PALAEOWEATHERING, PALAEOSURFACES AND RELATED CONTINENTAL DEP OSIT S
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
S PE C I A L PUB L I CAT I O N NUM B E R 27 O F T HE I N TE RN AT I O N A L A S S O C I AT I O N O F SED I ME N T O L OGI S T S
Palaeoweathering, Palaeosurfaces
and Related Continental Deposits EDITED BY MEDARD THIRY AND REGINE SIMON-COIN<;ON
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© 1999 The International Association of Sedimentologists
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Contents
vu
ix
3
Preface
Acknowledgements
Introduction Problems, progress and future research concerning palaeoweathering and palaeosurfaces
M. Thiry, J-M. Schmitt & R. Simon-Coim;on
21
Geochemistry and isotopes Weathering, rainwater and atmosphere chemistry: example and modelling of granite weathering in present conditions in a C02-rich, and in an anoxic palaeoatmosphere J-M. Schmitt
43
Stable carbon isotopes in palaeosol carbonates
61
Palaeoenvironment, palaeoclimate and stable carbon isotopes of Palaeozoic red-bed
TE. Cerling
palaeosols,Appalachian Basin, USA and Canada
C.!. Mora & S.C. Driese
87
Peculiar palaeoweathering types Diversity of continental silicification features: examples from the Cenozoic deposits in the Paris Basin and neighbouring basement
M. Thiry 129
Authigenic clay minerals in continental evaporitic environments
153
Saprolite-bauxite facies of ferralitic duricrusts on palaeosurfaces of former Pangaea
189
Karst bauxites: interfingering of deposition and palaeoweathering
J.P Calvo, M.M. Blanc- Valleron, JP Rodriguez-Arandia, JM. Rouchy & M.E. Sanz I. Valeton
G. Bardossy & P-J Combes
v
207
225 245
Precambrian palaeosols: a view from the Canadian shield Q. Gall
Regional palaeosurface and palaeoweathering reconstructions Palaeolandscape reconstruction of the south-western Massif Central (France) R. Simon-Coim;on Lateritization, geomorphology and geodynamics of a passive continental margin: the Konkan and Kanara coastal lowlands of western peninsular India
M. Widdowson & Y Gunnell
275
Relief features and palaeoweathering remnants in formerly glaciated Scandinavian basement areas
K. Lidmar-Bergstrom, S. Olsson & E. Roaldset
303 323
Palaeosol sequences in floodplain environments: a hierarchical approach M.J Kraus & A. Aslan Carbonate-rich palaeosols in the Late Cretaceous-Early Palaeogene series of the Provence Basin (France)
I. Cojan 337
Sedimentary infillings and development of major Tertiary palaeodrainage systems of south-central Australia NFAlley, JD.A. Clarke, M. MacPhail & E.M. Truswell
367
Weathering surfaces, laterite-derived sediments and associated mineral deposits in north east Africa
T. Schwarz & K. Germann
391
Index Colour plate section appears facing p.158
Preface
future to set-up mass balances between erosion and deposition, landmasses and basins. Several regional studies in Australia, India, France and Scandinavia draw up these geometrical aspects and their relation to crustal tectonics. The geochemical aspects are mainly bound to the palaeoclimatic studies and to global changes. Several chapters show how pecu liar palaeoweathering types allow us to reconstruct ancient climates and how this has recorded important changes in the Earth's environment. Isotopic studies, used either as palaeoenvironmental indicators or as dating tools, are particularly invaluable for such reconstructions. This book also shows how geochemi cal modelling can help in testing hypotheses and in simulating weathering conditions that differ from the modern ones. It must also be noted that the study of palaeo weathering is not a simple transposition to geologi cal deposits of results established by soil scientists in modern soils. The geological approach to ancient weathering features often reveals weathering profiles that are much thicker and have unusual geochemical signatures in comparison to present landscapes. Thus, studies of palaeoweathering can lead to concepts, processes or models that are novel to soil scientists. The strength of this book is in the variety of its techniques and in the pluridisciplinary approaches, which have brought together geologists, geomor phologists and geochemists working on the recon struction of ancient continents. This collection of papers should illustrate the range of palaeoweath ering occurrences and the great variety of their geochemistry and environmental significance. It should serve as an introduction to the reconstruction of ancient continental environments and will be of use to geologists, soil scientists and geomorphologists as well.
Most previous lAS Special Publications have dealt with marine deposits, with only a few devoted to con tinental deposits (i.e. Alluvial Sedimentation, Aeolian Sediments, Lacustrine Facies Analysis and Calcretes), and none has been dedicated to continental palaeosurfaces and palaeoweathering. To complete the panel of the topics covered by the Special Publi cations, Michael Talbot asked us to prepare a volume on problems related to the reconstruction of ancient continents. Moreover, there has been a recent growth of interest in the study of ancient continents. As most basin deposits have a land-based origin it appears important to reconstruct successive weatherings, landscapes and environments that have prevailed on ancient continents. Palaeoweathering is, in this respect, of prime importance: it is the means by which basins are fed with soluble and solid compounds, it forms a useful basis from which to reconstruct ancient environments, palaeoclimates and palaeo geographies and it also acts as an exchange interface with the atmosphere and thus participates in global changes. Palaeoweathering studies also contribute to the correlation of deposits and establish rate and timing of uplift and erosion. The core of this volume is a contribution from the International Geological Correlation Program IGCP 317 Paleoweathering Records and Paleosurfaces. Several meetings and field trips, that were the occa sion of numerous official presentations and informal discussions from which emerged a number of ideas leading to papers in this volume, have marked this International UNESCO-lUGS Program. Two main aspects of the palaeolandscape studies are tackled within this volume, namely the physical and the chemical approaches. Palaeoweathering pro duces invaluable markers of continental palaeosur faces, allowing us to reconstruct palaeolandscapes, but also to assess the volume of denudation or sta bility in palaeosurfaces. It will contribute in the near
Medard Thiry Regine Simon-Coin<;on
vii
Acknowledgements
Alain Godard, Universite Paris 1-Sorbonne, Meudon, France; Heinrich D. Holland, Harvard University, Cam bridge, USA; Dale Leckie, Geological Survey of Canada, Calgary, Canada; Earle F. McBride,University of Texas, Austin,USA; Arthur H. Palmer, State University College, Oneonta, USA; Nigel H. Platt, Ranger Oil Limited, Guildford, United Kingdom; Jean-Michel Schmitt, Ecole des Mines de Paris, Fontainebleau, France; Yves Tardy, Universite Louis Pasteur, Strasbourg, France; M.F. Thomas, University of Stirling, Stirling, United Kingdom; Cesar Viseras, Universitad de Granada, Granada, Spain; James C.G. Walker, University of Michigan, Ann Arbor,USA; Michael Widdowson, The Open University, Milton Keynes, United Kingdom; Paul Wright,University of Reading, Reading, United Kingdom; Dan Yaalon, The Hebrew University, Jerusalem, Israel; Richard Yuretich, University of Massachusetts, Amherst, USA.
In addition to the editors, who served as reviewers of all manuscripts, numerous outside referees, selected for their experience in various aspects of palaeoweathering, have given freely of their time and expertise for their conscientious and constructive perusal of the manuscript. They contributed to the quality of this volume and the editors wish to thank them here. Yvonne Battiau-Queney, Univ. Sci. Techn. de Lille, Villeneuve-d 'Ascq, France; Michael Bird, Australian National University, Can berra, Australia; Robert P. Bourman, The University of South Australia,Underdale, Australia; Emery Cleaves, Maryland Geological Survey, Balti more,USA; Isabelle Cojan, Ecole des Mines de Paris, Fontainebleau, France; Avigator Cohen, Geological Survey Israel, Maale Adumim, Israel; J.P Calvo,Universidad Complutense, Madrid, Spain; K.G. Cox, University of Oxford, Oxford, United Kingdom Bruno D'Argenio, Geomare Sud, Napoli, ltalia; Gregg Davidson, University of Mississippi, Missis sippi,USA; T.C. Devaraju, Karnatak University,Dharwad, India; Jane E. Francis, University of Leeds, Leeds, United Kingdom; R.J. Gilkes, University of Western Australia, N edlands, Australia; Klaus German, Technishe Universitaet Berlin, Berlin, Germany;
Special thanks to Bob Bourman, Klaus German, Michael Talbot and Mike Widdowson who helped in language correction of some papers.
ix
Introduction
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
Spec. Pubis int. Ass. Sediment. (1999) 27,3-17
Problems, progress and future research concerning palaeoweathering and palaeosurfaces
M . TH I RY , J . - M . S C H M I TT andR. S I M O N - C O I N<;:O N C. I. G., Ecole Nationale Superieure des Mines de Paris, 35 rue SLHonore, 77305 Fontainebleau Cedex, France and CNRS - UMR SISYPHE C 7619 Structure et fonctionnement des systemes hydriques continentaux
INTR O D U CTION
deposits appeared in the second half of the 19th century. The'superficial' hypothesis, based on the dis solution power of rainwater, developed through the impetus provided by chemists. For example, Ebelmen (1851) reported the dissolution of the limestones by rainwater, Del esse (1853) showed the transformation of granite into sand and of orthoclase into kaolinite, Delanoue (1854) described the dissolution of lime stones and the oxidation of the organic matter they contain and Friedel (1876) described the weathering of agate and flint. It was van den Broeck (1881) who first applied these observations in order to explain the genesis of particular geological formations. He spoke of'hydro chemical metamorphism' in the development of karst and its infill material. Dieulafait (1885) used a similar approach with regard to phosphorite deposits in the karst of south-west France. He used geometric argu ments to demonstrate that cave development and the associated phosphate deposits are due to down ward flushing by water and that rainwater containing carbonic acid can explain the dissolution of the limestone. Fleury (1909), in his work on the Swiss Siderolithic Formation, was the first to demonstrate clearly that these sequences are the result of the weathering of rock, ditferential leaching of various chemical elements and the transformation of the dissolution residues. He made a distinction between the periods of formation and reworking of the weathering products, and emphasized the role of palaeoclimate. For the first time the notion of palaeo weathering had appeared in geology. Afterwards, until the 1960s, the study of ancient weathering remained confined to the study of palaeoweathering resulting from leaching processes under humid tropical climates, with the development of kaolinite-rich residues, commonly called laterites.
Continents form wide areas o f Earth's surface, in fact much wider areas than the limited marine areas where most sediments accumulate. Despite this, geological records of the ancient continents are scarce, with some records perhaps totally missing. When present they are often puzzling owing to their successional evolution. Nevertheless, continen tal palaeolandscape and palaeoweathering recon structions are an important aspect of global palaeoenvironment understanding. Often, continents preserve only scanty testimony of their evolution; they offer evidence of only a patchwork of landforms and weathering horizons of different age and origin, discontinuous in time and space. It also has to be noted that the study of palaeo weathering is not simply the transposition to geologi cal deposits of results established by pedologists in modern soils. The geological approach to ancient weathering features involves descriptions and study of complete geological formations, rather than just superficial levels, and frequently reveals weathering profiles that are much thicker and which have unusual geochemical signatures in comparison with present landscapes. Thus, studies of palaeoweather ing can lead to concepts of processes and models that are novel to pedologists. Palaeoweathering studies contribute to palaeo environmental and palaeogeographical reconstruc tions, global correlation of deposits, records of global change, rate and timing of uplift and erosion, and an inventory of ore resources.
T H E E A R LY S T U D I E S
The concept of 'meteoric' alteration o f rocks and its importance in the development of some sedimentary 3
4
M. Thiry et al.
Bauxite is the palaeoweathering example quoted most frequently for this period of study. Even if some silicification or development of silcretes were described as being formed during specific periods under surficial conditions, they generally were not recognized and discussed as true weathering products. Similarly, the recognition of the palaeosur faces was most often restricted to their geometrical recognition, often by construction from isolated and distant topographical points. Since the 1960s, the recognition of the ancient land scapes has been refined considerably. This has been due in part to the contribution made by new analyti cal techniques, but above all it results from a better knowledge of modern weathering processes. Pedology provided the main 'tools' for analysing palaeoweathering. Micrographic techniques that were developed during the 1960s provided the specific criteria for soil structure recognition: illuvia tion, cappings, distinction of illuviated and eluviated horizons, diverse nodules and cutans, for example. Furthermore, studies of modern weathering covered a wide variety of parent rocks and morphoclimato logical conditions. Recognition of weathering profiles characterized by accumulation processes, such as fer ricretes, calcretes, gypcretes, vertisols, and so on, added to the range of weathering types characterized by leaching processes already known at that time. In the meantime, pedologists introduced the concept of catenas, with diverse weathering features of the same age present at a variety of topographical levels in the landscape. This notion of various weathering horizons in a landscape was an essential step in study of the ancient landscapes, it allowed the linking of different palaeoweathering features and from here the reconstruction of palaeolandscapes. The development of analytical techniques, such as scanning electron microscopy, palaeomagnetism, thermoluminescence, and especially the develop ment of isotope geochemistry, provided new 'tools' for studying palaeoweathering and palaeosurfaces. Two main categories of 'tools' can be distinguished: those that provide information on palaeoenviron mental conditions (composition of the palaeoat mosphere, concentration or composition of hydrous solutions, etc.) and those that allow dating of the palaeoprofiles themselves or the associated azoic continental deposits. Knowledge of the palaeoenvi ronment is invaluable, but from the geologist's point of view dating is even more important. Indeed, the difficulty, or even the impossibility, of dating most palaeoweathering horizons was always a major
handicap to the reconstruction of ancient landscapes: dating the horizons, correlating with confidence, and estimating the duration of formation of some pro-files, has entirely revised our ideas about some palaeoweathering horizons or major palaeosurfaces.
W H E R E TO L O O K F O R PA L A E O W E ATH E R I N G
Weathering is active everywhere on the continents. Its intensity depends mainly on geomorphology, on climate and on the time span during which weather ing processes have been operating. The best way for palaeoweathering phenomena to be recorded in geological materials is by their preservation as palaeoprofiles in the stratigraphical column. On upland areas, weathering may act continually for long periods, but the weathered material frequently, if not regularly, will be eroded and stripped off, so that much of it will not be preserved. In basinal areas, weathering is active only during short periods of emergence: palaeoprofiles remain incipient because they have little time to develop, although they are preserved more easily between successive sedimen tary episodes. Therefore, mature and deep weathering profiles have the best chance to be preserved and become palaeoprofiles on the margins of basins. This espe cially is true along passive margins, which are charac terized by crustal flexure, with inland uplift and subsidence of the basin. Here, the land surface is rela tively stable with regard to regional base level, the rate of erosion is low, and sediment deposition occurs only from time to time, allowing weathering to develop and providing a protective cover to weather ing features. Such sedimentary covers also provide valuable dating criteria. They are excellent record ers of vertical movements and of palaeolandscape changes. Within this specific geodynamic framework, geometric relationships between successive palaeo profiles and palaeosurfaces can be clarified (Fig. 1 ) . I n the basin, the accumulation o f sediments leads to superimposition of successive palaeosurfaces. Uplands, on the other hand, are affected by successive stages of downcutting, which will lead to preservation of the older palaeosurfaces on high plateaux and of the younger ones in lower positions. Plana tion surfaces provide essential markers. The use of the correlative deposits and palaeoweathering features can explain how planation surfaces were generated.
Problems, progress and future research 1 2 3
+
+
+
+ +
+
-l+
+
+ +
+
+
+ +
+
+
+ +
+
+
+ +
+
+
+ +
Fig.l. Mature and deep weathering profiles have the best
chance of preservation and of becoming palaeoprofiles when they are situated along the passive margins of basins. On the continent, weathering may occur, but the weathered material frequently, if not regularly, will be eroded and stripped off. In basinal areas, weathering is active only during short periods of emergence and therefore palaeoprofiles remain incipient.
The major Tertiary palaeodrainage systems of south-central Australia are good examples of the complicated geometric relationships between succes sive stratigraphical units (Alley et al., this volume, pp. 337-366). Marine influence extended several hundred kilometres up the major channels. Very diverse palaeoenvironments developed along these narrow strips, such as gallery forest, swamps, dry lakes, etc., allowing fruitful palaeogeographical and palaeoclimatological reconstructions. Nevertheless, at times of overall rising sea-level and aggradational sedimentary systems, the relative dating of palaeosurfaces may be more complicated. The lowest parts of the landscape are buried first and the higher parts remain exposed for longer. In this way the oldest palaeosurfaces will be preserved in lower positions and the youngest ones on the high plateaux. To be useful in palaeoreconstructions, however, the sedimentary systems must be aggrada tional so that parts of the palaeosols and palaeosur faces are buried and preserved.
DAT I N G
Independent dating o f palaeosurfaces and associated palaeoweathering horizons has been a major concern for palaeolandscape and palaeogeographical recon structions. Development of isotope geochemistry has opened new pathways in the study of palaeowea thering horizons. First there are the direct methods of dating by radiometric analysis, such as K-Ar dating, which has been applied to alunite and also to man ganese oxides that contain potassium (cryptomelane
5
and hollandite). Nuclides formed cosmogenically promise an alternative solution to the problem of surface dating, by allowing estimates of the duration of exposure from the presence of radionuclides (lOBe, 2 6Al) and stable nuclides (3H, 21Ne) that have formed by the interaction of cosmic radiation with atoms exposed at the rock or sediment surface. Similarly, quartz and feldspar exposure to sunlight provides the basis of thermoluminescence dating. Variation of stable isotopes (13C, 180, D) can be used as an 'indirect' method for dating regoliths based upon changing composition of the atmosphere, either global or in response to climatic variation, or to latitu dinal drift. Establishing the oxygen-isotope composition of various regolith profiles and kaolinitic sediments from across Australia has allowed the distinction of profiles formed in the Late Palaeozoic and late Mesozoic to early Tertiary from profiles formed in mid-Tertiary times during the northward drift of the continent (Bird & Chivas 1988, 1993). An alternative way of using stable isotopes for dating is to link the isotopic variations to a global reference curve. It works especially well when the global curve dis plays distinct variations, such as the ol3C excursion around the Palaeocene-Eocene boundary. The 1)13C organic record from Palaeocene-Eocene deposits and palaeosols of the Paris B asin shows several 'unique' events. These events can be used as geo chemical fingerprints to correlate specific horizons from section to section, and moreover to relate them to the deep-sea reference records (Sinha et al. 1 995; Thiry & Dupuis 1998). Dating also may be carried out by using the record of major changes in the ecosys tem, such as the change from C -dominated to C43 dominated biomass during the Neogene (Cerling et al. 1993; this volume, pp. 43-60) . Development of new techniques, such as the laser microprobe, allows the application of isotopic methods at a microscopic scale, in situ, without the need to purify or separate mineral phases. Thus the field of investigation has been enlarged considerably. Such refined dating allows us to capture the progres sion of an oxidation front during weathering and to propose a 'land-based stratigraphy' of past climatic and geomorphological events. Samples of void-filling banded and colloform manganese oxide from vertical profiles reveal remarkably extended periods related to episodic mineral precipitation, corresponding probably to the alternation of more humid and drier climatic periods (Vasconcelos et al. 1 994a,b; Dammer 1995) (Fig. 2). Moreover, microsampling of thin over-
M. Thiry et a!.
6
hand specimen of banded cryptomelane, Australia
Fig. 2. Relative probability plots
of K-Ar and laser microprobe Ar/39Ar ages from Australia and Brazil ( after Vasconcelos et al. 1994b & Dammer 1995). Good correlation of the data from several localities and of texturally distinct features in subsamples show that the formation of K-Mn oxides, and thus the period of intense weathering during the Tertiary, occurred during many relative short intervals, which alternated with periods when weathering was less intense. The periods of intense weathering proposed in Australia are comparable with similar periods identified in Brazil, thus suggesting that factors triggering intense weathering have widespread significance. 40
MIOCENE
0 2
5
(Ma)
OLIGOCENE
24
I
I 38
growths even allows the calculation of colloform manganese oxide concretion growth rates (Dammer 1 995). It has been shown that some weathering features from the Brazilian, African and Australian regoliths extended over the whole Tertiary, with specific age clusters that can be correlated across the continents (Vasconcelos et al. 1994a,b; Dammer 1995) (Fig. 3).
S T R ATI G R A P H I C A L R E C O R D
Palaeoweathering horizons and palaeosurfaces form an important part of the geological record because they represent fundamental events (climatic, tectonic, etc.) that aid in stratigraphical correlation and division. The formation of large-scale palaeosurfaces and extensive palaeoweathering horizons often will have influenced sedimentation in adjacent marine basins through changes in the rate, nature and composition of sediment supply. Intense weathering acting on stable continents imprints the landscape with strong mineralogical and geochemical fingerprints in the form of bauxites, laterites, ferricretes, silcretes, cal-
I 55
I 60
cretes, etc. These geochemical signatures recorded in palaeoweathering profiles and their correlative conti nental deposits, can provide the basis for large-scale correlation in azoic sequences. Weathering surfaces of north-east Africa and the sediments derived from these surfaces provide good examples of such stratigraphical correlations (Schwarz & German, this volume, pp. 367-390). The development of major palaeosurfaces may be recognized from changes in the pattern of correlative sedimentation in neigh bouring basins. Alternatively, palaeoweathering horizons can be used to identify palaeosurfaces that may be traced across ancient shields into marginal basin sequences, where they are marked by unconformities. Ancient stable continental surfaces, which have experienced little or no deposition and erosion, may be veneered by discontinuous but locally thick palaeosols. In a stratigraphical sequence such surfaces represent periods for which the geological record is missing, they are unconformities or diastems of variable dura tion. Numerous palaeosols and palaeokarst features line major sedimentary unconformities and point to stratigraphical gaps of eustatic or tectonic origin. Recognition of such weathered surfaces may help in
7
Problems, progress and future research EPOCH
MA
-
USA Nevada
PLIOCENE
10
Brazil Atlant. - Amazon
MIOCENE
20
c:::::J,c::::J
� �
OLIGOCENE
-
.0.
40
'Q"
50-
BEl c:J
......
E3
30-
EOCENE
Africa western
Australia western central
! c:Jc:J
::::J c:::c: ::J � �
-
Chile
-
c:J c:J
jarosite 40Ar/39Ar jarosite alunite K-Ar Mn-oxides 40Ar/39Ar Mn-oxides K-Ar
c:J c:JE3c:::::J l=l �c:::::�
QEJ
c:Jc::: E3E3::J
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El ...... l=l c:J
Fig. 3. Diagrammatic illustration correlating weathering ages obtained by dating of alunite,jarosite and manganese oxides from weathering profiles in South America, North America, Africa and Australia. The dates show that humid climates, favourable for manganese oxide formation, prevailed at least from early Tertiary until mid-Miocene times, then development of a major desiccation event affected the deep weathering profiles and led to the formation of jarosite and alunite in the deepest parts of the profiles. Data from Alpers & Brimhall 1988; Bird et a/. 1990; Vasconcelos et a/. 1994a,b; Dammer 1 995.
tracing such unconformities, from the basin on to the continent, and provide arguments that help to decipher the sequence of incision/erosion related to sea-level change or tectonic uplift. The global record reveals a series of specific time intervals in Earth history during which deep weather ing profiles were established and have been widely preserved. It is well known that intense bauxite formation occurred during specific periods: the Carboniferous, from Mid- to Late Cretaceous until the Palaeocene-Eocene and during the Miocene and Pliocene (Bardossy & Aleva 1990; Bardossy & Combes, this volume, pp. 189-206; Valeton, this volume, pp. 153-188). Pedogenic silcretes also form under specific morphoclimatic conditions and thus may be useful for tracing stratigraphical correlations from basin towards the continent, for instance in central Australia (Simon-Coin9on et al. 1 996a) and in the Paris Basin during mid- to late Eocene time (Thiry, this volume, pp. 96-127). A pervasive mid- to late Miocene oxidation event marked by sulphide-bearing rocks is recorded
in weathering profiles in the western USA and Africa (Vasconcelos et al. 1994a), Brazil (Vasconcelos et al. 1994b ), Chile (Alpers & Brimhall 1988) and Australia (Bird et al. 1990; Dammer 1 995) (Fig. 3). Dating of supergene manganese oxides and sulphates (alunite-jarosite) allowed the progression of the oxi dation front during this chemical weathering to be timed. This weathering process indicates dry climatic conditions promoting the lowering of the water table (Anderson 1981; Alpers & Brimhall 1988; McArthur et al. 1 991 ; Thiry et al. 1995). Moreover, in Brazil, Aus tralia and western Africa, sulphide oxidation often occurs at the base of laterite profiles, indicating that the onset of the lateritization processes pre-dates the middle Miocene. Evidence from dating of K-Mn oxides from similar laterites indicates that lateritiza tion started during the late Cretaceous/early Eocene, and persisted episodically throughout the Tertiary, until late Miocene times. The dry Miocene climate was conducive to world-wide development of deep oxidation sequences. Subsequent weathering processes have not been as pervasive as the late
M. Thiry et al.
8
Miocene event, indicating a general trans1t1on to cooler, drier climates in most of these areas.
D U R AT I O N
Life span of a palaeosurface and the associated palaeoweathering profiles is another important aspect directly linked to the dating problem. It is a particularly crucial question for establishing the rates of weathering and landscape evolution processes. In aggradational situations, buried soil profiles can be treated as traditional sedimentary facies. For instance, the degree of development of palaeosol profiles or their frequent alternation with flood deposits allows a reconstruction of the 'architecture' of fluvial deposits, which reveals the evolving palaeo geography of floodplains (Cojan, this volume, pp. 323-335; Kraus and Asian, this volume, pp. 303321 ). In the meantime, the maturity of the palaeosols can be used to analyse in detail the succession of flooding events. The distinction of short-lived expo sure of the flooded land surface from more mature soils formed on unconformities allows the recon struction of sedimentary processes, such as the shifting of position of crevasse splay events, or the identification of tectonic activity and eustatic sea level changes. Accurate dating of the rocks immediately above and below a palaeoweathering profile provides an estimate of the life span of a palaeoweathering profile and, thus the precise determination of periods of palaeolandscape stability or even spatial variations in the subsidence of a basin. Such an approach has been applied successfully by studying the 81 3 C record of organic carbon from the Palaeocene-Eocene flood-plain deposits and palaeosols of the Paris Basin. The accuracy of the isotopic dating by refer ence to the deep-sea 81 3 C record allows the identification of stratigraphical gaps of about 0.5 Myr, corresponding to the development of a mature nodular calcrete horizon record (Stott et al. 1996; Thiry & Dupuis 1 998) (Fig. 4 ).
E N V I R O N MENTAL AND GLOBAL CHAN GES
Recent advances in the understanding of palaeo weathering features have been the natural con sequence of new instruments, techniques and approaches. Much research involving isotope geo-
chemistry has given precise information on palaeoen vironmental conditions. Oxygen isotopes provide invaluable information on environmental conditions during the formation of soil minerals. The systematic variations in oxygen isotope composition of kaolinites from Australian regolith profiles show that the composition of meteoric water in equilibrium with these kaolinites changed through the Tertiary, thereby allowing the drift of the continent from high to low latitudes and changes in global climate to be traced (Bird & Chivas 1989). Moreover, the stable-isotope composition reveals that much of the Australian regolith formed during much earlier geological periods (early or mid Mesozoic) under comparatively cold conditions, sug gesting that, contrary to traditional interpretations, lateritization and deep weathering phenomena are not solely the result of weathering in tropical or sub tropical climates. The possibility of deep kaolinitic profiles formed in cold climates may have import ant implications for palaeoclimatic reconstructions, which may even feed back to an understanding of present-day soil formation processes. Oxygen isotope ratios of other clay minerals, soil carbonates, iron or aluminium oxides, and even mammal teeth and bones (Koch et at. 1991 ), are commonly used for such palaeoclimatic interpretations. Modification of the isotopic composition of mete oric waters by evaporation can be highlighted by hydrogen isotopes. For example, the hydrogen iso topic composition of the alunite-group minerals at the base of the Australian regolith profiles suggests that these minerals formed during the gradual 'drying out' of the regolith profiles by evaporation after periods of active weathering (Bird et al. 1989). The stable carbon isotopes provide another powerful tool for palaeoenvironmental reconstruc tions. Numerous palaeoenvironmental interpreta tions based on the concentrations of carbon isotopes in palaeosols have been made in order to estimate palaeo-pC02 1evels of the atmosphere (Cerling 1991, and discussion in this volume, pp. 4 3-60; Mora & Driese 1993, and this volume, pp. 61-84 ; Sinha & Stott 1994 ; Mora et al. 1996; Yapp & Poths 1996). The resulting palaeoatmosphere pC02 curves show values up to 100 times the present-day value (Berner 1 994 ). In addition to its effect on climate, such high C02 contents have considerable effects on weather ing geochemistry (Schmitt, this volume, pp. 21-4 1 ) . Palaeoenvironmental interpretations o f carbon iso topes in palaeosols also include estimates of the fraction of C4 biomass in soils, which is very useful
9
Problems, progress and future research s:13
lithology
u
-26
c orgamc matter .
-24
-22
s:13
u
c orgamc matter
-26
.
-24
-22
Ma
Fausses G/aises
dark clays 54.0
mottled red and grey clays
8
mature nodular calcrete
54.5
mottled clays with carbonate granules sandy calcrete
55.0
4
beige clays with carbonate granules
55.5
colluvium of weathered Chalk
0
Cretaceous
� calcrete
�
burrows
LIMAY SECTION
PARIS BASIN REFERENCE
Fig. 4. Variation of (}13C values representing organic matter preserved in the Argiles Plastiques floodplain deposits is
characterized by a sharp negative peak followed by three positive peaks, which are bracketed between 55.5 and 53.5 Ma by reference to the deep-sea carbon isotopic record. The absence of the second peak (P2) in the Limay section is linked to the well-developed nodular calcrete. The mature calcrete most probably developed over about 0.5 Myr, pointing to a period of non-deposition in the Limay area at that time (Thiry & Dupuis 1998).
for ecosystem reconstructions (grasses, shrubs, trees, etc.) (Quade et al. 1 989; Cerling, this volume, pp. 4 3- 60; Mora & Driese, this volume, pp. 61 -84 ) . Lastly, the sedimentary environment o f pedogenesis, whether inland alluvial or coastal margin, wet-dry cycles, or depth of groundwater table, leave its mark on the isotopic composition of pedogenic carbonates and in turn can be used in palaeoenvironmental interpretations (Mora & Driese, this volume, pp. 61 -84 ).
W E AT H E R I N G P R O FI L E S P A S T A N D P R E S E NT?
Owing to the obvious relationships between ob served (functional) weathering profiles and present
morphoclimatic environments, the palaeoweathering record is naturally regarded as a straightforward testimony of climatic conditions and hence of global change. The basis for this approach is a direct com parison between palaeoprofiles preserved in the geological record and present-day weathering fea tures, with some limitations related to conditions of preservation. Some difficulties may arise for those weathering processes that have a very low alteration rate or which develop at depth, below the surficial soil hori zons commonly studied by pedologists. In such cases, one may never be certain that the profiles are still operating or have just been fossilized. This is the case for silcretes. Silcretes have long been thought to form in dry environments, by refer ence to the silcrete exposures in the dry landscapes of
10
M. Thiry et a!.
central Australia and South Africa. Recent studies have shown that numerous silcretes are not of pedo genic origin, but result from groundwater processes (Thiry & Milnes 1991 ; Thiry, this volume, pp. 87-127). Furthermore, pedogenic silcretes show numerous micromorphological features related to water perco lation through well-differentiated soil horizons and it can be proved easily that the geochemical processes at work also require significant amounts of per colating water to leach all the foreign cations, including aluminium. Only incipient development of silicification is known in present-day soils (Netter berg 1974; Hay & Wiggins 1980; Arakel et al. 1989; Blank & Fosberg 1 991; B albir Singh & Gilkes 1993; Dubroeucq & Thiry 1994). These recent silicifi cation developments cannot really be compared to the thick and massive silcretes known from older formations. Bauxite development is another case of the dif ficulties faced when comparing present-day weather ing profiles to palaeoprofiles. Lateritic bauxites have long been regarded as being formed on stable and deeply dissected plateaux, such as those on which they are exposed today in western Africa, Australia · and Brazil. This association has resulted in the concept of bauxite development being confined to well-drained environments, which has been ques tioned by recent studies. For instance, the micromor phological organization of bauxites, especially the pisolithic structures, involves a hydromorphic environment during profile development (Boulange 1984). The relief of the pre-, syn- and post-bauxitic palaeolandscapes more likely indicates tectonic instability, short periods of bauxite formation, and a low-elevation topography (Valeton, this volume, pp. 1 53-188) . These concepts that are emerging from the study of fossilized bauxite occurrences allow the establishment of a parallel between lateritic bauxites and karst bauxites. Karst bauxite also develops mainly on coastal lowlands, in tectonically active areas, where sea-level oscillations controlled the karst formation, with subsequent preservation of the bauxite beneath sedimentary deposits (Combes 1 990; Bardossy & Combes, this volume, pp. 1 89-206). The method of comparing ancient and recent mor phoclimatic environments has nevertheless proved profitable on the whole, especially for younger (Cenozoic) profiles. The same approach could, for several reasons, meet with other types of difficulties when applied to increasingly older weathering profiles. In the first place, with climate being the complex result of many factors (astronomical vari-
abies, total rainfall and yearly distribution of precipi tation, wind pattern, vegetation cover, exposure and altitude, distribution of land masses and so on), there is little doubt that presently known climates repre sent only a very restricted range of all past terrestrial climatic conditions, and this is especially true for the more ancient climates. We have of course no present day equivalent of a weathering profile corresponding to these extinct conditions and will have trouble in deciphering them. Weathering, being the effect of the exposure of fresh rocks to atmospheric agents, is also strongly dependent on rainwater chemistry and hence on atmospheric composition, as shown in the paper by Schmitt (this volume, pp. 21-41 ). Many recent studies have shown that several critical components of the atmosphere, such as oxygen and carbon dioxide con tents, have varied significantly during Earth's history. These variations, and possibly the varying abundance of other gases, could be responsible for dramatic changes in the intensity and even in the very nature of weathering processes. The very unusual iron depleted pyrite-bearing Huronian palaeoprofiles of Canada (Roscoe 1 969; Prasad & Roscoe 1996) are probably the most dramatic example of this atmos pheric imprint on weathering profiles, as discussed below (Schmitt, this volume, pp. 21-41) . ODD AND UNCOMMON P A L A E O P R O FI L E S
Apart from peculiar climatic o r atmospheric con texts, groundwater of 'exotic' composition may lead to profiles with unusual or misleading characteristics. In southern France, many Triassic palaeoprofiles exhibit albitization phenomena that have passed unnoticed or have been ascribed to a later diage netic/metamorphic event. A careful study (Schmitt 1992) has shown that low-temperature neomorphic albite had actually formed in the lower pink-coloured horizon of these profiles, even when their upper argillized and ferruginized horizons are apparently of normal appearance and composition. Systematic mapping has proved that these albitized profiles were exclusively located in the low-lying coastal areas of the Triassic landscape, strongly suggesting that saline waters could have played a leading role in their for mation. Geochemical modelling and simulation of this unusual weathering by NaCl-enriched ground water has confirmed the validity of this interpretation (Schmitt 1 994) .
11
Problems, progress and future research
chemical modelling has confirmed that these profiles most probably formed during a single weathering phase from the action of acidic sulphate-rich phreatic waters (Thiry et a/. 1995). There are no present-day equivalents to these two examples of palaeoprofiles, either because the weathering action of such unusual ground waters has remained unnoticed, or more probably because they are linked to very peculiar continental palaeoen vironments. Once buried under younger sediments, the preser vation of the weathering profiles is still not guaran teed, and significant changes may alter their original character. In the stratigraphical column, weathered horizons always represent a contrasting-and thus generally unstable-geochemical feature. Diagenesis will tend to slowly re-equilibrate the local (rock pile) system. In the process it will alter the initial mineral ogy and geochemistry of the weathered horizons, although their major structure may be retained. This
In inland Australia, deep weathering profiles linked to the Tertiary palaeosurface were established on a variety of parent rocks, especially dark-coloured Cretaceous shales. Typical profiles (Thiry et al. 1 995) show a thick horizon of bleached shales composed of relict quartz, kaolinite and opal. Several types of silicified facies are present and the profile is topped with a pedogenic columnar silcrete (Fig. 5). Gypsum is common in all profiles, and alunite [KAI 3 (S04 )z(OH) 6] is found throughout the bleached horizon as centimetre-size nodules and larger lenses. The coexistence of mineral phases of apparently opposite geochemical signature led early authors to believe that these profiles were the result of two successive weathering phases: an initial humid phase of kaolinization; and a later arid phase during which silicification and sulphates may have formed. Recent data and various field observations have shown, however, that the hypothesis of two succes sive weathering phases was at least unlikely, and geo-
LITHOLOGY
pedogenic silcrete
WHOLE ROCK
CLAY MINERALS
lithology
whole rock
� termite burrow � gypsum in joint
[MJ CJ CJ � lmii!J
ground-water silcrete bleached Cretaceous
alunite nodules
joint with precious opal Cretaceous shales
clay minerals
CJ illite 1±:::±] kaolinite
calcite quartz clay minerals opaiA opal CT
E72J smectite
minerals (log moles) 0
quartz
-1
2
-3
-
4
::-:-:-,
1
---� 1muscovite....- ..'
I
I ..
/
'
:1 r1
� �e !---'\. \...__ ••
i""'--...
/
nontronite
,
-
'
kaolinite I I 1
gypsum
--
'
-
" alunite
haematite
�k �� feld� � ,o � ar ______________________________�
yrite
���������;;���h172���??r-��
mineralogical composition(%)
100
simulation progress � Fig. 5. Lithology and mineralogy of a typical bleached profile of inland Australia. Groundwater silicification penetrates into the bleached formations along deep vertical cracks. Alunite occurs in nodules at the base of the section, where precious opal veins have been mined. Gypsum has not been plotted on the mineralogical log. The simulation of the flushing of the Cretaceous shale by sulphate-rich brine (similar to those of saline lakes) shows a mineralogical assemblage similar to that of the bleached profile. Opal does not appear in the simulation owing to problems related to the introduction of amorphous silica solubility and/or precipitation kinetics into the model. Thus, the quartz increase seen during the simulation has to be transformed into opal.
12
M. Thiry et al.
is of course especially sensitive in the case of the older profiles, such as the Huronian profiles discussed by Prasad & Roscoe (1996), Gall ( 1994, and this volume, pp. 207-221) and Schmitt (this volume, pp. 21-41) . The detailed study o f unconformity related uranium deposits in Saskatchewan (Prasad & Roscoe 1996; Fedo et al. 1997) has also shown that in some cases circulation of hydrothermal fluids may erase the initial mineralogy and zonation of palaeoweath ering profiles more or less completely. High-grade metamorphism ultimately will transform both the mineralogy and structure of the weathered material totally, turning, for instance, bauxitic profiles into emery (corundum-bearing) rocks.
