K. Hamblin). Soc. econ. Paleont. Miner., Tulsa, 16, 108- 145. PIPER, J.D.A., BRIDEN, J.C. & LOMAX, K. ( 1 973) Precambrian Africa and South America as a single continent. Nature, 245, 244-248. REEVES, C . C . ( 1 968) Introduction to Paleolimnology. Developments In Sedimentology, 1 1 . Elsevier Publishing Co., Amsterdam. REINECI;., H.E. & SINGH, l. B .( l 973) Depositional Sedimentary Environments. Springer-Verlag, Berlin. ROWLANDS, N.J. ( 1974) The gitology of some Adeladian copper deposits. In: Gisements Stratiformes et Provinces Cupriferes (Ed. by P. Bartholome), pp. 41 9-427. Societe Geologique de Belgique, Liege. ScHWARTZ, H., EINSELE, G. & HERM, D. ( 1 975) Quartz-sandy, grazing-contoured stromatolites from coastal embayments of Mauritania, West Africa. Sedimentology, 22, 539-5 6 1 .
Spec. Pubis int. Ass. Sediment. ( 1 978) 2, 279-290
Economic significance of playa lake deposits
C. C. REEVES Jr Department of Geosciences, Texas Tech University, Lubbock, TX 79409
AB STRACT Playa lake deposits accumulate in the topographic lows of arid and semi-arid localities. Regardless of latitude or elevation, the sedimentary history of any playa basin reflects periods of intermittent flooding interspersed with periods of desiccation. The lithology of the sedimentary fill is determined by topography, climate and geologic history. Other factors which may have local to regional influence are post-depositional faults, availability of aeolian debris, spring activity and nearness of active volcanic centres. Playa lake sediments range lithologically from boulder gravels to silts and clays, and may contain commercially valuable mineral deposits of various evaporites and uranium minerals. Future demands on minerals commonly found in playa lake deposits are bound to increase, and the more exotic mineral occurrences, such as those containing lithium, are expected to be utilized in the near future by expanding technology.
DEFINITION
The widespread and variable use of the term playa (Reeves, 1 968) requires a definition of the word as used in this paper. Gilbert ( 1 875) and Russell ( 1 883) were apparently the first geologists to use the term playa (which actually translates beach or shore) for the low central areas of intermontane basins (actually the barria[). Later Davis ( 1 905) and Tolman ( 1 909) persisted in wrongly calling the barrial the playa, so that full rhetorical entrenchment occurred. Lahee ( 1 9 1 6) then implied that playa lakes formed intermittently in the wind-scoured basins common to desert areas, however, this implicit genetic origin has been discounted in recent years. The term playa is used here to refer to the intermittently flooded low area of a basin, not necessarily restricting the term to only basins in classic intermontane troughs. There is no climatic or genetic restriction implied, although such basins obviously exist principally in semi-ahd lands.
Modern and Ancient Lake Sediments Edited by Albert Matter and Maurice E. Tucker © 1978 The International Association of Sedimentologists. ISBN: 978-0-632-00234-4
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C. C. Reeves
Jr
I M PORTANCE O F PLAYA LAK E S E D I M ENTS Playa lake sed i ments often contain not only a host of u n u s ual evaporitic m i nerals, but other m i n erals w h ic h until recently had little econo m i c or industrial im portance. B orax, soda ash, and sodi u m su lphate have been produced from playa lake deposits fo r many years, but it is only recently that ancient lake deposits have become a prime u ra n i u m exploration target. Present k n own s u p plies of lithi u m are totally inadequ ate to su pport pending technological developments for electrically-powered veh icles and thermonuclear power plants ( V ine, 1 9 7 6 ) . Zeolites, w h ic h are ubiquitous constituents of alkaline playa lake sediments, are also experiencing increasing com mercial u t ilization as are certa in types of playa lake clays. Playa lake sediments are also o f great academic i n terest. Not only do they m i rror the depositional history o f the su rround ing area, but they often contain su pplementary palaeoclimatic ev idence in t he fo rm of fossil re mains of flora and fa u n a . Early man sites, as at Lake R udolph, Africa, and t h roughout the western U n ited States, are associated with ancient lake basins and their deposits, and even earth quakes leave their signature in the sediments. Although the use of the term playa lake sediment may unconscio usly
lead
one
to
visu alize
rather
thicknesses of lacustrine deposits, much o f
meager accum ulations,
s u rprising
which acc u m ulated under interm ittent
Jake conditions, exist. For exam ple, 5,000 m of lacustrine strata occur i n R idge B asin, California ( Eaton, 1 9 3 9 : L i n k
&
Osborne, 1 9 7 8 ) and basins with 1 000 m or more of
lacustrine fill are too common to designate ( Reeves, 1 97 7 ) . The questional aspect of such thick lacustrine sections is to determine how much represents peren nial lake conditions and how much actually represents playa lake deposition, the best clue being the common presence of evapo rite m i nerals in the sediments deposited under intermittent lake cond itions.