R O L E O F P R E - E XI S TI N G L A N D S C A P E S A N D E X H U M ATI O N S
Palaeolandscapes frequently are polycyclic and display features of various ages corresponding to combinations of successive changes in tectonic and climatic conditions. Whatever the weathering processes may be, they act on an existing framework developed during the previous former stages of land scape evolution. Of noteworthy importance in this respect are the nature of the substratum, the presence of abundant pre-existing weathered material, and the existence of tectonic and morphostructural frameworks with well-developed surfaces and specific landforms. The nature of the parent rock commonly is consid ered to be a major factor in landscape evolution: whereas granitic areas evolve to glacis landscapes and inselbergs, limestone platforms may give rise to wide karstic landscapes. Weathered material and sed iments will of course greatly differ in each case. The presence of a thick and widespread mantle of pre existing alterites will have almost the same effect as a specific parent rock. In western Europe, as well as in North America, the huge kaolinitic mantle inherited from the humid Cretaceous climatic phase has strongly influenced the shaping of subsequent land scapes and affected sedimentation during the whole Tertiary (Sigleo & Reinhardt 1 988; Blanc-Valleron & Thiry 1 997) . Well-developed pre-existing palaeosurfaces, espe cially when established on resistant rocks, will remain as important features of subsequent landscapes. Throughout the Tertiary the palaeochannels of Aus tralia that were initiated during the earliest Creta-
ceous were being infilled with sediments between periods of weathering, which led to the development of ferricretes, calcretes and silcretes (Alley et al., this volume, pp. 337-366). The earliest incision steered the Tertiary morphodynamic evolution of large parts of Australia. In the south-west Massif Central (France) ancient surfaces (infra-Liassic or Cenomanian) were partly exhumed during the Tertiary and largely preserved from weathering as a result of increasing aridity (Simon-Coin�on et al. 1996b; Simon-Coin�on, this volume, pp. 225-243) . Elsewhere, removal of the deep Cretaceous weathering cover has exposed the undulating etched surface of the crystalline uplands, approximately reproducing the ancient topography. Smaller landforms also may be long-lived features. Present landscapes of the south-western Massif Central likewise still exhibit many remnants of a long and complex continental history that traces back to the early Tertiary and beyond. For example, along the Lot Valley various remnants of palaeolandscapes are combined into the present-day landscapes (Fig. 6). The highest basement outcrops show remnants of the Triassic albitization (Simon-Coin�on this volume, pp. 225-243), flanks of the Lot Valley uncover sections of the Jurassic dolomitized palaeosurface, and else where kaolinitic palaeoweathering profiles and deposits mark out the Tertiary palaeosurface on basement and Jurassic limestones. Finally, Plio Quaternary lava flows fossilize the former exhumed palaeosurface patchwork. Likewise, relief features and saprolite remnants on Scandinavian basement areas show that surfaces that have been exposed for long periods can largely retain the original morphology but the associated saprolites might have gone (Lidmar-Bergstri:im et al., this volume, pp. 275-301 ) . Thus saprolites of wide!y differ ent ages can occur on ancient surfaces that have been exposed for a long period.
M I S S I N G I N F O R M AT I O N
Palaeoweathering profiles are indeed a significant part of the geological record and a direct witness of the palaeoenvironments that prevailed on conti nents. Clay minerals are indirect tracers of the palaeoenvironments and are linked to the pal aeoweathering history. In the geological record, however, numerous components are systematically missing, others are side-stepped, and the picture pro-
13
Problems, progress and future research s
N
m asl
Lot Riv.
+ +
+
+ + paleosu rfaces
& paleoweathering
� Triassic � Lower J u rassic � U pper Cretaceous-Tertiary
1 000
stratigraphy
D lo 0 ol
E:±J
;-
+
+ +
;-
+
+ +
+
+
+ +
basement
Trias sandstone Permian sandstone
+
+
+ +
+
+
+ +
500
+
+
+
+
+ +
+
+ � + + 1 km + �
�
+
Ill Q uaternary lava flow D Tertiary sandstone rn J u rassic limestone
Fig. 6. Schematic section showing the patchwork of palaeosurfaces integrated in the present-day landscape of the south
western border of the Massif Central. Large parts of the present-day landscape are controlled by exhumed palaeosurfaces.
vided in reconstructions is distorted and incomplete. We have to be aware of these difficulties. The various components of the palaeorelief are always fossilized unequally. Elevated parts of the landscape are seldom preserved, except in the case of tectonic or karst collapse. In contrast, lowland areas near the edge of the basin are buried by sedimentary deposits very frequently. The same goes for pal aeoweathering profiles. Profiles developed on the uplands are always eroded and reworked, only weathering profiles developed in the lowlands will be recorded. Most frequently, only the deepest horizons of the palaeoprofiles are preserved, the upper ones having been stripped off. The general lack of pedogenic and biological structures does not allow us to link most palaeoprofiles to modern soil taxonomies and thus to precise environments. Moreover, only thick and highly differentiated soils are recorded, thin soil profiles, such as those developing under temperate climates, are rarely, if ever, preserved beneath sedi mentary deposits. The hardness of the weathering products is a main factor in the differential preservation of palaeo profiles. Duricrusts, like ferricretes, silcretes and cal cretes are described very often from continental series. The soft weathering products are missing and/or not recognized. Numerous series are related to alternating dry and wet climates, with widespread formation of smectites, obviously resulting from development of vertisols. Such palaeosols, however, or palaeosols with slickensided structures, are seldom described. They are easily reworked and/or specific
pedogenic structures are destroyed by burial compaction. Soluble compounds are generally lacking in the palaeoprofiles. In modern soils, most of them are present in the surficial horizons and so they are stripped off with the uppermost horizons during sedi mentary or erosional processes. They are also washed out during sedimentation or dissolved during the course of diagenesis. This is the case of halite and gypsum impregnations in soils from arid environ ments. These soluble compounds generally accumu late in ancient soils and in lowland areas. Thus we will systematically miss the latest evolutionary stage of this kind of environmental evolution. Relicts of old weathering surfaces generally exhibit a long polyphasic history associated with geochemi cal re-equilibration of the minerals. Numerous examples can be quoted; we may emphasize the long history of the successive palaeosurfaces or 'glacis' of western Africa (Michel 1978; Boulange 1984) and eastern Africa (Schwarz & German, this volume, pp. 367-390). Much of the primary bauxite minerals of the highest (oldest) surfaces have been transformed into kaolinite during subsequent processes of resili cification. Caution has to be used in any palaeocli matic interpretation based upon indicator minerals, in order to check whether palaeoweathering pro files have not been deeply affected by successive alterations. Very often continental sedimentary successions show a general geochemical and palaeoenviron mental sequence related to palaeoclimatic evolution from wet to dry. This may be an incomplete sequence.
14
M. Thiry et a!.
Under dry conditions, weathering processes are less aggressive than under wet conditions. Thus, evolution to a dryer and more arid climate preserves for a long time the weathering products formed during wet periods. In contrast, evolution to wetter climates will destroy the former weathering products and mark the soils with its own fingerprint. Little evidence will be then preserved of an evolution toward wetter climates. The geological record clearly favours the preservation of mineralogical and pedological sequences indicating the evolution of a wet to dry palaeoenvironment. These are major distortions that affect the view of past continental environments indicated in the palaeoweathering horizon of the geological record. Several other distortions exist. They explain why kaolinitic palaeoweathering profiles of a warm and wet climate are recorded so frequently in continental deposits and why hydromorphic soils are described so often. This does not mean that the study of palaeo weathering is of little help to palaeoenvironmental reconstruction, but the missing members of the sequence always have to be kept in mind and sought for more carefully.
F U T U R E D E V E L O P M E N T S IN P A L A E O W E AT H E RI N G S T U D I E S
Isotope geochemistry will most probably b e a major contributor to palaeoweathering studies in the future. With further instrumental progress, these techniques will be applied to a larger range of miner als and so will allow the investigation of very differ ent types of palaeoweathering phenomena and hence a large variety of palaeoenvironments of all ages. Dating will be of prime importance for palaeo geographical and palaeoclimatic reconstructions. The more precise the dating, the more accurate will be the environmental reconstructions (palaeoatmosphere composition, palaeotemperature, etc.). Quantitative records of global change will be the challenge for palaeoweathering studies in the next decade. Geochemical modelling will probably make a major contribution to palaeoweathering studies in the future. As more accurate data are obtained on the palaeoenvironments, geochemical modelling will enable the simulation of palaeoweathering processes under various atmospheric compositions, tempera ture, rainfall regimes, etc. and hence most probably will be able to discriminate successive weathering processes superimposed on old regoliths.
Global circulation models and palaeoclimate mod elling are outstanding techniques (Barron et al. 1 993; Pollard & Schultz 1995). At present such models may be too broad to be used to explain and understand specific palaeoweathering features. With more accu rate models in the future, climatic reconstructions may be useful because they will be able to provide invaluable information on wind regimes, rainfall peri odicity, effect of altitudes, etc. (Moore et al. 1992). Palaeosurface modelling is another emerging discipline. Computers now allow new and rapid appraisal of spatial data obtained from remote sensing or GIS (McAlister & Smith 1 997; Ringrose & Migon 1997; White et al. 1 997). By modelling succes sive loading and unloading of the palaeosurfaces, endogenic phenomena, such as tectonics and crustal studies, may be linked to exogenic processes (Wyns 1991a,b; Kooi & Beaumont 1994; Tucker & Slinger land 1994; Quesnel 1997; Widdowson & Gunnell, this volume, pp. 245-274). These advances in understand ing of the relationships between geomorphological processes, the sedimentary record and tectonics will provide tools for evaluating models of landscape evolution. Clearly, sophisticated instruments, techniques and modelling will lead to major progress in palaeo weathering studies. It is hoped, however, that this 'high-tech' approach will not become divorced from the more general techniques, including field mapping and petrography. Although perhaps obvious, a multi scaled and interdisciplinary evaluation of specific palaeosurfaces may improve our overall understand ing of their origin (Alley et al., this volume, pp. 337-366; Cojan, this volume, pp. 323-335; Calvo et al., this volume, pp. 129-151; Mora & Driese, this volume, pp. 61-84; Simon-Coinc;on, this volume, pp. 225-243; Widdowson & Gunnell, this volume, pp. 245-274). Defining the processes that have created the major palaeosurfaces is not a simple matter. In addition to the careful geochemical and petrographical studies necessary for the improved understanding of palaeo surfaces and palaeoweathering, however, we suggest that the more geological aspects of these features (including basin history, tectonics, eustatic sea-level changes, etc.) also be pursued.
ACKNOWLED GE MENTS
The core o f this volume i s a contribution t o IGCP 317 'Palaeoweathering Records and Palaeosurfaces'. Several meetings and field trips that were the occa-
Problems, progress and future research sion of numerous official presentations and informal discussions, from which emerged a number of ideas leading to papers in this volume, have marked this international UNESCO-lUGS programme. We are indebted to all the people who made a great job orga nizing the meetings and trips and would especially acknowledge herein: Patricia Zalba, Renato Andreis and Mario Iniguez for the 1992 meeting in La Plata (Argentina) on 'Landscape Reconstruction' and the field trips to Sierras de Tandilia and Patagonia; Tony Milnes and Malcom Wright for the 1993 meeting in Adelaide on 'Clay Minerals for Palaeolandscape Reconstructions' and the marvellous camping field trip to show soils, landscapes and palaeoweathering features of inland Australia; Janet Springer, Quentin Gall, Nirankar Prasad, James A. Robertson, Gerry Bennett and Mireille Bouchard for another 1993 meeting held in Hamilton on 'Palaeoweathering and Palaeolandforms' and field trips on the Precambrian palaeoweathering profiles in Canada; Maria de Grac;:a de V.X. Ferreira and Jannes M. Mabesoone for organizing a 1994 meeting in Recife and the field trip to palaeolandscapes and palaeoweathering profiles in north-east Brazil; Mike Widdowson for organizing a 1 994 workshop-style meeting at Sheffield (UK) on ' Tertiary and Pre- Tertiary Palaeosurfaces ' and B.J. Smith and P.A. Warke who organized another meeting with field trips in 1994 in Belfast (Northern Ireland) on 'Laterites, Palaeoweathering and Palaeo surfaces'. To complete the information about meet ings of IGCP 317, two other meetings and field trips have to be mentioned that were held in France in 1991, with a workshop on 'Surficial Silicifications ' and on 'Mineralogical and Geochemical Records of Palaeosurfaces', followed by a field trip in the Paris Basin, and in 1995, with a meeting on 'Stability and Evolution of Continental Palaeolandscapes' and a field trip 'From Inland Palaeosurfaces towards Sedi mentary Basins ' in the south-western French Massif Central and the nearby Aquitaine B asin. The editors thank Michael Talbot for offering us the opportunity to prepare this volume and in addi tion for editing the English language of the introduc tory paper.
REFERENCES ALPERS, C.N. & BRIMHALL, G.H. ( 1988 ) Middle Miocene climatic change in Atacama Desert, northern Chile; evi dence from supergene mineralization at La Esconbida. Geol. Soc. Am. Bull. , 100, 1640-1656.
15
ANDERSON, J.A. ( 1981 ) Characteristics of leached capping and techniques of appraisal. In: Advances in Geology of the Porphyry Copper Deposits, Southwestern North America (Ed. Titley, S.R.), pp. 245-287. University of Arizona Press. ARAKEL, A.V. , JACOBSON, G., SALEHI, M. & HILL, C.M. ( 1989 ) Silicification of calcrete in paleodrainage basins of the Australian arid zone. Aust. J Earth Sci. , 36, 73-89. BALBIR SINGH, & GILKES, R.J. ( 1993 ) The recognition of amorphous silica in indurated soil profiles. Clay Minerals, 28, 461--474. BARDOSSY, G. & A LEVA, G.J.J. ( 1990) Lateritic Bauxites. Else vier Science Publishers, Amsterdam. BARRON, E.J., FAWCEIT, P.J., POLLARD, D. & THOMPSON, S.L. ( 1993) Model simulation of Cretaceous climates.The role of geography and carbon dioxide. Philos. Trans. R. Soc. London, Ser B, 341, 307-315. BERNER, R.A. ( 1994 ) 3GEOCARB-II -a revised model of atmospheric C02 over Phanerozoic time. Am. J. Sci. , 294, 56-91. BIRD, M.I. & CHIVAS, A.R. ( 1988 ) Oxygen isotope dating of the Australia regolith. Nature, 331 (6156 ) , 513-516. BIRD, M.I. & CHIVAS, A.R. ( 1989 ) Stable-isotope geochronology of the Australian regolith. Geochim. Cos mochim. Acta, 53, 3239-3256. BIRD, M.I. & CHIVAS, A.R. ( 1993 ) Geomorphic and palaeo climatic implications of an oxygen-isotope chronology for Australian deeply weathered profiles. Aust. J Earth Sci. , 40, 345-358. BIRD, M.L, ANDREw , A.S. , CHrvAs , A. R . & LocK, D.E. ( 1989 ) An isotopic study of surficial alunite in Australia: 1 . Hydrogen and sulphur isotopes. Geochim. Cosmochim. Acta, 53, 3223-3237. BIRD, M.I., CHIVAS, A.R. & Mc DouGLAS, I. ( 1990 ) An iso topic study of surficial alunite in Australia. Chem. Geol. , 80, 133-145. BLANC-VALLERON, M.M. & THIRY, M. ( 1997 ) Clay minerals, paleoweathering, paleolandscapes and climatic sequences: the paleogene continental deposits in France. In: Soils and Sediments: Mineralogy and Geochemistry (Eds Paquet, H. & Clauer, N.), pp. 223-247. Springer Verlag. BLANK, R.R. & FOSBERG, M.A. ( 1991 ) Duripans of the Owyhee Plateau region of Idaho: genesis of opal and sepi olite. Soil Sci. , 152/2, 1 1 6-133. BouLANGE, B. ( 1984 ) Les formations bauxitiques lateritiques de Cote d'Ivoire. Travaux et Documents de l'ORSOM, Paris, 175, 3 1 1 , pp. CERLING, T.E. ( 1991 ) Carbon dioxide in the atmosphere: evi dence from Cenozoic and Mesozoic paleosols. Am. J Sci. , 291, 377--400. CERLING,T.E., WANG, Y. & QuADE, J. ( 1993 ) Expansion of C4 ecosystems as an indicator of global ecological change in the late Miocene. Nature, 361, 344-345. CoMBES, P.J. ( 1990 ) Typologie, cadre geodynamique et genese des bauxites fran<;aises. Geodynam. Acta, Paris, 4, 91-109. DAMMER, D. ( 1995 ) Geochronology of chemical weathering processes in the northern and western Australian regolith. PhD thesis, Australian National University, Canberra. 214 pp. DELANOUE, J. ( 1854 ) Des sources sulfurees et des eaux ordi naires. Bull. Soc geol. France, 2 ( 1 1 ) , 569-574.
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DELESSE (1853) Sur Ia transformation du granite en ar(me et en kaolin. Bull. Soc geol. France, 2(10), 256-266. DIEULAFAIT, L. (1885) Origine et mode de formation des phosphates de chaux en amas dans les terrains sedimen taires. Leur liaison avec les minerais de fer et les argiles des terrains siderolithiques. Ann. Chim. Phys. , 6(5), 204-240 . DUBROEUCQ, D. & THIRY, M. ( 1994) Indurations siliceuses dans des sols volcaniques. Comparaison avec des silcretes anciens. 15th World Congress of Soil Sciences, July 10-16, Acapulco, Mexico, Transactions, 6a, 445-459. EBELMEN (1851) Alteration des roches stratifiees sous !'influence des agents atmospheriques et des eaux d'infiltration. C.R. Acad. Sci. Paris, 33, 678-682. FEDo, C.M., YouNG, G.M., NESBITI, H.W. & HANCHAR, J.M. (1997) Potassic and sodic metasomatism in the Southern Province of the Canadian Shield: evidence from the Pal eoproterozoic Serpent Formation, Huronian Supergroup, Canada. Precam. Res. , 84(1-2), 17-36. FLEURY, E. (1909) Le Siderolithique Suisse. Contribution a Ia connaissance des pheomenes d'alteration superficielle des sediments. Mem. Soc. fribourgeoise Sci. Nat., 6, 260, pp. FRIEDEL, C. (1876) Sur !'alteration des agates et des silex. Ann. Chim. Phys. , 5(7), 540-546. GALL, Q. (1994) Proterozoic paleoweathering and diagene sis, The/on Basin, Northwest Territories, Canada. PhD Thesis, Carleton University, Ottawa, Ontario. HAY, R.L. & WIGGINS, B. (1980) Pellets, ooids, sepilolite and silica in three calcretes of the southwestern United States. Sedimentology, 27, 559-576. KocH, P.L., ZACHOS, J.C. & GINGERICH, P.D. (1991) Paleo cene-Eocene climates in Western North America: infer ences from oxygen isotope analysis of mammal teeth and soil carbonates. Geol. Soc. Am. Abstr. Program, 23(5), 301. KoOJ, H. & BEAUMONT, C. (1994) Escarpment evolution on high-elevation rifted margins: insights derived from a surface processes model that combines diffusion, advec tion, and reaction. J geophys. Res. , 99/86, 12191-12209. Mc ALISTER, J.J. & SMITH, B.J. (1997) Geochemical trends in Early Tertiary palaeosols from northeast Ireland: a statis tical approach to assess element behaviour during weath ering. In. Palaeosurfaces: Recognition, Reconstruction and Palaeoenvironmental Interpretation (Ed. Widdowson, M.), Spec. Pub!. Geol. Soz. London, No. 120, 57-65. Geologinal Society of London, Bath. McARTHUR, J.M., TuRNER, J.V., LYONS, W.B., OssoRN,A.O. & THIRLWALL, M.F. (1991) Hydrochemistry on the Yilgarn Block, Western Australia: ferrolysis and mineralisation in acidic brines. Geochim. Cosmochim. Acta, 55, 12731288. MICHEL, P. (1978) Cuirasses bauxitiques et ferrugineuses d'Afrique Occidentale. Apen;u chronologique. In: Geo morphologie des reliefs cuirasses dans les pays tropicaux chauds et humides Travaux Doc Geogr. Trap. (CEGET, Talence), 33, 9-32. MooRE, G.T., SLOAN, L.C., HAYASHIDA, D.N. & UMRIGAR, N.P. (1992) Paleoclimate of the Kimmeridgian/Tithonian (Late Jurassic) world: II. Sensitivity tests comparing three different paleotopographic settings. Palaeogeogr. Palaeoclimatol. Palaeoecol. , 95, 229-252. MoRA, C.l. & DRIESE, S.G. (1993) A steep, mid- to late Paleozoic decline in atmospheric C02; evidence from
the soil carbonate C02 paleobarometer. Chern. Geol. , 107 217 219 . MoRA, C.I., DRIESE, S.G. & CoLARusso, L.A. (1996) Middle to late Paleozoic atmospheric C02 levels from soil car bonate and organic matter. Science, 271, 1105-1 107. NETIERBERG, F. ( 1974) Calcretes and silcretes at Sambio, Okavangoland, South West Africa. S. Afr. archaeol. Bull. , 29, 83-88. PoLLARD, D. & ScHULTZ, M. (1995) A model for the potential locations of Triassic evaporite basins driven by paleocli matic GCM simulations. Global planet. Change, 9, 233-349. PRASAD, N. & RoscoE, J.A. ( 1996) Evidence of anoxic to oxic atmospheric change during 2.45-2.22 Ga · from lower and upper sub-Huronian paleosols, Canada. Catena, 27, 105-121. QUADE, J., CERLING, T.E. & BOWMAN, J.R. (1989) Develop ment of the Asian monsoon revealed by marked ecologi cal shift during the latest Miocene in northern Pakistan. Nature, 342, 163-166. QuESNEL, F. (1997) Cartographie numerique en geologie de surface. Application aux alterites a silex de l'Ou.e st du Bassin de Paris. These, University of Rouen., 256 pp. RING ROSE, P.S. & MIGON, P. (1997) Analysis of digital eleva tion data for the Scottish Highlands and recognition of pre-Quaternary elevated surfaces. ln. Palaeosurfaces: Recognition, Reconstruction and Palaeoenvironmental Interpretation (Ed. Widdowson, M.), Spec. Pub!. geol. Soc. London, No. 120, 25-35. Geological Society of London, Bath. RoscoE, S.M. (1969) Huronian rocks and uraniferous con glomerates. Can. Geol. Surv. Pap. , 68-40, 205 pp. ScHMITI, J.M. (1992) Triassic albitization in southern France: an unusual mineralogical record from a major continental paleosurface. In: Mineralogical and Geochemical Records of Paleo weathering, IGCP 317 (Eds Schmitt, J.M. & Gall, Q.). E.NS.M.P Mem. Sci. Terre, 18, 1 1 5-132. ScHMITI, J.M. (1994) Geochemical modeling and origin of the Triassic albitized regolith in Southern France. 14th International Sedimentology Congress 20-26 Aug. Recife, Brazil, Abstracts SS, 19-21. SJGLEO, W.R. & REINHARDT, J. (1988) Paleosols from some Cretaceous environments in the southeastern United States. In. Paleosols and Weathering Through Geologic Time (Eds Reinhardt, J. & Sigleo, W.). Geol. Soc. Am. Spec. Pap., 216, 123-142. SJMON-COINc;:ON, R., MILNES,A.R., THIRY, M. & WRIGHT, M.J. (1996a) Evolution of landscapes in northern South Australia in relation to the distribution and formation of silcretes. J geol. Soc. London, 153, 467-480. SIMON-COINc;:oN, R., THIRY, M. & ScHMITI, J.M. (1996b) Variety and relationships of weathering features along the early Tertiary palaeosurface in the south-western French Massif Central and the nearby Aquitaine Basin. Palaeogeogr. Palaeoclimat. Palaeoecol. , 129, 5 1-79. SINHA, A. & STOTI, L.D. (1994) New atmospheric p C0 2 esti mates from paleosols during the late Paleocene/early Eocene global warming interval. Global planet. Change, 9, 297-307. SiNHA, A., AUBRY, M.P., STOTI, L.D., THIRY, M. & BERGGREN, W.A. (1995) Chemostratigraphy of the 'lower' Sparna cian deposits (Argiles Plastiques bariolees) of the Paris B asin. Isr. 1. Earth Sci. , 44, 223-237. ,
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Problems, progress and future research STOTT, L.D., SINHA, A., ThiRY, M., AUBRY, M.P. & BERGGREN, W.A. ( r 996) Global o13C changes across the Paleocene/ Eocene boundary: criteria for terrestrial-marine correla tions. In: Correlation of the Early Paleogene in Northwest Europe (Eds Knox, R.W.O'B., Corfield, R.M. & Dunay, R.E.) Spec. Pub!. geol. Soc. London, No. 101, pp. 381-399. Geological Society of London, Bath. ThiRY, M. & DuPUIS, C. (1998) Chemostratigraphy of the Sparnacian deposits and correlation to deep-sea record. In: Paleocene/Eocene Boundary in Paris Basin: the Sparnacian Deposits Field Trip Guide (Eds Thiry, M. & Dupuis, C.). E.N.S.M.P Mem. Sci. Terre, 34,39-47. ThiRY, M. & MILNEs, A.R. (1991) Pedogenic and groundwa ter silcretes at Stuart Creek Opal Field, South Australia. J sediment. Petro/. , 61, 1 1 1-127. ThiRY, M., SCHMITT, J.M., RAYOT, V. & MILNES, A. (1995) Geochimie des alterations des profils blanchis du regolithe tertiaire de l'interieur de I' Australie. C.R. A cad. Sci. Paris, 320(ser. Ila), 279-285. TucKER, G.E. & SLINGERLAND, R.L. (1994) Erosional dyn amics, flexural isostasy, and long-lived escarpments: a numerical modeling study. J. geophys. Res. , 99/86 , 12 224-12243. VAN DEN BROECK, E. (1881) Memoire sur les phenomenes d'alteration des eaux meteoriques etudies dans leurs rapports avec Ia geologie stratigraphique. Mem. Acad. Sci. Lett. Bx-Arts Belgique, 44, 180 pp.
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VASCONCELOS, P.M., BRIMHALL, G.H., BECKER, T.A. & RENNE, P.R. (1994a) 40Arf39Ar analysis of supergene jarosite and alunite: implications to paleoweathering history of western USA and West Africa. Geochim. Cosmochim. Acta, 58, 401-420. VASCONCELOS, P.M., RENNE, P.R., BRIMHALL, G.H. & BECKER, T.A. (1994b) Direct dating of weathering phenomena by 40A/39Ar and K-Ar analysis of supergene K-Mn oxides. Geochim. Cosmochim. Acta, 58, 1635-1665. WHITE, K., DRAKE, N. & WALDEN, J. (1997) Remote sensing for mapping palaeosurfaces on the basis of surficial chemistry: a mixed pixel approach. In: Palaeosurfaces: Recognition, Reconstruction and Palaeoenvironmental Interpretation (Ed.Widdowson, M.), Spec. Pub!. geol. Soc. London, No. 120, 283-293. Geological Society of London, Bath. WYNS, R. (1991a) Evolution tectonique du bati armoricain oriental au Cenozo!que d'apres !'analyse des paleosur faces continentales et des formations geologiques asso ciees. Ceo!. France, 3, 1 1-42. WYNS, R. (1991b) L'utilisation des paleosurfaces conti nentales en cartographie thematique probabiliste. Geol. France, 3, 3-9. YAPP, C.J. & PoTHS, H. (1996) Carbon isotopes in continental weathering environments and variations in ancient atmospheric C02 pressure. Earth planet. Sci. Lett. , 137, 71-82.
G eochemistry and isotop e s
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
Spec. Pubis int. Ass. Sediment. (1999) 27, 21-41
Weathering, rainwater and atmosphere chemistry: example and modelling of granite weathering in present conditions in a C02-rich, and in an anoxic palaeoatmosphere
J . - M . S CH M ITT C. I. G., Ecole Nationale Superieure des Mines de Paris, 3 5 rue St-Honore, 77305 Fontainebleau Cedex, France; and IGCP 317 Palaeoweathering records andpalaeosurfaces and CNRS- UMR SISYPHE C 7619 Structure et fonctionnement des systemes hydriques continentaux
A B S T R AC T Chemical weathering refers t o the chemical and mineral transformations induced by rainwater percola tion through rocks at Earth's surface. The main types of weathering reactions are oxidation and carbona tion, which are caused by the oxygen and carbon dioxide dissolved in rainwater. The fugacities of these two gases in rainwater are imposed by their abundance in the atmosphere, which in this way governs the chem istry and intensity of weathering. There is very strong evidence, however, that atmospheric abundance of both 02 and C02 have varied significantly during Earth's history. In this paper two striking examples are provided of palaeoprofiles formed, respectively, in a C02-rich atmosphere and in an anoxic atmosphere, and the conditions of their development are discussed with the aid of a geochemical modelling code. Recent data and models of the variation of atmospheric C02 show that the Middle Cretaceous corresponds to a well-marked high, with values approaching 10 times that observed at present. Modelling shows that such C02 abundance highly accelerates the formation of deep kaolinitic profiles, without the need for weathering duration as long, nor climate as hot and humid, as thought previously. Palaeoweathering profiles fossilized below the lower Huronian strata in Canada are characterized by an unusual depletion in total iron, and the preservation of pyrite. These characters are considered as strong indicators of reducing atmospheric conditions c. 2.4 Ga. Geochemical modelling enables us to reproduce the major features of these profiles for 02 and C02 partial pressures compatible with current estimates for this period. These results also are used to discuss the anoxic character as well as the potential H2S content of the Huronian atmosphere. The main conclusion that can be drawn from this review and from these examples is that palaeoweath ering profiles have registered the fluctuations in the atmosphere's chemistry and climatic conditions, and that consequently they are one of the main indicators of Earth's past environments.
I NT R O D U C T I O N
The characteristics and rapidity of rock weathering depend on many factors, including the nature and properties of the parent rock, the local drainage conditions, and the geomorphological and climatic context. These many factors naturally result in a great variety of weathering profiles around the world, and many types of classification (climatic, chemical, genetic or morphogenetic) have been proposed to conveniently describe their major features.
Under similar settings and on a given rock type, the nature of observed profiles can be considered as characteristic of the climatic zone. For instance (Fig. la) the weathering of granitic rocks today results in the development of lateritic profiles under equatorial and humid tropical climates, of lithosolic profiles under desert conditions, and of podzolic or isohumic soils and profiles in the temperate zone.
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
21
22
1. M. Schmitt Table 1. Partial pressures of main atmospheric gases (After
Hem, 1 985)
Gas
Partial pressure ( atm)
N2 02 H20 CO, CH� CO S02 Nz O
0.78 0.21 0.001-0.23 0.0003 1 .5 X 10-6 (0.06-1.0) X 10-6 1.0 x 1 0-6 5 x 10-7
pared most often to present-day features in order to infer the type of climate under which they formed. Although reasonable, this actualistic approach may meet with serious difficulties, particularly when dealing with pre-Quaternary features (Fig. lb ) Two main cases can be highlighted: (i) some profile types are apparently ubiquitous at certain epochs, regard less of their latitudinal setting; and (ii) some others appear to have no present-day equivalent. In this paper we will try to review the main geo chemical aspects of weathering, and to highlight the intricate links between weathering, rainwater and atmosphere. Geochemical modelling will be used to discuss the unusual features and the origin of two noteworthy types of palaeoweathering profiles. These results will provide an illustration of the utility of weathering in constraining Earth's past atmospheric conditions, and will show how weather ing itself plays a crucial part in regulating these conditions. .
4
R O C K W E ATH E R I N G : A G E O C H E M I C A L A P P R OA C H
Fig. I. Present and past distribution of weathering profile
types around the world. The present distribution (a) reflects mainly the climatic zonation, with for instance: lateritic profiles (1) in the humid tropics; lithosolic profiles (2) in desert zones; podzolic or isohumic soils (3) in the temperate zone. In the geological record (b), some profile types ( 4) appear ubiquitous for certain epochs, and some others (5) are without present-day equivalents.
When we study palaeoweathering profiles pre served in the geological record, we are inclined to interpret them in the same way, and palaeoprofiles as well as palaeoweathered material have been com-
Origin and nature of weathering
Weathering commonly refers to the different (and often destructive) processes affecting rocks exposed to atmospheric agents. In its more restrictive sense, or what could be referred to as chemical weathering, it is the process of chemical and mineral transformation induced by rainwater percolation through rocks near Earth's surface. Rainwater chemistry is characterized mainly by its 02 and C02 fugacities, which are imposed by the abundance of these two gases in the atmosphere. Given the present composition of the atmosphere (Table 1) this confers to rainwater a strongly oxidizing
23
Weathering, rainwater and atmosphere Table 2. Chemical models of rainwater (all concentrations in mg L- I )
Station
Temperature (0C) pH Eh (mV) ClNa+
so4-
Mg++ Ca++ K+ HC03 Si02(aq) 02(aq)
Kiruna, Sweden*
Massif Central, Francet
Monsoon rain, Delhi, India:j:
Cyclonal rain Reunion Island§
c. 8 5.60 910
12 6.37 814
c. 25 7.14 796
c. 25 5.46 896
0.40 0.30 2.02 0.12 0.49 0.20 1 .26 0.002 1 1 .48
1.10 0.80 1.00 0.50 0.20 1 .82 0.001 10.57
0.99 0.69 0.38 0.10 1.16 0.27 4.76 0.005 8.08
1 1.2 3.00 <0.01 0.98 1 .88 0.27 0.77 2.40 8.07
* After Granat (1972), in Appelo & Postma (1993). t Unpublished personal data. :j: After Sequeira & Kelkar (1978). § After Grunberger (1989).
and in most cases a moderately acidic character (see examples in Table 2). Most fresh rocks, on the other hand, are in a reduced state (mainly because they contain ferrous iron and often disseminated sul phides) and may be considered alkaline (because they are composed of carbonate minerals and/or alkaline alumino-silicates). The contrasting chemistry of rain water with respect to fresh rocks could thus be regarded as the actual motor of weathering. The weathering of a silicate such as anorthite to form kaolinite may be written: 2 CaA12Si208 + H20 + 2H+ � Ca + + Al2Si205(0H)4 If we assume that rainwater acidity is mainly derived from its dissolved C02 following the reaction: C02(aq)
+
H20
�
H+ + HC03
combining these reactions shows that silicate weath ering is in fact a carbonation reaction: CaA12Si208 + 2C02(aq) + 3H20 � Ca2+ + 2HC03 + Al2Si205(0H)4
(1)
Holland (1984), after Goldich (1938), has thus shown that during weathering, three types of chemi cal reactions affect the rock-forming minerals: car bonation, oxidation (reaction with rainwater oxygen) and hydration (although some reactions occurring in weathering profiles may be of mixed character). Here are simple examples of the last two types of reaction:
1 oxidation reaction, e.g. magnetite to hematite conversion
(2) 2
hydration reaction, e.g. transformation of anhy drite into gypsum (3)
Hydration reactions affect rocks everywhere in the hydrosphere, but carbonation and oxidation reactions (1) and (2) are typical of weathering, and control the development of the weathering profile. Geochemistry of the weathering profile
From the geochemist's point of view, a weathering profile can be defined as a reaction zone between two media (Fig. 2): 1 the atmosphere, which comprises both a gaseous phase (the air) and a liquid phase (rainwater); 2 the parent rock, which is another two-phase medium, with the mineral assemblage of the unweathered rock as the solid phase and the rock's interstitial water (i.e. the ground water) as the liquid phase. In a simplified approach, we can assume reason ably that these two media are in thermodynamic equilibrium, i.e. that rainwater is in equilibrium with the air and that the unweathered rock's interstitial
J M. Schmitt
24 PHYSICAL
PEDOLOGIC
P E DOGEN IC
PHASE
WEATH E R I N G POTEN TIAL
MEDIUM
HORIZONS
FEATURES
ASSEMBLAGE
O F P ERCOLAT I N G WATE R
horizon surface•---- ("topsoil")
ATMOS P H E R E
biologic activity
WEAT H E R I N G
plant rooting
PROF ILE
humic material
ZONE OF AERATION
B -
horizon
accumulation concretion illuviation pedogenic structures
C -
ZONE OF SATURATION
AIR
oxidizing power
A -
horizon
saprolite incipient weathering relict primary structures
carbonating
RAINWA TER SOIL'S A TMOSPHERE PERGOLA TING WA TER WEA THERED MINERAL ASSEMBLAGE
INTERSTITIAL WA TER WEA THERED MINERAL ASSEMBLAGE
core-stones features
PARENT ROCK
fresh rock
GROUNDWA TER
primary minerals
INITIAL MINERAL
and structures
ASSEMBLAGE
Fig. 2. Main features and geochemical characteristics of a weathering profile. The different pedologic horizons correspond to
specific phase assemblages. The oxidizing and carbonating potential of the percolating water is more or less maintained in the upper aeration zone; it diminishes rapidly towards the base, and is practically exhausted at the weathering front.
water is in equilibrium with the initial mineral assemblage. The soil surface marks the boundary between the top of the profile and the atmospheric medium, whereas the weathering front (Fig. 2) separates the base of the profile from the unweathered parent rock. As rainwater percolates the profile, its 02 and C02 are consumed progressively by reaction with the solid phase, giving rise to successive pedologic hori zons, in which mineral assemblages are in equilibrium with decreasingly aggressive solutions. The weather ing front may be regarded as the front under which the oxidizing and carbonating potential of rainwater, with respect to the parent-rock mineral assemblage, has been practically exhausted (Fig. 2). Within the profile, the water-table level constitutes a major division. It separates the zone of saturation comprising only the interstitial liquid phase and the weathered solid phase, from the zone of aeration (vadose zone). In this part of the profile, the weath ered mineral phase is accompanied by a gaseous phase (the soil's atmosphere), and generally by a liquid phase as well, present as thin capillary films on mineral grains. At least for the biotic era, and owing to the respiration of microbes and plant roots, the soil's atmosphere composition generally differs from
that of the air, being notably richer (10-100 times) in C02 (Holland 1 984; Yolk 1987; Ceding 1 991), and slightly depleted in 02. The presence of the soil's atmosphere nevertheless forms a gaseous buffer, and may allow replenishment of the 02 and C02 content of the percolating water, thereby speeding up weath ering reactions. The mineral assemblage present at ·the base of the profile thus reacts with a small amount of rainwater borne 02 and C02, whereas the maximum amount of the two gases has been used in weathering reactions in the upper horizons and especially in the unsatu rated zone. For a given type of rock, the mineralogy and geo chemistry of the profile depend primarily on the chemistry of rainwater and the atmosphere, which will be discussed now. Chemistry and chemical controls of rainwater
The simplest chemical model of rainwater is given by pure water (condensed water vapour) that has equili brated with atmospheric 02 and C02. Although there is a slight shift in temperature as a result of decreas ing gas solubilities, the approximate pH of this solu tion will be 5.65 (at 25°C). Reported pH values for
Weathering, rainwater and atmosphere modern rainwater vary, however, largely between 4.0 and more than 7 .0. Varying pH values result from the presence of other ions that have been acquired by water during its transport through the atmosphere. These ions are mainly of three types (Appelo & Postma 1993). 1 Na+ and Cl- (and some trace elements) are essen tially of marine origin (Holland 1978; Grunberger 1989). The concentration ratios between these ele ments are nearly equal to those in sea water. They derive from marine aerosols diluted in rainwater, and their concentration usually diminishes rapidly (from more than 30 to less than 1 mg L-1 for Cl-) away from coastal areas and with elevation above sea-level (Meybecq 1984). 2 The anions S04 and N03 originate from the reac tion of rainwater with S02, H2S and NO, gases. The natural source for the first two gases is volcanic activity and to a lesser extent natural forest fires. The NOx gases are formed mainly during biological denitrification of soils (Bouwman 1990), but also can result from lightning discharges in the atmosphere (Holland 1984). These gases oxidize in the atmos phere and produce H2S04 and HN03 acids, thus lowering the pH of rainwater. In modern urban areas the main sources of the above gases are industrial and traffic fumes, which produce the acid rain phenomenon. 3 Ca++, Mg++ and K+, for their major part, as well as the totality of dissolved silica, have a land-based source. They are acquired by rainwater through reac tion with airborne continental dust (and eventually volcanic ashes, Grunberger 1989) . The concentration of Ca++ is of special interest, because it traces the reaction with calcite particles. This reaction results in partial neutralization of rainwater and ultimately controls its pH. Calcite-containing dust especially forms over continental regions that experience arid climates, and is much less abundant in humid cli mates. Climate has in this way a certain control on the pH and chemistry of rainwater. These different contributions and their effect are illustrated in Table 2, which lists the chemical compo sitions of four types of rainwater representing differ ent parts of the world. The analyses are more or less arranged in order of increasing marine contribution. Sample 1 from Sweden records an acid-rain effect, as shown by its S04- content and low pH. Sample 4 has a major marine component, but also exhibits some anomalies (high silica content, excess CJ-, abnormally low pH) that probably result from contamination by volcanic ash (Grunberger 1989). Samples 2 and 3
25
have a rather similar (but minor) marine contribu tion, but the pH of the first sample is notably more acidic than the second. There are two reasons for this: the acid rain component is somewhat higher in the rain over France, whereas the composition of the rain over Dehli reveals a higher dust-induced neutralization. Although appreciable, these chemical variations of rainwater that result from the presence of foreign ions, are of limited extent, with the possible exception of stations near the coast. We may in any case con sider that the marine and land-based (dust-induced) contributions are more or less constant for a given geographical location. As stated before, the overall chemistry of rainwater, and especially its weathering potential (i.e. oxidizing and carbonating potential), depends primarily on its dissolved 02 and C02 contents, which are imposed by the atmosphere chemistry. As we will see in the following section, there are very strong indications that atmosphere chemistry has experienced notable variations on a geological time-scale. Examples of modern rainwaters in Table 2 will help us in constructing and discussing models of rainwater for various atmospheric compositions, before we use them in weathering simulation.
C H E M I S T RY A N D E V O L U T I O N O F E A RT H ' S ATM O S P H E R E Origin and chemical control of Earth's atmosphere
Current knowledge about the formation of the Solar System tells us that like other terrestrial planets, Earth formed about 4.6 billion years ago from the accretion of planetesimals (i.e. small condensed solid bodies) within the solar nebula. Earth's early atmosphere is thought to have formed through the degassing of H20, C02, CO, and N2 (as well as other minor gases) from impac ting bodies. This degassing took place mainly at an advanced state of the accretion process (Kasting et al. 1988; Pollack 1990). Once the main accretionary phase had ended, most of the H20 condensed to form the oceans (before 3.8 Ga). The main components of this early atmosphere at that time were N2 and C02, with estimated partial pressures for both gases near 1 bar. From then on, the composition of the atmosphere was controlled and evolved through a variety of processes which are still active today (Fig. 3).
26
J M. Schmitt
Fig. 3. Sources and sinks of atmospheric gases, and main controls of the atmosphere's chemistry, including the carbonate-silicate cycle. (Modified after Arthur et a/. 1985; Kasting et a/. 1988; Pollack 1990.)
1 New gases (H20, N2, C02, together with minor H2, CO, H2S and rare gases) are introduced in the atmosphere by volcanic outgassing. We must note that owing to variations in volcanic activity, this out gassing has varied markedly, and was probably much more important in Earth's youth than it is today. 2 Both H20 and C02 are depleted from the atmos phere, the former through condensation in meteoric and ocean water, the latter through the formation of carbonate rocks and carbon fossil deposits. 3 � is formed through various photochemical reac tions in the upper atmosphere, but escapes to space together with other light gases. These different processes have resulted in accumu lation of N2 and rare gases over time, and to a certain regulation of the atmospheric abundance of some gases, and especially of 02 and C02. During Earth's history, oxygen and carbon dioxide levels have never theless undergone important fluctuations, as detailed below.
Variation of the oxygen atmospheric level
It is admitted generally (Holland 1984; Gillett 1985; Kasting 1993) that the early atmosphere was totally devoid of free molecular 02, and weakly to strongly reducing because of the composition of volcanic out gassing. As a whole, the atmosphere slowly became
more hydrogen-poor and C02-rich through the pho tochemical dissociation of water and the escape of H2 to space (Fig. 3). Large quantities of molecular oxygen appeared only as a by-product of life (Cloud 1 972), resulting from the conversion by plants of water and C02 into food (sugars) (Lovelock & Whitfield 1982). The rise of free molecular oxygen in the atmos phere has been discussed by many authors (Cloud 1972; Holland 1 984; Gillett 1 985; Kasting et al. 1 988; Pollack 1990) and modelled for the Precambrian era by Kasting ( 1993) both from geological evidence and astronomical arguments. Three stages are distin guished. 1 The early stage (stage I), during which reducing conditions prevailed in the atmosphere, surface ocean and deep ocean, with 02 partial pressure cer tainly below 10-4 bar. This stage ended between 2.4 and 2.0 Ga. 2 An intermediate oxidizing stage (stage II) during which the atmosphere and surface ocean may have become oxidizing, although with 02 atmospheric abundance not above 0.03 PAL (present atmospheric level), whereas the deep ocean remained reducing. 3 A third aerobic (stage III) where both the atmos phere and hydrosphere contained abundant free 02 as on modern Earth. The beginning of this stage was marked by a major increase in 02 levels around 1 .85
27
Weathering, rainwater and atmosphere Ga, with early concentration probably well below that at present but exceeding 0.002 PAL. A final major episode in the rise of the 02 level is thought to have occurred at the end of the Proterozoic (c. 550-600 Ma) and may have triggered the global diversification of organisms during the Precambrian. What caused the initial appearance of free 02 at the beginning of the intermediate oxidizing stage remains unclear, but all authors agree that from then on, the 02 level rose as a consequence of photosyn thesis and organic carbon burial. Contrasting with the picture for carbon dioxide (as we will see below), the Phanerozoic evolution of the atmospheric 02 level is not well documented, although minimal values can be inferred indirectly from the study of some palaeosols (Holland 1984; Feakes et al. 1989; Holland et al. 1989). Some fluctuations could well have occurred, but we may nevertheless consider that since the beginning of the Palaeozoic era the 02 abundance has not been significantly different from that at present, at least for its effect on rock weathering. Variation of atmospheric C02
Pollack (1990) has shown that in comparison with other terrestrial planets, the present-day distribution of Earth's carbon dioxide endowment is unusual (Table 3). The main C02 sink lies in sedimentary rocks, both in carbonate rocks and as fossil organic carbon. Oceans contain a mere 0.05 % of total C02 and the atmosphere less than 0.001 % , representing a partial pressure of only 0.0003 bar. This low atmospheric abundance of C02 results from a balance between its production by volcanism, and its withdrawal through organic carbon and car bonate sedimentation. In fact C02 is part of a rather complex cycle, known as the carbonate cycle, or the carbonate-silicate cycle, which has been investigated and detailed more or less independently by several authors (Holland 1978; Walker et al. 1981; Berner et al. 1983; Yolk 1987). The main parts of this geo chemical cycle (see Fig. 3), as summarized by Kasting et al. (1988), are as follows: 1 atmospheric carbon dioxide dissolves in rainwater and forms carbonic acid; 2 rainwater weathers continental rocks, especially calcium-bearing silicates, and releases Ca++ and HC03 ions into groundwater; 3 groundwater is transported to streams, rivers and the ocean;
Table 3. Today's Earth carbon dioxide endowment and distribution (Modified after Pollack, 1990)
Atmosphere Oceans Surface organic matter Organic carbon in sedimentary rocks Carbonate rocks Total
4
5.6 x 1016 mol (0.001 % ) 3.2 x 1 018 mol (0.05 % ) 2 x 1017 mol (0.003 % ) 1 x 102 1 mol (16.66%) 5 x 102 1 mol (83.29 % ) 6.003 x 1021 mol (100%)
in the ocean, dissolved ions are incorporated by organisms in shells and form carbonate sediments; 5 during subduction, sediments are subjected to increased temperature and pressure and calcium car bonate reacts with quartz, forming new calcium bearing silicates and releasing carbon dioxide; 6 the C02 ultimately returns to the atmosphere through volcanic outgassing; The global budget of atmospheric C02 depends on the different fluxes in this complex cycle and conse quently on many factors, such as climate, continental land area, sea-floor spreading rate, and so on. It is of course complicated by the presence of the biomass and by biological activity. Considering the very small part of the atmosphere in Earth's total C02 endow ment, marked variations in the C02 atmospheric level at a geological time-scale appear likely. The possibility of tracing past C02 atmospheric abundance through isotopic studies of pedogenic carbonate concretions has been demonstrated by Cerling (1991 ; this volume, pp. 43-60). Below 20 em depth, the carbon isotopic composition of soil carbon ate appears constant. Carbonate is formed in isotopic equilibrium with soil C02. In the absence of C4 plants (that is before late Miocene), the isotopic composi tion (oBC value) of soil C02 in turn depends on the concentration of C02 in the atmosphere. This model has been applied by Cerling himself on Cenozoic and Mesozoic palaeosols and by other authors (see e.g. Mora etal. 1991 ; Driese & Mora 1993; this volume, pp. 61-84, and other references therein) on older palaeosols. Similar oBC isotopic data also have been obtained recently on a Fe(C03 )0H com ponent of pedogenic goethite (Yapp & Paths 1996). These data have shown that the C02 level has fluctu ated effectively between one and 20 times PAL during the Palaeozoic-Miocene interval. Walker et al. (1981) and Berner et al. (1983) have quantified the different terms of the global carbon ate-silicate cycle, enabling them to construct a
28
J. M. Schmitt
C02 x TODA Y
1 slf1]�����===i;;;;�b�ee:s�ot:-��iat�c�u�re�v�e�==� 1 46 �����illit;==��-��i�sr�or;to�rpi�r�cn�mg�e�asur�e 1 2 �����; � 1 0 -Hl ·• ·• • • • "N 1 64 2 8
Fig. 4. Variation of the atmospheric C02 level for the Phanerozoic (after the GEOCARB I model of Berner 1991 ) Isotopic measures by Cerling (1991), Driese & Mora (1993) and other data in Appenzeller (1993). .
actually warmer than today. This problem is known as the faint young Sun or cool Sun paradox (Sagan & Mullen 1 972). Because C02 is the main greenhouse gas, the solution to this apparent paradox requires C02 partial pressures high enough to maintain a suffici ent surface temperature (Walker 1982). The pC02 required is estimated approximately at 1 bar at 4 Ga, 0.1 bar at 2.5 Ga and around 0.001 bar in the Late Precambrian (Kasting 1 993) . I n the following sections w e will first see how geo chemical modelling may be used to simulate granite weathering under present atmospheric conditions, and then we will show two examples of how different compositions of the atmosphere may dramatically change the conditions of weathering and the result ing profiles.
S I M P L I F I E D S I M U LATI O N O F G R A N I T E W E AT H E R I N G Modelling code and its use
computer model of the C02 atmospheric level. The early (Berner et at. 1 983) and more recent (Berner 1 991, 1 994) dynamic models of Berner are sufficiently complete to incorporate all the main aspects of the carbonate cycle, and tentatively trace C02 atmos pheric abundance back to the Early Palaeozoic. These results (Fig. 4) agree well with isotopic data from pedogenic carbonates (Appenzeller 1993; Driese & Mora 1993) and show a rather spectacular succession of highs and lows: several C02 peaks occurred during the Palaeozoic, a well-marked one during the Cretaceous, and a pronounced minimum during the Carboniferous-Permian. Up to now, Berner's model has not been applied to the Protero zoic. The reason is that it requires reliable data on palaeogeography, palaeoclimatology and palaeogeo dynamics, which are increasingly difficult to obtain for remote periods of Earth's history. The early abundance of atmospheric C02 has, however, been approached on the following astro nomical and climatological grounds. According to the physics of stellar evolution, there is little doubt that solar luminosity during the Archaean was significantly fainter (about 30%) than at present. Given the present C02 level, this normally would result in an ice-covered Earth. The geological record, however, shows that liquid oceans have existed since at least 3.8 Ga, and suggests that early Earth was
Various geochemical modelling codes now enable us to simulate the reaction between a fluid and a solid phase. The React code is part of a group of interac tive programs, the Geochemist's Workbench™ by Bethke (1994, 1 996). The code describes a geochemi cal system (given the appropriate constraints) in terms of physico-chemical parameters (pH, Eh, ionic strength, etc.), composition and speciation of the aqueous phase, minerals saturation index, gas fugaci ties, and mineral composition of the solid phase. The code allows us to trace the progress of reaction processes (reaction path modelling) involving solu tions, minerals and gases. The accompanying program Gtplot plots and prints the results along the simu lated reaction path. Geochemical modelling codes, whether simple or not, have of course been used to simulate weathering reactions (see e.g. Fritz 1981; Steinmann et a!. 1 994; Bethke 1996). A weathering profile, however, is an open system and simulation would normally require the use of coupled or continuum models, which are difficult to set up and compute (B ethke 1996). Con tinuum models nevertheless predict that reactions occur in the same sequence as reaction path models (Lichtner 1988). In this way, the React code may be used to simulate rock weathering, and to some extent the develop ment of a weathering profile (Schmitt 1994; Schmitt
29
Weathering, rainwater and atmosphere & Thiry 1994). We can simulate the reaction of an initial mineral assemblage and accompanying aqueous phase with an adjustable volume of a reac tant fluid. In the flush option of the code the reactant fluid displaces at each step an equal volume of the interstitial water. If we use a model of the parent rock for the initial system, and a model of rainwater for the flushing solution, the reaction path simulates the successive weathering stages of the initial mineral assemblage as a result of the oxidation and carbona tion reactions induced by the percolating water. Keeping in mind that these reactions are only incipi ent in the bottom horizon and maximum at the top of the weathering profile (Fig. 2), we can consider that the flushing reaction path simulates the succession of mineral assemblages (i.e. of horizons) met with from the unweathered parent rock upward. We describe below the assumptions made in this simulation and the results obtained for a complete weathering of granite under present conditions. Initial system representing the parent rock
The initial system [solid phase + fluid phase] must be in thermodynamic equilibrium. This means in par ticular that the mineral assemblage of the rock will control the interstitial fluid chemistry. Most of the components of the initial fluid will thus be fixed through the equilibrium with one of the minerals. For instance, quartz may be used to fix the concentration of dissolved silica species -namely, quartz is swapped for Si02(aq), muscovite for K+ species, and so on. A more complete discussion of this technique is given in Bethke (1996) . The granite model used in the simulation is an assemblage of 40 % quartz, 25% K-feldspar, 20 % plagioclase (0.9 albite + 0.1 anorthite), 10% biotite (0.7 annite + 0.3 phengite), 5% muscovite, and trace amounts of pyrite. The initial system contains 1 g of this granite.The present simulation, as well as most of those that follow, has been conducted at a constant temperature of 25°C. At this temperature anorthite is unstable and the program readily converts it to prehnite during the initial equilibrium calculation, before the simulated reaction begins. Rainwater model used as flushing fluid
The flushing fluid used in the simulation is a model based on sample 1 from Table 2. It is a moderately acidic (pH = 6.4) water that contains small amounts of N a+, Cl-, Ca2+, K+, SO�- , and Si02( aq). Equilibrium
with the present atmosphere is assumed, and hence the fugacities of gases are fixed to f02 = 0.2, and fC02 10-3 .5. In some runs, for calculation require ments (coherence of composition of the two reac tants) minute amounts of other elements (iron, aluminium, etc.) also must be included in this rain water model. When the simulation is completed (1.0 reaction 120 L) of rainwater have progress) 120 kg (i.e. been flushed through 1 g of the granitic fresh rock. =
=
Results of weathering simulation
This simplified simulation of granite weathering under the present conditions has been performed and is discussed below as an example and reference for the palaeoprofiles that will be examined later. The results may be visualized conveniently as a plot giving the mineral composition of the rock versus reaction progress (Fig. Sa). This plot shows the fol lowing steps in the weathering process: 1 initial stage of alteration -the plagioclase (albite and anorthite), as well as the annite component of biotite, and pyrite, are readily dissolved and replaced by muscovite and iron-bearing smectites (nontronite); 2 from 0 to 0.12 reaction progress, K-feldspar and phengite are destroyed and replaced mainly by muscovite; 3 from 0.12 to 0.42 progress, quartz begins to dis solve, and muscovite is weathered to kaolinite; 4 from 0.42 to 0.74 progress, clay minerals (kaolinite and nontronite) remain stable whereas quartz continues dissolving; 5 from 0.74 to 0.85 progress, quartz is no longer present, kaolinite remains stable, but nontronite is destroyed and iron is retained as haematite; 6 in the final stage, haematite remains stable, whereas kaolinite weathers to gibbsite. The simulation also allows us to verify the behaviour of chemical elements. Sulphur and sodium are leached as soon as weathering begins, Ca, Mg and K are partly retained in clay minerals until smectites disappear, silica is slowly but continually leached away throughout, and iron and aluminium remain practically immobile. As there is a volume and mass loss of about two-thirds of the solid phase during the simulated reaction, this actually means that the rela tive content in both aluminium and iron is strongly enriched. The minerals-reaction progress plot also may be considered as a simulated weathering profile (Fig. 5b)
30
J. M. Schmitt
Minerals in system (log grams)
results above is a slight shift in the appearance of hematite. We can thus claim that this simulated complete flushing of a granite down to the [gibbsite + hematite] stage represents well the succession of mineral assemblages and horizons observed in lateritic profiles under present hot and humid climates.
EXAMPLE AND M O D E LLING O F W E AT H E R I N G I N A C 0 2 - R I C H ATM O S P H E R E D U R I N G T H E C R E TAC E O U S
0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 Reaction progress (a) Land surface
0.8
0.9
--------�
kaolinite + haematite ± gibbsite topsoil ®
kaolinite + quartz + smectite horizon
@
quartz + kaolinite + mica + smectite horizon
G)
quartz + K-feldspar + mica + smectite horizon unweathered granite
(b) Fig. 5. Granite weathering simulation for present
atmospheric conditions: (a) minerals (log moles) versus reaction progress (at the end of the simulation 120 kg of rainwater have been flushed through 1 g of granite); (b) corresponding simulated profile. Numbers 1 through to 4 refer to the successive mineral assemblages in (a).
comprising: (i) a lower [quartz + K-feldspar + illite/ mica + smectite] horizon; (ii) a [quartz + kaolinite + smectite + illite/mica] horizon; (iii) a [kaolinite + quartz + smectite] horizon; (iv) a [kaolinite + haematite ± gibbsite] topsoil. In this simple simulation we have not taken into account the existence of an unsaturated (or aeration) zone, that could buffer 02 and C02 fugacities during the weathering process at the top of the profile. We have, however, verified in different modelling runs that the main potential effect of this buffering on the
The singular Cretaceous weathering record
Unusually deep and leached profiles are often pre served in the geological record, even at extratropical palaeolatitudes (Bird et al. 1990). This is particularly true for the Cretaceous Period, for which deep kaolinitic or bauxitic profiles of lateritic type are known and kaolinite deposits are ubiquitous (Millot 1970). At present, bauxites form only under hot and humid equatorial climates (Tardy et al. 1990) and the conditions of laterite formation are currently satisfied only between 30°N and 30°S (Bardossy 1981). For the Middle Cretaceous, palaeogeographi cal reconstructions (Scotese et a!. 1981) show that the palaeoequatorial zone corresponds to northernmost South America, western Africa, northern Central Africa and Ethiopia, Indonesia and Borneo. Creta ceous bauxitic formations are effectively present in these zones, e.g. on the Guiana Shield (Aleva 1981), in the Amazon Basin, Guinea, Ivory Coast, Nigeria and Cameroon (Tardy et al. 1990). The Middle Cretaceous, however, shows a large expan sion of bauxites in the Mediterranean region (Bardossy 1981; Bardossy and Combes, this volume pp. 189-206), in southern France and north-eastern Spain (Combes 1969), probably as far north (30--40°N) as Middle Asia (Bardossy 1981), and as far south (35--45°S palaeolatitude) as Jammu province, India (Mohan et a/. 1981). At the same time, deep kaolinitic weathering profiles (or laterites, or ferralitic soils) formed at still higher latitudes. In northern America, such profiles and sometimes extensive kaolinitic regoliths have been described, for instance in south-western Wisconsin (Dury & Knox 1971), Minnesota (Goldich 1938;Austin 1970; Parham 1970), and in other parts of the south-eastern USA (Sigleo & Reinhardt 1988). In
Weathering, rainwater and atmosphere Europe, Dupuis (1992) has described giant (up t o 70 m thick) kaolinitized regolith remnants of Early to Middle Cretaceous age in Belgium, northern France and The Netherlands. Intensive lateritic weathering also developed during Cretaceous times in the South Bohemian Basin (Slanska 1976) and in the Urals (Sapojnikov 1981). The erosion of most of these profiles has given rise to very important Upper Cretaceous and Lower Tertiary sedimentary kaolin ite deposits in the south-eastern USA (Austin 1970), in the London Basin and Belgium (Wealden facies, Millot 1970; Sladen 1983), in the Paris Basin (Thiry 1981), and in the Bohemian Basin (Klikov Formation, Slanska 1976). In the Southern Hemisphere, laterites, and kaolini tized profiles as well, formed during Cretaceous times at high latitudes (30-45°S) in India (Subramanian & Mani 1981), Argentina (Cravero & Dominguez 1992) and probably northern Australia (Schmidt & Oilier 1988). This striking expansion of intensive weathering at extratropical latitudes has been interpreted as the result of an anomalous climatic zonation during the Cretaceous. The Cretaceous climate
The Cretaceous palaeoclimate has been discussed in connection with a variety of geological evidence, in addition to the palaeoweathering record. The main features that have been considered are: distribution of diagnostic rocks (e.g. coals, evaporites, reefs); palaeobiogeography, both terrestrial (palaeofl.oristic provinces) and marine (distribution of Boreal versus Tethyan faunas); and marine palaeotemperatures (from oxygen isotopes data, mainly on benthonic and planktonic Foraminifera). These data have been dis cussed and synthesized by Frakes (1979), Arthur et al. (1985) and Kemper (1987). These authors agree that all data available suggest warm mid-latitude and polar conditions during much of the Cretaceous, and the absence of mid-winter freezing at middle to high latitudes. Although numer ous short cold periods as a result of climatic cyclicity apparently occurred (Kemper 1987), the mean global temperature seems to have undergone a slow rise during the Cretaceous, with a maximum value in the late Middle Cretaceous. Temperatures were then probably some 10oc warmer than today at mid latitudes. A marked and rapid cooling followed at the end of the Cretaceous. As a whole, the Cretace ous stratigraphical record indicates (i) that global eli-
31
mates were considerably warmer than at present, and (ii) that latitudinal temperature gradients were considerably less than at present. Consequently it is widely assumed that the climate was warm and equable during most of the Cretaceous. Various reasons have been put forward to explain this abnormally warm and uniform global climate: changes in the solar constant (Frakes 1979), transi tory anomaly in Earth's orbit, peculiar configuration of the continents and greenhouse effect. The role of these different factors has been tested using general circulation models (GCMs) and discussed in several papers (Barron & Washington 1982, 1985; Schneider et al. 1985). Peculiar oceanic circulations, with con tinents clustered in a great land mass, appear to have contributed to high poleward heat transport, and to have resulted effectively in more equable and warmer climates. General circulation models also indicate that mid-latitude rainfall in the Cretaceous was greater than at present (Barron & Washington 1982). Even considering various favourable parameters (including astronomical ones), global circulation models constrained by the Cretaceous palaeogeo graphy do not succeed, however, in reproducing the observed temperatures. This indicates a probable greenhouse warming by atmospheric C02. Sensitivity experiments by Barron & Washington (1985) and Barron et al. (1993) have shown that a two- to 10fold increase in pC02 is required to reproduce the Cretaceous temperatures, especially at mid-latitude. Schneider et al. (1985) also have shown that similar high atmospheric C02 concentrations are needed to avoid mid-winter freezing at middle to high latitudes. The reality of this increased C02 level during the Cretaceous (four to nine times PAL), has been sub stantiated by 813 C isotope data on pedogenic carbon ate and pedogenic goethite (Cerling 1991; Driese & Mora 1993; Yapp & Poths 1996). Similar carbon isotope data from Andrews et al. (1995) on Late Cretaceous palaeosols in India have shown that by this time atmospheric C02 had probably dropped again below 1300 p.p.m. (that is around four times PAL). These high values of Cretaceous atmospheric C02 are also reproduced in the study of Volk (1987) , and b y the dynamic models o f Berner (1991, 1994), which show (Fig. 4) a C02 peak of some eight to 10 times PAL in the Mid-Cretaceous. Both the peculiar palaeogeography and the green house warming induced by a high atmospheric C02 level, therefore appear as the likely reasons for the warm and equable Cretaceous climate. Warmer
32
J M. Schmitt
tropical-like climates at higher latitudes could well account for the unusual palaeoweathering record of the Cretaceous, but Bird et al. (1990) have shown that the o lSQ composition of kaolinite from some deeply weathered profiles formed at high latitudes reveal cool to cold conditions of formation, rather than a tropical climate. Consequently, the same authors have suggested that an elevated pC02 may well be an important factor in the formation of deeply weathered profiles at extratropical latitudes. We have made the same suggestion independently (Schmitt & Thiry 1994) and we discuss it below after the results of another geochemical simulation of granite weathering. Simulation of granite weathering in the high-C02 Cretaceous atmosphere
After Berner's model (Fig. 4), the C02 atmospheric level during the Mid-Cretaceous possibly peaked at about 10 times PAL. A model of rainwater in equilib rium with such an atmosphere has been constructed and used for granite weathering simulation. To allow easy comparison with the previous simulation (in present atmospheric conditions), the rainwater model has been modified only for its imposed C02 2 fugacity (10- .5 bar). Chemical composition of this Cretaceous model rain is given in Table 4. As in the simulation described in the preceding section, 120 kg of the Cretaceous rain are flushed through 1 g of the same granite model. The simula tion here is again conducted at 25°C constant temper ature. The results are visualized in Fig. 6, on the same plot giving rock composition versus reaction progress.
The different weathering phases are very similar to those of the previous weathering simulation. (Fig. Sa), but the following main differences can be observed: 1 most of the initial mineral assemblage, except for quartz and muscovite, disappears as early as 0.04 of the reaction progress. 2 kaolinite appears at this point, i.e. three times sooner than in the previous simulation, and domi nates the weathered mineral assemblage until the end of the simulation. 3 hematite and gibbsite are formed, respectively, from 0.76 and from 0.85 of the reaction progress, i.e. more or less at the same stage of flushing as in the previous run. When compared with the simulation made under present atmospheric conditions, it should be ap parent that the [kaolinite + quartz (+ smectite)] assemblage is reached about three times sooner (i.e. with three times less rainwater flushed through the granite) when C02 abundance is increased 10 times. It means that under similar rainfall and temperature conditions granite would be weathered and profiles would deepen three times faster than under present atmospheric conditions. It is obvious, however, from comparison of Figs 5 & 6, that the increased pC02 has no direct effect on the appearance of gibbsite (which occurs in both cases near 0.85 of the reaction progress), and hence on
Minerals i n system (log grams)
25°C
I_L I _L I -L������L-r O -r��������L
Table 4. Models of rainwater under present atmospheric
conditions (1), and under the C02-rich Cretaceous atmosphere (2) (all concentrations in mg L- 1) Model
[C02 (g) fOz (g) pH Eh (volts) Cl-
so4-
Na+ K+ Ca++ Mg++ Si0 2(aq)
Model l : Present atmosphere 10-3.5 0.2 6.37 0.842 1.10 1 .00 0.80 0.20 0.50 0.10 0.001
Model 2: Cretaceous atmosphere 10-2 .5 0.2 5.52 0.892 1 .10 1.00 0.80 0.20 0.50 0.10 0.001
-2
0
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 Reaction progress
Fig. 6. Granite weathering simulation for Cretaceous
atmospheric conditions at 25°C. Same plot as in Fig. 5. Numbers 1 through to 4 refer to the same mineral assemblages as above. Note the early appearance of kaolinite, and the unmodified one for gibbsite.
Weathering, rainwater and atmosphere
Minerals in system (log grams)
-
1
33
likely origin of the expansion of bauxites during the Cretaceous. The facility of both kaolinization and bauxitization at extratropical latitudes probably vanished at the end of the Cretaceous, as a result of the important drop in atmospheric C02 (Andrews et al. 1995), and the concomitant return to a more usual climatic zonation (Frakes 1979).
-2 E X A M P L E A N D M O D E L LI N G O F G R A N I T E W E AT H E R I N G U N D E R A N O X I C C O N D I T I O N S c. 2 . 4 G A
0
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 Reaction progress
Fig. 7. Granite weathering simulation for Cretaceous
atmospheric conditions at 35°C. Mineral assemblages 1 through to 4 as before. Note this time the early appearance of gibbsite.
bauxite formation. Faster bauxitization may result only from increased rainfall, or from higher mean temperature. This latter effect is best illustrated by a different simulation of granite weathering at a warmer temperature. Figure 7 illustrates the result obtained at 35°C (instead of 25°C) . Gibbsite here appears as early as 0.60 of the reaction progress, as a result of the important increase (some 40% ) in silica solubility between 25 and 35°C. This means that for a 10°C warming, the total rainfall needed for bauxitiza tion is some 30% less than at a cooler temperature. Conclusion on the effect of the Cretaceous atmosphere and climate on granite weathering
Modelling shows that high atmospheric C02 values greatly accelerate the formation of deep kaolinitic profiles, without needing periods as long, nor climates as warm and humid as in present conditions. This explains why deep kaolinitized profiles, and kaolinite deposits have been widespread during the Cretaceous, even at extratropical latitudes, and under cool, moderately humid climates. There is no direct effect of the increased C02 atmospheric abundance on the rapidity of bauxitiza tion, but we know that the induced greenhouse effect, and the peculiar Cretaceous palaeogeography, have resulted simultaneously in an increase in rainfall and in an important warming at mid-latitudes. Modelling shows that the conjunction of these two factors is the
Palaeosols and sedimentary formations from the Elliot Lake area, Ontario, Canada
The Blind River and Elliot Lake area of the Canadian shield (see for instance Roscoe 1957, 1969) is famous for its detrital uranium deposits in quartz pebble conglomerates, the only parallel of which are the uranium-gold ores of the Witwatersrand Basin in South Africa. The area also is known for some of the oldest palaeosols of the geological record (Roscoe 1957; Gay & Grandstaff 1980; Prasad et al. 1993). These profiles show typical pedogenic features (core stones, argillized horizon, cutan structures, etc.) but they also exhibit very unusual characteristics that are considered evidence of reducing or anoxic conditions of the atmosphere at this time. The profiles are mainly developed on Archaean rocks, and are overlain by the detrital lower Huronian Matinenda Formation, which contains authigenic pyrite as well as detrital pyrite and uraninite (in the mineralized quartz pebble conglom erates). Younger palaeoweathering profiles also have been found and studied in the Upper Huronian, but they do not display the unusual characteristics of the older ones, and hematite develops in their upper saprolite zone, like in present-day profiles. The Upper Huronian Cobalt Supergroup moreover con tains the first red beds known from the geological record. The above succession has been considered by Prasad & Roscoe (1991, 1996) as an effect of the oxy atmoversion or transition from an early anoxic to the oxic atmosphere. The time of this transition is constrained by field relationships and regional chronology to the 2.45-2.22 Ga interval, the authors' estimate being 2.4 Ga. In Kasting's (1993) scheme of atmosphere evolution (see above), this date would correspond more or less precisely to the end of stage
34
J. M. Schmitt
I, with maximum 02 partial pressures around 10-4 bar. The sub-Huronian palaeoweathering profiles
Palaeoweathering profiles fossilized beneath the lower Huronian strata have been described on various parent rocks, such as granite (Pronto palaeosol), greenstone (Denison palaeosol) and basalt (Quirke II mine) (Roscoe 1957, 1969; Gay & Grandstaff 1980; Prasad & Roscoe 1991), and more recently on sandstones (Sutton & Maynard 1993). For profiles established on metabasalts at Quirke II mine, Prasad & Roscoe (1991) have distinguished: massive metabasalt; increasingly (upward) altered metabasalt, highly altered metabasalt with calcite cement and minor pyrite, and palaeosol argillite. On granite at Pronto and, above all, Panel Mine (Prasad et al. 1 993) the following succession is observed (Fig. 8): unweathered homogeneous granite; granite with corestone facies; gradually argillized horizon with uppermost argillaceous cap containing floating quartz grains. The lower zone of initial alteration may exhibit carbonate enrichment (with calcite neo genesis) and Fe, Mn, Mg and K depletion. The upper, argillite-like, intensely weathered section is com posed mainly of sericite, with minor chlorite and quartz.
land surface
Most authors have shown that, unlike more re cent profiles, these palaeosols are characterized by upward depletion of Fetotai and Fe3+ in the whole rock. The same depletion is observed in chlorites and micas as well (Prasad & Roscoe 1996). The fact that the decrease in Fetotai is not accompanied by the for mation of ferric minerals (goethite, haematite) is con sidered by most authors a major indicator of anoxic conditions. All these profiles are also characterized by the preservation of Fe2+ minerals and especially of pyrite. Pyrite is present in several morphological forms, par ticularly in the top part of many profiles. Pyritic crusts also are present on the top of some profiles, and are well preserved in underground workings (Gay & Grandstaff 1980; Prasad & Roscoe 1996). As stated previously, pyrite is present together with uraninite (and with brannerite, a Ti + U ± Ca silicate) as a detri tal and authigenic mineral in the overlying Mati nenda Formation. This means that uraninite also has been preserved sufficiently well in weathering profiles, unlike in present profiles developed in oxic conditions. The origin and significance of these unusual palaeoweathering profiles have been much dis cussed. Several authors have wondered for instance if the formation and preservation of pyrite at the top of the profiles should imply that H2S was present and that truly reducing conditions prevailed in the
observed {a)
argillaceous topsoil sericite, floating q u a rtz g rains pyrite
simulated { b)
land surface
kaolinite to
argillaceous topsoil
��
kaolinite + q u artz ± smectite + m in o r pyrite
argillized h o rizon q u a rtz + m uscovite + s mectite
g ra d u a lly arg illized horizon
�1±±flli§q_ q ua rtz · a u th ig en ic pyri te � ·
sericite, chlo rite
chlorite
Fe, Mn,Mg , K depletion core-stones featu res h o m ogeneous g ranite
�����[_���������:::::_ J ± kaolinite
sericite, m in o r ch lorite,
pa
f2\ �
__
Q uartz + K-feldspar + m uscovite
+ biotite ± calcite
Likely diagenetic reactions
Fig. 8. Schematic profiles of Lower Huronian palaeosols: (a) main features as observed by Prasad et al. (1993) and Prasad &
Roscoe (1996); (b) simulated profile corresponding to the plot in Fig. 9. The discrepancies between (a) and (b) are thought to result from diagenetic reactions indicated in the column at centre.
35
Weathering, rainwater and atmosphere atmosphere. The interpretation generally accepted has been disputed by some authors, including Rain bird et al. (1990) who argued that the presence of illite/sericite in the uppermost horizon (instead of kaolinite) could result from a later event of K metasomatism and that part of the early Huronian profiles unusual characteristics could be due to diage netic transformations. More recently the overall reducing character of the early Proterozoic atmos phere also has been disputed by Ohmoto (1996). Geochemical simulation of weathering can help when discussing these problems, but to model these Huronian palaeoprofiles we first need to model the Huronian rainwater. Modelling the early Huronian rainwater
As we know little about the early Proterozoic atmos phere, and less about the Huronian rainwater, its modelling may appear a hopeless enterprise. We can, however, make some reasonable assumptions about these conditions, and see how the Huronian palaeo profiles help to explain them. We know first, as imposed by the solution of the faint young Sun paradox (see above) that the C02 partial pressure at this time was about 0.1 bar, that is about 300 times PAL. As for the oxygen partial pressure, the model of Kasting (1993) indicates a maximum value of some 10-4 bar (i.e. 0.5%o of PAL) . We also must consider the possibility of a significant H2S partial pressure in the Huronian atmosphere. The coexistence of this gas with only minor amounts of free oxygen would result in a sulphuric acid rain, with a very low pH, but this fact appears most
unlikely at this time (Holland 1984). We may thus consider two main alternative types of rainwater models: one corresponding to an oxygen-poor (anoxic) atmosphere totally devoid of free H2S; and the other corresponding to a truly reducing atmos phere containing appreciable traces of this gas. We have of course no clear ideas of the dissolved ions, either of marine or land-based origin, contained in the Huronian rainwater. We have seen that such ions may somewhat buffer the rain's pH, but we have no valid reasons to suppose that the buffering was significantly higher at that time than it is now. For this reason, and for easier comparison with other models, we have thus retained values of dissolved ions similar to those in average present rainwater. Table 5 lists the chemical data for our basic model of the early Huronian rainwater (based on Kasting's model), as well as for several alternative models with varied p02 pC02 and pH2S. These models will be used and eventually adjusted for the simulation of Huronian palaeoprofiles. Simulation of granite weathering in the early Huronian atmosphere
The React code has been used to model early Huronian weathering, in the same way as we simu lated previously granite weathering in present condi tions. The granite model, representing the initial system, has been modified slightly in order to incor porate trace amounts (0.5 wt % ) of uraninite and so follow the behaviour of uranium minerals during weathering. We will first describe the results of the weathering
Table 5. Main model types of the Huronian rainwater used in weathering simulation (all concentrations in mg L- 1)
Model [C02 (g) [0 2 (g) [H2S (g) pH Eh (volts) Cl-
so4-
Na+ K+ Ca++ Mg++ Si02 (aq)
Model l : Kasting's (1993) 2.4 Ga atmosphere
Model 2: higher C02 atmosphere
Model 3: higher 02 atmosphere
10-1 10-4
10-15 10-4
10-1 10-2.5
4.72 0.891 -tid.* -tid. -tid. -tid. -tid. -tid. -tid.
4.44 0.930 -tid. -tid. -tid. -tid. -tid. -tid. -tid.
4.44 0.907 1.10 1.00 0.80 0.20 0.20 0.10 0.001
* -tid., is the same value as Model l .
Model 4: reducing H2S-rich atmosphere 10-1 10- 1 10-5 4.45 -0.022 -tid. -tid. -tid. -tid. -tid. -tid. -tid.