PLAYA LAK E S E D I M ENTS Playa lake sed i ments and sediments deposited under perenn ial lake conditions may appear stri k ingly s i m ilar or exhibit sig n i fi cant d i ffe rences. Whereas genesis of a lake basin has relatively li ttle to do with sediment deposition within the basin, climatic parameters ( p rec i pitation, evaporation, prevailing w i nds, w i nd velocity, temperature and tem perat u re range ) exert a maj o r i n fl u ence on the s u rrounding sediment source area and thus on the lithology o f the lacustrine fill. Coarse clastics as well as fine clastics ( fi ne sands and clays) are deposited in lake basins i n all climatic regions, whereas evaporitic deposits occur i n basins only in the sem i-arid to arid areas. Twe n h o fel's ( 1 9 3 2 ) schematic of sediment distribution in a lake basin, ranging fro m the outer peripheral ring of coarse-grained sandstone around sandstone which su rrounds the more central m u dstones may exist as a gross s u rficial generalization ( H ardie et a/. , 1 9 7 8 ) , but d rilling often shows disruptions. For example, h ighland areas su rround ing a small playa basin w ill produce a dis proportionate amount of large clastics w hich w i ll be w idely s p read over playa clays by intermittent thu nderstorm runoff. O n the other hand, larger playa lake basin fills may exhi bit, in both a horizontal and vertical perspective, variable lithologic 'rings' o f sediment. The lithologic changes reflect local i n fl u e nces ( s u c h as lithology o f s u rrounding highlands, composition of
Economic significance ofplaya lake deposits
28 1
local spring waters, mineralogy of surrounding soils) and may, in both small and large basins, undergo rapid variations. Interbedding of facies is common. Playa lake basins are predominantly closed, thus chemical precipitants in the form of evaporites are common (see Hardie et a!., 1 978) in basins surrounded by.highland bedrock source beds. However, basins surrounded by lowland sedimentary strata with well-developed soils commonly exhibit sand or clay fills. For example, Jones ( 1 965), working in Deep Springs Lake basin, California, which is surrounded by highlands, mapped a playa zonation of saline minerals as well as a vertical succession through surficial salt polygons. However, playas of the same relative size in West Texas exhibit little more than peripheral sand lenses and central clay zones. Table 1 presents an outline of the common playa lake sediments, those having particular economic significance marked with an asterisk. Although many ofthe listed minerals may form under more than one depositional environment a singular listing is used for simplicity and brevity.