36
J M. Schmitt
simulation using our basic model of Huronian rain water (model 1, Table 5), assuming the atmosphere composition proposed by Kasting (1993). In a further step, we will compare these results with the actual palaeoprofiles, and show how the main alternative models of rainwater (and of atmosphere composi tion) also may satisfy the field data, or not. The result of the simulation is again visualized (Fig. 9) as a plot showing the mineral assemblage as flushing progresses. When the reaction is compl eted here, only 3 kg (i.e. "' 3 L) of rainwater have been flushed through 1 g of granitic fresh rock. The final [quartz + kaolinite] assemblage is obtained very
Minerals in system (log grams)
(a)
(b)
0
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 Reaction progress Elemental rock composition (log grams)
0
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 Reaction progress
Fig. 9. Simulation of granite weathering using the early Huronian rainwater model l (in Table 5): (a) Minerals versus reaction progress plot, numbers 1 through to 4 refer to the main successive mineral assemblages described in the text, and to the different horizons in Fig. S(b ); (b) evolution of the rock elemental composition during the simulation.
rapidly (compared with Fig. 5) because of the much higher weathering (i.e. carbonating) power of this high C02 rainwater model. The leaching out of quartz and the formation of gibbsite would occur at much higher flushing ratios (100-120 kg rainwater per gram of granite) The overall characteristics of early Huronian profiles and especially the constant pres ence of relict quartz grains in their upper horizons, however, rule out such high weathering ratios. The plot (Fig. 9a) enables us to distinguish the fol lowing steps of weathering: 1 Initial stage of alteration (0-0.05 of reaction progress): the plagioclase (albite), as well as the annite component of biotite, and a small part of the K-feldspar are destroyed and replaced by quartz, muscovite and iron-bearing smectite (nontronite) . 2 From 0.05 to 0.22 of reaction progress, K-feldspar and phengite first remain stable and then are weath ered to muscovite and quartz. Calcite neogenesis (dashed curve, Fig. 9) will occur at the beginning of this step if a calcic plagioclase is present in the parent rock. 3 From 0.22 to 0.26 of reaction progress the as semblage [quartz + muscovite + nontronite] remains stable. 4 From 0.26 to 0.47 of reaction progress, part of the nontronite is destroyed and siderite temporarily forms, and muscovite is weathered to kaolinite. 5 From 0.47 to 0.70 of reaction progress the remain ing nontronite is in turn weathered to [kaolinite + quartz]. 6 From then on, the [quartz + kaolinite] assemblage remains stable until the end of the simulation. Pyrite, whether present in appreciable quantities or only in trace amounts in the parent rock, is preserved throughout the weathering process. In any case, authigenic pyrite is formed during the first half of the simulation (0.06-0.47 reaction progress). Uraninite is preserved largely during the initial weathering, well into the kaolinization stage (up to reaction progress 0.58) . It slowly weathers to the hydrated Ca-silicate uranophane (Ca(H20h(U02h(Si02h(OH)6) which remains stable until the end of the simulation. The individual behaviour of chemical elements during the simulation can be summarized on a plot showing the elemental rock compositions versus reaction progress (Fig. 9b). Silicon and aluminium remain stable throughout the simulation, as well as uranium. Sodium, potassium, and magnesium are successively leached out, as is calcium if it is abundant initially in the parent rock. If uraninite is present
Weathering, rainwater and atmosphere initially, however, calcium is partly preserved from leaching in uranophane. At the end of the simulation iron is still present and remains immobile but is very distinctly depleted (there is no significant volume change here) during all of the major part of the simu lated weathering, unlike in present profiles. Comparison with observed profiles
The result of this modelling (Fig. 9a) may be regarded as a simulated weathering profile comprising the fol lowing horizons (Fig. Sb) : 1 partially weathered granite; 2&3 middle gradually argillized [smectite + mus covite + kaolinite] horizon, with authigenic pyrite, locally containing siderite, and calcite at its base; 4 an upper argillized [kaolinite + quartz + pyrite] horizon. The comparison with the observed palaeoprofiles (Fig. Sa) is striking because the simulated profile reproduces the overall characters of horizons, and most of the peculiar features of the lower Huronian weathering: pyrite authigenesis and preservation, calcite precipitation at the profile's base, as well as the behaviour of chemical elements and especially the depletion of iron in the lower horizons. If present, uraninite also will be preserved, well into the argillized horizons (unlike in present profiles where it would disappear at the base of the profile). Uraninite as well as pyrite therefore would be easily available for reworking and preservation in overlying fluvial (unaerated) sediments. Uranophane is not a rnaj or ore mineral in the Elliot Lake area, but it may be regarded as an equivalent or a possible precursor for brannerite. A major discrepancy, however, appears between the observed and the simulated profiles, in the fact that sericite is found instead of kaolinite in the upper argillized horizons. This anomaly has been pointed out and discussed previously by several authors (Gay & Grandstaff 1980; Rainbird et a!. 1990; Roscoe et at. 1992; Prasad et a!. 1993; Schmitt & Thiry 1994), and seems actually to be a common feature of Protero zoic, and of some Palaeozoic, palaeosols (Feakes et a!. 1989; Holland et a!. 1989). The most likely explana tion of this problem is that kaolinite has formed in these profiles, but has been turned back to potassium bearing clays or micas during a later event. The nature of this event may be a K-metasomatic phase, as proposed by Rainbird et a!. (1990), but we suggest a simpler geochemical process. In the presence of water, kaolinite and K-feldspar form an unstable mineral assemblage that turns . to
37
[K-feldspar + muscovite] , or [muscovite + kaolinite]. Given the abundance of K-feldspar both in the Archaean parent rocks, and in the Huronian igneous rocks surrounding the weathered profiles, we think it is highly probable that a sufficiently long evolution of the system formed by the kaolinitized horizons, the K-feldspar bearing rocks, and the ground water has led to its re-equilibration and hence to sericitization of kaolinite. A second discrepancy is caused by the presence of siderite in the simulated profile, whereas it is absent from the actual profiles. This could be the result of a different effect of diagenesis on these old profiles. Geochemical modelling shows that the [siderite + kaolinite + quartz] assemblage is unstable at temper atures higher than 40°C, and that above 100°C it will produce chlorite following a reaction of the type: 5 siderite + kaolinite + quartz + 2 H20 --7 Fe5Al2Si3 0 10(0H)8 + 5 C02 Fe-chlorite
(4)
It is thus probable that the only notable discre pancies between simulated profiles and real sub Huronian profiles result from a rather simple diagenetic evolution (Fig. 8). This means that the atmosphere composition proposed by Kasting (1993) that we used in building rainwater model 1 (Table 5), accounts well for the observed weathering. Addi tional simulations have been performed in order to determine the range of atmospheric compositions that would still produce identical profiles. Constraints on early Huronian atmosphere's chemistry
The other main models of early Huronian rainwater and atmospheric composition that have been tested are given in Table 5. Model 2 enables us to test the effect of lower pC02s on simulated profiles. The C02 partial pressure has little effect on the stability of uraninite, but a direct effect on the rapidity of silicate weathering. From this simulation it follows that partial pressures of carbon dioxide below 0.1 bar are unlikely in the early Huronian because uraninite will not remain stable in the partly argillized horizons of the profile, making its reworking improbable in over lying sediments. Partial C02 pressures of 10-05 bar and higher are ruled out, because they will not permit calcite formation and preservation even in the lowest weathering horizons. They would also lead to a rather improbable acidity of rainwater and rapidity of weathering at this epoch.
38
J M. Schmitt
The oxygen content and reducing character of the atmosphere may be estimated from comparison of the simulation results obtained with rainwater models 1, 3 and 4 (Table 5). Atmospheres with pC02 around 0.1 bar, and somewhat richer in oxygen 2 (p02 10- .5 bar, model 3) lead to simulated profiles still close to those observed. Higher p02s are unlikely because they result in poorer stability of uraninite, and above all, in a rapid destruction of pyrite in the upper weathering horizons. For truly reducing atmospheres containing sizeable amounts of free H2S (model 4), similar profiles will be formed too. Uraninite will not be much better preserved (it still weathers slowly to uranophane unless rain water is practically free of calcium and silica). Pyrite on the other hand will become increasingly stable and with pH2S greater than 10-4 bar (Schmitt & Thiry 1994), will accumulate everywhere near the surface, a situation that is not really supported by field data. The formation of the early Huronian weathering profiles of Canada therefore agrees with a pC02 value close to 0.1 bar, which is consistent with the astro-climatic model of Kasting (1993). The con straints on the p02, however, are not so clear, because 2 oxygen-poor (p02 < 10- .5 bar, i.e. < 1 . 5 % of PAL) to truly reducing atmospheres apparently may result in similar profiles. Additional constraints of other types therefore are needed to better evaluate this crucial parameter of the atmosphere's evolution. We can conclude that the geochemical simulation of early Huronian weathering strongly supports recent models of Earth's atmosphere evolution, with early Proterozoic pC02 near 0.1 bar, and a low oxygen content (< 1.5% of PAL) . Although weakly constrained, the range of admissible p02s neverthe less agrees with the proposed transition from a reduc ing to an oxidizing atmosphere at that time (Kasting 1993), and still supports the general interpretation of the Huronian cycle given by Prasad & Roscoe (1991). =
CONCLUSIONS
I n this paper w e have tried t o summarize, and t o high light with a few examples, some of the strong links that bind together weathering, rainwater and Earth's atmosphere. We have focused on simple geo- and hydrochemistry and are well aware of the many aspects that have been neglected in this elementary approach.
The crude tentative attempt at modelling palaeo weathering profiles that we have described above shows the potential importance of chemical varia tions of the atmosphere for the development of weathering profiles. The existence of these variations, as well as their important amplitude, is now rea sonably established and there is no doubt that atmospheric compositions significantly differing from the present one have resulted in unusual (or unusually distributed) weathering profiles. This sensitivity of chemical weathering to atmos pheric conditions results from the existence in the weathering profile of several major chemical reaction fronts that behave more or less independently, as shown in the different simulations conducted above: 1 the weathering front (where only the most un stable minerals are destroyed) and the kaolinization front, the positions of which depend primarily on the carbonating power of rainwater, are hence essentially governed by the level of atmospheric C02; 2 the redox front of pyrite dissolution, the position of which depends on the 02 fugacity of rainwater, there fore is influenced directly by the oxygen content of the atmosphere; 3 the bauxitization front appears largely indepen dent of the atmosphere's chemistry, but is sensitive to both rainfall and temperature. These main reaction fronts (as well as minor ones corresponding to the stability limits of other miner als), migrate downwards at different speeds, thereby producing various types of weathering profiles. In this way, palaeoweathering profiles constitute a major record of Earth's surficial conditions. We should not overlook the fact that some palaeoprofiles may have registered only local anomalies (e.g. in drainage conditions, in groundwater chemistry, and so on), and that the more ancient ones may have been altered significantly during diagenesis or metamor phism. The weathering record (together with the sed imentary record and the palaeontological record) is nevertheless one of the best indicators of Earth's past environments. As shown briefly here, and demon strated by other papers in this volume, palaeogeo graphical distribution, petrography, chemistry and stable isotope chemistry of pedogenic features, are able to record information on palaeotemperatures, palaeoclimate, palaeoabundance of atmospheric gases and so on. A final aspect of the links between weathering and the atmosphere that may be suggested from the pre ceding simulations, is the part played by weathering in regulating Earth's surficial conditions. We have
Weathering, rainwater and atmosphere seen, for instance, from the Cretaceous example that increased atmospheric C02 results in increased weathering, especially at higher latitudes. As silicate weathering essentially is a carbonation reaction, we can conclude that the higher the C02 atmospheric level, the higher its consumption by weathering. This regulating process, or negative feedback (Yolk 1987), is well illustrated in Berner's model of the C02 cycle (Berner 1991, 1994). Weathering also has a negative feedback on mean global temperature (Walker et al. 1981). If for example the mean surface temperature falls sufficiently, the atmospheric humidity and pre cipitation consequently will drop. The weathering rate of surface rocks will slow down correspondingly, and hence so too will the uptake of C02 from the atmosphere (Brady 1994; Blum & White 1995). As C02 is the main greenhouse gas, the following increase in the C02 atmospheric level would, in turn, increase the mean surface temperature. Weathering thus may be considered as one of the major regulat ing agents of both atmospheric composition and global mean temperature. Not forgetting the many other aspects of weather ing (and especially its complex links with the terrestrial biota), we shall maintain that the palaeo weathering record is one of our main keys to past environments. Further progress in the study of palaeoweathering features, and in modelling the peculiar geochemical systems that weathering profiles represent, will certainly bring more con straints and valuable landmarks to our understanding of the history and evolution of our planet.
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AUSTIN, G.S. (1970) Weathering of the Sioux quartzite near New Ulm, Minnesota, as related to Cretaceous climates. J. sediment. Petrol., 40(1), 184-193. BARDOSSY, Gv. (1981) Paleoenvironments of laterites and lateritic bauxites - effect of global tectonism on bauxite formation. In: Lateritisation Processes (Project IGCP129) Proceedings of the International Seminar on Lateriti sation Processes, 1979, Trivandrum, India, pp. 287-294. Balkema, Rotterdam. BARRON, E.J. & WASHINGTON, W.M. ( 1982) Cretaceous climate: a comparison of atmospheric simulations with the geologic record. Palaeogeogr. Palaeoclimatol. Palaeoecol. , 40, 103-133. BARRON, E.J. & WASHINGTON, W.M. (1985) Warm Cretaceous climates: high atmospheric C02 as a plausible mecha nism. In: The Carbon Cycle and Atmospheric CO,: Natural Variations Archean to Recent (Eds Sundquist, E.T. & Broeker, W.S. ) . Geophys. Monogr., Am. geophys. Union, 32, 546-553. BARRON, E.J., FAWCETT, P.J., POLLARD, D. & THOMPSON, S.L. (1993) Model simulations of Cretaceous climates: the role of geography and carbon dioxide. Philos. Trans. R. Soc. London, Ser B, 341, 307-315. BERNER, R.A. (1991) A model for atmospheric C02 over Phanerozoic time. Am. J. Sci., 291, 339-376. BERNER, R.A. (1994) GEOCARB II: a revised model of atmospheric C02 over Phanerozoic time. Am. J. Sci., 294, 56-91. BERNER, R.A., LASAGA, A.C. & GARRELS, R.M. (1983) The carbonate-silicate geochemical cycle and its effect on atmospheric carbon dioxide over the past 100 million years. Am. !. Sci., 283 , 641-683 . B ETHKE, C. (1994) The Geochemist's WorkbenchTM. A User's Guide to Rxn, Act2, Tact, React and Gtplot. University of Illinois, Urbana-Champaign. BETHKE, C. (1996) Geochemical Reaction Modelling: Concept and Applications. Oxford University Press, New York. BIRD, M.l., CHJVAS, A.R., FYFE, w.s. & LONGSTAFFE, F.J. (1990) Deep-weathering at extra-tropical latitudes: a response to increased atmospheric C0 2. In: Soils and the Greenhouse Effect (Ed. Bouwman, A.F. ) , pp. 383-389. John Wiley & Sons, Chichester. BLUM,A.E. & WHITE,A.F. (1995) The effect of global climate on silicate weathering rates; impact of present day green house warming and Cretaceous climatic changes. Geol. Soc. Am. Abstr. Program , 21(6) , 237 . BouwMAN, A.F. (1990) Exchange of greenhouse gases between terrestrial ecosystems and the atmosphere. In: Soils and the Greenhouse Effect (Ed. Bouwman,A.F. ) , pp. 61-127. John Wiley & Sons, Chichester. BRADY, P.V. (1994) Looking behind the wizard's curtain: mineral surface controls on long term climate. Geol. Soc. Am. Abstr. Program, 26, a-287. CERLING, T.E. (1991) Carbon dioxide in the atmosphere: evi dence from Cenozoic and Mesozoic paleosols. Am. J. Sci., . 291, 377-400. CLOUD, P. (1972) A working model of the primitive Earth. Am. J. Sci., 272, 537-548. CoMBES, P. (1969) Recherches sur Ia genese des bauxites dans le nord-est de l'Espagne, le Languedoc et 1'Ariege (France). Mem. Centre d'Etudes Recherch. Geol. Hydro geol., m-IV, 335.
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CRAVERO, M.F. & DoMINGUEZ, E.J. (1992) Evidencias de una paleosuperficie de meteorizacion durante el Berriasiano Aptiano en Ia Patagonia Argentina. Cuarta Reun. Argent. Sediment., ID, 185-192. DRIESE. S.G. & MoRA, l.M. (1993) Introduction and primary objectives of field excursion. In: Paleosols, Paleoclimate, and Paleoatmospheric CO"· Paleozoic Paleosols ofCentral Pennsylvania (Ed. Driese, S.G. ). Univ. Tenn. Dept Ceo!. Sci., Stud. Ceo!., 22, 22-29. DuPUIS, CH. (1992) Mesozoic kaolinized giant regolith and Neogene halloysitic cryptokarst: two striking paleoweathering types in Belgium. In: Mineralogical and Geochemical Records of Paleoweathering (Eds Schmitt, J.M . & Gall, Q. ). Ecole Nationale Superieure des Mines de Paris Mem. Sci. Terre, 18, 61-68. DuRY, G.H. & KNox, J.C. (1971) Duricrusts and deep weathering profiles in southwestern Wisconsin. Science, 174, 291-292. FEAKES, C.R., HOLLAND, H.D. & ZBINDEN, E.A. (1989) Ordovician paleosols at Arisaig, Nova Scotia, and the evolution of the atmosphere. In: Paleopedology: Nature, and Application of Paleosols (Eds Bronder, A. & Catt, J.A. ). Catena Suppl., 16, 207-232. FRAKES, L.A. (1979) Climates Through Geologic Time. Else vier,Amsterdam. FRITZ, B. (1981) Etude thermodynamique et modelisation des reactions hydrothermales et diagenetiques. Memoires Sci ences Geologiques, Universite Louis Pasteur, Strasbourg, France, 65-197. GAY, A.L. & GRANDSTAFF, D.E. (1980) Chemistry and miner alogy of Precambrian paleosols at Elliot Lake, Ontario, Canada. Precam. Res., 12, 349-373. GILLETT, S.L. (1985) The rise and fall of the early reducing atmosphere. Astronomy, 66-71. GOLDICH, S.S. (1938) A study in rock weathering. !. Ceo!., 46, 17-58. GRANAT, L. (1972) On the relation between pH and the chemical composition in atmospheric precipitation. Tellus, 24, 550-560. GRUNBERGER, 0. (1989) Etude geochimique et isotopique de !'infiltration sous climat tropical contraste -Massif du Piton des Neiges -Ile de Ia Reunion. These, Doctoral es Sciences, Universite Paris XI, Orsay. HEM, J.D. (1985) Study and Interpretation of the Chemical Characteristics of Natural Water, 3rd edn. U.S.geol.Surv. Water-Supply Pap., 2254, 263. HoLLAND, H.D. (1978) The Chemistry of the Atmosphere and Oceans. Wiley Interscience, New York. HoLLAND, H.D. (1984) The Chemical Evolution of the Atmosphere and Oceans. Princeton University Press, Princeton. HoLLAND, H.D., FEAKES, C.R. & ZBINDEN, E.A. (1989) The Flin Flon Paleosol and the composition of the atmos phere 1.8 by b. p. Am. J Sci., 289, 362-389. KASTING, J. (1993) Earth's early atmosphere. Science, 259, 920-956. KASTING, J.F., TooN, O.B. & PoLLACK, J.B. (1988) How climate evolved on the terrestrial planets. Sci. Am .,258(2), 46-53. KEMPER, E. (1987) Das Klima der Kreide Zeit. Ceo!. Jahrb., Reihe A, 96,399. LICHTNER, PC. (1988) The quasi-stationary state approxima tion to coupled mass transport and fluid-rock interaction
in a porous medium. Geochim. Cosmochim. Acta, 52, 143-165. LOVELOCK, J.E. & WHITFIELD, M. (1982) Lifespan in the bios phere. Nature, 296, 561-563. MEYBECQ, M. (1984) Les Fleuves et le Cycle Geochimique Des Elements. These, Doctoral es Sciences, Universite Paris VI, Paris. MILLOT, G. (1970) Geology of Clays: Weathering, Sedimen tology, Geochemistry. Springer-Verlag, New York. MoHAN, L. JAMWAL, J.S. & NANDA, M.M. (1981) Bauxite deposits of Jammu, India. In: Lateritisation Processes (Project IGCP-129) Proceedings of the International Seminar on Lateritisation Processes, Trivandrum, India, pp. 190-192. Balkema, Rotterdam. MORA, C.J., DRIESE, A.G. & SEAGER, PG. (1991) Carbon dioxide in the Paleozoic atmosphere: evidence from carbon-isotope compositions of pedogenic carbonate. Geology, 19, 1017-1020. OHMOTO, H. (1996) Evidence in pre-2.2 Ga paleosols for the early evolution of atmospheric oxygen and terrestrial biota. Geology, 24(12), 1 135-1138. PARHAM, WE. (1970) Clay Mineralogy and Geology of Min nesota's Kaolin Clays. Minnesota Geological Survey, Min neapolis. POLLACK, J.B. (1990) Atmospheres of the terrestrial planets. In: The New Solar System , 3rd edn (Eds KELLY Beatty, I & Chaikin, A.), pp. 91-106. Sky Publishing Corporation, Cambridge. PRASAD, N. & RoscoE, S.M. (1991) Profiles of altered zones at ca 2.45 Ga unconformities beneath Huronian strata, Elliot Lake, Ontario: evidence for early Aphebian weath ering under anoxic conditions. In: Current Research, Part C , Paper 91-1 C, pp. 43-54. Geological Survey of Canada, Ottawa. PRASAD, N. & RoscoE, S.M. (1996) Evidence of anoxic to oxic atmospheric change during 2.45-2.22 Ga from lower and upper sub-Huronian paleosols, Canada. Catena, 27, 105-121. PRASAD, N., ROBERTSON, J.A. & BENNETT, G. (1993) Paleo weathering, Paleosurfaces and Precambrian Stratigraphy, Elliot Lake-Thessalon Area, Ontario. IGCP 317 Field Trip Guide Book. Third International Geomorphology Conference, Hamilton, Canada. RAINBIRD, R.H., NESBITT, H.W. & DoNALDSON, J.A. (1990) Formation and diagenesis of a sub-Huronian saprolith; comparison with a modern weathering profile. J Ceo!. , 98, 801-821 . RoscoE, S.M. (1957) Geology and uranium deposits, Quirke Lake-Elliot Lake, Blind River area, Ontario. Ceo!. Surv. Can. Pap. , 56-7. RoscoE, S.M. (1969) Huronian rocks and uraniferous con glomerates. Ceo!. Surv. Can. Pap., 68-40. RoscoE, S.M., THIERAULT, R.J. & PRASAD, N. (1992) Circa 1.7 Ga Rb-Sr re-setting in two Huronian paleosols, Elliot Lake, Ontario and Ville Marie, Quebec. In: Radiogenic Age and Isotopic Studies Report 6, Paper 92-2, pp. 119-124. Geological Survey of Canada, Ottawa. SAGAN, C. & MULLEN, G. (1972) Earth and Mars: evolution of atmospheres and surface temperatures. Science, 177(4043), 52-56. SAPOJNIKOV, D.G. (1981) Lateritic formations of the U.S.S.R. In: Lateritisation Processes (Project IGCP-129) Pro ceedings of the International Seminar on Lateritisation
Weathering, rainwater and atmosphere Processes, Trivandrum, India, pp. 185-189. Balkema, Rotterdam. ScHMIDT, P.W. & 0LLIER, C.D. (1988) Palaeomagnetic dating of late Cretaceous to early Tertiary weathering in New England, N.S.W.,Australia. Earth Sci. Rev., 25, 363-371 . ScHMITT, J.M. ( 1994) Geochemical modeling and origin of the Triassic albitized regolith in Southern France. 14th International Sedimentology Congress, Aug 20-28, Recife, Brazil, Abstracts S8, 19-21. ScHMITT, J.M. & THIRY, M. (1994) Simulation of granite weathering for various compositions of the atmosphere and importance for paleoweathering interpretation. 14th International Sedimentology Congress, Aug 20-28, Recife, Brazil, Abstracts S8, 21-23. SCHNEIDER, S.H., THOMPSON, S.L. & BARRON, E.J. (1985) Mid Cretaceous continental surface temperatures: are high C02 concentrations needed to simulate above-freezing winter conditions? In: The Carbon Cycle and Atmos pheric C02: Natural Variations Archean to Recent (Eds Sundquist, E.T. & Broeker, W.S.). Geophys. Monogr., Am. geophys. Union, 32, 554-559. ScoTESE, C.R., VAN DER Voo, R. & Ross, W.C. (1981) Meso zoic and Cenozoic base maps (abstract). Bull. Am. Assoc. petrol. Geol., 65,989. SEQUEIRA, R. & KELKAR, D. (1978) Geochemical implica tions of summer monsoonal rainwater composition over India. J. Appl. Meteorol., 17, 1390-1396. SIGLEO, W.R. & REINHARDT, J. (1988) Paleosols from some Cretaceous environments in the southeastern United States. In: Paleosols and Weathering Through Geologic Time: Principles and Applications (Eds Sigleo, W.R. & Reinhardt, I ). Geol. Soc. Am. Spec. Pap., 216, 123-142. SLADEN, C.P. (1983) Trends in Early Cretaceous clay miner alogy in NW Europe. Zitteliana, 10, 349-357. SLANSKA, J. (1976) A red-bed formation in the South Bohemian B asins, Czechoslovakia. Sediment. Geol., 15, 135-164.
41
STEINMANN, P., LICHTNER, P.C. & SHOTYK, W. (1994) Reaction path approach to mineral weathering reactions. Clays Clay Mineral., 42(2) , 197-206. SUBRAMANIAN, K . S. & MAN ! , G. (1981) Genetic and geomor phic aspects of laterites on high and low landforms in part of Tamil Nadu, India. In: Lateritisation Processes (Project JGCP-129) Proceedings of the International Seminar on Lateritisation Processes, Trivandrum, India, pp. 237-245. Balkema, Rotterdam. SuTTON, S.J. & MAYNARD, J.B. (1993) Sediment- and basalt hosted regoliths in the Huronian Supergroup: role of parent lithology in middle Precambrian weathering profiles. Can. J. Earth Sci., 30, 60-76. TARDY, Y., KoBILSEK, B., RoQUIN, C. & PAQUET, H. (1990) Influence of Periatlantic climates and paleoclimates on the distribution and mineralogic composition of bauxites and ferricretes. Chem. Geol., 84, 179-182. THIRY, M. (1981) Sedimentation continentale et alterations associees: calcitisation, ferruginisation, et silicification. Les Argiles Plastiques du Sparnacien du Bassin de Paris. Memoires Sciences Geologiques, Universite Louis Pasteur, Strasbourg, France 64, 173. YOLK, T. (1987) Feedbacks between weathering and atmos pheric C02 over the last 100 million years. Am. J. Sci., 287, 763-779. WALKER, J.C.G. (1982) Climatic factors on the Archean Earth. Palaeogeogr. Palaeoclimatol. Palaeoecol., 49, 111. WALKER, J. C.G. , HAYS, P.B. & KASTING, J. F. (1981) A negative feedback mechanism for the long-term stabilization of Earth's surface temperature. J. geophys. Res., 86(C10) , 9776-9782. YAPP, C.J. & POTHS, H. (1996) Carbon isotopes in continen tal weathering environments and variations in ancient atmospheric C02 pressure. Earth planet. Sci. Lett., 137, 71-82.
Spec. Pubis int. Ass. Sediment. (1999) 27, 43-60
Stable carbon isotopes in palaeosol carbonates
T. E . C E R L I N G Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112, USA
A B S T R ACT The stable carbon isotope composition of carbonates formed in soils can be modelled using a steady-state diffusion-production equation. This model predicts an enrichment of 13C in soil C02 compared with soil respired C02, which results from the difference in diffusion coefficients of l2C02 and 13C02 and from the influence of the atmosphere. Carbon isotope studies of modern soils show that the specific predictions of the diffusion-production model are fulfilled, giving confidence to predictions made by the model that are not readily testable. The diffusion-production model was developed for soils where mass transport is controlled by gaseous diffusion and should not be applied to other conditions. It assumes isotopic equilibrium between all oxi dized carbon species, which means it cannot be applied to soils (or palaeosols) where there is an inherited detrital component. For modern soils, the conditions of diffusion-controlled mass transport are met in those soils with high free-air porosity but not with those that are saturated with water. Identification of palaeosols meeting these conditions is more difficult than simply identifying palaeosols, so that care must be taken in establishing the character of palaeosols where this model is used. The palaeoenvironmental interpretations of carbon isotopes in palaeosols include estimates of the fraction of C4 biomass in soils, which is very useful in the Neogene. An example from Pakistan shows the change from a C3-dominated to a C4-dominated biomass between about 7 and 5 Ma. Another important application of carbon isotopes in palaeosols is to the study of paleo-pC02 levels of the atmosphere. Because the solution to the diffusion-production model includes atmospheric C02 as one of the boundary conditions, studies of palaeosol carbonate can give estimates of ancient atmospheric C0 2 levels. Prelimi nary studies of palaeosols indicate that atmospheric C0 2 levels in the Mesozoic were between 2000 and 3000 p.p.m.
I NT R O D U CT I O N
carbonate is a measure of soil productivity and of the fraction of C4 biomass in soils. C4 plants, which pri marily are grasses and sedges (although some forbs and shrubs use this photosynthetic pathway), have a 8I3C value of about -1 1 to -13%o (Deines, 1980) and generally are restricted to regions that are hot during the growing season (e.g. temperate to tropical grass lands and savannas to deserts). In contrast, C3 plants, which include cool-season grasses and most shrubs and trees, have 8BC values between about -24 and -30%o (Deines, 1980) and make up diverse ecosys tems, from the tropics to the high latitudes (e.g., all forests, Mediterranean climates, cool-temperate to boreal grasslands). It appears that C4 biomass has been significant in ecosystems only in the Neogene (Cerling et al. , 1993), where soil carbonates have been
The stable carbon isotope composition o f paleosols is a powerful tool for palaeoenvironmental studies. It is useful in determining the amount of recrystallization of carbonate soils in very early soil formation, it can be used to estimate the fraction of C3 versus C4 biomass in fossil soils, it is sensitive to the total pro ductivity of soils and so may indicate aridity, and it can be used as a barometer of pC02 in pre-Neogene soils. Carbon isotopic ratios in soil carbonates are deter mined by the fraction of C4 biomass in the local ecosystem, by the influence of atmosphere on C02 in soil gas and by temperature (Cerling, 1984, 1991; Cerling & Quade, 1993). For periods when the pC02 level of the atmosphere is relatively low (less than 1000p.p.m.) the carbon isotopic composition of soil
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
43
44
T E. Cerling
used to estimate the fraction of C4 biomass in a variety of places (Quade et al. , 1989a; Cerling 1992a; Mack et al. , 1994; Cole & Monger, 1994). The atmospheric influence on total soil C02 is impor tant because the atmospheric component can be a relatively high proportion of total C02 in low productivity soils or in soils formed during periods of high atmospheric C02 levels (> 1000 p.p.m.; Cerling, 1991c). Cerling (1984) derived a diffusion production model to describe the isotopic composi tion of soil C02 and the isotopic composition of pedogenic carbonate precipitated in isotopic equilib rium with soil C02. This model has been used to study low-productivity desert soils (e.g., Amundson et at. , 1988a,b; Quade et al. , 1989b; Pendall e t al. , 1994) and to estimate the pC02 content of the Palaeozoic and Mesozoic atmosphere (Cerling, 1991c; Mora et al. , 1991, 1996; Yapp & Poths, 1992, 1994, 1996). The diffusion-production model was also used to distin guish pedogenic carbonate from spring carbonate (Quade & Cerling, 1990). It is the purpose of this paper to outline the use of carbon isotopes in palaeosols. I will first discuss some early models of carbon isotopes in soils and the theory of gas transport in soils, show a diffusion production model for C02 transport in soils, use modern soils to show how this model works, and discuss implications and limitations of using carbon isotopes for palaeoenvironmental studies. I will show examples of carbon isotope studies of palaeosols in studying the early history of carbonate remobiliza tion is soils, in reconstruction of the fraction of C4 biomass, in studies of arid soils and its use in studying the history of atmospheric pC02 levels.
P R E C I P I TAT I O N O F P E D O G E N I C C A R B O NATE
The carbon isotopic composition of soils was first modelled (Salomons et al. , 1976) using the model for carbonate formation resulting from C02 degassing that was developed by Hendy (1971) for speleothem formation, where the system is characterized by the reaction: (1) The carbon is derived from the rock and from biolog ical C02, with the solubility product: Kcal ci te -
a a a c c •' � o '-
pco,
(2)
where Kcalci te' aca+2, a HCO) ' and P co, are the solubility product of calcite, the activity of the Ca+2 ion, the activity of the HC03 ion, and pC02, respectively. Reaction (1) was later used to determine the amount of recrystallization in soils developed on a carbonate substrate by estimating the fraction of pedogenic versus detrital carbonate (Magaritz & Amiel, 1980). With the realization, however, that the bio logical component of C02 in soils was much larger than the amount of detrital carbonate dissolved on soil-formation time-scales (C02(biological) » CaC03 ( detrital), in the order of 100 to 1000 times) it was clear that the soil C02-H20-Ca0 system for the isotopic species should be modelled using diffusion theory (Cerling, 1984). The diffusion theory for C02 transport in soils is well established (Baver et al. , 1972; Kirkham & Powers, 1972) and the diffusion model is extended to include transport of 1 2 C02 and 1 3C02 in the soil atmosphere. It is clear from reactions (1) and (2) that soil car bonate can form in several ways. 1 Degassing of the soil solution. Soil pC02 profiles generally increase with depth so that downward movement of soil solutions increases solubility rather than decreases solubility. Upward movement of soil solutions is important in groundwater discharge zones, such as prairie pothole soils in glacial terrains, or by capillary rise. Degassing also occurs as soils decrease respiration rates seasonally as a result of decreased temperature or soil moisture. 2 Increase in ion activity as a result of evaporation or evapotranspiration. Direct evaporation of soil solu tions is important in soils with bare ground exposed. Soils with 100% vegetation cover, however, do not always show evaporative enrichment of stable oxygen isotopes in soil waters (Hsieh et al. , 1998). Evapotranspiration (water uptake by roots with evaporation taking place at the leaf surface to the free atmosphere) is the dominant form of water loss at depth in most soils. This results in an increase in ion 2 activity of Ca+ and HC03 and leads to carbonate for mation. 3 Increase in ion activity as a result of ion exclusion during freezing. Ion exclusion of salts is a well-known property of water during freezing. It does not appear to be a major contributor to soil carbonate formation. 4 Retrograde solubility of calcite. Calcite is more soluble at low temperatures than high temperatures. Therefore, changes in temperature result in changes in the solubility of calcite. Although C02 degassing is an important mecha nism for carbonate precipitation in some environ-
45
Stable carbon isotopes ments, it is Jess important than evapotranspiration in soils. Observations on modern soils show that the zone of carbonate accumulation generally is deep (> 30 em) within the soil. Soil solutions moving down ward encounter higher levels of C02 at depth. Thus, C02 degassing is less important than evapotranspira tion. Water loss across the cell membrane, with selec 2 tion against Ca+ and carbonate species, causes local increases in the ion activities of Ca+2 and COi which leads to carbonate precipitation in the growing season when pC02 levels are high, rather than in the autumn or winter when pC02 levels decrease. The rates of carbonate precipitation are low and the rates of exchange of carbon-bearing dissolved species 2 (C02, H2C03 , HC03, and C03 ) are high, so it is expected that there is isotopic equilibrium between the carbon-bearing species. Formation of pedogenic carbonate around roots and in areas of high root mass is compatible with evapotranspiration driving carbonate precipitation in soils. Figure 1 shows the isotopic composition of four carbonate soils developed on Holocene parent mate rial in central North America; the stable carbon iso topic composition at depth in these soils has a small range, although near the Earth's surface it is a strong function of depth because of diffusion processes. It has been argued previously that inheritance of BC from carbonate dissolution of detrital carbonate should buffer the oBC of pedogenic carbonate. The annual flux of biological C02, however, is of the order 2 of 10-3 mol cm- yr-1, which is much higher than rates of carbonate accumulation, typically 10-6 to 10-5mol cm-2 yr-1 (Cerling, 1984). Quade et al. (1989b) tested this idea by analysing oBC from pedo genic carbonate along two elevation transects in western North America; one had parent material derived from Palaeozoic limestones, whereas the other was derived from Tertiary volcanics. Inheri tance of detrital carbonate would attenuate the isotope signal because of addition of a carbonate fraction of uniform isotopic composition. Quade et al. (1 989b ), however, found that the oBC of pedogenic carbonate had the same oBC m-1 gradient for both suites of soils over an elevation difference of almost 2500 m, indicating that inheritance of carbon isotopes during the dissolution of detrital carbonate in the soil zone does not occur for these soils. Marion et al. (1991) found that pedogenic carbonate formed very quickly in Alaskan soils (less than a few hundred years) and the isotopic values of the new carbonate did not show evidence of inheritance of carbon from dissolution of the detrital phase. Pendall et al. (1994)
0 ao 00 0
50
8
0
E
__.._
u
..c:
'--'
OlD
1 00
••
0.. 0.)
.t.Saskatchewan
a
•
(
• Kansas
! 50
• •
e !owa ONevada
200 -15
-10
-5
0
o 1 3 C (pedogenic CaC0 3 )
5
Fig. I. (ii 3 C of pedogenic CaC0 from four modern soils 3 in North America. All soils were developed on Holocene to late pleistocene parent material. Uppermost point for each profile is the first carbonate encountered in the profile. The profiles show fairly constant values at depth, but the one profile from Nevada shows a strong (ii 3 C gradient in the upper 30 em of the profile.
dated incipient pedogenic carbonate with ages less than 1000yr, also arguing against inheritance of carbon during dissolution-precipitation. Detrital car bonate that is not dissolved but is occluded in pedo genic carbonate, or that dissolves in a system where the carbon budget is not overwhelmed by the biologi cal signal, however, are important inheritance prob lems. Likewise, the misidentification of detrital carbonate as pedogenic carbonate could lead to erro neous conclusions.