Table
l. Playa lake sediments
Clastics Silts, clays* Sands, gravels Volcanics Evaporites Carbonates - limestone, dolomite, natron * , trona* Sulphates - gypsum*, thenardite*, glauberite* ,epsomite*, bloedite* Halides - halite*, sylvite* Borates and Nitrates - borax*, searlesite*, colemanite*, nitre and soda nitre Silicates - magadiite, gels Lithium* Uranium* Authigenics (common) Pyrite, marcasite, goethite, sulphur, gypsum, banded iron (?) Glauconitic mica Silicates - kenyaite, bedded chert (?), quartz-chalcedony or opal Zeolites•, erionite, analcime, clinoptilolite Potassium feldspar (also calcite, dolomite, clays, salts) *Having particular economic significance
CLASTICS Silts and clays
The very fine grained clastics, which consist predominantly of calcium carbonate, gypsum, quartz, biogenic debr�s and clays, are found throughout playa lake basins. In the more distal parts of large lake basins the sediments often consist of nothing but fine clays (perhaps with some salts) whereas in near-shore areas the clays are mixed and interbedded with silts, sands, and gravels. The near-shore and surficial lithology of
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typical playa lake basins is thoroughly presented by Motts ( 1970). Ancient near- shore playa lake deposits can therefore be distinguished in cores by such sedimentary successions. Lacustrine clay sections may exceed several hundred metres in thickness, perhaps because they often represent the deepest parts of the ancient basins. Coates ( 1 952) reported at least 250 m oflacustrine clays in the Gila River Basin, Arizona, and Allison ( 1 945) mentioned 400 m of ' . . . soft blue lake muds . . . 'at Sumner Lake, Oregon. Thicknesses of clays in many of the deep troughs of the Basin and Range, U.S., undoubtedly exceed the previous figures. However, it is conjectural just how much of such sections represent playa as opposed to perennial lacustrine conditions. There is generally little correlation between clay mineralogy and palaeoclimates, the surrounding source area apparently exerting more influence on the clay mineralogy than the climatic regime. However, in some areas, such as West Texas (Parry & Reeves, 1 968), lacustrine clays exhibit increasing kaolinitic content on approaching a classic Pleistocene pluvial depositional boundary. On the other hand, clays deposited under semi-arid to arid conditions (regardless of mineralogy), usually contain salt beds or crystals (gypsum, halite), colloidal-sized calcite (Kerr & Langer, 1965), and are thinly and irregularly bedded.
Economic significance
Clays, usually of the bentonite and sepiolite type, are utilized as fillers in paper and paint manufacturing and as an anti-settling agent in paints. Sepiolite clays can be used as special drilling muds (temperature stability to 400'C) whereas the swelling montmorillonites are used as drilling muds or sealers for ponds or trenches. Hectorite clays, which are lithium-rich (see later section) are used in cosmetics, pharmaceuticals and various organophyllic compounds. High magnesia montmorillonite (saporite) is now being used for water treatment and in construction materials.
Sands and gravels
Sands and gravels accumulate principally along playa lake basin shores as gravel beaches or spits, and where canyons or alluvial fans intermittently disgorge large volumes of storm runoff. Excellent examples can be seen in nearly any intermontane valley in the western United States, and Pleistocene examples are well exposed along the eastern shorelines of ancient Lake Bonneville, Utah. Exploratory drill holes in some of the deep grabens of the B asin and Range have cut ancient (probably early to mid-Tertiary) gravels well removed from the present flanking mountain fronts. In some cases such gravels represent ancient stream channels but in other cases they apparently result from step-faulted mountain flanks which during earlier lacustrine stages, were much closer together (Reeves, 1 977).
Economic significance
Lacustrine sands and gravels have little widespread economic importance unless they happen to occur near urban centres where they can be commercially utilized as road metal, fill, or in concrete mix.
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Volcanics
Volcanic debris, in the form of ash, cinders, basalt flow fragments, bombs and pumice, has accumulated in many playa lake basins. Those lake basins fringed by volcanoes active during times of flooding (as near Elephant Mountain, Texas) may also contain wedging pillow flows, yet the most widespread volcanic sediment is undoubtedly volcanic ash. Volcanic ash deposited in playa lakes surrounded by low-lying sedimentary source rock (as throughout the Southern High Plains, Texas) usually undergoes little alteration and, due to runoff, may accumulate exaggerated thicknesses. However, in highly alkaline playa lake water, as once existed in the Big Sandy River valley, Arizona (Sheppard & Gude, 1 973), the volcanic ash rapidly alters to bentonitic clays, zeolites, alkali feldspars and silicates. Economic significance
Unaltered volcanic ash and pumice has been mined and used for several years as a cheap oil adsorptive in garages, truck and automotive terminals, loading docks, and factories (Reeves, 1 962). However, the most important economic products are derived from altered ash (discussed under Authigenics).
EVAPORITES
The evaporitic minerals which collect in a playa lake basin reflect the surrounding rock types, the local climatic regime, and atmospheric contributions, yet as Hardie & Eugster ( 1 970) wrote ' . . . the composition of the brines of modern continental evaporites exhibits a bewildering range, lacking any obvious correlation with respect to geography or drainage basin rock types . . . '. Krauskopf ( 1 967) believed the composition of basin tributaries, although initially dependent on local lithology, is quickly altered by climate, but Jones ( 1 966) stated that the primary anion in closed basin waters is determined by lithology of the drainage basin. Thus, although there is a diversity of opinion, playa basin waters, on evaporation, tend to occur as one of four principal brine types: Na-C03-S04-Cl, Na2 S04-NaCl, Na-Mg-Ca-Cl, and Na-Mg S04-Cl (Hardie & Eugster, 1 970). For a more recent discussion see Hardie et a!. ( 1 978). Carbonates
The most common lacustrine carbonates are calcite, aragonite and dolomite, with calcite and aragonite always the first minerals to appear on evaporation of carbonate rich water. The occurrence of dolomite is determined by the presence of high Mg/Ca ratios (Parry & Reeves, 1 965; Muller, Irion & Forstner, 1972). Ripple-marked lentils of limestone and/or dolomite are therefore common in playa basin silts and clays, particularly where the basins are not surrounded by highlands. Near highland areas the coarser clastics, as along the western side of Guzman Playa, Chihuahua, Mexico, are coated with calcium carbonate. In many playa basins in West Texas the limestone/dolomite zones are near-shore with the central playa areas exhibiting a gypsum mush. However, local areas around springs may show no calcium carbonate
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due to the acidic nature of the local water created by decay of local organic debris. Irregular lenses of carbonate, as well as other evaporites, are common in playa lake basin sediments due to deposition in restricted lake areas (Bates, 1 968, Bradley, 1 966). When playa lake alkaline brines have a pH > 9 trona may be deposited, providing the brine contains little sulphate and is rich in Na2 + , HC03- and C032-(Bradley, 1 964; Baker, 1 958; Eugster, 1 970, 1 97 1). Precipitation results from a substantial loss of both C02 and H 2 0 (Eugster, 1 97 1 ). Economic significance
To date lacustrine limestone and dolomite have no economic significance. Trona and other carbonates of soda (soda ash, natron) are basic industrial chemicals, but do not commonly occur in commercially-sized deposits in many playa lake basins (Smith et at. , 1 973). Most of the world's demand for trona is supplied from lacustrine deposits at Searles Lake, California, in the Green River Basin (Green River beds), and from Lake Magadi, Kenya, Africa. Sulphates
The principal sulphates found in playa lake deposits are bloedite, epsomite, gypsum, glauberite and thenardite, with gypsum being the most widespread. Lacustrine gypsums range from the very fine grained 'gypsum mush' so often found at playa surfaces to gigantic, exaggerated blades (Reeves, 1 968) produced by oxidation of sulphides. Playa surfaces commonly contain appreciable quantities of salt-resistant vegetation or vegetative debris which blows onto the surface, much of which is incorporated into the playa sediments. Sulphides, such as pyrite and hydrogen sulphide, are then produced by reduction during bacterial decay. In turn the oxidation of the sulphides provides the sulphate which combines with available carbonate to form the gypsum. How much gypsum in any one zone may be produced by bacterial action is not easily determinable, however, sulphur oxidizing bacteria T denitrificans, novellus, coproliticus, thioparus, and thiooxidans are common in gypsum brines. Thenardite (sodium sulphate), epsomite (magnesium sulphate), glau berite (calcium sodium sulphate) and bloedite (magnesium sodium sulphate) frequently occur with trona salts or mixed with halite and gypsum. Deposits of these sulphates are widespread, forming in both large playa environments such as Great Salt Lake, Utah (Cohenour, 1 966) as well as in small basins. Economic significance
Bloedite, epsomite, glauberite and thenardite are basic industrial salts, being used by many companies for production of everyday essentials (explosives, clothing, fertilizers, soaps/ detergents, glass, pharmaceuticals), thus even relatively small deposits, if strategically located, may be profitably utilized. Deposits of these sulphates, of considerable magnitude, are known in remote areas (for example, in northwestern Chihuahua, Mexico) which are not yet exploited. Thus no apparent shortages exist. Gypsum, used by the construction industry (cement use, plaster, sheetrock) and agriculture, is supplied principally by marine deposits.