M O D E L F O R M A S S T R A N S P O RT I N D I F F U S I O N - C O NT R O L L E D SYSTEM C02 in soils
The isotope model for C02 transport in soils has been described previously (Cerling, 1984; Quade et al. , 1989b; Cerling & Quade, 1993) and has been extended to include terms for radioactive decay (Wang et al. , 1 994). Mass transport of C02 in soils is described by (Baver et al. , 1972; Kirkham & Powers, 1972):
T E. Cerling
46
(3) where c: is the concentration of soil C02 without iso topic distinction (mol cm-3), t is time (s), D';' is the dif fusion coefficient for C02 in soils (cm2 s-1), and
air, £ is the free-air porosity, and p is a tortuosity factor (typically 0.8 to 0.6 for soil porosities from 0.5 to 0.25). Using a Dco,-air of 0.14 cm-2 s-1, £ of 0.25 and 2 p of 0.6 gives u;' 0.02 cm- s-1. The production rate of C02 is taken to be proportional to the organic matter concentration in soils because C02 is pro duced by microbial oxidation of organic carbon and by root respiration. Organic matter in soils typically decreases exponentially in soils: =
(5) where z is the characteristic depth of C02 produc tion. The characteristic depth of C02 production has been measured in situ (Gaudry et al. , 1990; Dorr & Mtinnich, 1990) and is of the order of 10-20 cm for forest soils; the organic matter concentration in soils decreases exponentially with depth with characteris. tic depths of the order of 10-20 cm. The characteristic response time: * -z 2 (6) = *
'ts
boundary. Most soil C02 profiles show the character istic profile that approaches (aC�/az) 0 (e.g. Fig. 2). Exceptions to this shape of profile can be found in soils with very deep unsaturated zones (> c. 20 m); in these soils significant degassing takes place in the non-growing season so that in the early part of the growing season a local high concentration at shallow depths occurs, which dissipates as the growing season progresses (Reardon et al. , 1979). Using these bound ary conditions the solution to the diffusion equation is of the form: =
(10) where S(z) is the specific solution to the diffusion equation. For the case of an exponentially decreasing C02 production function:
S(z) is: S (z)
=
Equation (10) has some interesting implications concerning modern soils and ancient soils. In the steady-state case all soils have increasing C02 with depth. Figure 3 shows C02 profiles for soils with different respiration rates, ranging from 2 to 2 8 mmol m- h-1; porosities ranging from 10 to 40 % ,
20
aJr 0
soil
s
(7) with the boundary conditions:
c: c:' at z =
=
o
-20
E
-40
0..
-60
,--.._
u
.c
'-./
-80
(8)
and (9) (act ja z) 0 at the bottom of the soil where c; is the concentration of C02 in the air. These =
assumptions have the practical meaning that the con centration of C02 is continuous across the soil-air interface, and that the lower boundary is a no-flux
(12)
s
D
is of the order of hours (where D'; is c. 0.02cms-1). Therefore, we use the steady-state case for C02 diffu sion in soils:
(11)
=
· 1 00 100
Brighton, Utah 4.1 mmole/m 2 /hr
9
Sept. 84
0.5 porosity
0
model measured
1000
1 0000
C02 (ppmY) Fig. 2. Soil C02 profile diffusion control of mass transport. The boundary conditions c;'= qir at z 0 and (Cl C;'/Clz) 0 at the bottom of the soil appear to be met in this soil. Data from Solomon & Cerling, 1987. =
=
Stable carbon isotopes and characteristic depths of C02 production from 1 to 40 cm. Soil respiration rates during the growing season generally range from about 5 to 10 mmol m-2 h-1, with occasional higher values being mea sured in agricultural areas, and lower rates being measured in water-stressed regions such as deserts and in the winter or non-growing season (Singh & Gupta, 1977; Schlesinger, 1977; Parker et al. , 1983; Dorr & Mtinnich, 1987; Gaudry et al. , 1990; Solomon & Cerling, 1987; Quade et al. , 1989b). At low soil res piration rates the C02 concentration in soils is quite low, as shown in Fig. 3. The effect of porosity is very important and low free-air porosities, such as occur in
( 0 )
0
soils with high water content, have very high pC02 levels, leading to significant oxygen depletion and often to iron reduction. The characteristic depth of C02 production is important in considering the con ditions that could lead to weathering of silicate min erals, which today is driven in part by the high soil pC02 levels, before the advent of vascular plants with rooting systems. If the soil respiration is limited to the upper few centimetres of the soil then the contribu tion of the biological component of soil C02 (the S( z) term) is very small (Fig. 3). It is also significant to note that the specific form of the production function (e.g., exponentially decreasing with depth, linear decrease
( b )
z = 20 em
E
50
20%
0.. �
100
Cl
2
-.....£
a. 0)
4
40%
150
._ .... .... . ....___. _._ ._ ._ ---1 .. _...._... _ ,._ ___. _ .... .... ._ .. _._
1 0000
5000
100
Cl
8
150
200
E' u
Respiration Rate
; (mmolfm2fhr)
,...._
..c::
0
1 0°1i
50
u -... ..-
47
pC0 2 ppmV
200
; 20000
40000
=
z
8 mmol/m 2/hr =
60000
20 em 80000
1 00000
pC02 ppmV
; = 8 mmol/m2/hr £ = 40%
50
E
,...._
u -... ..£
a. �
Cl
100
10
20
z
(em)
40
1 50
5000
1 0000
pC0 2 ppmV
1 5000
20000
Fig. 3. Variations in pC02 concentration for soils of
differing porosity, soil respiration rates and characteristic depth of C02 production.
48
T E. Cerling
with depth, etc.) has little effect on the C02 profiles for equations with the same average depth of C02 production (Solomon & Cerling, 1987); for the exponentially decreasing case, Z average = z/0.693. For well aerated non-agricultural soils,pC02 levels in soils generally reach values between about 5000 and 9000 p.p.m. in temperate to subtropical soils (e.g., Solomon & Cerling, 1987; Brook et al. , 1983). The pC02 levels in desert soils tend to be lower, often less than 5000p.p.m. during the growing season (e.g., Quade et al. , 1989b). In water-saturated soils, pC02 levels can be very high with oxygen depletion and iron reduction, as mentioned above. It is clear from this modelling exercise that pC02 levels in soils can be quite variable from soil to soil depending on depth in the soil, seasonality, soil texture including porosity, soil moisture, soil productivity, and distribution of organic matter in the soil. Fortunately, except for waterlogged soils, some of the factors governing soil C02 levels tend to work against each other. For example, high productivity levels deplete the soil in available moisture, causing an increase in porosity and a decrease in productivity. Because of this self regulating property of soil, in the ensuing discussion we will consider that soils have growing season bio logical component (S ( z)) values between about 5000 and 9000 p.p.m. at depths > 50 em in high productivity soils and between 3000 and 5000 p.p.m. in low-pro ductivity soils, unless specific conditions suggest oth erwise. The concentration of soil C02 during the period of pedogenic carbonate formation, however, has not been determined for any soils. An interesting thought exercise is to consider soils in the world before vascular plants with well devel oped rooting systems. Such soils may have had a veg etative cover similar to the modem cryptogamic soils that stabilize many arid and semi-arid landscapes. Cryptogamic soils have most of their organic matter in the upper few centimetres of the soil, so that the characteristic depth of C02 production is small, perhaps of the order of 1 em or so. Figure 3 shows that C02 production at such a shallow depth does not lead to a significant increase in C02 with depth in the soil. Yapp & Poths (1994), however, show evidence that the biological component of C02 in Ordovician (pre vascular plants) soils was high. This presents some thing of an enigma that still needs to be explained. Carbon isotopes in soils
The principles of C02 transport in soils also applies to 2 the different species of C02, namely 1 C02, 13C02
and 14C02 (Cerling, 1984; Cerling & Quade, 1993; Wang et al. , 1994). The equation for the stable carbon species is:
a cns . a2 c_ sn + <)> n ( z) = D ·I· s a z2 s at
__
_ _
(13)
and for 14C is:
aCJ 4 at
_s_
=
. a2 C 14
D "· __ s_ + <)> 14(z ) - A,1 4C1 4 s s s a z2
(13a)
where the superscript n refers to the 12 C02, 13C02 or 14C02 species, and 1..1 4 is the decay constant of 14C. In a stable ecological setting, the stable isotopes have the convenient property of having a constant isotope production ratio PfF- Once again, this allows the convenient assumption of a steady-state condition, so that
a(cp ;cp ) =0 at
(14)
This is not the case, however, for the radioisotope 14C02 because the 14Cf12C ratio of organic matter in the soils increases with time, at least in the early stages of the soil. The steady-state assumption is suit able to describe the short-term 14C02 distribution in soil, but is not suitable to describe the distribution of 14CaC03 in soils, which is important for the applica tion of 14C dating to pedogenic carbonate (see Wang et al. , 1994, 1996). The evolution of 14C in organic matter in soils has been modelled as (Trumbore et al. , 1990; Amundson et al. , 1994; Wang et al. , 1996):
:. c* at a q� = "'14 14 + A.14)COM 'i'OM - ( kOM at
� - *OM - kOMc*OM
(15) (16)
where C()M and qjM are the carbon and 14C content of soil organic matter (mol cm-3), k0M is the first order oxidation rate of organic matter (s-1), and <)>()M and <)>�M are the production rates of organic matter and of 14C in soils (mol cm-3 s-1). The soil organic matter concentration reaches steady-state rapidly in soils, but the C�M does not reach steady state quickly. This approach to modelling the age of pedo genic carbonate and the occluded organic matter in soils, however, is a very fruitful approach to dating Holocene and late Pleistocene palaeosols (Wang et al. 1996). In the remainder of this paper I will confine my remarks to the stable carbon isotopic composition of soils.
Stable carbon isotopes The stable isotopes are related by: 813Ci =
(___&___ ] RPDB
- 1 1000
49
eter for the different isotope species is constant, so that:
where Ri - ( B Cj 12C)i (17)
and where i is the sample and PDB is the isotope ref erence standard for carbon. Solution of the appro 2 priate diffusion equation for the 1 3C02 and 1 C02 species gives ( Cerling, 1984; Cerling & Quade, 1 993):
8 13C5 (z) =
[( -J 1
RPDB
(18)
where
(19)
and 8i is the isotopic composition of phase i. lt is clear from equation (18) that the biological (S(z)) and atmospheric ( C�) components make direct contribu tions to the isotopic composition of soil C02 and that the term
(22) which means that the 1 2C02 species has a diffusion coefficient 4.4%o greater than that of 13C02 ( Craig, 1 953). The 4.4%o difference in diffusion coefficients causes soil C02 to be enriched in BC by at least 4.3%o compared with soil-respired C02 (for the natural case of the atmosphere being c. 5-20%o enriched in BC compared with soil respired C02) . The value 2 is not exactly 4.4%o because the diffusion of 1 C02 and 13C02 is related to their respective gradients in the soil as well as their diffusion coefficients. For example, Davidson (1995) showed that L\-$ (85 - 8$) can be as low as 4.2 for the case of a soil with very depleted soil organic matter ( -36%o ) , where the upper boundary is the average atmosphere. This is a special case, however, because such depleted 13 C organic matter forms under a closed canopy as a result of the canopy atmosphere being depleted in BC (Medina & Minchin, 1980; Medina et al. , 1 986) because of poor exchange with the troposphere under the canopy; in such a setting the upper bound ary at the soil-air interface is unlikely to have the same concentration and isotopic composition as the average troposphere. Likewise, when soil organic matter approaches the isotopic composition of the atmosphere:
L'l.(air - organic matter) = 8 air - 8organic matter < 4.4%o
(20)
is simply (13C/12C) soil C02. The Stefan-Maxwell relationship for binary diffu sion is:
_[
[J:__ ] J
+ J:__ 2 kT D = 1 Y� n:<J �n n; my m� �
I/2 (21)
where D Y � is the diffusion coefficient of species y in medium �, <JY� is the collision diameter between the y and �, k is the gas constant, n is the number of mole cules, T is temperature, and my and m� are the atomic masses of y and � , respectively. For the case of the dif 2 fusion of 1 C02 and BC02 in air, the collision diam-
the limiting value of enrichment is the difference between the air and respired C02 (Davidson, 1995). In the discussion below we will consider that the isotopic composition of organic matter is depleted in BC by 5-20%o compared with the overlying atmos phere. In these soils the theoretical minimum differ ence L'l.s-$(85 - 8<1,) ranges from 4.4 to 4.3%o. Figure 4 illustrates the diffusion effect where the 8BC value of soil C02 is, at all depths in the soil, greater than the respired value of -27.7%o. Therefore the distinction must be made between the isotopic composition of soil C02 (concentration ) and soil-respired C02 (flux) (Dorr & Mi.innich, 1980, 1986; Cerling et al. , 1991b). This theoretical basis for the isotopic composition of C02 in soils has important implications concerning the 8BC values of carbonate in pedogenic carbonate. It makes the following predictions about the carbon
T E. Cerling
50 0
-20
0 uI
'---'
-60
.fi
0..
� ·n �
9
;;;
7
'B' c
respirmion nne
(mmoles/m 2/hr)
-80
one$
-2 5
-20
-
15
0
-27.7%o
=
oi3Cair
-100
=
D •
D
Mook el al ( 1 974) Turner ( 1 98 2 )
Romanek et al ( 1 992)
-10
6
-S.Oo/oo -5
0
Fig. 4 .
Isotopic composition o f soil C02 for different respiration rates using the model described in the text. Soil parameters are ()1 3 Cair -8%o at 350 p.p.m., porosity £ =0.5, tortuosity p 0.6, T = 15°C, with a characteristic depth z = 25 em for C0 2 production assuming an exponential decrease in soil respiration. All curves intersect the atmospheric value at the soil-air interface, have the steepest gradient just below the soil-air interface, and approach a constant ()1 3 C value at depth. Note that the limiting ()13C value is enriched in 13C by several parts per thousand compared with ()13 C�, which results from differences in the diffusion coefficient for 1 3 C02 and 1 2C02 (see text). Figure modified from Cerling & Quade (1993). =
=
isotopic composition of soil C02 and soil carbonate in isotopic equilibrium with soil C02: 1 the isotopic profiles are in isotopic equilibrium with the atmosphere at the soil-air interface; 2 the one values decrease rapidly in the upper few centimetres and reach a constant value at depth, usually below 30-50 cm; 3 the oBC of soil C02 is always 4.3%o or more enriched in BC compared with soil-respired C02; 4 for modern soils, the difference between the soil respired C02 and soil carbonate Ll( ocaicite - 8$) should be between about 13%o and 16.5%o for high-produc tivity soils, assuming the fractionation factors dis cussed in the paragraph below. The equilibrium fractionation factor: ,-
•
u
0
a. co
+ Emr ic h et a! ( 1 970)
'N 1 0
-40
E u
......._
J03Jna = 1 1 .709 - 0. 1 1 6(T)+2 . 1 6x I0-4 (T)2
II
R eo _ _ _ _ , cac o, RCaC03
1000 + 0 co ,
1000 + 0 cacn
'-'3
(23)
is temperature dependent and has been reevaluated recently by Romanek et al. (1992), who found that the previous empirical estimate of Deines et al. (1974) needed revision. Romanek et al. (1992) point out that
10
20
30
40
50
60
70
Fractionation factor 103 ln a of calcite and C02 during equilibrium formation of calcite at low temperatures. Data from Emrich et al. (1970), Mook et a/. (1974), Turner (1982) and Romanek et al. (1992). Fig. 5.
the difference in their work from previous work is in the characterization of the carbonate phase, which can be either calcite or aragonite. In this paper, we use the Romanek et al. (1992) results, which agree with the determinations of Mook et al. (1974), Emrich et al. (1970) and Turner (1982). This leads to 103 ln Uco caco values of 1 1 .7 and 8.4 at 0° and 30°C, r 3 respectively (Fig. 5). Field validation of diffusion model
A number of field sites have been studied to test the validity of the diffusion-production model and its application to carbon isotope studies of pedogenic carbonate. We discuss here three tests of the model: 1 the 4.4%o diffusion effect and the difference between the soil-C02 end-member value and the soil-respired C02 value; 2 the shape of the isotope profile; 3 the 13-16.5%o difference between respired C02 and pedogenic calcite. The discussion below is related to results reported previously by Cerling (1984 ), Quade et al. (1989b ), Cerling et al. (1991b) and Cerling & Quade (1993). 4.4%o diffusion effect
Cerling et al. (1991b) reported the carbon isotopic composition of C02 in a montane soil at Brighton,
51
Stable carbon isotopes -5 .-------� Air ,.......__
8
'-"
(.)
-15
- 25
•
0
•
•
•
�
0.. v
•
Konza
•
Wasatch
0
1000
1 /C0 2
40
...c:
•
-20
20
0
Little Bluestem
2000
3000
Fig. 6. Isotopic composition of soil C02 from three soils in North America. Solid lines represent modelled /il 3 C respired values given in Table 1 and using an atmospheric C0 2 value of 350 p.p.m. with /il3C -8%o (the atmospheric value at the time the measurements were made). Modified from Cerling & Quade (1993). =
60
e SM-2b --- model
80
]()()
-5
5
0
8 1 3 C soil carbonate Fig. 7. Isotopic composition of soil carbonate in a desert soil near Las Vegas, Nevada, USA. /i1 3 C values near the soil-air interface are in isotopic equilibrium with the atmosphere. Solid line uses the model described in the text for soil carbonate assuming a I)BC (soil-respired) value of -23.4%o and a low soil respiration rate (see Quade et al. , 1989b, for details). Figure is modified from Quade et al. (1989b ).
Utah, which had been studied previously by Solomon & Ceding (1987). This soil system was chosen because it had a deep snow cover in the winter so that the diffusion profile was observed across the additional 2+m of snow cover, instead of only the upper 30 em of the soil. Figure 6 shows measurements of ()13C and l/pC02 of the soil-snow-atmosphere system and two other soils. The end-member {)13C value for C02 is -23.4%o using a modelled o1 3Crespired -27.7%o, whereas the measured {)13C of soil-respired C02 was -27.5%o, which is close to the predicted enrichment due to diffusion effects. Figure 6 also shows measurements and modelling results for two other North American soils. Table 1 shows that the isotopic composition of soil-respired C02 is com patible with organic matter {)13C values observed for these soils.
Fig. 4, which have {)13C values with an atmospheric value at the soil-air interface, a steep gradient near the top of the soil, and a fiat gradient deep in the soil. Figure 7 shows an example of one of these soils measured by Quade et al. (1989b). Recent developments in mass spectrometry (i.e., using gas-chromatography flow-through stable isotope ratio mass spectrometry (GC-IRMS)), enable analyses of very small gas samples, and allow direct collection and measurement of soil C02 in the upper few centimetres of the soil profile. This will allow these observations on diffusion profiles to be extended to other soils that do not have pedogenic carbonate preserving the {)13C of the soil C02.
Shape ofthe carbon isotope profile in soils
L1(0calcite - 0¢)
=
Quade et al. (1989b) , Pendall et al. (1994) and Wang et al. (1996) have measured detailed profiles of pedogenic carbonate in desert soils, where CaC0 is 3 precipitated at the soil-air interface. In these soils pedogenic carbonate is assumed to form in isotopic equilibrium with the oxidized carbon species (C02 H2C03 , HC03, COi, CaC03 ) in the soil. Soils i� these studies show the characteristic shape shown in
The model described above predicts a difference:
�( CaC03 - respired C02 ) {)13 Cca co - {)13C$ =
=
,
13 to 16.5%o
using the fractionation factor ex(calcite - C02) of Romanek et al. (1992) for soils with relatively high respiration rates. Soils with low respiration rates would have � values greater than discussed here, as
52
T E. Cerling
Table 1. Isotopic composition of soil C02, soil-respired C02, soil organic matter and the modelled soil-respired C02 used in
Fig. 6. Data from Cerling eta/. (1991b) and Cerling & Quade (1993) o13C soil C02 (1/C02 intercept) Konza Little Bluestem Wasatch
ol3C respired modelled
-23.4
would soil carbonate collected high in the soil profile. Again, isotopic equilibrium is assumed between all oxidized carbon species. Cerling et al. (1989) and Cerling & Quade (1993) report the results of 34 modern soils developed on Holocene or late Pleistocene parent material, collected with the intention of minimizing the effects of climate and vegetation change. Soils were devel oped under varied vegetation (forest, savanna, shrub land, grassland), climate (Mediterranean, tropical, monsoon, semi-desert, boreal), and from six con tinents (all except Antarctica). All localities were presently in apparently stable vegetation settings, and not situated near ecologic boundaries. Soil car bonate was measured at depths greater than 30cm for all the soils. Figure 8 shows that the isotopic com position of soil organic matter, taken as a proxy of the isotopic composition of soil-respired C02, and pedo genic carbonate does indeed differ by the amount predicted by the model described in the text. These field results give confidence that the diffusion-production model is valid for describing C02 transport in soils. One of the boundary condi tions of this model is the atmospheric C02 concentra tion; if this model is valid, it implies that stable carbon isotopes in paleosols can be used to estimate pC02 of the palaeoatmosphere.
S O ILS AND PALAE O S O L S : C H A RACTE R I S T I C S A N D R E C O G N ITI O N
I have outlined a model for soil C02 and pedogenic carbonate that is very useful for palaeoenvironmen tal studies. Here, however, I make some remarks concerning the characteristics of soils, and the recognition of palaeosols and pedogenic carbonate. Soils are developed either on bedrock or on reworked sediments. Sequences of soils are almost
-14.9 -17.8
-15.2 -19.1 -27.7
-10.8 -14.7
...
�
... c:
s
- 10
0I3C organic matter
o13C respired measured
-26.7
-27.5
Modern soils model: 30°C - - model:
u
·a c:
0
0 -
/
/
/
/
�
0�
·c; "'
u
"'
00
-10
0
o l 3 C pedogenic calcite
10
Fig. 8.
Isotopic composition of pedogenic carbonate and soil organic matter for 3 4 Holocene soils. Pedogenic carbonate from below 30cm. Solid lines are using the model described in the text for S(z) values of 8000 p.p.m. at ooc and 30°C. Figure modified from Cerling & Quade ( 1 993).
invariably developed on reworked sediments, whereas bedrock soils are isolated soils. In the case of bedrock-developed soils and palaeosols, the unifor mity of the parent material and the change in soil structure are generally sufficient to make a clear identification of palaeosols. In many cases the unifor mity of the parent material can be used to determine the amount of weathering that has taken place in the soil using mass balance considerations (e.g., Brimhall & Dietrich, 1987); in certain cases this is even true in alluvial soils and palaeosols. For palaeosols developed in alluvial systems, it is often necessary to recognize a repetitive pattern characteristic of soil development.
Stable carbon isotopes 'Pedogenic carbonate' versus 'carbonate in soils'
Carbonate in soils can be either inherited from the parent material (either as bedrock or as detrital grains), or it can precipitate as a result of soil-forming processes. The diffusion-production model described above is for the pedogenic component, that is, the car bonate formed during soil formation. Several studies of soil carbonate developed on limestone bedrock have been published where it is clear that complete dissolution and reprecipitation has resulted in the biological signal being preserved in the soils (e.g., Rossinsky & Swart, 1993). In others, it is not clear that identification of an exposure surface is sufficient for application of the 8BC model described above. Marine or lacustrine exposure surfaces must have undergone sufficient pedogenesis to allow for separa tion between the detrital (or bedrock) component and the pedogenic component (Cerling, 1992b,c; Wright & Vanstone, 1991, 1992). For carbonate formation to take place in soils, there must be sufficient weathering of calcium bearing minerals. The easiest mineral to weather in most soils is detrital (or bedrock) calcite. Marion et a!. (1991) show an example in Alaska where a carbonate leached zone of soils is recognizable (but not fully developed) in a period of a few hundred years, showing how rapidly pedogenic carbonate can form from the weathering of carbonate minerals. In the absence of carbonate minerals available for weather ing other calcium-bearing minerals, such as silicates, must weather before carbonate formation can occur in soils. In typical soils where carbonate is present as a detrital phase, such as in alluvial sediments, glacial deposits, or loess, a leached zone develops in the upper part of the soil. Pedogenic carbonate can form both in the leached zone, and in the lower unleached part of the soil. In the former case, the carbonate is unlikely to be contaminated by detrital carbonate, but in the latter case contamination is potentially a serious problem. Recognition of palaeosols
Modern soils are characterized by their intense biological activity, which results in organic matter accumulation high in the profile, bioturbation that homogenizes the original rock or sediment texture, enhanced weathering and translocation of mobile elements and even minerals. The consideration of modern soil features is very useful in identifying
53
paleosols. For the purpose of geochemical arguments, especially concerning carbon isotopes in carbonates, it is better to be somewhat cautious in the identification and interpretation of palaeosols. In any sequence of palaeosols, such as is found in sedimen tary deposits, it is useful to establish the pattern that results from soil-forming processes. The presence of preserved organic matter in palaeosols is a very useful criterion in identifying palaeosols. Organic matter, however, although common in Holocene and some Pleistocene soils, is less abundant in increasingly older palaeosols because of the oxidation of organic matter through time. Cerling (1991) and Mora et al. (1996), however, have found trace amounts of organic matter in Palaeozoic and Mesozoic palaeosols. In alluvial soils the destruction of the original bedding is an important clue to identifying palaeosols. Complete homogenization of the original bedding indicates significant bioturbation. Incom plete bioturbation, where evidence of bedding can be identified in the upper part of the soil profile, indi cates only incipient soil development. Many palaeosols have a Bt horizon (t denotes illu viated clay) and the development of ped structure. Peds are generally discrete, subrounded, with an outer layer of oriented clay particles, generally a few millimetres in diameter. Peds often survive diagene sis and are readily recognizable in many Cenozoic, Mesozoic and Palaeozoic palaeosols (Retallack, 1990a,b). Translocation of iron minerals occurs in soil horizons and is often recognized by the presence of orange to reddish orange colour banding. The colour preserved in palaeosols is not the original colour, but is generally enhanced from the original because of the oxidation of organic matter, which darkens the soil or palaeosol. Iron or manganese nodules also are found in the B-horizons of many soils and palaeosols. These often can be found reworked in channel conglomerates. The development of leached zones in the upper part of the soil is also a key towards identifying soils. The leached zone develops where there is a net flushing of ions from the soil. In vadose soils of North America, the leached zone ranges from 0 to > 2 m in regions of high rainfall (> c. 1 m yr-1 ). Pedogenic car bonate often forms in the leached part of the soil. It occurs as discrete nodules, as rhizoliths associated with roots, as pendants beneath large clasts, and as semi-continuous to continuous horizons. Pedogenic carbonate can occur as massive accumulations in
54
T. E.
regions of groundwater discharge, sometimes to the soil-air interface, although it may be difficult to identify such accumulations in the fossil record. Reworked nodules in basal conglomerates in sand stone channels provide evidence that the nodule for mation was early. Quade & Cerling (1995) found that reworked nodules in such conglomerates had the same isotopic composition as nodules in leached horizons, and differed from late-stage veins or detri tal carbonate. They also found that carbonate precipi tated in the unleached zone was contaminated by detrital carbonate. Groundwater calcretes, which have an extensive literature in Australia, are not suitable for this isotope model because carbonate is precipitated in the satu rated zone, where gaseous diffusion is not the domi nant mechanism for mass transfer. In some cases it can be very difficult to distinguish between ground water calcretes and those formed in the unsaturated zone. In summary, accurate identification of palaeosol carbonate is a very important aspect of using their isotopic composition as a palaeoenvironmental indi cator. Occurrences of calcretes or other carbonates that do not have a suite of other evidence to confirm their pedogenic origin should be suspect. Detrital contamination, or carbonate formation in a non-soil environment, precludes the use of the model described here for palaeoenvironmental studies in the context of palaeosols.
A P P L I C ATI O N S : C I S O T O P E S IN P A L A E O S O L S
Carbon isotopes i n palaeosols have been used i n two principal ways: in the study of the fraction of C3 and C4 plants in a local ecosystem, and in estimating the pC02 of the atmosphere. I will give examples of each of these applications below. C3/C4 ecosystem changes
The relative contributions of C3 and C4 biomass to an ecosystem can be estimated using the 8BC of pedogenic carbonate. This has been used in a number of studies (e.g., Cerling & Hay, 1986; Cerling et al. , 1988, 1991a; Cerling, 1992a; Mack et al. 1994; Quade et al. , 1989a, 1994; Koch et al. , 1995; Quade & Cerling, 1995; Slate et al. , 1996) on Tertiary and Quaternary sediments. Figure 9 shows the results of about 120 palaeosols from the Siwalik sequence in Pakistan,
Cerling 0
Pakistan Siwaliks
�
5 ---
..
�
:E
..._.,
Q) 01)
10 .
--<
15
20
-15
o0 co0 os o c? o oora� 00 0 -10
% 0-
0 a> o 0 (Y 0 O B �0 0 0o o 0
�
-5
0
5
() 1 3 C pedogenic CaC0 3 Fig. 9.
Carbon isotopic composition of pedogenic CaC03 from the Siwalik sequence in Pakistan. Figure modified from Quade & Cerling (1995).
studied by Quade & Cerling (1995; see also Quade et al. , 1989a). The Siwalik sequence is particularly instructive because it is an almost complete Neogene section with an exceptional preservation of palaeosols (see also Retallack, 1990a). The palaeosols in this section illustrate the points made above in the section on the recognition of palaeosols: the parent material was originally bedded and calcare ous; bioturbation and leaching are readily observ able; colour differentiation is well developed; ped structure is well developed; and carbonate nodules are found in both the leached and the unleached soils. The range of 1)13C values in soils older than 8 Ma (Fig. 9) may represent either a very small fraction of C4 biomass (> 10%) or the variation could be the result of changes in the 813C of organic matter of the ecosystem; under conditions of high light irradi ance and water stress the carbon isotope discrimi nation is decreased and the 1)13C of C3 plants can be enriched in BC by several parts per thousand (Ehleringer et al. , 1986; Ehleringer & Cooper, 1988). Thus the 1)13C (CaC03 ) values of -9%o could be the result of either a small fraction of C4 biomass or that the 813C of the C3 plants was shifted to about -23%o. In addition, a change in the isotopic composition of
55
Stable carbon isotopes the atmosphere would also cause a shift in the 8BC of plants. Starting about 7 Ma the Siwalik palaeosols shift to more enriched BC values and by 5 Ma reach values as high as +3%o, indicating a pure C4 ecosystem. The pure C4 ecosystem of the Siwaliks was almost cer tainly a C4 grassland. The lower 813C values in the post-5 Ma palaeosols, ranging to -3%o, indicate that mixed C3/C4 ecosystems were present as well as the pure C4 ecosystems. Studies of palaeosols in Africa (Ceding & Hay, 1986; Cerling, 1992a) do not show evidence for very much C4 biomass before 7 Ma, although some samples indicate that C4 plants were present. The palaeosol evidence from the Olduvai Gorge region and from the Turkana and Baringo basins do not show the abrupt change seen in Pakistan, but rather show a gradual increase in C4 biomass starting about 7 Ma, with the modern ecosystem having a high C4 biomass compared with most of the previous 7 Ma. Studies of fossil tooth enamel, however, show that there was a rapid expansion of C4 biomass in Asia, Africa, South America and North America starting about 7 Ma (Cerling et al. , 1993; Morgan et al. , 1994; MacFadden et al. , 1996). Quade & Cerling (1990) used the carbon and oxygen isotope composition of calcite and its occluded carbon to demonstrate the pedogenic origin of shallow carbonate in a fault trace near a proposed nuclear waste repository. They showed changes in vegetation from late Pleistocene conditions com pared with the present ecosystem. Studies of atmospheric pC02
Equation (18) shows that the isotopic composition of soil C02 is related directly to the concentration of C02 in the atmosphere. The form of eqn (18) is useful in evaluating the different parameters contributing to the isotopic profiles in soils, but a more convenient expression for palaeo-pC02 studies can be derived. The S(z) term is related to total soil productivity, porosity, depth in the soil and the characteristic depth of C02 production; however, it reaches a constant value at depth in vadose soils, usually between about 5000 and 8000p.p.m. (Brook et at. , 1983) for soils likely to produce soil carbonate. If we consider the carbonate that forms at depth in the soils, below c. 50 em, the S (z) term can be considered constant. Recall ing that: q (z ) = S i (z ) + q
and Csi = S i i*D ':' + Cai D i
(24)
2 C1 85 = 1.00448$ + �2 ( 8 a - 1 .00448 $ - 4.4) + 4.4 Cs
(25)
c* - a* cs1 2 cs
(26)
where i can be the bulk C02, 1 2 C02 or 13C02, eqns (17), (22) and (24) can be expressed as (Davidson, 1995):
Because 1 2C is the dominant isotope in C02 (99 % ) w e can make the approximation that Cla
2
_
and recalling that C�(z) = S(z) + c:ir' eqn (25) then can be recast: -
c.* - s (z)
(8 s - 1 .00448<1> - 4.4) (8 a - 8 s )
(27)
·.
p(C02) of the atmosphe�e is seen to be related to the biological component of C02 in soil, and the isotopic composition of soil C02, soil-respired C02 and the air (S(z), 85, 84>, and 83, respectively). To use palaeosols to estimate pC02 of palaeoatmospheres it is necessary to estimate S(z), 85, 84> and 83. Yapp & Paths (1994) have studied the carbon isotopic com position of goethite and found that the co3 concen tration of goethite is related to q. With pedogenic carbonate, however, that is not the case and an edu cated guess must be made. If pedogenic carbonate forms at depth within a soil (c. > 50 cm) then the C02 gradient is essentially constant and it is reasonable to assume values similar to that observed in modern soils (Brook et al. , 1983); for high productivity soils, S(z) probably is in the range 5000-8000 p.p.m., and for arid zone soils S(z) probably is in the range 3000-5000 p.p.m. Figure 10 shows the relationship of atmospheric pC02 to S( z) and to !). (a - <1>) = 83 - 84>, which is depletion in 13C in C3 plants compared with the air in which they grew. It is clear from this figure that a significant uncertainty is present in atmospheric pC02 estimates using palaeosols. However, the palaeosol pC02 barometer, in spite of this uncertainty, is more promising than searching for Mesozoic ice! The 85 value is obtained from the 8l3C of pedogenic carbonate, assuming isotopic equilibrium between pedogenic calcite and soil C02 (from Fig. 5):
56
T. E. Cerling 5000
> 8
.----
0.. 0.. '-'
N
0 u 0... II
u"'
5000
�'>a-(j> S(z) 1 9.5 17.5 1 9.5 17.5
4000 3000
9000 9000 6000 6000
/
2000 1 000 0
// 0
�
//
/ / / / //
/ /
/ /
/ /
� I �
/
4000
/
> E
0..
3000
&
N
0 u 0...
2000
�
1 000
5
15
10
500
400
300
200
5000 4000 3000
�a �
2000 1000
100
Age (Million years) Fig. lO. Relationship between pC0 2 and L\.5_$ using model described in text.
103 ln <Xco -caco , ,
+
=
1 1.709 - 0. 1 1 6 T 2.14x 10-4 T 2 (28)
Figure 5 shows the temperature dependence of the fractionation factor between calcite and C02, which has a temperature relationship of about 1%o 10°C-1, so the uncertainty in the temperature estimate is very important. The respired C02 value, Oq,, is assumed to be the same as the oBC value for soil organic matter; unfortunately it is not known if the palaeosol organic matter is shifted significantly during diagenesis. Vari able 03 is the least sensitive value that needs to be estimated. The pre-industrial value of 03 was about 6.5%o but it is now about -8%o; the history of 03 is not well constrained. Taken together, the difference between o, and 84>: (29) is a measure of pC02, with a slope of greater than 1 %o 1000 p.p.m.-1 C02 (Fig. 10). Therefore, the tempera ture estimate is significant because it has a slope of about 0.1 %o °C-1 and an error in the temperature esti mate of 10°C is an error of more than 1000 p.p.m. In spite of the problems in estimating the parameters for the study of the history of atmospheric pC02, this method has given a number of estimates of C02 that agree with other geological considerations and models. Cerling (1991, 1992b,c), Mora et al. (1991 , 1996),
Fig. ll.
Estimates of atmospheric C02 using palaeosol carbonate (Cerling, 1992c, and unpublished data; Mora et al. , 1996) and the model described in the paper.
Mora & Driese (1993) and Yapp & Poths (1992, 1994, 1996) have used this model to estimate pC02 levels of the Phanerozoic using palaeosol carbonate and goethite. Table 2 shows the results for some modern soils dominated by C3 plants; o, values are calculated using the growing season temperature for each locality. It is useful to note that these modern C3dominated ecosystems all have cool growing seasons; C4-dominated ecosystems (e.g., examples in Cerling & Quade, 1993) all have warmer growing seasons. Table 2 shows that modern soils give an estimate for modern atmospheric pC02 of 510 ± 370 p.p.m. This estimate gives an idea of the uncertainty under the best of conditions: of the order of± 500 p.p.m. Figure 11 and Table 3 show estimates of palaeo atmospheric pC02 from palaeosols. The general pattern is one of low atmospheric C02 during the Tertiary (less than 1000 p.p.m.), higher during most of the Mesozoic (c. 2000-3000 p.p.m.), low in the Permian (about 1000 p.p.m.), and high in the late Palaeozoic. This agrees with the results of the model of Berner (1991, 1994).
C O N C L U D I N G S TAT E M E NT S
I have discussed the development in using stable carbon isotopes to study soils and palaeosols over the last two decades, with emphasis on the application to studies of palaeosols in the geological record. Stable
57
Stable carbon isotopes
Table 2. Parameters from modern C3-dominated soils used to calculate pC02• Range of values shows that the soil carbonate
barometer has an uncertainty of the order of ±500 p.p.m. Temperature is the estimated soil temperature at 50 em during the growing season (lower temperatures would result in a lower pC02 estimate). occ > o,, o., are the oBC values for average pedogenic carbonate, for calculated soil C02 using eqn (27), and for the pre-industrial atmosphere, respectively. o�, the respired component, is taken to be average soil organic matter for modern soils. �s-am and �a-om are the differences (o, - 00111 ) and (o. - o0n,), respectively. pC02 is calculated from eqn (26) assuming that S(z) 5000 p.p.m. for all soils except aridisols, where S(z) 4000 p.p.m. =
=
New York Nevada Nevada Nevada Bolivia Bolivia Utah Utah Utah Saskatchewan Saskatchewan Saskatchewan Greece Greece Turkey Turkey France
Alfisol Aridisol Aridisol Aridisol Aridisol Aridisol Aridisol Aridisol Aridisol Mollisol Mollisol Mollisol Mollisol Mollisol Vertisol Vertisol Alfisol
T
0cc
0$
o,
o.
�s-�
6,a-om
pC0 2
15 15 13 11 10 10 10 10 10 15 15 15 15 15 15 15 15
-9.4 -6.8 -8.5 -8.5 -7.3 -8.5 -7.4 -8.8 -7.5 -7.9 -8.4 -6.3 -7.5 -9.3 -10.0 -10.3 -10.0
-25.6 -23.4 -23.7 -23.9 -22.8 -23.3 -24.5 -23.8 -24.4 -24.2 -24.1 -22.1 -23.7 -25.7 -24.5 -24.5 -25.1
-19.3 -16.7 -18.6 -18.8 -17.7 -18.9 -17.8 -19.2 -17.9 -17.8 -18.3 -16.2 -17.4 -19.2 -19.9 -20.2 -19.9
-6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5 -6.5
6.3 6.7 5.1 5.1 5.1 4.4 6.7 4.6 6.5 6.4 5.8 5.9 6.3 6.5 4.6 4.3 -5.2
19.1 16.9 17.2 17.4 16.3 16.8 18.0 17.3 17.9 17.7 17.6 15.6 17.2 19.2 18.0 18.0 19.6
800 940 270 260 270 30 840 90 760 940 650 820 920 880 130 20 330
5.6
17.4
510
Average (1cr 0.9)
isotope studies in palaeosols have great potential towards further understanding the history of global ecosystem changes and global atmospheric chem istry, as well as shedding light on the systematics of soil behaviour and its role in modifying global climate. The dynamics of soil processes, however, are still very poorly understood, as is the diagenesis of soil carbon. This still limits the usefulness of isotopes in palaeosol studies and their interpretation. One of the next steps for the study of palaeosols and soils is to develop the use of 14C isotopes in soils in order to understand the rates of soil formation and diagenesis (see Amundson et al. , 1994; Wang et al. , 1996). This step i s considerably more complicated for modelling because steady-state conditions are not reached for thousands of years for the 14C isotbpe system, compared with tens of hours or less for the stable isotopes. Unlike the stable carbon isotopes, however, 14C has a built-in clock and can address other problems of time that the stable isotopes cannot. Understanding the 14C input to the soil system will be very important in studies of soil devel opment, turnover of carbon in soils and the interac tion of soils in the global carbon cycle.