Economic significance ofplaya lake deposits
285
Halides
Lacustrine deposits frequently contain significant deposits of halite (e.g. Lake Eyre, Australia), whereas most of the potassic chlorides, like sylvite, stay in solution or the miniscule amounts precipitated are rapidly removed by deflation. However, under certain environmental/depositional conditions which are largely unknown, sylvite lenses may occur with associated halite beds (Smith et a!., 1 973).
Economic significance
Playa lacustrine sediments contain little potassic halides and attempts to recover potassic salts from such deposits, except as by-products, have not been widely successful. However, sodium chloride is mined, as at Saltair, Utah (Cohenour, 1 966), for commercial, industrial and domestic use. Playa-lake basin brines, from California, Utah, and the Dead Sea, Israel, do yield potassium compounds (Smith et al. , 1 973).
Borates and nitrates
The borates (borax, kernite, searleasite, colemanite) and nitrates (soda nitre and nitre are usually found in most playa deposits formed in basins surrounded by volcanic terrains, yet commercially-sized deposits of the borates and nitrates are not common. It is suspected that this is mainly the result of the past reluctance to drill playas, combined with the ready availability of such minerals from such classic areas as Death Valley and Searles Lake, California. Borax and kernite are mined from lacustrine deposits at Boron, California and lacustrine deposits yield calcium borates near Death Valley, California (Smith et al. , 1 973). However, borate production from Searles Lake originates from the highly alkaline lacustrine brines.
Economic significance
Borax is a basic industrial chemical, being used in a multiplicity of products (Bateman, 1 956), from soaps and detergents in the chemical industry to use as a chemical in gasoline additives and fire-retardants. Nitrates are not mined deliberately from lacustrine deposits.
Silicates
Playa lake basins in semi-arid to arid areas commonly contain silicates in their sediments from the devitrification of volcanic ash in the highly alkaline environments. However, the original silicate may have actually been a precipitate like the hydrous sodium silicate magadiite (Eugster, 1 967) or perhaps a siliceous gel. Chert and opal lenses are also often found in lacustrine beds (as at Virgin Valley, Nevada), Eugster ( 1 967) showing that bedded chert could originate from kenyaite which originated from magadiite.
Economic significance
No economic significance is presently known for lacustrine silicates.