Another important step is to quantify some of the parameters in the soil diffusion-production model for application to palaeosols. For instance, the soil C02-production value (S(z)) for carbonate precipita tion is not known for any modern soil, let alone fossil soils. Is it the maximum C02 value attained in soils, or some value intermediate between the maximum achieved in the growing season and the minimum found in the non-growing season? Annual soil tem perature ranges can easily be 20°C or higher. At what average soil temperature does pedogenic carbonate form: is it the maximum soil temperature, is it the temperature during the period of maximum soil res piration, or is it the soil temperature related to some other process? In addition, how would this be esti mated for palaeosols? We are left also with the problem of diagenesis. What is the best estimate of the original 813C value of soil respired C02? Many workers have measured the 813C of organic carbon preserved in palaeosols (e.g., Cerling, 1 991; Cerling, 1992b; Mora et al., 1996) and have estimated that the 813C of soil-respired C02 is the same as that preserved as organic carbon pre served in palaeosols. Balesdent and Mariotti (1996),
T E. Cerling
58
Table 3. Calculated pC02 for the Phanerozic Eon using pedogenic carbonate. ow 8$ , o,, oa , t.,-1>' and L':.a--i> are the ot 3 C values for pedogenic carbonate, for palaeosol respired C02, the calculated oBC of the soil atmosphere, the estimated o13C of the atmosphere and the difference between o, - O$ and oa - 8$ , respectively. Respired C02 (8$) was assumed to be 1 %o depleted in BC compared with the measured oom · o, was calculated at 25°C except for modern soils. pC02 values are calculated for two S(z) values, 5000 and 8000 p.p.m.
Soil Modern§ Francett Francett Pakistan� Fort Ternanll Willwoodll India** Ephramtt Proctor Lake [[ , tt Dolorestt Chinlett Dunkard:j::j: Conemaugh:j::j: Hinton:j::j: Pennington:j::j: Mauch Chunk:j::j: Maccrady:j::j: Catskill:j::j: Catskill:j::j: Bloomsburg:j::j:
Age (Ma)
0cc
0$*
o,
pC02t
pC02:J:
0 1 4 8 14 51 70 100 110 220 230 285 305 334 334 339 351 364 367 412
-9.8 -10.0 -10.6 -11.9 -10.6 -10.6 -6.7 -6.5 -6.3 -6.9 -8.7 -7.2 -7.6 -7.0 -7.0 -7.6 -9.8 -9.0 -5.3
-25.0 -25.4 -25.0 -28. 1 -25.6 -27.0 -28.1 -25.1 -24.7 -25.0 -24.5 -24.6 -24.8 -24.8 -24.8 -24.8 -28.2 -28.2 -28.4
-18.6 -18.8 -19.4 -20.7 -19.4 -19.4 -15.5 -15.3 -15.1 -15.7 -17.5 -16.0 -16.4 -15.8 -15.9 -16.4 -18.6 -17.8 -14.2
510 860 930 504 1 100 740 960 4580 2580 3040 2690 1240 2250 2060 2540 2490 2090 2200 2740 6510
885 1380 1490 810 1760 1180 1540 7330 4130 4870 4300 1980 3600 3290 4070 3980 3350 3520 4380 1 1 400
�
* is estimated to be 1 %o depleted in BC compared with measured organic carbon in the palaeosol. t Calculated from eqn (28) assuming S(z) 6000 p.p.m. :j: Calculated from eqn (28) assuming S(z) 8000 p.p.m. § From Table 2. � Data from Quade & Ceding (1995). II Data from Ceding (1992c). Data from Tandon et al. (1995) and unpublished data. tt Unpublished data. :j::j: Data from Mora et al. (1996). =
=
••
however, studied a modern soil that had been cleared of vegetation for over 60yr and found that the resid ual BC increased by 1 .6%o during this 60-yr interval. More studies of changes in the isotopic composition of soil carbon are needed before one can confidently asign o$ values from residual soil organic matter. In summary, stable isotope studies of palaeosols have great potential for helping to understand the history of global climates and ecosystems. Continued studies of modern soils and their counterparts, palaeosols, are necessary to be able to fully realize that potential. A C K N OW L E D G E M E N T S
This work was supported over many years b y the Research Corporation, the National Science Founda-
tion, Mifflin and Associates, and most importantly BLS (Boot Leg Science). J.R. Ehleringer, J. Quade, D.K. Solomon and Y. Wang contributed to various parts of this subject. This paper benefited from the reviews by M.I. Bird, G.R. Davidson and M. Thiry. This paper was written while the author was a visitor at the California Institute of Technology.
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Cerling ROMANEK, C.S., GROSSMAN, E.L. & MORSE, J.W. (1992) Carbon isotopic fractionation in synthetic aragonite and calcite: effects of temperature and precipitation rate. Geochim. Cosmochim. Acta , 56,419-430. RossiNSKY, V. & SwART, P.K. (1993) Influence of climate on the formation and isotopic composition of calcretes. In: Climate Change in Continental Isotopic Records (Eds Swart, P., Lohmann, K.C., McKenzie, J.A. & Savin, S.M.). Geophys. Monogr. Am. Geophys. Union, 78, 6775. SALOMONS, W., GooDIE, A. & MooK, W.G. (1976) Isotopic composition of calcrete deposits from Europe, Africa, and India. Earth Swf Processes, 3, 43-57. ScHLESINGER, W.H. ( 1 997 ) Carbon balance in terrestrial detritus. Ann. Rev. ecol. Syst. , 8, 5 1-81. SINGH, J.S. & GUPTA, S.R. (1977) Plant decomposition and soil respiration in terrestrial ecosystems. Bot. Rev. , 43, 449-528. SLATE, J.L., SMITH, G.A., WANG, Y. & CERLING, T.E (1996) Carbonate-paleosol genesis in the Plio-Pleistocene St. David Formation, southeastern Arizona. J. Sediment. Res. , 66, 85 94 . SOLOMON, D.K. & CERLING, T.E. ( 1 987) The annual carbon dioxide cycle in a montane soil: observations, modeling, and implications for weathering. Water Resour. Res. , 23, 2257-2265. TANDON, S.K., SooD, A., ANDREws, I.E. & DENNIS, P.F. (1995) Palaeoenvironments of the dinosaur-bearing Lameta Beds (Maastrichian), Narmada Valley, Central India. Palaeogeogr. Palaeoclim. Palaeoecol. , 117, 15 3- 184 . TRUMBORE, S.H., BoNANI, G. & WoLFLI , W. (1990) The rate of carbon cycling in several soils from AMS 14C measure ments of fractionated soil organic matter. In: Soils and the Greenhouse Effect (Ed. Bouwman, A.), pp. 407-414. Wiley, New York. TuRNER, J.V. ( 1 982 ) Kinetic fractionation of carbon-13 during calcium carbonate precipitation. Geochim. Cos mochim. Acta, 46, 1183-1191. WANG, Y.,AMUNDSON, R. & TRUMBORE, S. (1994) A model for soil i4C02 and its implications for using I4C to date pedo genic carbonate. Geochim. Cosmochim. Acta, 58, 393399. WANG, Y., McDoNALD, E., AMUNDSON, R., McFADDEN, L. & CHADWICK, 0. (1996) An isotopic study of soils in chrono logical sequences of alluvial deposits, Providence Moun tains, California. Ceo!. Soc. Am. Bull. , 108, 379-391. WRIGHT, V.P. & VANSTONE, S.D. ( 1 991) Assessing the carbon dioxide content of ancient atmospheres using palaeo calcretes: theoretical and empirical constraints. J. geol. Soc. London, 148, 945-947. WRIGHT, V.P. & VANSTONE, S.D. (1992) Further comments on using carbon isotopes in paleosols to estimate the C02 content of the atmosphere. J. geol. Soc. London, 149, 675676. YAPP, C.J. & PaTHS, H. (1992) Ancient atmospheric C0 2 pressures inferred from natural goethites. Nature, 355, 342-347. YAPP, C.J. & PaTHS, H. ( 1 994 ) Productivity of pre-vascular continental biota inferred from the Fe(C03 ) content of goethite. Nature, 368, 49-51 . YAPP, C.J. & PaTHS, H. (1996) Carbon isotopes in continen tal weathering environments and variations in ancient atmospheric C02 pressure. Earth Planet. Sci. Lett. , 137, 71-82. -
Spec. Pubis int. Ass. Sediment. (1999) 27, 61-84
Palaeoenvironment, palaeoclimate and stable carbon isotopes of Palaeozoic red-bed palaeosols, Appalachian Basin, USA and Canada
C . I . M O RA and S . G. D R I E S E Department of Geological Sciences, University ofTennessee-Knoxville, Knoxville, TN 37996-1410, USA
A B S T R AC T Palaeosols with vertic (Vertisol-like) features occur in the upper clay-rich parts o f upward-fining sequences in Palaeozoic red-bed successions ranging from Ordovician to Permian age within the Appalachian Basin, USA and Canada. Occurrences of vertic features in nearly all of the claystone palaeosols indicate persistence of a seasonally wet-dry palaeoclimate and smectitic clay sources in the Appalachian region for nearly 180 Myr, over palaeolatitudes ranging from 0 to 30° south. Palaeosols are developed in both allocyclic, marginal-marine deposits and autogenic, alluvial-plain deposits, and are char acterized by very weak horizonation, abundant pedogenic slickensides, and a micromorphology domi nated by sepic-plasmic fabrics and peds bounded by stress cutans, hence they broadly are analogous to USDA vertic Entisols and Inceptisols. Pedogenic carbonate is generally abundant, and consists of calcite nodules and rhizoliths. Variations in palaeosol morphology and stable isotope geochemistry are attributed to: 1 differences in the pedogenic and geomorphic environments, whether coastal margin or inland alluvial; 2 differences in the evolutionary state of the soil ecosystem, in particular, the presence of vascular plants, with or without deep root systems. Consideration of these controls permits interpretation of the carbon isotope compositions of pedogenic carbonate as a proxy for Palaeozoic atmospheric C02 levels. Our results suggest a steep decrease in atmos pheric C02 levels between the late Silurian (3200-5200 p.p.m.) and early Permian ( 150-200 p.p.m.), which was associated with the rapid evolution and diversification of vascular land plants and global climate change, leading to the extensive Permo-Carboniferous glaciation.
INTR O D U CTION
geochemistry of Palaeozoic Appalachian red-bed palaeosols, drawing attention to the many character istics common to all of the palaeosols, as well as the important differences that can be ascribed to the pedogenic palaeoenvironment and/or evolutionary advances in the soil biomass.
Palaeosols crop out extensively i n the Appalachian region of the eastern USA and maritime Canada (Fig. 1), occurring predominantly in terrigenous clastic red-bed deposits ranging in age from Ashgillian (Upper Ordovician) to Lower Permian (Fig. 2; Table 1). The wide stratigraphical occurrence of these palaeosols spans periods of rapid evolution and diversification of the terrestrial ecosystem and of long-term climatic change. Because the palaeosols formed under relatively constant source-area and pedogenic conditions, they share generally uniform physical and chemical properties and are thus suit able for investigations of the influences of long-term changes in variables such as soil ecology (Table 1), palaeoclimate or palaeoatmospheric p(C0 2). This study summarizes the morphology and stable isotope
G E N E R A L A S P E CT S O F PALA E O Z O I C R E D - B E D PA L A E O S O L S Geographical and stratigraphical distribution
Palaeosols crop out extensively in terrigenous clastic red-bed deposits of the Appalachian Foreland Basin (Fig. 1), which extends from the Canadian Maritime
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
61
62
C. I. Mora and S. G. Driese
MAUCH CHUNK
PALAEOZOIC
CATSKILL
SUCCESSION Central Pennsylvania BLOOMSBURG
JUNIATA
Fig. l. Map showing distribution of Palaeozoic red-bed palaeosols in Appalachian region of eastern USA and maritime Canada. Palaeosol locality numbers are keyed to Table 1 .
Provinces southward t o the Tennessee-Alabama border (Fig. 1) along the western side of the Appalachian Orogen. Basin formation was initiated by lithospheric loading of a passive margin by Taconic (Middle Ordovician) thrust sheets (Quinlan & Beaumont 1984; Tankard 1986; Beaumont et al. 1988), with each of the three major episodes of Palaeozoic orogeny (Taconic, Acadian, Alleghanian) resulting in deposition of a major clastic wedge. The molasse phase of deposition for each clastic wedge provided abundant parent material for the various depositional systems within which the palaeosols developed. (For in-depth reviews of the tectonic and stratigraphical history of the Appalachian Orogen and Foreland Basin, see Colton (1970), Meckel (1970), Thomas (1977) and Williams & Hatcher (1983), amongst others.) . Depositional setting o f parent material
For nearly 200Myr, until the Appalachian Foreland Basin was deformed by compression during the Alleghanian orogeny (Upper Carboniferous Permian) , a depositional pattern was established that
Fig. 2 . Generalized stratigraphical column for Palaeozoic rocks exposed in central Pennsylvania, Appalachian Foreland Basin, USA. Red-bed formations containing vertic claystone palaeosols are named. See discussion in text. persisted through the development of three major Palaeozoic clastic wedges. A spectrum of deposi tional environments extended away from the foothills of linear highland uplifts towards the west and the north-west into the Appalachian Foreland Basin. Piedmont alluvial fans graded downslope to a broad alluvial plain, which in turn led to low-gradient delta-plain and coastal mud-fiat environments at the interface with a shallow-marine system. Proximal, higher gradient alluvial facies are largely coarser grained, light-coloured sandstones and conglomer ates. Red-bed deposits with palaeosols, consisting largely of upward-fining sequences of red channel sandstone overlain by red shale and siltstone, were deposited lower on the alluvial and deltaic plain, and in coastal-margin mudflat environments (Table 1 ) . As a result of the relative constancy of depositional processes, Appalachian palaeosol-bearing deposits are all red beds with grossly similiar physical and chemical attributes (Figs 3, 4 & 5).
63
Palaeozoic red-bed palaeosols
Table 1. Characteristics of red-bed palaeosols in the northern, central and southern Appalachian Basin region*. Carbonate
morphologies include: R, rhizoliths; RC, rhizoconcretions; N, nodules; L, lacustrine; E, evaporites; B, animal burrows. Localities in USA: KY, Kentucky; NY, New York; OH, Ohio; PA, Pennsylvania;TN, Tennessee; VA, Virginia; WV, West Virginia. Localities in Canada: NS, Nova Scotia; Q, Quebec Age
Orogenic events
Marine parented coastal soils
Permian
Non-marine parented inland alluvial soils
Organic advances16
Carbonate morphology
Dunkard Group (WV, OH)15
R,RC,N
Monongahela Group (KY,WV,OH)14 Conemaugh Group (KY,WV,OH)13
R,RC,N,L R,RC,N,L
Alleghanian Pennsylvanian
Orogeny
Widespread peat swamps Pennington Formation (TN)12 Maccrady Formation (WV)9
Mississippian
Devonian
Acadian Orogeny
Mauch Chunk Formation (PA)11 Hinton Formation (VA,WV) !O Catskill Formation (PA,NY)S Malbaie Formation (Q)6
Catskill Formation (NY)7
R,RC,N R,RC,N R,RC,N,E Large, deep root systems Spiders Arborescence, insects Moderate root size, depth
R,RC,N
Earliest shallow roots Centipedes, millipedes No true roots (rhizomes) First vascular plants First soil animals
N,B N,B
Land plant spores Soil animal traces (?)
B
R,N,L R,N,B
B attery Point Formation (Q)5 Silurian
Ordovician
Moydart Formation (NS)4 Bloomsburg Formation (PA) 3 Taconic Orogeny
Juniata Formation (TN,VA)1
Juniata Formation (PA) 2
* References: Algeo et al. (1995)16; Banks et al. (1985)16; B arlow (1975)15; Blodgett (1985)1 3-15; Boucot et al. (1974)4; Bridge & Gordon (1985)8; Bridge & Willis (1994)7; Cant & Walker (1976)5; Caudill et al. (1992a)12 , (1992b )1 3 , (1996)1 2; Diemer ( 1992)8; DiMichelle & Hook (1992)16; Dineley (1963)4; Driese & Foreman (1991,1992)1; Driese & Mora (1993a)8; Driese et al. (1992)3, (1993b)11; Edmunds et al. (1979) 1 1 ; Fastovsky et al. (1995)15; Feakes & Retallack (1988) 2; Gensel & Andrews (1984, 1987) 1 6; Gordon & Bridge (1987)8; Gray & Shear (1992)16; Hoskins (1961)3; Jaeckel (1995)13; Lawrence & Rust (1988)5,6; Milici (1974)1 2; Milici & Wedow (1977)1; Mora et al. (1996) 3 ,7-13,15; Neal (1995)1 0; Rahmanian (1979)8; Retallack (1986) 2,3,8, (1993)2 ; Retallack & Feakes (1987) 2, 16; Rust (1984)6; Sevon (1985)8: Stefaniak et al. (1993)9; Stewart (1983)16; Strother (1988)3; Thompson (1970)1; Walker & Harms (1971)8; Warne (1990)9; Woodrow et al. ( 1973)7,8.
Coastal-margin environments and palaeosols Coastal-margin environments encompass a wide spectrum, ranging from delta-plain to coastal mudflat systems. During Late Ordovician, Late Silurian, and Late Devonian times, low-gradient braided and meandering rivers terminated and graded seaward (north-westward to westward) into low-energy, coastal mudflat and tidal-flat environments (Table 1 ) .
Milankovitch-scale sea-level changes resulted i n sub aerial exposure and pedogenesis of coastal-margin deposits to form vertic palaeosols (Fig. 3a-c), fol lowed by marine transgression and drowning of the palaeosols (Walker & Harms 1 971; Driese & Foreman 1992; Driese et al. 1992; Cotter & Driese, 1998). Later Mississippian-Pennsylvanian (Car boniferous) coastal-margin environments were more characteristically delta-plain, having formed in
�
Juniata
(U.
@]
Catskill - Irish Valley Mbr.
Ord.)
(U.
Dev.)
cross-bedded N4 sandstone
fine sandstone
t O R 3/4 1aminated
t OR 3/4 smstone
t OYR 612 smstone
smstone
5Y 5/2 claystone 5R 412 claysione
·5R 4/2 claystone
5R 412 silty claystone
'I
�
,- ,.., 5Y 6/4 med. sandstone
-
5R 412 clayshale, fissile
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5R 412 silty
claystone
claystone
0
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(U.
•
5R 4/2 silty
t O R 3/4 si�stone
t O R 3/4 si�stone
claystone
t OY 6/2 to 5Y 512 silty
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reduction mottles Fe glaebules/
•
concretions
I A. 9
ee
�
-
'7
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granular peds
�
-
�
angular blocky peds platy peds
Catskill - Duncannon Mbr.
cross-bedded 5R 4/2 clayshale, fissile
0
[]
•
�
Skolithos
@]
5R 412 silty
desiccation cracks dolom�e/calc�e nodules
rhizol�hslrhizocretions Lingulid brachiopods articulate brachiopods general bioturbation framboidal pyr�e
(L.
Carb.)
cross-bedded
Pennington
claystone
burrows
root traces
Mauch Chunk
claystone
0\ ..,..
Legend
(L.
N4 sandstone
Carb.)
5R 412 to N4 clayshale,
fissile
t OYR 212 claystone 5R 4/2 silty
t OR 3/4 claystone
claystone
t O R 3/4
t OR 3/4 claystone
l 1d::::<£) :: -
-
1
-
lc-=--...c-> .q
t OR 3/4 wave-rippled si�stone to very fine sandstone
em
current-rippled fine sandstone
5GY 4/t clayshale, fissile
Fig. 3. Representative columns describing features of Palaeozoic vertic claystone palaeosols, Appalachian Foreland Basin, USA. (a-c) Palaeosols formed in coastal margin palaeoenvironments, whereas in ( d-f) they have developed in alluvial-plain palaeoenvironments.
0 !"--<
�
;::;
!':> ;,s
!':>... )-'l
0 tl
....
�·
"'
Palaeozoic red-bed palaeosols association with peat swamps and mires (Donaldson 1974; Cecil et al. 1985; Caudill et al. 1992a; Joeckel 1995). Late Ordovician vertic palaeosols of the Juniata Formation in Tennessee consist of slickensided marginal-marine claystone containing inarticulate brachiopods (Fig. 3a; Driese & Foreman 1991, 1992). One palaeosol was bioturbated by marine invertebrates during transgression and submer gence, as manifested by Skolithos burrows with prominent reduction haloes and pynt1zation, which penetrate the top (Fig. 3a). Juniata Formation vertic palaeosol chemistry indicates salinization, enrichment of phosphate and some marine trace ele ments, and localized iron reduction towards the top of the palaeosol, which are all associated with marine transgression and flooding (Fig. 4a). The marine burrows and soil fractures served as permeable flow paths for marine fluids (Driese & Foreman 1991, 1992). Further evidence of marine modification of coastal-margin vertic claystone palaeosols is pro vided from the Bloomsburg Formation (Upper Silurian) of central Pennsylvania (Driese et al. 1992; 1993a). Palaeosols show extensive slickensides and developed from coastal mudflat deposits containing a marine fauna that includes articulated brachiopods (Fig. 3b ). After flooding and submergence of the palaeosols during metre-scale transgressions, the palaeosols experienced salinization, calcification, enrichment of phosphate and some marine trace ele ments, and localized iron reduction (Fig. 4b ), changes similar to those documented in the Juniata Forma tion by Driese & Foreman (1991, 1992); such chemi cal modification of the palaeosol by marine fluids was termed marine hydromorphism by Driese et al. (1992). Late Mississippian (late Early Carboniferous) vertic claystone palaeosols occurring in the Penning ton Formation of Tennessee palaeosols are slicken sided and formed from lagoonal and coastal mudflat deposits exposed during metre-scale sea-level drops (Fig. 3c; Caudill et al. 1992b, 1996). Pedogenic carbon ate is present in many of the palaeosols, and is chiefly dolomite that was probably precipitated during or just immediately after pedogenesis (Fig. 4c; Caudill et al. 1992b ). One palaeosol, interpreted as a palaeo Vertisol, has a complete profile preserved as a result of the fortuitous precipitation of a dolomite phosphate crust, which armoured the top of the palaeosol and precluded significant erosion upon submergence and burial (Fig. 4c; Caudill et al. 1996);
65
this palaeosol exhibits granular peds at the top and a pedogenic calcite horizon at depth that permits a palaeoprecipitation estimate of 648 ± 141 mm yr-1 (Fig. 3c; Caudill et al. 1996). High-sinuosity, alluvial channel-floodplain environments and palaeosols High-sinuosity alluvial channel and floodplain deposits comprise a major part of the post-Silurian, Palaeozoic molasse of the Appalachian Basin (Fig. 2; Table 1 ) . The architecture of alluvial deposits typi cally consists of repetitively stacked, upward-fining sequences 1-5 m thick, with palaeosols chiefly formed within the upper, clay-rich, floodplain portions of each sequence (Fig. 3d-f). Pre-Devonian alluvial deposits were low-sinuosity and dominated by braided patterns (Schumm 1968; Cotter 1978), even where deposited in a lower alluvial plain setting (e.g. Juniata Formation (Upper Ordovician) of central Pennsylvania), Table 1; Cotter 1978; Thompson & Sevon 1982). Some Juniata streams could be inter preted as meandering, however, based on the pres ence of upward-fining sequences 1-5 m thick in mudrock-rich parts of the Formation, which are capped by palaeosols (Thompson & Sevon 1982; Feakes & Retallack 1988). Vertic claystone palaeosols analogous to vertic Entisols and Inceptisols (Soil Survey Staff 1990) are especially abundant in the Catskill Formation (Upper Devonian) alluvial succession (Cotter et al. 1993; Driese & Mora 1993a; Capelle & Driese 1995). These palaeosols are extensively slickensided and formed on thick deposits of overbank and floodplain alluvium. Palaeosols lacking pedogenic carbonate development (Fig. 3d) presumably represent shorter durations of pedogenesis and/or poor soil drainage, whereas those with extensive pedogenic carbonate development (Fig. 3e) probably represent longer durations of pedogenesis and/or better soil drainage. Many profiles are extraordinarily thick (up to 5 m), and must represent cumulative profiles in which there were constant additions of sediment to the soil surface coincident with pedogenesis (Fig. 3e; Driese & Mora 1993a) or compound palaeosols consisting of several stacked soils welded together by pedogenic processes. The whole-rock chemistry of these palaeosols, although showing weak evidence for leaching, seems more strongly influenced by original depositional texture, as well as by the presence or absence of pedogenic carbonate (calcification, Fig. 4d & e) .
66
C. I. Mora and S. G. Driese
"E
o -2o
0.1
Juniata Palaeosol ( U . Ord.) 1
m
l l l l l lll
-40 .!::: -60 g. -80 0 -100 - 1 20 -���� .s
----
IN
--���
---------
Molecular ratio leaching --+-- Ba/Sr base loss ---18-- AI203/Ca0 + MgO + Na20 + K20 salinization --...6-- Na20/K20 oxidation --0- Fe203/AI203 AI203/Si02 clayeyness --¢-calcification -4>- CaO + Mg0/AI203 (a) Bloomsburg Palaeosol (U. Sil.) 10 1 0.1 0 -40 -80 -120 "E .s -160 .!::: 15. -200 Q) 0 -240 -280 -320 -360 Molecular ratio (b) 0.1
"E
.s .!::: 15. Q) 0
10 -10 -30 -50 -70 -90 -1 10 -130 -150 (c)
Pennington Palaeosol (L. Carb.) 1
0.1
10
"E
.s .!:::
15.
Q) 0
0 -40 -80 -120 -160 -200 -240 (d)
Catskill (Irish Valley) ( U . Dev.) 1
Molecular ratio Catskill (Duncannon) ( U . Dev.) 1
"E
.s .!::: 15. Q) 0
-50 -100 -150 -200 -250 -300 -350 -400 -450 -500 (e)
10
Molecular ratio 0.1
10
10
Mauch Chunk Palaeosol (L. Carb.) 1
10
0 -50 "E
-100
Q) 0
-1 50
.s .!::: 15.
-200 Molecular ratio
-250 (f)
Molecular ratio
Whole-rock XRF data, expressed a s molecular ratios (see Retallack 1990), for palaeosols depicted i n Fig. 3. (a-c). Data are from palaeosols formed in coastal-margin palaeoenvironments, whereas (d-f) data are from palaeosols developed in alluvial-plain palaeoenvironments. See discussion in text. Fig. 4 .
Palaeozoic red-bed palaeosols Late Mississippian (late Early Carboniferous) vertic claystone palaeosols occur in the upper parts of Mauch Chunk Formation upward-fining alluvial deposits (Driese et al. 1993b; Fastovsky et al. 1993). Palaeosols are slickensided and exhibit varying degrees of pedogenic carbonate development, with some horizons thick and massive enough to qualify as K horizons (Fig. 3f). As was the case for the Catskill (Upper Devonian) palaeosols, Mauch Chunk palae osol whole-rock chemistry is largely inherited from the parent material and shows little variation that can be attributed to pedogenesis, except for that related to the presence or absence of pedogenic carbonate (calcification, Fig. 4f) . Late Pennsylvanian (late Late Carboniferous) and early Permian vertic claystone palaeosols occur in the Conemaugh and Dunkard Group, respectively (Caudill et al. 1992a; Fastovsky et al. 1995; Jaeckel 1995; Caudill 1996; Caudill & Driese, submitted). Conemaugh Group palaeosols exhibit striking colour variations, with chromas ranging from < 2 to > 6; low chroma portions of palaeosols apparently formed by groundwater pseudo-gley as water tables perched on top, and within, low-permeability horizons, some of which include lacustrine and palustrine limestones (Fig. 5; Caudill et al. 1992a; Caudill 1996). Although the occurrence of abundant red, oxidized, vertic claystone palaeosols in direct juxtaposition with superjacent coals (palaeo-Histosols) may seem con tradictory, it apparently relates to a progressive dete rioration of soil drainage conditions preceding peat mire development that was possibly transgression driven (Caudill & Driese, submitted). Palaeoclimate information
The palaeogeographical reconstructions of Ziegler et al. (1979) and Scotese et al. (1979) placed the Appalachian Foreland Basin region at about 20°-30° south palaeolatitude during Late Ordovician and Silurian times, with the palaeoequator trending N-S (present orientation) through the centre of the Laurentian continent. As the Laurentian continent progressively rotated counterclockwise, by Late Devonian to Mississippian (Early Carboniferous) time the Appalachian region was located at about 4°-10° south palaeolatitude (Van der Voo et al. 1979; Kent 1985). Pennsylvanian (Late Carboniferous) and Permian reconstructions place the Appalachian Foreland Basin more or less astride the palaeoequa tor; most palaeogeographical models place the region within 5° north or south of the palaeoequator during
67
the Carboniferous and Permian Periods (Heckel 1980;Witzke 1990; Crowley & Baum 1991). Palaeoclimatic models for Late Ordovician to Silurian times predict warm, moist winters and hot, dry summers (Ziegler et al. 1977). The Devonian palaeoclimate was subtropical to tropical and strongly controlled by the orographic effects of the Acadian orogen, which would have blocked south easterly trade winds, resulting in a seasonally wet and dry (monsoonal) pattern of precipitation (Woodrow et al. 1973; Woodrow 1985). The general post Devonian palaeoclimate was 'megamonsoonal' and strongly influenced by the Appalachian Orogen, which acted as an orographic barrier (Kutzbach & Gallimore 1989). The palaeoclimate varied from drier during the Early Carboniferous (Mississippian) to wetter during the early Late Carboniferous (Early and Middle Pennsylvanian), to drier once again during latest Late Carboniferous (Late Pennsylvan ian) and Permian times (Cecil 1990; Heckel 1995), based upon abundance and distribution of coals, and the colour of associated palaeosols. The abundance of vertic (shrink-swell) features preserved in most of the red-bed claystone palaeosols (Figs 3, 5 & 6a,b; see also descriptions in subsequent section) reinforces palaeoclimatic models predicting strong seasonality of precipitation (Soil Survey Staff 1990). The climate necessary for development of Holocene Vertisols (and vertic fea tures in other soil types and vertic intergrades) must be seasonally moist and typically tropical to warm temperate, with typically 4-8 dry months each year (Ahmad 1983; Ductal & Eswaran 1988). Such a sea sonal wet-dry palaeoclimate can be inferred for the Appalachian Foreland Basin region throughout the Palaeozoic Era, based on the widespread distribution of red-bed palaeosols with vertic features (Figs 2, 3 & 5) and the existence of favourable palaeolatitudes and palaeogeography, as discussed previously. That pedogenic carbonate deposits occur in nearly all of the red-bed palaeosols in the Appalachian Foreland Basin succession also ha palaeoclimatic · significance (Figs 3, 5 & 6c,d; Table 1 ). Pedogenic car bonate horizons form in Quaternary soils under conditions of low mean annual precipitation (< 50 em yr-1), or under higher precipitation where there is a significant moisture deficit as a result of high evapo ration or evapotranspiration (Goudie 1983; Cecil 1990). The abundance of pedogenic carbonate in the red-bed palaeosols is compatible with a warm tropi cal to subtropical palaeoclimate and seasonal mois ture deficit.
68
C. I. Mora and S. G. Driese
2428
2552
Legend
5G4/1
Om
0
0 0 0
5Y5/1 5GY4/1 5Y4/2 5GY4/1
Om
3 4 5
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slickensides� J. (calcite-lined) _:;;? clastic dikes
�
rhizocretions � carbonate nodules 19 haematite nodules Ell burrows 1 ostracods (g) marine fossils 0 pyrite 0 siderite 181 vein-network (siderite or haematite) � root mottles 1; root traces (carbonized) t1 colour (chroma) mottles C:::::, ' 0 . blocky peds weak oo moderate go stong BB granular peds 00 platy peds black/greyish black • dark to med. grey • chroma s 2 chroma > 2 s 4 D chroma > 4 s 6 D chroma s 6 D
Fig. 5.
Schematic profiles of colour-mottled Upper Pennsylvanian (upper Upper Carboniferous), sub-Ames (Conemaugh Group) palaeosol complex formed at sites interpreted to have been well-drained (from Caudill 1996).
C H A R AC T E R I S T I C P H Y S I C A L , CHEMICAL AND BIO LO GICAL F E AT U R E S Claystone matrix
Vertic features Prominent slickensides occur in all Palaeozoic clay rich palaeosols (Figs 3, 5 & 6a,b ); these 'pedogenic slickensides' (Gray & Nickelsen 1989) are orientated randomly and locally form pseudo-anticlines (cf.
Driese & Foreman 1992; Driese et al. 1992; Driese & Mora 1993a; Caudill et al. 1996). Pedogenic slicken. sides form in clay-rich soils when swelling pressures exceed shear strength at depths where vertical movement is confined and may result in development of surface hummock-and-swale structure (gilgai) expressed as pseudo-anticlines in the subsurface (Watts 1977; Yaalon & Kalmar 1978; Knight 1980; Wilding & Tessier 1988). Pedogenic slickensides are orientated randomly, in contrast to tectonic slickensides, which generally are aligned pre ferentially (relative to a stress field) in response
Palaeozoic red-bed palaeosols
69
(a)
(c)
(b)
(d) Fig. 6.
Examples of vertic features and pedogenic carbonate deposits in Catskill Formation (Upper Devonian) vertic claystone palaeosols depicted in Fig. 3( e). (a) Pedogenic slickensides intersecting to form pseudoanticline (above 15 em scale card). (b) Pedogenic slickenside surfaces (smooth) with random orientations. Note also the well-developed medium angular blocky ped fabric. (c) Calcite rhizoliths (white) in palaeosol claystone. Lens cap is 5.5 cm in diameter. (d) Calcite nodules (white) in palaeosol claystone.
to structural deformation (see Driese & Foreman 1992). Well-developed sepic-plasmic (bright-clay) micro fabrics consisting of subangular to angular blocky aggregates of sand, silt and clay bounded by orien tated clay minerals with bright interference colours are interpreted as 'peds' formed by differential shear ing (Fig. 7a); the bright-clay coatings on the peds, or 'stress cutans' (Brewer 1976), are associated with wetting and drying cycles (Fig. 7a-c; Nettleton & Sleeman 1985; Wilding & Tessier 1988; Blokhuis et al. 1990). Many of the claystone palaeosols (Figs 3 & 5) therefore have been interpreted previously by us as being analogous to Holocene Vertisols, vertic Enti sols, and vertic Inceptisols (Soil Survey Staff 1990) based on the abundant vertic (i.e. Vertisol-like) macro- and microfeatures.
Clay content and mineralogy Parent material of Holocene Vertisols typically has a high clay content (> 30% ), consisting predomi nantly of expandable smectite mineralogies possess ing a high shrink-swell potential (Ahmad 1983; Ductal & Eswaran 1988; Soil Survey Staff 1990). The requirement of a high clay content is consistent with our observation that most of the Appalachian Foreland Basin palaeosols occur in the upper, clay-rich parts of sedimentary upward-fining sequences (Figs 3 & 5). Progressive burial diagenesis, however, altered the clay mineralogies of all Appalachian palaeosols examined to predominantly well-ordered illites and Fe chlorites, with no preservation of original expandable clays (Fig. 8; cf. Gray & Nickelsen 1989; Driese & Foreman
70
C. I. Mora and S. G. Driese
(b)
(a)
(c)
0 . 5 mm Fig. 7. Micromorphology o f vertic claystone palaeosols. (a-c) These are under crossed-polarizers, whereas ( d)-(h) these are in plane-polarized light. Parts (a)-(c) are from Juniata Formation (Upper Ordovician) palaeosol depicted in Fig. 3(a). (d), (e) and (h) are from Catskill Formation (Upper Devonian) palaeosol depicted in Fig. 3(e). (f) and (g) are from Mauch Chunk Formation (Upper Mississippian, upper Lower Carboniferous) palaeosol depicted in Fig. 3(f). (a) Angular blocky ped (p) encircled by birefringent clays (bright white). (b) Reworked ped (pedorelict?,p) bounded by birefringent clays (bright white). (c) Stress-orientated clays (bright white) aligned along pedogenic slickenside surface. Dark grain at bottom centre is haemetite glaebule. (d) Vertical root traces (r) lined with clay and Fe oxide hypocoatings. (e) Micrite nodule (n) cross-cut by sparry calcite cement (s ) filling septarian shrinkage void. (f) Dense micrite showing incipient pisoid grain development; note coatings on grain (arrows); sparry calcite cement (s ) occludes interpisoid porosity. (g) Longitudinal cut through rhizolith (RH) rimmed by micrite and with centre (arrow) infilled with Fe oxides, clays and detrital quartz grains. (h) Axial cut through rhizolith (RH) rimmed by micrite and with centre (arrow) infilled with Fe oxides, clays and detrital quartz grains.
(d)
Palaeozoic red-bed palaeosols
71
(e)
(g)
(f)
Fig. 7.
Continued.
1992; Driese et al. 1992; Driese & Mora 1993b; Sheldon 1995; Mora et al., 1998). In one of the least-buried vertic claystone palaeosols there is preservation of kaolinite, which increases upward at the expense of illite (Fig. 9); this relationship appar ently relates to greater intensity of weathering towards the top of the palaeosol (Caudill et al. 1992b; Sheldon 1995). Biological features
Major diversification and adaptive radiation of land plants occurred during the Palaeozoic Era, with the zenith of land plant evolution occurring during the Early and Middle Devonian (Table 1). Colonization of the terrestrial environment by plants followed a major colonization by land animals (Gray & Shear 1992). The presence or absence of soil macrofiora and macrobiota predictably would be manifest by atten dant changes in soil morphology and soil chemistry.
These relationships are summarized in Driese & Mora (in press). Animal burrows Animal burrows are the characteristic macroscale biological features of Ordovician and Silurian red bed palaeosols (Table 1). Retallack & Feakes (1987) reported animal burrows with meniscate structures occurring in alluvial palaeosols in the Juniata Forma . tion (Upper Ordovician) of central Pennsylvania, and interpreted them as the earliest evidence of dry soil animals. Driese & Foreman (1991, 1992) also described vertical animal burrows occurring in pedogenically modified tidal-fiat deposits in the Juniata Formation of eastern Tennessee, but these burrows are associated with marine flooding surfaces and clearly post-date palaeosol formation. Large burrows, 1-3 cm in diameter and up to 30 cm long dominate palaeosols formed in more proximal parts
C. I. Mora and S. G. Driese
72
--' "'
�
:::! d
r-- o 1'-
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:::!