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Lithium
Lithium in the lacustrine environment is so soluble that it concentrates during the late evaporative stages, with the potassium, boron, chloride or sulphate brines (Cannon, Harms & Hamilton, 1975; Vine, 1 975) and with the montmorillonite clays. Lithium-bearing clays (smectites), which represent the only lithium-rich playa lake basin sediment known, are commonly associated with calcite, dolomite, sepiolite and attapulgite. Several small occurrences oflithium-bearing clays from West Texas playa sediments are similar to hectorite, a lithium-rich clay from hydrothermal deposits near Hector, California. The West Texas lithium-rich clays, however, contain somewhat more Mg2 + and Fe2 + and about 0·8% less lithium (Goolsby, Reeves & Lee, 1 977). Lithium (as a phosphate) is produced as a by-product of the salt production at Searles Lake, California, and lithium (as a carbonate) is produced from playa lake brines in Clayton Valley, Nevada (Vine, 1 975). A major lithium occurrence in the Salar de Atacama, Chile, was also recently reported (Vine, 1 975) Economic significance
Lithium is used (Vine, 1 975) by several commercial industries (glass-making, ceramics, industrial chemicals, dry-cell batteries), perhaps the most notable being its use as lithium carbonate in the medical field to stabilize manic-depressive behaviour. However, the development of electric vehicles requires the lithium battery which could, by the early 1 980s, double lithium demand. A second technological development which may soon spur lithium demand would be the development of thermonuclear fusion plants which require a lithium blanket around the reactor unit (Bogart, 197 5). Uranium
Radioactive anomalies are common to lacustrine sediments and particularly to those deposited in closed basins. Such occurrences are known from Poland, Czechoslovakia, Turkey (Fakili), Japan (Mingyo-toge), France (Loveve), the Sila Plateau, Italy, western Australia and throughout the western United States. This propensity results from the extreme mobility of the uranyl ion under alkaline conditions, the availability of suitable reductants in the lacustrine environment, and the natural trapping mechanism of a closed basin. Frequently the uranium minerals will be closely associated with opaline beds (Virgin Valley, Nevada, or the Southern High Plains, Texas), with calcareous zones (Yeelirrie, western Australia or western United States), with zeolites (Tono Mine and Oochi deposits, Japan), or with volcanic ash (western United States and southern Texas). Although some authigenic uranium may occur in playa lake sediments, most appears to have been transported by meteoric water until encountering a reducing environment to cause precipitation of the uranyl carbonate complex. Uranium minerals encountered below the water table in playa lake sediments will usually occur as uraninite or coffinite whereas that above the water table will usually occur as carnotite, tyuyamunite or autunite. Uranium in lacustrine sediments is always in close proximity to igneous bedrock and adjacent deposits of arkosic sands and volcanic ash, but which acted as the original source of the uranium has long been debated (Grutt, 1 972; Rosholt & B artel, 1 969).
Economic significance of playa lake deposits
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Although we have now identified additional causative factors (such as structure, fault control, spring locations), the role of the playa lake environment (alkaline pore water, plant material, humic acids) can hardly be overestimated.
Economic significance
By the year 2000 about 60% of the nation's electric energy is expected to be produced by nuclear power plants, requiring at least ten times the present annual domestic uranium requirement. This realization has resulted in an approximate ten-fold increase in the price of U308 in the last few years, creating an exploration 'boom' of classic proportions. The discovery of a commercially-sized uranium deposit in playa lake sediments (Jones, 1 970) therefore focused exploratory attention on playa lake environments throughout the world. It is expected that there will be several significant discoveries of massive uranium deposits from both ancient and modern playa lake basins during the next decade.
A UT H I G E N I C S
Authigenic lacustrine sediments, by definition, cannot form a primary deposit, yet the relative amounts in select zones may be uncommonly high. Certainly the recognition of over seventy authigenic minerals in the Green River Formation, Wyoming (Milton, 197 1 ) illustrates that fantastic mineralogic variations may occur in the playa depositional environment.
Zeolites
Zeolites, which are common in alkaline playa lake sediments, occur as clinoptilolite, erionite and phillipsite with some chabozite, mordenite and analcime. All apparently form from volcanic ash and tuff which accumulate in the basins. Although the progressive sequence of volcanic glass to alkalic zeolite to analcime to potassium feldspar is well documented, the particular alkalic zeolite which may form in a local lacustrine environment is not predictable (Surdam & Eugster, 1 976). What has been shown is that volcanic glass to alkalic zeolites, with the exception of analcime, form in the peripheral (freshest) lake areas. Analcime then occurs further basinward, and finally potassium feldspars form in the most alkaline/saline environments (Sheppard & Gude, 1 973). Erionite and phillipsite form mainly in saline lacustrine deposits while clinoptilolite and mordenite may form in either fresh or saline environments (Hay, 1 964, 1 966). The presence of zeolites in playa lake sediments then supplies important depositional environmental clues. However, perhaps the most important aspect is the possible relationship between zeolite formation and release of uranium from the volcanic glass. Although Rosholt, Prij ana & Noble ( 1 97 1 ) do not believe volcanic ash can supply the amounts of urarnum required, the massive amounts of volcanic ash in many basins and the close mineralogic association of uranium minerals with remnant ash shards suggests we may not know all the answers here.