:::!
Fig. 8.
66.66 33.33 1 5.6 19.0 22.4 25.8 29.2 32.6 36.0 Two - theta (degrees) d
1�
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-,-,-
-,-,--,-c�r-,--,-,--�,
66.66 33.33 36.0 39.4 42.8 46.2 49.6 53.0 56.4 59.8 63.2 66.6 70.0 Two - theta (degrees)
0
E (,)
.c c. C1) "'0
0
50
1 00
-40
-80
- 1 20 - 1 60
-200
-240 Fig. 9.
Clay mineralogy as a function of depth for Pennington Formation (Upper Mississippian, upper Lower Carboniferous) palaeosol depicted in Fig. 3(c). Note upward increase in kaolinite and upward decrease in illite; chlorite appears invariant.
of the Bloomsburg Formation (Upper Silurian) red bed succession in central Pennsylvania and south eastern New York (Retallack 1986; Driese & Mora, in press). Animal burrows are less important biotur bation features in Middle Devonian and younger
Whole-rock XRD spectra for three red claystone palaeosols showing similar mineralogical composition despite greatly differing geological age. Lower curve: Bloomsburg Formation (Upper Silurian) palaeosol depicted in Fig. 3(b ), sampled at -180 em level. Middle curve: Catskill Formation (Upper Devonian) palaeosol depicted in Fig. 3( d), sampled at -200 cm level. Upper curve: Mauch Chunk Formation (Upper Mississippian, upper Lower Carboniferous) palaeosol depicted in Fig. 3(f), sampled at -120cm level. Minerals identified include: Fe chlorite (CH), 2M 1 , 1M and lMd illites (IL), quartz (Q), haematite (H), and calcite (C). Bloomsburg sample contains more calcite because parent material was a marine claystone.
palaeosols, in which turbation is dominated by plant root systems. Root traces Silurian and Early Devonian plant commumtJes dominantly occupied wet coastal-margin environ ments owing to lack of true root systems and poor reproductive strategies (Gensel & Andrews 1984, 1987). By Middle Devonian time, plants occupied exposed parts of the coastal-alluvial plain between the river courses, and by Late Devonian time rela tively diverse communities of shrubs and trees, par ticularly vascular plants with well-developed root systems, had begun to spread into continental inter iors. Although concentrated on levees and near standing bodies of water, the plants expanded into drier interfiuve areas (Banks et al. 1985). The earliest unequivocal root traces with rhi zomous habit in red-bed palaeosols of the Appalachian Foreland Basin occur in the upper part of the Bloomsburg Formation (Upper Silurian) at Danville, and Port Clinton, Pennsylvania (Driese et al. 1992; Driese & Mora, in press). The traces are preserved as anastomosing, micrite-lined, 0.5-2mm
Palaeozoic red-bed palaeosols
diameter tubes that display a decidely horizontal, rhi zomous mat-like aspect (cf. Driese et al. 1992; 1993a); central pore spaces are typically occluded by one or more generations of calcite spar cement. Similar rhizomous structures occur within large micrite nodules found in Late Silurian (Ludlovian Pridolian) palaeosols in the Moydart Formation of Nova Scotia (Driese & Mora, in press) . Late Silurian rhizomous structures probably represent simple below-soil extensions of above-soil plants (rhynio phytes, probably Cooksonia ) that were of relatively small stature (perhaps 2-3 em tall) living in coastal mudflat and marsh environments. Alternatively, they may represent calcified nematophyte thalli, possibly of the non-vascular plant Nematothallus, which has been reported from the Bloomsburg Formation (Strother 1988). Root structures occurring in late Early Devonian (Emsian) palaeosols of the Battery Point Formation of Quebec are considerably larger and less cryptic than those of the Silurian, and are the earliest exam ples of large root traces exhibiting a dominantly ver tical disposition (Driese & Mora, in press; Elick et al., 1998) . Root systems from Middle Devonian (Givet ian) palaeosols containing large (up to 55 em diame ter and 1 0 m tall) in situ stumps of arborescent cladoxylaleans (Eospermatoperis) are considerably larger ( 1-2 em diameter, up to 15 em long), although the plants still occupied wet, marsh-type settings con stituting the famous 'Gilboa Forest' of New York (Bridge & Willis 1994; Driese et al., 1997) . By Late Devonian (Famennian) time large (up to 1 8 m tall) arborescent progymnosperms (Archaeopteris) occu pied well-drained, upland alluvial environments characterized by seasonal moisture variations (Banks et al. 1985; Gensel & Andrews 1987) . Archaeopteris root systems are larger (10-15 em diameter and up to 1 .5 m deep) , and exhibit a more 'modern' morphol ogy characterized by deep vertical tap roots and large secondary roots that branch from primary tap roots (Driese et al., 1997) . Most root traces we observe in Appalachian Foreland Basin palaeosols, however, are more commonly of relatively small stature, ranging from a few millimetres to a few centimetres in diameter (e.g. Fig. 7d) . Carbonate Morphology
Pedogenic carbonate occurs m nearly every Appalachian red-bed formation that contains palaeosols (Table 1) . Carbonate morphology is rela tively simple in Ordovician and Silurian palaeosols,
73
and shows a trend towards increasing diversity in the Middle to Late Palaeozoic that parallels land plant evolution and diversification (Table 1; Driese & Mora, in press) . Nodules
The earliest unequivocal pedogenic carbonate in the Appalachian Foreland Basin succession occurs in Late Silurian palaeosols of the Bloomsburg and Moydart Formations as micrite nodules, which range from 0.5 to 5 em in diameter (Table 1 ; Mora et al. 1991; Driese et al. 1992). Nodular pedogenic carbonate exhibits two characteristic morphologies: spherical to prolate micrite nodules, which are most common below the lowest occurrence of pedogenic slickensides within any individual palaeosol (Figs 6d & 7e); and large compound macronodules, up to 15 em in diameter, which consist of cemented aggregates of smaller diameter micrite nodules (Fig. 7f; Driese & Mora 1993a). Neither morphology resembles that of a 'rhizolith', calcite precipitated in association with plant roots (Klappa 1980). Micro scopically, the micrite is very finely crystalline, includes floating sediment grains, and exhibits a uniform dull luminescence under cathodolumines cence (CL) conditions. The micrite therefore resem bles the 'alpha-type' calcrete microstructure defined by Wright (1990) , for which precipitation is largely abiological. Pedogenic carbonate nodules are dolomitized in Mississippian (Lower Carboniferous) coastal-margin successions such as the Pennington and Maccrady Formations (Caudill et al. 1992b, 1996; Stefaniak et al. 1993; Driese, unpublished data) . The dolomite is extremely finely crystalline and replacement is inter preted as very early. Evaporite pseudomorphs occur in association with dolomitized carbonate nodules in the Maccrady Formation, which also contains evap orite dissolution breccias. Rhizoliths/rhizoconcretions
Pedogenic carbonate exhibiting a rhizolith morphol ogy first occurs in th� Appalachian Foreland Basin succession in the Battery Point Formation (Lower Devonian) as downward-branching, micrite-lined, 1-5 mm diameter features (Table 1; Driese and Mora, in press). Rhizoliths in Middle Devonian and younger palaeosols appear as highly branching micritic cylinders that are millimetres to tens of centimetres in diameter (Figs 6c & 7g,h; Mora et al.
74
C. I. Mora and S. G. Driese
1991; Driese & Mora 1993a, in press). In thin-section, these rhizoliths consist of an outer region of dense micrite containing organic carbon inclusions, rare microbial tubes, alveolar-septal fabric, and calcified faecal pellets. Such features, in addition to the general rhizolith morphologies, result from largely biologi cally induced calcification, characteristic of 'beta type' calcretes (Wright 1990). Many rhizoliths also exhibit an inner region (pore) that is occluded with clear, calcite spar cement. The root-marginal micrite displays a uniform dull luminescence, whereas former root voids are filled by an early dull lumines cent calcite spar cement, followed by a very brightly luminescent calcite spar cement (cf. Driese & Mora 1993a). Rhizoconcretions occur as vertically stacked or coalesced micrite nodules that exhibit an overall root morphology (Mora et al. 1991; Caudill et al. 1992a,b, 1996; Driese et al. 1993b; Fastovsky et a/. 1993). Rhizo concretions probably represent early formed rhi zoliths in which calcite precipitation was enhanced or supplemented by groundwater input.
The stable carbon and oxygen isotope compositions of soil-formed minerals, including carbonates, clays and oxides, may preserve information about the pedogenic environment, including palaeoclimate, palaeoecology and palaeoatmospheric levels of C02 (Cerling 1984, 199 1 ; Bird & Chivas 1988; Yapp & Poths 1992; Mora et al. 1996). Studies of modern soil systems (Cerling 1984; Quade et al. 1989) have eluci dated the important controls on carbon and oxygen isotopes in soil carbonate; these studies provide useful guidelines for the evaluation of isotopic com positions in Palaeozoic red-bed palaeosols. Some of the challenges associated with interpretation of iso topes in ancient palaeosols have been discussed in previous studies (Cerling 1991, 1992a; Wright & Vanstone 1991, 1992; Mora et al. 1993, 1996). Questions raised by these studies, and other specific issues, are addressed below.
Microcodium!cyanobacterial spherulites
Composition of soil organic matter
Soil microorganisms (bacteria, fungi) may affect carbonate precipitation in a number of ways, includ ing microbial uptake or respiration of C02, degrada 2 tion of organic matter, or excretion of Ca + (Klappa 1978; Monger et al. 1991; Wright & Tucker 1991; Verrecchia et al. 1995). Petrographic evidence for microbial action includes the occurrence of spherulites (biomineralization associated with cyanobacteria; Verrecchia et al. 1995) or Microco dium ('rosette' or 'corn-cob' like aggregations of calcite associated with rhizoliths that are thought to be the product of a symbiotic root-fungal associa tion; Klappa 1978; Freytet & Plaziat 1982; among others). The oldest known Microcodium rosettes in the Appalachian Foreland Basin succession occur in the Catskill Formation (Upper Devonian) of central Pennsylvania (see Fig. 6a in Driese & Mora 1993a). The rosettes consist of individual bladed crystals, 100 11m wide by 500-800 11-m long, arranged in a radi ating manner around a hollow axis. Smaller scale, spherulitic structures observed in the Catskill Forma tion, as well as in younger Palaeozoic palaeosols (Maccrady and Pennington Formations), are inter preted as spherulites formed by cyanobacteria, in accordance with recent work by Verrecchia et al. (1995).
One of the most important controls on the carbon isotope composition of pedogenic carbonate is the composition of the soil organic matter (Cerling 1984), the isotopic composition of which depends largely on the photosynthetic pathway(s) utilized by the soil biomass. Although it has been proposed that the C4 (Hatch-Slack) pathway arose several times in the geological past (Spicer 1989), and may have constituted a significant proportion of the biomass in some Carboniferous palaeosols (Wright & Vanstone 1991, 1992), the fossil evidence for plants likely to have utilized the C4 pathway extends only to the Miocene (Thomasson et al. 1988). In the absence of isotopic evidence for C4 vegetation, studies of Palaeozoic palaeosols may reasonably assume an exclusively C3 (Calvin cycle) flora. The isotopic composition of C3 plants in the geological record, as represented by compositions of coals, lignin, and palaeosol organic matter appears to be comparable to modern C3 plants (Degens 1969; Popp et al. 1989; Cerling 1992b ). Terrestrial plants utilize atmospheric C02 during photosynthesis and it is expected there fore that the isotopic composition of vascular plants tracked secular changes that are inferred for the 8BC of Palaeozoic atmospheric C02 (Fig. 10). The isotopic composition of atmospheric C02 is regulated on a
S TA B L E I S O T O P E G E O C H E M I S TRY OF P A L A E O Z O I C R E D B E D PALAEO S O LS
75
Palaeozoic red-bed palaeosols
1 0 .-------�
0
r------ ---- -----�
5
m 0 D..
()
C")
......
c.o
-10
_20
-25
paleosol om -
-1 0
-_
-1 5
. -350
-300
-250
Age (Ma) Fig. lO.
Secular trend in the carbon isotope compositions (±1 cr SD from the mean; Mora et al. 1996) of Appalachian palaeosol bulk organic matter (om) compared with trends established for shallow marine inorganic carbon (I C) (after Lindh 1983) and atmospheric C02 in equilibrium with marine I C.
long-term basis by equilibrium with marine inorganic carbon (marine carbonate rocks), which increased from O%o in the late Silurian to + 6%o in the early Permian (Lindh 1 983; Popp et at. 1986; Veizer et at. 1986). A number of studies (Degens 1969; Dean et at. 1986; Popp et at. 1989) suggest that the ()13C of terres trial organic matter remained relatively constant through the Phanerozoic Eon, despite inferred changes in ()BC of atmospheric C02, possibly as a result of physiological or biochemical regulation of plant internal p C02 and ()13C (Farquhar et at. 1982; Popp et at. 1989). These studies are at odds with isotopic studies of Permian coals (Jeffery et at. 1955; Compston 1960) and therapsid tooth enamel (Thackeray et at. 1990), which suggest significant changes (3-5%o) in ()BC between the early and late Permian. Palaeosols offer an inviting alternative to coal as a rock type that may record the terrestrial organic carbon record. Although organic matter is scarce in most Palaeozoic red-bed palaeosols, it is possible to concentrate sufficient bulk organic carbon for iso topic analysis in some palaeosols (Mora et at. 1996). Results from Appalachian Foreland Basin palaeosols (18 samples) indicate that the ()BC value for bulk organic matter increased from -27.4%o in late Silurian palaeosols to -23.5%o in early Permian palaeosols
1
-10
3 80
{PDB)
-5
oloo
0
Fig. ll.
Stable isotope compositions of pedogenic micritic rhizoliths (wlid symbol) or nodules (open symbol) for Late Mississippian (late Early Carboniferous) palaeosols that experienced different depths of burial: 7-8 km (11 ), 4 km (9,10), 1.5-3 km (12); palaeosol numbers are keyed to Table 1. Shown for comparison are other Carboniferous pedogenic carbonates reported by (G) Goldstein (1991) and (WV) Wright & Vanstone (1991).
(Fig. 10). These values track secular changes in the isotopic composition of middle to late Palaeozoic atmospheric C02, with a relatively constant isotopic fractionation of 21-22.5%o. Significantly, the compositions of soil organic matter in the Carboniferous palaeosols we examined do not require a component of C4 flora; therefore other explanations must be considered for the iso topically heavy pedogenic carbonate sometimes measured in palaeosols of this age (Fig. 1 1 ) (Gold stein 1991; Wright & Vanstone 1991). These explana tions may include factors affecting C02 diffusion in soils, including shallower depths of rooting, wetter soil conditions, or different vegetation types and den sities that affect the soil productivity, or contamina tion of pedogenic carbonate by the marine carbonate substrate of these palaeosols. The possible impact of animal-related organic matter (Table 1) on the carbon isotopic composition of bulk soil organic carbon was evaluated in five samples from a 30-50 cm organic-rich lens within a Silurian red-bed, palaeosol-bearing succession (described by Strother 1988). The samples had quali tatively different proportions of plant-derived (pos sibly Nematothallus) and animal-derived (cuticular fragments) organic matter, and ranged in ()13C from -27.0 to -28.3%o, respectively. Only the sample most
C. I. Mora and S. G. Driese
76
INLAND - AI..LWIAL PALEOSOLS: •<> Parbville, NY • oNewport, PA .6. �Sunbury, PA e 0 Trout Run, PA COASTAL - MARGIN PALEOSOLS:
g
-
0
.Q
+ liE Selinsgrove, PA
0
m -a 0 a..
-
('t)
0 -1 0
..... c.o
* -18
-16
-14
18 o 0
-12
•
(PDB) o/oo
-10
-8
Fig. 12. Stable isotope compositions of rhizoliths (solid symbols and crosses) and nodules (open symbol and stars) from Late Devonian palaeosols of the Catskill Formation (location 8,Table l , Fig. l). Note the greater variability of values for pedogenic carbonate developed in a coastal margin environment. Cyanobacteria may have precipitated much of the carbonate comprising nodules in one of the Selinsgrove, Pennsylvania palaeosols (stars).
dominated by cuticular fragments was lighter than -27.6%o, suggesting that even significant amounts of admixed animal organic matter do not significantly alter the bulk oBC of palaeosol organic matter. Little has been established about the influence of microbial action on the isotopi'c composition of pedogenic carbonate or soil C02. Carbonate inter preted as spherulitic is common in a coastal-margin palaeosol in the lower Sherman Creek Member of the Catskill Formation near Selinsgrove, Pennsylva nia (Fig. 12). These carbonate samples have ()13C values in the range established for cyanobacteria (Hoefs 1987) and are approximately 2-4%o lighter than pedogenic micritic rhizoliths from a more inland palaeosol occurring c.SOO m up-section. The influence of microbial action on oBC ultimately may depend on whether the microbes attack pre-existing pedo genic carbonate or precipitate a non-pedogenic car bonate phase that may later be admixed with (or misidentified as) pedogenic carbonate (cf. Verrecchia et al. 1995). Burial diagenesis
The potential for diagenetic modification of pedo genic carbonate chemistry has been viewed as a major obstacle to isotopic studies of Palaeozoic
palaeosols. Isotope exchange occurs during recrystal lization of pedogenic carbonate and clays in the palaeosol matrix. Because of the water-rich nature of many diagenetic fluids, the probablility for alteration of pedogei).ic oxygen isotope ratios during dia genesis is great. Alteration of ()1 80 values may occur without concomitant alteration of oBC values if dia genetic fluids are relatively carbon-poor or diagen etic water/rock ratios remain low. For example, Late Mississippian (late Early Carboniferous) palaeosols in the Mauch Chunk, Hinton and Pennington Forma tions (Table 1 ; Fig. 1) experienced burial conditions of 1 .5-3 km, 3-4 km and 7-8 km, respectively (Sheldon 1995; Mora et al., 1998). Despite oxygen isotope com positions that range from -3.2 to -15.9%o Pee Dee Belemnite (PDB) (Fig. l l ) , the average oBC values of pedogenic micrite in these palaeosols are -7.0, -7.6 and -7 .O%o, respectively. In general, more deeply buried palaeosols record lower ()18 0 values (Fig. l l ) , consistent with exchange between pedogenic carbon ates and 1 80-enriched diagenetic fluids that are strongly influenced by the transformation of smectite to illite. In certain circumstances (e.g. Pennington Formation) illitization, driven by wet-dry cycling, may proceed under pedogenic or very shallow burial conditions without significant influence on the ()1 80 value of pedogenic micrite (Mora et al., 1998). The preservation of pedogenic carbon isotope ratios through burial diagenesis suggests that o13C values of Palaeozoic pedogenic carbonates likely preserve meaningful information on the pedogenic environment, including atmospheric levels of C02. Variability in ()ISO values, however, is most likely the result of diagenetic overprinting, rather than an inter pretable pedogenic signal and, accordingly, we do not further discuss oxygen isotope ratios in this paper. Influence of the pedogenic palaeoenvironment
Palaeosols developed in both coastal-margin and inland alluvial channel-floodplain sediments (Table 1; Figs 3 & 4), and it is important to evaluate how the isotopic composition of the pedogenic calcite might be affected by pedogenesis in these different envi ronments. Perhaps the most predictable problem with pedogenesis in coastal-margin environments is the possible introduction of heavy marine carbon through detrital marine carbonate grains, or marine ground water or spray, leading to an apparent enrich ment in the isotopic composition of soil carbonate. Numerous pedogenic nodules in the Bloomsburg Formation (Upper Silurian) palaeosols contain finely
77
Palaeozoic red-bed palaeosols
comminuted marine skeletal grains (see Fig. 6c in Driese et al. 1992). The isotopic compositions of these largely pedogenic nodules (average 8BC=-1 .9%o) are clearly distinct from pedogenic micritic rhi zomous mats (average -5.32%o), which are devoid of detrital carbonate (see Fig. 13 in Driese et al. 1992). Carbon isotope compositions of rhizoliths from Late Devonian coastal-margin palaeosols are more variable and enriched in BC compared with rhi zoliths from four inland alluvial Catskill Formation palaeosols (Fig. 12). Similar systematic morphologi cal and stable isotopic differences between fluvial and supratidal caliche are also reported in the Permian Abo Formation, New Mexico (Mack et al. 1991). Several factors may contribute to these dif ferences. Supratidal carbonate morphologies are typically dominated by small nodules, horizontal rhi zoliths or thin, coalesced horizontal root mats, char acteristic of calcareous palaeosols formed in areas with high water tables (Cohen 1982; Mount & Cohen 1984). Gleying of the palaeosol clay matrix and relatively good preservation of organic matter are further evidence of a high water table in the suprati dal environment (Buurman 1980; Fastovsky & McSweeney 1987). By comparison, carbonate in fluvial palaeosols commonly is characterized by larger nodules and numerous, downward-tapering rhizoliths (Driese & Mora 1993a, in press). Even if the isotopic compositions of the soil biomass are equivalent in the two environments (i.e. both C3), dif ferences in the depth of rooting and soil moisture will affect diffusion of soil C02 (Ceding 1984), leading to more positive oBC for soil carbonates in coastal margin palaeosols. It was noted in an earlier section that some coastal margin palaeosols contain a significant amount of carbonate spherulites, which were likely associated with cyanobacterial mats. Misidentification of these surface-precipitated spher ulites (Verrecchia et al. 1995) as pedogenic calcite would further obfuscate interpretation of the isotopic compositions. Detailed examination of Late Mississippian (late Early Carboniferous) palaeosols in the Pennington Formation elucidate another common feature of coastal-margin palaeosols: the mineralogy and iso topic compositions of pedogenic calcite in these palaeosols have been variably altered by dolomitiza tion. Petrographic evidence for early post-pedogenic dolomitization of pedogenic calcite nodules includes retention of pedogenic fabrics, including disorthic nodules, root tubules and Microcodium. Microcrys talline dolomite is constrained to nodules; the clay
matrix essentially is devoid of carbonate. The carbon isotope ratios of dolomite nodules are significantly depleted in BC relative to Mississippian marine car bonates (Fig. 13), most likely reflecting inheritance of light soil carbon from the pre-existing pedogenic calcite. The 8180 values of dolomite nodules are significantly enriched in 180 relative to pedogenic calcite nodules and are similar to compositions of thin, supratidal dolomite beds interbedded with shale/palaeosol units in the lower portions of the section (Fig. 13). The isotopic compositions of calcite and dolomite in these palaeosols fall along arrays suggesting alteration of pedogenic calcite by a modifed marine fluid (Fig. 13). The arrays for each palaeosol have different slopes, suggesting that dolomitization occurred in the very early post pedogenic environment, prior to the deposition of the immediately overlying clastic material. The extreme variability of dolomitized pedogenic carbon ate may reflect different degrees of fluid-rock inter action accompanying dolomitization. We hypothesize that dolomitization of the pedogenic calcite resulted from minor, local transgressions which superimposed a highly evaporative, supratidal environment on
Miss. marine carbonate
-
m Q a..
-
u
('f) .....
c.o
-5
0 Stable isotope compositions of pedogenic carbonate in Late Mississippian (late Early Carboniferous) palaeosols within the Pennington Formation (location 12, Table 1, Fig. 1 ) . Early post-pedogenic dolomitization has altered pedogenic compositions along arrays between a pedogenic calcite end-member (box) and an evaporated, modified marine fluid. Many dolomitized nodules are similar in composition to a thin supratidal dolomite overlying palaeosol pLl . Fig. 13.
78
C. I. Mora and S. G. Driese
vertic soils developed in an emergent, marginal marine setting (see also Caudill et al. 1992b, 1996). Thus, the sedimentary environment of pedogene sis, whether inland alluvial or coastal margin, may contribute to significant differences in the isotopic compositions of pedogenic carbonate, as a result of different types of vegetation, shallower rooting depths or groundwater tables, and likely early post pedogenic processes that may alter the original isotopic compositions. These differences may compli cate (or make impossible) the interpretation of certain palaeoenvironmental variables, such as palaeoatmospheric C02 levels, but do not negate the use of these data to constrain general aspects of the palaeoclimate and palaeoenvironment. Mack et al. (1991) noted that regardless of the absolute carbon isotope composition, isotopic trends in both alluvial and supratidal pedogenic carbonate reflected changes in Permian palaeoclimate to drier and warmer conditions. Influence of physical pedogenic processes
Specific pedogenic processes, such as the wet-dry cycling of vertic soils, may leave their mark on the iso topic composition of pedogenic carbonate. In our studies of Palaeozoic vertic, red-bed palaeosols, we note that post-Silurian pedogenic carbonate assumes two general morphologies (spherical micrite nodules and micritic rhizoliths; Fig. 6c & d; Table 1), each occurring in identifiable zones within the palaeosol. Nodules are found higher in the soil profile, within the zone of soil shrink-swell, as indicated by numer ous pedogenic slickensides in the palaeosol matrix. Rhizoliths typically are found at slightly deeper levels, below the lowermost occurrence of pedo genic slickensides, and have a distinctly biogenic micromorphology. Details of nodule and rhizolith micromorphology, cathodoluminescence (CL) and chemistry are given elsewhere (Driese & Mora 1993a). The carbon isotopic compositions of the nodules are invariably enriched in 13C relative to rhizoliths in the same profile (Fig. 12), most likely reflecting an increased concentration of isotopically heavy, atmospheric C02 in the upper portions of the soil; vertic soils may crack open to depths of 1 m or more during the dry season (Dudal & Eswaran 1988; Wilding & Tessier 1988). Compositions in rhizoliths are lighter and less variable (Fig. 12), yielding the most consistent, minimum estimates of atmospheric C02 when using the soil carbonate palaeobarometer of Cerling (1991) . We have noted the isotopic distinc-
tion between nodules and rhizoliths in more than 20 Appalachian vertic palaeosols, ranging in age from Late Silurian through to Permian. Pedogenic carbon ate formed in other types of soils, subject to a differ ent set of physical processes, may not show this distinction. This example illustrates the need to eval uate the possible influences of physical pedogenic processes on the chemical and isotopic composition of soil-formed minerals. Soil C02 production in Late Silurian soils
Very old palaeosols (e.g. Silurian) pose a special problem for the soil C02 palaeobarometer. The prim itive plants that colonized these soils were small and lacked true roots (Table 1; Stewart 1983; Gensel & Andrews 1987; Driese & Mora, in press) and Late Silurian plant cover may have been sparse compared with more robust Devonian and Carboniferous ecosystems (Table 1; Gensel & Andrews 1984). These factors suggest a shallow mean production depth for C02 in these soils (possibly « 5 em), as well as a low rate of soil respiration; both factors would lead to lower estimates of atmospheric C02 for a given car bonate ()13C value (Cerling 1991). Possible mediating factors, such as incorporation of organic matter into deeper levels of Silurian soils as a result of pedotur bation or preferential plant colonization of moister soils (i.e. lowland or near-channel environments), would affect other soil factors, such as porosity, per meability or average depth of soil C02 production, thereby effectively increasing the organically derived proportion of soil C02. At present, quantifying these factors is very difficult and the errors in p(C02) esti mates for this time period must be considered quite large. Estimation of palaeoatmospheric p(C02)
We have used the isotopic compositions of pedogenic carbonate from red-bed vertic palaeosols in nine stratigraphical successions, ranging in age from late Silurian to early Permian, to estimate palaeoatmos pheric levels of C02 using the soil carbonate palaeo barometer of Cerling (1991). Our estimates are summarized in Fig. 14. To make these calculations, we used average o13C (to -1 a) of pedogenic micritic rhizoliths, o13 C for atmospheric C02 calculated from the time-appropriate estimate of marine inorganic carbon (Fig. 10) assuming a fractionation of -7%o, o13C of organic matter from the studied, or a time equivalent, palaeosol, and soil C02 = 3000-7000
79
Palaeozoic red-bed palaeosols
�
CD .r:.
� 15 0
1000
atmospheric p(C02) (ppmV)
.5
� 10 0
..... 5 <1
S(z)
= 5000 ppm V T = 25 C
14
16 18 20 22 24 1 ..i 3c (carbonate - o.m.)
26
Fig. 14.
Estimates of Palaeozoic atmospheric C02 levels using the soil carbonate palaeobarometer of Cerling (1991); palaeosols keyed to Table 1. See text for details of the calculations. Atmospheric C02 levels dip sharply between the late Silurian and early Permian, approximately coincident with afforestation of the land surface and global climate change leading to the extensive Permo Carboniferous glaciation.
p.p.m. (S(z) = 5000 p.p.m. in Fig.14), a range typical of many tropical to temperate soils (Cerling 1991); these data are tabulated in Mora et al. (1996). In contrast to the nearly constant carbon isotope fractionation between soil organic carbon and atmos pheric C02 (Figs 10 & 14), differences in the fraction ation between pedogenic carbonate and soil organic carbon range from -21%o in late Silurian to -14%o in early Permian palaeosols. Fractionations in the range 14-16%o are observed in modern, semi-arid, C3dominated Holocene soils (Quade et al. 1989) and indicate the absence of a significant input of atmos pheric C02; significantly larger fractionations are interpreted as indicating higher atmospheric levels of C02 in the mid-Palaeozoic. Estimates of atmospheric C02 levels from geographically separated, time equivalent palaeosols are consistent, suggesting that a coherent record of changing atmospheric chemistry is preserved in the ancient soil record. Estimates of Carboniferous p(C02) reported here and in Mora et al. (1996) are significantly lower than those we reported previously (Mora et al. 1991), as a result of heavier organic matter than we modelled originally ( -23.8%o instead of -26%o ) , and a lower range of esti mated soil C02 (3000-5000 p.p.m., instead of 5000-
1 0 000 p.p.m.); this range is more consistent with the more arid early Carboniferous palaeoclimate (Cecil 1990). Our results suggest that atmospheric C02 levels declined from 3200 to 5200 p.p.m. in the late Silurian to 150-200 p.p.m. in the early Permian, closely following a decline predicted by the theoreti cal carbon mass-balance model of Berner (1994). The most significant decrease, between the late Silurian and late Devonian, coincides with a period of rapid evolution and diversification of the terrestrial ecosys tem (Table 1; Stewart 1 983; DiMichele & Hook 1992; Driese & Mora, in press).
S U M M A RY A N D C O N C L U S I O N S
First-order controls on the morphology and stable isotope chemistry of Appalachian Foreland Basin red-bed claystone palaeosols include: 1 variations in geomorphological and pedogenic environments and soil parent material (coastal versus inland); 2 soil development in the presence or absence of vas cular plants with significant root systems. The morphology and stable isotope chemistry of Ordovician and Silurian palaeosols differ substan tially from that of Devonian, Carboniferous and Permian palaeosols partly because of differences in these variables. There is great potential in the use of Cerling's (1991) soil carbonate palaeobarometer, but careful consideration must be given to the many factors that may influence the isotopic compositions in the pedogenic environment.
AC K N OW L E D G E M E NT S
This research was supported b y PRF Grant 25678AC8-C-SF94, and by NSF Grants EAR-9206540 and EAR-9418183 awarded to Mora and Driese.
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Peculiar palaeowe athering typ e s
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
Spec. Pubis int. Ass. Sediment. (1999) 27, 87-127
Diversity of continental silicification features: examples from the Cenozoic deposits in the Paris Basin and neighbouring basement
M . T H I RY C. I. G., Ecole Nationale Superieure des Mines de Paris, 35 rue SLHonore, 77305 Fontainebleau Cedex, France and CNRS, UMR SISYPHE C 7619, Structure et fonctionnement des systemes hydriques continentaux
A B S T R AC T Silicification occurs i n almost all the Tertiary formations o f the Paris Basin and can b e grouped into two main types: pedogenic silcretes and groundwater silcretes. Pedogenic silcretes are linked directly with the surface. They display specific structures and fabrics, such as illuviations and cappings, and are well-differentiated between the top and the base of the sections. There is always only one single silicified horizon. On the edge of the Paris Basin, hardened red sandstone or hardpan developed on Tertiary deposits blanketing the Massif Central basement. The red sandstone surrounds the basement palaeo hills and marks out a buried palaeolandscape. Well-structured palaeosols with biological activity were developed in the transitional zone between uplands and lowlands. Silica hardening was superimposed on the palaeosol catena. Upstream profiles show a characteristic columnar structure and have been plugged by silica illuvi ation structures. Downstream profiles have a planar structure and contain opal deposited in subhorizontal fractures, indicating a prevailing lateral water flow. Concretionary deposits of purer opal formed further inside the Paris Basin. A silcrete armour caps the lower Tertiary formations and the palaeoweathering profiles of the Chalk. The silicified pans crown kaolinite-rich clastic sediments and cover a wide palaeosurface extending over the whole southern part of the Paris Basin. They consist of hard quartzite with typical columnar and large illuviation features. Several horizons can be identified in which the grain size of the illuviated material decreases from top to bottom of the profiles. There is abundant evidence of silica dissolution in the upper horizons and silica deposits in the lower ones. The fact that these silicified horizons have been reworked in late Eocene and Oligocene formations shows their early age. Products of silicification by ground water retain the primary structures of the host material and are always linked to an absolute silica accumulation, either by silica deposits in pores or by matrix replace ments. Tiley are found in clastic rocks as well as in limestone and evaporitic rocks. Superimposed flat-lying, very tightly cemented quartzite lenses occur in several sand formations of the Paris Basin. The quartzite is composed of quartz grains with overgrowths. Exposures in valley-side out crops of quartzite lenses developed in Fontainebleau Sand (Oligocene) show that they do not extend more than a few hundred metres beneath the limestone-covered plateaux. The quartzite is thus linked to the present or recent morphology and silicification occurred relatively recently in outcrop zones of the Fontainebleau Sand. The quartzite lenses are thought to relate to silica precipitation controlled by ground water flow. Almost all of the continental limestones in the Paris Basin contain silicified facies.TI1e silicified facies are distributed irregularly in the limestone and their sizes vary from millimetre-sized dots to units several tenths of metres long. They are discordant on the sedimentary structures, are of irregular shape, massive or have numerous inclusions, almost exclusively formed of quartz. Silica occurs as deposits in voids and replacement of the Limestone. Because there is always a strong relationship between voids and silicification, and because the limestone is pure, without clay or sand layers, the silica also is thought to be related to groundwater flow. Peculiar siliceous and cavernous rocks, called meulieres in the Paris Basin, result from the weathering of the silicified lacustrine limestones, which are embedded in a residual clay formation, the 'Argiles a Meulieres'. These weathering horizons mark out the plateau surfaces of the southern Paris Basin. The meulieres display a large variety of facies, which develop from the most massive facies to highly porous Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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M. Thiry 'cellular' ones and are finally weathered into sand-like materials. Silica redistribution plays a major role in the development of the meulieres facies. Weathering causes dissolution of the weakest crystallized silica phases in the upper horizons of the profiles and deposits of secondary silica phases in the lower ones. Pedogenic silcretes are linked directly to surface conditions. They relate either to a relative accumula tion of silica with leaching of the other elements or to an absolute silica accumulation. Pedogenic silcretes generally provide good markers for the reconstruction of palaeosurfaces. Moreover, they indicate long periods of palaeolandscape stability and alternating wet and dry climates. Ground water silcretes are not linked directly to the surface and soils. They develop at depth in relation to groundwater pathways and water-table levels. Groundwater silicification features relate to absolute silica accumulation processes. They can develop in superimposed layers in relation to water-table fluctuations and/or downcutting of the landscape. From this standpoint they could be valuable guides to the geodynamic evolution of a basin.
I NT R O D U C T I O N
Silicification features are widespread i n continental environments. Although silicification in supergene conditions has been recognized for a long time (Gosselet 1 888; Kalkowsky 1 901; Lamplugh 1907; Woolnough 1927; Stortz 1928), these materials have, until recently, mostly been called silcretes in general descriptions, without further details being provided. The term 'silcrete' is widely used to refer to silicified materials of different structures and fabrics. Their origin usually is thought to be diagenetic processes above the phreatic zone or in environments associ ated with soil formation. During the last 20 years, interest in supergene silicification has been reviewed and new ideas have appeared. New studies have taken advantage of improved knowledge of the geochemistry in surface environments and new techniques. In particular, the interpretation of the micromorphological structures of soils has been deci sive for the understanding of pedogenic silcretes. A more quantitative approach, formulating questions in terms of mass balance, has led to different approaches to the questions of silicification in con nection with ground water. Silcretes are widely recognized features of land scapes in Australia (Woolnough 1 927; Stephens 197 1 ; Wopfner 1978; Callen 1983; Milnes e t al. 1991; Simon Coingon et al. 1996), in southern Africa (Goudie 1973; Summerfield 1982), in parts of western Europe (Summerfield & Whalley 1980; Blanco & Cantano 1983; Thiry et al. 1983; Wopfner 1983; Meyer & Pena dos Reis 1985), in parts of North and South America (Goudie 1973), in northern Africa (Millot et al. 1959; Conrad 1969; Thiry & Ben Brahim 1 990) and in the Arabian Gulf (Khalaf 1988). Note that all the above continental silicification features occur in relatively young deposits (mostly Tertiary) that have never undergone burial diagenesis. In older and deeper buried deposits, late diagenetic transformations have
occurred, making it difficult to recognize the features of early supergene silicification. Nevertheless, this process has been held responsible for large porosity variabilities in Triassic sandstone (Durand & Meyer 1982) and sometimes also in oil reservoirs. Contem porary materials resembling silcretes, however, have not been identified. Although there is evidence of cementation by silica in contemporary soil environ ments, for example the duripans of North America (Flach et al. 1969; Chadwick et al. 1989), tepetates in Mexico (Dubroeucq & Thiry 1994), dorbanks in southern Africa (Ellis & Schloms 1982) and red brown hardpans in Australia (Brewer et al. 1972; Chartres 1985), these materials themselves usually occur on remnants of former landscapes. There fore, by implication, silcrete duricrusts are probably formed over time spans that are considerably greater than the life span of an individual soil and require very stable geomorphological and geological condi tions to achieve silicification. Nevertheless, con temporary silica deposits are known in evaporitic environments, where they form chert-like beds (Peterson & von der Borch 1965; Eugster 1969; Maglione & Servant 1973). Previous studies of siliceous duricrusts have con sidered aspects such as field distribution and petrol ogy, geomorphological patterns and relationship to palaeolandscapes and mineralogical-geochemical characteristics. Silcretes take a wide variety of forms; detailed investigations of these forms from the field scale to that available by scanning microscopes can provide clues to their origin. A combination of studies of the complex textures and fabrics of the natural materials and data from experimental silica systems has greatly facilitated interpretations of the processes involved in the deposition of silica at low temperatures and assessments of the environments in which such processes were active. This chapter
Diversity ofcontinental silicifications
focuses on the most recent concepts relating to the formation of silcretes, and is essentially a review of work in progress. Relationships between processes involved in the origin, transport and deposition of silica are discussed, significant environments are suggested and lines of further research are identified.