C. C. Reeves Jr
288 Economic significance
Zeolites have become an important industrial resource in the last few years due to their unusual ion-exchange, dehydration, absorptive and catalytic properties. Of particular interest are the uses of zeolites for municipal sewage control, metal recovery and in pollution control. It therefore seems that the increasing importance of zeolites will lead to additional exploration of playa lake environments.
CONCLUSIONS
Even from this brief review it is apparent that the playa lacustrine environment is, and throughout geologic history has been, responsible for a magnificent variation of common and even uncommon sediments and minerals. What was once considered as 'worthless basin fill' is now the subject of an unusual concentration of geologic investigation. The recent need for several of the uncommon minerals produced from the lacustrine sediments will inevitably divest data which cannot, under ordinary investigative budgets, be secured. Thus, not only will geologists trained in working with the lacustrine facies be able to adequately contribute to exploration/exploitation ventures, but will be rewarded with those 'dry' academic morsels so necessary to future research.
REFERENCES ALLISON, l.S. ( 1 945) Pumice beds at Sumner Lake, Oregon. Bull. geol. Soc. Am. 56, 789-808. BAKER, B . H . ( 1 958) Geology of the Magadi area. Kenya geol. Surv. Rep. 42. BATEMAN, A.M. ( 1 956) Economic Mineral Deposits. John Wiley and Sons, New York. BATES, T.R. ( 1 968) Late Pleistocene geology ofpluvial Lake Mound, Lynn and Terry Counties, Texas. MS. Thesis, Texas Tech. Univ. BOGART, S.L. ( 1 975) Fusion power and the potential lithium requirement. Lithium Resources and Requirements by the Year 2000. Prof Pap. U.S. geol. Surv. 1005, 1 2-2 1 . BRADLEY, W . H ( 1 964) Geology o f the Green River Formation and Associated Eocene Rocks of Southwestern Wyoming and Adjacent Parts of Colorado and Utah. Prof Pap. U.S. geol. Surv. 4%-A. BRADLEY W . H . ( 1 966) Paleolimnology of the trona beds in the Green River Formation of Wyoming. 2nd Symp. Salt Ohio Geol. Soc. Cleveland, 1 60-1 64. CANNON, H .L., HARMS T.F. & HAMILTON J.C. ( 1 975) Lithium in unconsolidated sediments and plants of the Basin and Range province, southern California and Nevada. Prof Pap. U.S. geol. Surv. 918. COATES, D.R. ( 1 952) Gilda Bend basin, Maricopa County. In: Groundwater in the Gila River Basin and Adjacent A reas, A rizona. A Summary (Ed. by L.C. Halpenny et a/.), pp. l 7 1 - 1 76. U.S. geol. Surv. open file report. COHENOUR, R.E. ( 1 966) Industrial development and potential of Great Salt Lake with notes on engineering geology and operational problems. In: Guidebook to the Geology of Utah (Ed. by W. L. Stokes), pp. 1 53- 163. Utah geol. Surv., Salt Lake City. DAVIS, W.M. ( 1 905) The geographical cycle in an arid climate. J. Geol. 13, 3 8 1 --407. EATON, J.E. ( 1 939) Ridge Basin California. Bull. Am. Ass. Petrol. Geol. 23, 5 1 7-558. EuGSTER, H.P. ( 1 967) Hydrous sodium silicates from Lake Magadi, Kenya: precursors of bedded chert. Science, 157, 1 1 77- 1 1 80. .
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