M A I N T Y P E S OF C O NT I N E NTA L S I L I C I FI C ATI O N
Most supergene silicification features described, both in ancient sedimentary series and in modern land scapes, can be grouped into three main types. 1 Pedogenic silicification. It occurs with the develop ment of specific structures and fabrics, characteristic of soils, that more or less completely obliterate the structures of the parent material. It develops in sand and clay materials. 2 Groundwater silicification. It usually maintains the primary structures of the host material and always relates to an absolute silica accumulation, either by silica deposits in pores, or by matrix replacement. It develops in clastic rocks, in limestones, or evaporites. 3 Silicification linked to evaporites. It is often difficult to be specific about the time when it devel oped, i.e. during early or late diagenesis. This type of silicification generally is defined by the silica mineralogy. Pedogenic silicification
Pedogenic silicification differs from the other types of silicification (groundwater silicification, burial dia genesis, etc.) mainly by its vertical arrangement, hori zons differentiated between the base and the top of the sections, and the presence of typical structures such as illuviation and capping features (Thiry 1 981 ; Milnes & Thiry 1992). These structures are related mainly to water seepage through soil horizons, with reworking and flushing of surface materials down into the subsoil, forming illuviation structures along the water pathway. The profile generally is character ized by a vertical columnar structure defined by verti cal and horizontal joints that obliterate the primary sedimentary structure of the materials. Pedogenic silicification also is characterized by its mineral com position. Often, the silicified layers show a specific distribution of secondary silica polymorphs within the profile, with amorphous and poorly crystallized phases of silica prevailing at its base, whereas micro crystalline and euhedral quartz develop higher up
89
and poorly crystallized phases even dissolve at the top. Two types of material can be distinguished, on one hand there are silcretes formed mainly of micro crystalline quartz, without clay minerals and iron oxides, but enriched in titania, and on the other hand there are soils and palaeosols called 'hardpans' or 'd1,1ripans', formed mainly of opal, that impreg nate the primary clayey and iron-oxide-containing materials. Groundwater silicification
In contrast to pedogenic silicification, groundwater silcretes develop at depths between 5 and 50-lOOm. Their main characteristics are superposed silicified lenses and preserved host-rock structures (strati fication, bioturbations, etc.). The silicification may take the form either of even silicified pans, extending over hundreds of metres, or irregularly shaped metre to centimetre-sized lenses or nodules. There are two main varieties: those that occur in sand and develop into massive sandstone facies by cementation of the pores and those that are created by replacement of limestone and claystone and therefore have more irregular and discontinuous shapes. Silicification associated with evaporites
Silicification associated with evaporites is complex and the time of silicification can be difficult to deter mine. Confusion has arisen because distinctive quartz forms such as quartzine (length-slow chalcedony), lutecite and 'flamboyant' or 'cubic' quartz often have been regarded as indicative of evaporitic environ ments (Munier-Chalmas 1890; Cayeux 1929; Folk & Pittman 1971; Arbey 1980). In modern evaporitic environments silicification occurs primarily through changes in the pH and salt concentration. These changes are both spatial and temporal, involving mixing of fresh continental water with brines or a succession of periods of desiccation and subsequent dilution by rain water. The mixing, either by surficial or ground water, occurs on borders of the evaporitic depressions. Irregularly shaped, chert-like nodules are the most common form of such silica deposits (Peterson & von der Borch 1965; Eugster 1969; Maglione & Servant 1973). Many of the evaporite silicifications that have been described, however, are formed mainly by replace ment, i.e. by diagenetic replacement following depo sition. Although some mineralogical forms of silica
90
M. Thiry
signal sulphate-rich environments, it does not neces sarily mean that they were formed in saline envi ronments. As an example, gypsum replacement by groundwater silicification (after deposition) occurs, at least at the replacement front, in the presence of high sulphate concentrations, but in spite of this silicification does not develop in saline environments. Interpretations of these replacements and of their palaeoenvironments therefore must be made with care. This kind of silicification will not be examined specifically in this paper.
G E O L O GY AND G E O M O RP H O L O G Y O F T H E PARIS BASIN
The Paris Basin i s an intracratonic sedimentary basin that has been subjected to successive transgressions and phases of deposition during the Mesozoic and Cenozoic (Fig. 1 ) . Its present-day morphology, com posed of superimposed limestone plateaux, results mainly from major uplift during the Pliocene and Quaternary, which caused downcutting and erosion of the Tertiary formations. Mesozoic deposits are formed mainly of three facies: hard chert-bearing Jurassic limestone; Lower
Cretaceous clay and sand; and flint-bearing Upper Cretaceous Chalk. During the Senonian-Maas trichian (Upper Cretaceous), the sea withdrew and left behind a platform of flint-bearing chalk and clayey sediments. Between the withdrawal of the Cretaceous sea and the early Eocene, at about 30 Ma, the limestone platform was weathered and thick flint bearing palaeoweathering profiles developed. During the Tertiary, the Paris Basin formed a wide shallow bay bordered by a fiat hinterland, where the boundary between the sea and the continent fluctu ated. In these conditions, small variations in sea-level led to broad transgressions and regressions. Conti nental and marine deposits alternated during the Ter tiary subera (Fig. 2). The most remarkable feature of the Tertiary sequence is the interbedding of thick sandstone formations with limestones and marls. These sandstones form the main aquifers of the Paris B asin. Until the Pliocene epoch, the centre of the Paris B asin was only slightly higher than the penecontem poraneous sea-level, and sloped gently towards the north-west. The main features of the present Paris Basin were established during the early Quaternary, following the cutting of the valleys and the large scale erosion of sand and marl above the main
THE CHAN N E L
Schematic geological map of the Paris Basin. On the western and southern borders the flint -bearing palaeoweathering formations have developed since the withdrawal of the Cretaceous sea. Tertiary deposits piled up in the centre of the basin and formed superimposed plateaux. The eastern border of the basin rose up and was eroded and dissected during the Pliocene and Quaternary Periods. Fig. l.
LJ
I2:QJ
basement
D
Trias
5]
Jurassic
flint bearing paleoweathering formations
� C retaceous D IS;;) m ain reg ional scarp
Tertiary 100 km
91
Diversity of continental silicifications
limestone formations. The result was a series of superimposed structural plateaux sustained by the major limestone formations, which are, from south to north: Calcaire de Beauce, Calcaire de Brie, Calcaire de St-Ouen and Calcaire grossier. Silicified facies occur in almost all the Tertiary for mations of the Paris Basin. Hardened red sandstone or hardpan developed on the Tertiary deposits blan keting the Massif Central basement. Silcrete armour caps the lower Tertiary formations and the palaeo weathering profiles of the Chalk in the southern and western parts of the basin. Tightly cemented quartzite lenses occur in all the sandy formations and silicified limestones are present in every lacustrine and brack ish limestone formation. Even the weathering mate rials topping the limestone plateaux contain large silicified features.
R E D H A R D PA N O F T H E S I D E R O LI T H I C FAC I E S
On the edge of the Paris Basin, in the northern French Massif Central, Tertiary deposits on the basement are scarce and occur mostly in areas of subdued relief and in grabens. These are mainly sandy and clayey, fluvio-lacustrine deposits of variegatedcolour. Car bonates are scarce, they are related to calcrete rather than to lacustrine limestone and are often silicified. There are no fossils, and based on lithostratigraphic correlations, the formations are generally assumed to be from the Eocene and/or Oligocene epochs. Hardened red sandstone surrounds the basement hills and passes laterally into Tertiary sediments. This sandstone belongs to palaeoweathering profiles of the siderolithic facies (Boulanger 1844; de Launay
North
South
PLIOCEN E MIOCENE
OLIGOCENE Argiles vertes
\=--::!Q
I �� ��· � � =�: ; :;;:p � •
Calcaire de Cha
.
'\
.p 1 ""1:& y ;,&
Calca�re de St-Ouen Calcaire Grassier Argiles Plastiques
=:::::::c
� I 't 'il .a.t �� &0 o 0� o o o o �
:::C:::: "
�.:..:YY�Y�YY�Y�YY.:.:. \
� T •I �P I ,� o o o o o o -=��������7r=�:a� � �0
a
'i • • •
� sandstone � limestone
Argile a silex
Cretaceous Chalk
CJ claystone/marl � gypsum
0/¥Y'f,Y¥Y'f,Y¥Y: �
I
- 50m
EOCENE
de Beauchamps
0
__ __
•
0
0
ble de Cuise
0
PALAEOCENE
Sable de Bracheux
� groundwater quartzites IIYYY\I pedogenic silcretes
IT •I discordant cherts
Fig. 2. Continental and marine deposits, interspersed with long periods of non-deposition and weathering, alternated during the Tertiary in the Paris Basin. The most noteworthy feature of the Tertiary sequence consists of successive thick sandstone formations that form extensive aquifers and which are interbedded within limestones and marls. Note the wide distribution of the silicified facies within deposits from all periods throughout the basin, in the clastic as well as in the limestone formations.
92
M. Thiry
1892-3; Deschamps 1973). These 'siderolithic' facies display features of buried palaeolandscapes (Thiry & Turland 1985). Palaeomorphology
The hardened red sandstone lies mainly at the edge of the Tertiary deposits and frequently masks the contact zones between the sedimentary deposits and the basement (Turland et al. 1989). They show four main structural types (Fig. 3). 1 The contact between basement and Tertiary deposits is sinuous: the indurated red sandstone rests on the basement and passes into the sedimentary deposits, and in places is interbedded within these deposits. 2 In places, wide stretches of the basement are still covered with saprolites, on which the red sandstone is preserved locally and forms small buttes a few metres high, which resemble armoured mesas. 3 In other areas the basement forms gently undulat ing plateaux and the red sandstone is regularly located in palaeovalleys through which the present day valleys are cut. The slopes are cut either into the red sandstone (in places over 10 m thick), or deeper into the basement; the red sandstone is then pre served like fluvial terraces. Remnants of saprolites remain on the adjacent plateaux.
4
Coarse red sandstone, contammg numerous quartz clasts (up to 1 0 cm in diameter) and even fresh basement fragments, closely surrounds steep palaeoreliefs of the basement and fault scarps, forming cones of fallen rock and slope breccia. In places, such coarse red sandstone can reach a thick ness of 30m. These relationships clearly show that the red sand stone is evidence of differentiated palaeotopogra phies, widely eroded and reused by the present-day hydrographic network. Detailed mapping of the red sandstone facies shows that they cap palaeoglacis (gently sloping piedmonts) cut into the Tertiary deposits as well as into the basement (Fig. 4), around the basement palaeorelief. Relationships between the red sandstone, the base ment and the Tertiary sediments can be defined by detailed studies of the sections around a granitic palaeohill at Perguines (north of Montlw;:on) (Thiry & Turland 1985). The macro- and micromorphologi cal structures and the mineralogical composition of these materials allow us to understand the mechan isms of their development and their geochemical and environmental significance. The facies
In the Montlw;:on-Perguines area, granite forms a small hill rising 30 m above the Tertiary deposits of the basin. Patches of red sandstone lie directly on the granite hill, but especially around its perimeter, where weathering overprints occur on the sedimen tary deposits, forming variegated sandstones. Jasper and chert boulders and layers, as well as limestone facies, stand against the greenish claystone of the basin. The weathered granite
On the hill, the granite is weathered and iron-stained down to a depth of at least 2-3 m. In places it has been weathered into a soft, clayey saprolite, whereas in others it appears hardened by silicification. In places at the edge of the hill, the granite fabric collapsed and turned progressively into a coarse red sandstone. D Tertiay deposits � fresh basement
�
rz::::LI
red sandstone
weathered granite
Hardened red sandstones related to palaeoweathering profiles of the siderolithic facies resting on various features of buried palaeolandscape at the edge of the Tertiary deposits.
Fig. 3.
The red columnar sandstone
The red sandstone duricrusts, 2-4 m thick, surround ing the granite hill are the most typical and the thickest palaeoweathering facies. The lower part of the duricrust is usually massive, whereas it becomes
Diversity ofcontinental silicifications
nodular toward the top with a pseudo-columnar structure (Figs 5 & 6a). The columns are formed of a red, hardened sandy clay, with angular quartz grains scattered in a red argillaceous matrix that has small irregularly shaped voids. The planar and vertical joints between the columns are filled with loosely cemented ochre sandstone composed of millimetre sized red pseudo-ooliths and display burrows, nodules and illuviation structures. The burrows average 1 . 5 cm in diameter and can be 3 0 cm long. They are mostly vertically elongated and have a wall 1-1 .5 mm thick 'constructed' of fine matrix material. The sandstone contains kaolinite. Illuviation struc tures, burrows, irregularly-shaped voids, small pipes (0.5-1 cm in diameter) and horizontal joints con tain white silica deposits, mostly in the middle of the profile. Burrows and nodules capped with fine illuvia tion laminae clearly show that the duricrust is related to a palaeoweathering profile. Beige to brown opal cutans with weak refraction and anisotropy occur in all the open pores of the sandstone (Fig. 6b ). The cutans are transected by thin cracks regarded as shrinkage cracks indicative of des iccation or gel-ageing phenomena. Mineralogical and microprobe analyses show that the cutans are com posed of a mixture of 50% opal and 50% kaolinite. Granular infilling in the joints is made up entirely of 'microgravels' or pseudo-ooliths of argillaceous and siliceous materials arranged in concentric layers
Fig. 4. Schematic geological map showing the layout of the Tertiary deposits and of the red hardened sandstone around the hill of Ste Agathe (15 km north-west of Montlw;on). Red sandstone armours a palaeoglacis (gently sloping piedmont) cut in the Tertiary deposits as well as in the basement.
93
around skeleton grains. They are cemented by brown opal similar to that of the illuviation cutans. This pseudo-oolithic structure is similar to the pellets observed in termite mounds and may result from biologic activity (Wielemaker 1984; Eschenbrenner 1986). The red duricrust is a kaolinitic ferruginous palaeosol with typical pedogenic features im pregnated with silica. The illuviation features and vertical joints indicate water percolation in a soil environment. The variegated planar sandstone
Downhill, towards the basin, the variegated sand stone is characterized by branching subhorizontal fractures that have produced a distinctive but irregu lar layering (Fig. 5). At the base of the section, the sandstone layers average 20 em in thickness. The hori zontal fractures show deposits (from 0.5 to 2 cm thick) of thin silica laminae that extend into the verti cal j oints intersecting the sandstone layers. The upper part of these profiles is formed of red-ochre sand stone nodules wrapped in undulating illuviation laminae. The nodules (from 0.5 to 4 cm in diameter) are composed of hardened ferruginous sandy clay with angular quartz capped with alternating millimetre-thick laminae of sandstone and fine grained silica products (Fig. 6c ). The caps can increase
D basement D weathering mantle D tertiary D red sandstone � escarpment � glacis � incised modern valley • 27 0 m a.s.l.
M. Thiry
94 DOWNSTREAM
variegated planar sandstone
red colu m nar sandstone
UPSTREAM
m col u m nar and nod u lar re sandstone
variegated planar sandstone with sil ica laminae
00 00 0 0 0 0 0 0
0 0
hardened red sandstone with joints fil led with loosely cemented pseudo-ooliths
0 0 0
0
beige sandstone
LITHOLOGY rn D sandstone � silica laminae �
bu rrow pseudo-ooliths
0 0 0 0 0 0
0 0 0 0 00 0 MINERALOGY tllll feldspar D q uartz � gffi goethite-hematite E/2
red mottled sandstone opal-cT & opal-A IIIJill illite smectite Jo 0oJ kaol i n ite
Fig. 5. Schematic sketch and mineralogical composition of the hardened red sandstone around the palaeo hill near Perguines. Upstream, against the granitic palaeo hill, red sandstone is characterized by a pseudo-columnar structure and kaolinite is almost the only clay mineral present. Downstream, interbedded in the Tertiary deposits, the variegated sandstone has a planar structure, opal becomes a significant compound, and smectite the main clay mineral.
in size and build up undulating illuviation laminae, creating a progressive transition to underlying thick silica laminae. The earlier silica deposits consist of illuviation structures of opal mixed with about 30% clay miner als. Later silica deposits are represented by clear isopachous opal concretions that develop preferen tially at the base of planar joints; pendant features are more scarce and thinner Fig. 6(d & f) . The silica laminae deposits show a succession of mammillary opal concretions and granular material illuviation structures (quartz grains, clay chips, silicified debris, microgravels) with erosive contact (Fig. 6e) . This con cretionary opal contains less alumina but has a rela tively higher potassium content than the early opal illuviations. The nodules of the upper horizon are formed either of ferruginous sandstone or micro gravels cemented by brown opal. The caps contain cross-bedded laminae of fine quartz and microgravels embedded in a matrix of brown opal. The succession of concretions and illuviations indi-
cates that water percolation occurred at the time of silicification and suggests that periods of percolation and saturation alternated. Here, planar structures are the main features of the profile, suggesting a prevail ing lateral water flow over vertical percolation in a downstream situation, close to the basin base level. Jasper and chert layers
In the basin, jasper and chert lenses with complex concretionary structures are associated with carbon ate beds. Downstream from the red sandstone, the first silica nodules to appear are red and contain red sandstone nodules, detrital quartz grains and silicified plant debris. The most typical facies are white or beige with undulating laminated structure. Some are more translucent with concretionary facies. Finally, some facies are white and porous with bluish translu cent opal nodules. The two main types of silica shown in thin-sections are nodules of pure opal and thick opal and chal-
95
Diversity of continental silicifications
(a)
(d)
(b)
(c)
(e)
(f)
Hardened red sandstone. (a) Columnar structure of a red hardpan profile (near Montlw;on). It is a ferruginous palaeosol similar to laterite and cemented by silica. (b) Illuviation cutans composed of a mixture of opal and kaolinite transected by shrinkage cracks in the columnar sandstone. Crossed polars: v, void; scale bar is 100 �m. (c) Hardened sandstone nodules capped with fine-grained illuviation laminae in the upper part of the variegated planar sandstone. Scale bar is 1 em. (d) Concretionary opal deposits developing at the base of the planar joints in the lower part of the variegated planar sandstone. Crossed polars at 80°. Scale bar is 3 mm. (e) Details of the opal concretions showing the intercalation of granular illuviations (arrow). Scale bar is 1 mm. (f) Sequence of concretionary opal deposits thinning out upwards. Scale bar is 1 mm. Fig. 6.
cedony concretions that cement together the opal nodules and the other compounds. The opal nodules range from 0.1 to 20 mm in diameter. They frequently have a network of radial and concentric cracks resembling shrinkage cracks indicative of desicca tion or gel-ageing phenomena. The concretionary deposits are thinly laminated and show successive
deposits of translucent opal and several varieties of chalcedony. The residual voids may be sealed by chal cedony sheaves. Chert lenses from the basin contain the purest varieties of opal, devoid of alumina but with significant potassium content. Small barite crys tals develop in some nodules. In the most translucent nodules and microlaminae, amoebic zones of micro-
96
M. Thiry
crystalline quartz and chalcedony appear. The rela tionships between opal and quartzose zones clearly indicate that the quartz formed by recrystallization of opal. The lack of illuviation cutans in these cherts and the concretions around the voids argue in favour of evolution in a water-saturated environment (lake or ground water). Significance and relationships of the structures
Two distinct types of evolution are superimposed in the sandstone studied: 1 'ferralitic' palaeoweathering that led to the devel opment of the red kaolinitic sandstone and is related to the major siderolithic weathering event; 2 silica hardening that impregnated the red sand stone and favoured its preservation. These successive weathering features show lateral variations according to their positions relative to the basin and are related to a palaeocatena; their main characteristics are summarized in Fig. 7. The ferralitic weathering profiles
The weathered basement rocks are found on the high points of the palaeolandscape. Weathering profiles are restricted to the lower saprolite horizon as a result of stripping by erosion. The red columnar sand stone forms the upper part of the catena, the vertical and eluvial structures suggest that downwards water percolation reworked the fine compounds of the soil. The materials are leached, kaolinite is the main clay mineral and the Si/Al ratio of the clayey matrix is around 1 :0.7. planar sandstones
structure
nodular
1-1-morphology
concretions
laminar IJ·Iaminar deposits
The variegated planar sandstone occurs down stream of the hill. Vertical percolation is restricted to the upper nodular horizons; at depth, the planar and laminated structures show signs of lateral water flow. The development of lateral flow and thick clayey illu viations probably is the result of downstream outlet plugging by clay deposits in the basin. Invasion of the upper nodular horizon probably resulted from a rising illuviated horizon caused either by pro gressive clogging of the profile or by a rising lake level. The materials being less leached, the better preservation of feldspars and micas, the presence of smectites as a main clay mineral in addition to kaolinite, and the higher Si/Al ratio of the clayey matrix (near 2) are signs of more confined environments. Micromorphological and geochemical characteris tics of these palaeosols make them comparable to soil catena in the tropical regions of West Africa, such as those described by Bocquier (1 973) and Boulet (1974). The palaeosols probably relate to palaeocli mates with 1 000-1200mm of rainfall during a relative short wet season (4-5 months) and a dry season with definite hydric deficit. Such palaeoclimatic conditions are also well known to have existed elsewhere in the Paris B asin during the early and middle Eocene (Thiry 1981 ). The palaeolandscape of this siderolithic catena was formed of basement hills rising out of a clayey, periodically flooded, plain. Silica hardening Nature and structure. Silica hardening followed the development of the ferralitic palaeoweathering profiles, but was superimposed on the ferralitic soil
columnar sandstones
columnar
preserved
i l l uviations
i l l uviations
cutans and
opal
75% opal
50% opal
20% Fe oxides
concretions
fibr. quartz
25% clay min.
50% clay min.
80% clay min.
Fig. 7. Pedogenic silicifications of the 'hardpan' type. Silicifications occur in catenas displaying distinctive structures that are a function of position in the palaeolandscape: siliceous illuviations, due to percolation, dominate upstream, whereas silica concretions due to lateral flow develop downstream in wide planar structures.
Diversity of continental silicifications
97
catena that remained functional. The upstream fer ralitic soils were progressively clogged by silica-rich illuviation deposits, which are composed of a mixture of opal and clay mip�rals. Shrinkage cracks suggest " -ageing phenomena. Down desiccation and/or g'M stream, the first s ilica-rich deposits were als o formed by illuviation of a mixture of opal and clay minerals enriched in silica. Later, there were concretionary deposits of purer opal. In the lacustrine realm, chert developed, which are composed of pure opal, without alumina, and of several varieties of crystallized silica phas es. There is a clear differentiation in the silicification of the palaeolandscape: upstream, silica deposits are mixed with clay minerals, whereas downstream, silica phases are purer and were precipitated in water logged environments.
structures may be compared to s ilica deposits in pres ent-day evaporitic environments (Peters on & von der Borch 1965; Eugster 1969; Surdam et al. 1972; Maglione & Servant 1973; Sheppard & Gude 1974). The palaeoclimate at the time of silicification is thought to have been relatively dry, with drought periods leading to evaporitic environments and periods of stronger rainfall causing a ris e in the lake level and flooding in the alluvial plains. Thes e palaeo climatic conditions correspond to the transition of the seas onally differentiated tropical palaeoclimate of the early Eocene to the warm and arid palaeo climate of the late Eocene in the Paris Basin (Chateauneuf 1980).
Geochemistry and environment. There is no alu minium and iron leaching during the silicification process. Silicification appears mainly as an absolute accumulation of silica in what must have been oxic and near-neutral pH environments. The fairly high potassium contents of silica phases and the pres ence of sulphates (gypsum moulds and barite crystals), particularly in the downstream domain, indicate a relatively confined, evaporitic environment. In the palaeos ols, silicification features correspond to periods of rising bas e levels, or flooding of the lake, alternating with periods of drought. It may appear contradictory that lake flooding, probably during periods of heavier rainfall, led to silica deposits. This can be compared to s imilar situations at the northern border of Lake Chad, where an abundance of diatomites corresponds to high water levels in the lake (Servant & Servant 1 970). During periods of low water there seems to be less silica available, although silica concentrations are higher. The s ilica may then have combined with other elements to form clay min erals, as noticed in other evaporitic environments (Gueddari et al. 1982). In the basin, the presence of nodular chert may have a different significance. With increased s alt concentration, the solubility of amorphous silica decreases cons istently (Marshall & Warakomski 1980; Marshall 1980). It decreas es by as much as three to ten times in a s olution saturated with salt. This may explain the behaviour of silica in the bas in. After dry periods, silica-loaded freshwater flooding into the basin mixes with brine and becomes strongly overs at urated, and the resultant mixture is therefore able to precipitate amorphous silica. Thes e silicification
The red and ochre mudstone relates to the reworking of kaolinite-rich palaeosols formed in the warm and humid climates that prevailed during the Eocene. They reflect lands cape instability and tectonic rejuvenation together with a change to a drier climate. The hardening of the red duricrust corre sponds to a period of calm and landscape stabiliza tion. Well-structured palaeosols with biological activity date from that time. The silica accumulation found in these palaeosols, especially in the transi tional zone between uplands and lowlands, marks a major climatic change from the previous more humid climates. The origin of the silica is not well established. There is no sign of dissolution of the silica phases (neither quartz nor clay minerals) and therefore the s ilica appears to have been added to the profiles. This could have happened either by leaching of the high points in the landscape or by s iliceous oozes depos ited at the top of the profiles during periods of flooding or high water levels in lakes. Whatever its s ource, silica must have been leached and the s olu tion concentrated to cause the deposition of silica gel and opal. Alternating wet and dry seasons or periods seems the most likely cause. The geochemistry of the system appears relatively simple. There is no dissolu tion of quartz, nor of iron oxides, not even of clay minerals. The environment was restricted and had a neutral pH. The formation of silica gels was probably favoured by the presence of brine, which lowered s ilica s olubility. . The red duricrusts from the northern Massif Central resemble the hardened s oils called red brown hardpans in Australia, which are composed
Palaeolandscape interpretation
M. Thiry
98
of a mixture of clay minerals, iron oxides and opal (Bettenay & Churchward 1974; Chartres 1985; Milnes etal.1991). Carbon ate and gypsum commonly are present in the hardpan profiles and barite has been observed occasionally (Wright et al. 1993). Duripans in the arid areas of the USA (Flach et al. 1974; Chadwick etal.1989) and dorbanks in southern Africa (Ellis & Schloms 1982) have similar character istics. Hardpans lie immediately beneath contempo rary s oils and are closely related to contemporary landscapes and some appear to be still developing. Bettenay & Churchward (1974) think that they develop in periodically flooded s ubdued drainage areas, with 250-300 mm annual rainfall. The Ameri can formations als o are related to alluvial and collu vial deposits in regions where the precipitation did not exceed 300mm during the Pleistocene (Harden etal.1991).
QUARTZOSE SILCRETE ARMOUR
Silicified pans crown kaolinite-rich clastic s ediments and cover a wide palaeosurface extending over the whole s outhern part of the Paris Basin. They consist of very hard quartzite in blocks of variable s ize and massive, columnar, bulbous an d mammillated shapes, coated with a thin s ilica layer that hides the internal structures. Geology
Tertiary deposits are scattered over the s outhern border of the basin and form a discontinuous blanket of palaeoweathering products recording the history of the long period of continental evolution following the withdrawal of the Cretaceous Sea. Fluviatile and lacus trine deposits related to Eocene times are confined mainly to several river valleys and grabens, among them the present-day Loire rift valley (Fig. 8). Toward the cen tre of the basin (Brie region), the clastic sediments deposited in wide alluvial plains and lakes during the early Eocene (Argiles Plas tiques Formation) were overlain by several limestone formations during the middle and late Eocene. After the withdrawal of the s ea, the Cretaceous deposits un derwent intense leaching: they were totally decalcified, flints partly dissolved and kaolin ite developed. The flint-bearing palaeoweathering profiles are formed of white, brown and red mottled clays, s andy in places, which con tain irregularly dis tributed flints. Flints are 5-20 cm in diameter and
� � � li!!lli!!J s::::::B
Upper Cretaceous Lower Cretaceous
CJ C=:J
Jurassic
�
Triassic
-
basement
Mio-Piiocene Oligocene Eocene fault
50 km
Fig. 8. Geological map of the southern Paris Basin. Eocene fluviatile flint-bearing deposits are mainly confined to grabens trending N-S along the Loire and Loing rivers. To the North, beyond the River Seine, the clay deposits of the Argiles Plastiques Formation develop in the Provins area. These Eocene clastic deposits are crowned frequently by pedogenic silcrete armour.
deeply corroded. These weathering formations s ubsist locally on the plateaux. In the south, there are s everal facies relating to fluvio-lacustrine deposits in the N-S-trending Loire
99
Diversity ofcontinental silicifications
rift valley. The coarse depos its at the base and the border of the grabens progressively evolve into s ilt and clay sediments and finally into palustrine deposits. Upstream, in the Loire gr aben, corroded and br oken flints have accumulated locally. The clay matrix between the flints shows numerous illuviation structures related to palaeosol features. They are s igns of a subaerial depos itional environment and of progressive accumulation in the gr aben of this thick (up to 50 m) flint breccia. Downstream, towards the bas in, the fluviatile deposits cover sever al areas that corr espond to palaeovalleys converging on the bas in. In the Loing valley, the deposits are restricted to a narrow strip and can exceed 40 m in thickness above the erosional gullies cut in the Cretaceous Chalk. The deposits ar e composed mainly of flint conglomerates and clayey and s andy lenses. The flint pebbles aver age 5-30cm in diameter and are well rounded within a s andy matrix. In the s outhern parts of the subsiding Paris Bas in, in the Brie region, the bas al Tertiary deposits, the Argiles Plastiques, are formed of a thick clay unit that varies between 5 and 1 0 m in thickness. Above a bas al colluvium of reworked weathered chalk products, the clay unit is composed of pure kaolinite without quartz and cr owned by a clayey s and that fills er o sional troughs. The quartzose silcrete duricrusts
Many Eocene clastic deposits are silicified. The early age of these silicification structures is shown by clasts having been reworked in the late Eocene limestone and Oligocene s andstone cover deposits. These silicifications have been known s ince the beginning of this century and their nodular structure and titania enrichment noticed (Janet 1903; Cayeux 1929; Vatan 1935; Jodot 1947). All the Eocene clastic deposits, with or without flints, fr om the s outhern bor der of the Paris Basin are affected by the silicification. Expo sures where the silicification can be seen in situ and over the whole profile are scarce. In the s outh there are frequently only residual boulders enclosed in the clayey flint-bearing formation of the plateaux; towards the subsiding basin the sections are more complete and preserved within the Eocene str ati graphical column. The silicifiedflint breccia of the Loire graben
The flint breccia is extens ively silicified in the Loire gr aben near Sancerre (Fig. 9), wher e the silicified unit
boulder horizon
m
% Ti02 .1
1
10
1
nodular structures dissolution features and large capping structures
silica illuviations
l'r'Oo-�a""-'-
14
�����
10
in the joints and above the blocks
clay with flints and. residual silicifications
fractured pan with silica rims on the joints
clay with flint breccia ----�0 ���--+
Fig. 9. Schematic diagram of the silicified flint breccia profile at Tracy near Sancerre. At the base of the section the silicification is massive, higher in the section the silcrete displays wide columnar structure, and the upper part of the profile contains boulders of complex nodular structure. Titania concentrations are clearly associated with the illuviation structures and locally with iron-rich cutans, and increase toward the top of the section.
reaches a thickness of 1 5 m (Thiry & Simon-Coin9on 1996). At the base of the section, the s ilicification is glassy and massive and rests on a loos e flint with a clay matrix. Flint clasts are embedded in a yellowish-grey matrix and are glassy when fractured. Horizontal and vertical joints intersect the har d silicified flint breccia, defining metric angular slabs. Most of the joints are unfilled and only a few vertical joints have a thin silica coating on the walls. Higher up in the section, the silcrete has a wide columnar structure (Fig. lOa). The silicified parts consist of flint clasts embedded in a yellowish-grey matrix (Fig. lOb & c). Finely laminated, yellowish, titania-rich s ilica illuvial deposits coat all the vertical and some of the horizontal joints. These illuviation
100
M. Thiry
(a)
(c)
(b) Fig. lO. Silicified flint breccia. (a) Columnar structure of a quartzose flint breccia profile (Tracy near Sancerre ). (b & c) Angular flint clasts embedded in a yellowish-grey quartzose matrix. Scale bars are 1 and 5 em.
deposits increase along the main vertical joints to form a centimetre-thick coating that includes small silcrete nodules. Capping the upper silcrete bodies, as well as the smaller nodules in the main joints, are alternating deposits of fine silica laminae and coarser layers of granules and small silcrete nodules. The upper part of the profile contains rounded sil crete boulders from 0.5 to 2m in diameter. The struc ture of the s ilcrete is more complex; fragmentation and nodulation of the material are very advanced and nodules are complex and polygenetic. Large parts of this facies consist of wide and thick illuviation struc tures. The matrix is, in places, variably coloured by iron oxides. Irregular voids are abundant and well developed around the nodules and the flint clasts, indicating loss of material by dissolution and eluviation. Thes e silicified breccias differ between the top and the base. Thin-section micrography shows the magnitude of the illuviation deposits during the silicification process. Moreover, the profile structure is accompanied by a chemical polarity: silica matrix and illuviation structures increase in titania content
towards the top of the profile (Fig. 9). These obviously are pedogenic silicifications that developed near the soil surface. The silicified flint conglomerate ofthe Laing channel
In the Loing valley, the silicified flint conglomerate crops out along a strip 4-Skm wide and 40 km long. It is about 5-15 m thick, and typically forms rock towers that are often hollow at the base and overlain by late Eocene limestone. These wide columns, up to 10m thick (Fig. ll), are silicified at the tops and outer walls, whereas the core remains unconsolidated. Several types of horizons are visible (Thiry 1981). At the very foot of the s ection, the conglomer ate has not been s ilicified: the flint pebbles are embedded in argillaceous s and. At the first stage of silicification, pebbles are topped by a glassy sandstone capping and the sand matrix remains unconsolidated. The geopetal pebble caps consist of alternating millimetre-thick laminae of sandstone and fine-grained s ilica with cross- and graded bedding (Fig. 12).
Diversity of continental silicifications Calcaire de Ch8teau-Landon calcrete invaded puddingstone
m 9
quartzitic puddingstone with dissolution features and overturned pebbles
lustreous puddingstone with silicified caps and matrix
unconsolidated conglomerate with silicified caps on the flint pebbles
flint conglomerate without silicification
Chalk
Schematic diagram of the silicified flint conglomerate of the Loing channel near Nemours. At the base of the profile the silicification is restricted to capping structures above the flint pebbles, higher in the section the silicification is pervasive, invading the sandy matrix between the flint pebbles. Lastly, at the upper part of the profile, numerous dissolution voids develop and flint pebbles are sometimes released and overturned.
Fig.ll.
3
101
Higher up in the section, the flint pebbles also have illuvial caps, but the matrix progressively turns into beige-ochre glassy quartzite. The matrix is either coarse-grained and homogeneous or fine grained and nodular. Fractures are glassy and cut evenly across the matrix and the flint pebbles. In the upper part of the silicified formation, the cement no longer shows a glassy appearance but has a dull granular fracture surface and a greyish colour. The matrix is quartzitic, with visible quartz grains. Voids are well developed around flint pebbles. Flint pebbles and their illuvial caps are etched and cut by the quartzitic matrix. Overturned flint pebbles (their caps are no longer in an apical position ) appear in the conglomeratic rock. The micromorphological organization of the silicified conglomerate is typical of pedogenic sil cretes, with silica illuviations and a matrix of inter locking fine-grained and microcrystalline quartz and brown titania pigment and/or larger leucoxene grains. The macro- and micromorphological charac teristics indicate that illuviation features developed at the base of the profile and dissolutions, eluvial structures and titania accumulation at the upper part of the profile. A concomitant mineralogical sequence of secondary silica phases also developed: opal and microcrystalline quartz at the base and larger crystals of clear quartz, whereas limpid overgrowths and euhedral quartz formed in voids at the top. The silicified clay deposits ofthe subsiding basin
On the border of the subsiding basin, the sandy clay of the Argiles Plastiques Formation frequently is crowned by a quartzitic pan with a columnar habit
(a)
(b) Fig. 12. Silicified flint conglomerate (near Nemours). Caps above the flint pebbles consist of millimetre-thick laminae of sandstone and fine-grained silica with cross- and graded bedding. (a) Polished slab. Scale bar is 1 em. (b) Thin-section. Scale bar is 1 em.
M. Thiry
102 HORIZONS sand pseudo-brecciated polygenic structures
MICROGRAPHY Ill .<:
m
"'
4%
c 0 0:: :I
"'
.<:
�"' al
I
:I CT Ill
I Qj
Cl "'
�
'iii
�
��
I
CJ
I
T
·�Ill
siliceous granules in a sandy clay
3
•�I .� •.g!a :I
columnar
clayey formation
2
·�II! �
nodular structure and thick illuviation caps
granular
TiOz
�
2
prismatic structure and thinly laminated illuviation cutans
ILLUVIATIONS
•::: 0
c. 0
ECJ ·e
"0
I
.
\.._. . ' ' . . . ..
..
I�
.<:
0
.l:! Ill
Schematic diagram of the pedogenic silcrete that develops at the top of the upper sandy clay of the Argiles Plastiques Formation near Provins. The profile is characterized by a vertical organization and large illuviation structures. Distribution of secondary silica shows a highly soluble silica phase (opal) at the base of the profile, more stable phases (microcrystalline and euhedral quartz) higher in the section, and finally dissolution features appear at the top of the profile. Titania concentrations are clearly associated with the illuviation structures and with the opal and altered clay minerals at the base of the profile.
Fig.13.
and large illuviation features (Thiry 1 981). Several horizons can be identified, and profile thickness varies between 0.5 and 2 m (Fig. 13). At the base, a granular horizon is formed of small siliceous granules (2-5 mm in diameter) embedded in yellow s andy clay. Towards the top of the horizon the granules become progressively larger and are welded into centimetre-s ized nodules by brown coloured opal that has replaced the clay matrix. Higher up in the profile, the granular horizon is overlain by a more massive slab with a columnar structure, formed by vertically elongated decimetre sized blocks that are delineated by vertical fractures and horizontal cracks (Fig. 14a). The interior of the columns has a granular arrangement similar to that of the nodules of the lower horizon and are formed of quartz grains with irregular overgrowth apophyses that grade into titania-enriched microcrystalline quartz. The columns themselves are capped by illu-
vial deposits that continue into vertical joints filled with siliceous granules embedded in thinly laminated deposits of opaline silica and titania (Fig. 14b ). The voids are arranged in typical patterns with clear illu viation structures at the base and dissolution struc tures at the top (Fig. 14c). The upper horizon is more uneven and has a pseudo-brecciated internal structure. The fragmen tation and nodulation of the material are very advanced. Nodules are complex and polygenetic, sur rounded by a titania-enriched rim and irregularly shaped voids. The largest elements are capped by very thick (up to 5 em) deposits of interstratified fine grained s ilica and coarse clastic materials (with frag ments reaching around 1 em). Most of these structures appear related to eluvial and illuvial organizations. The grain s ize of the reworked materials decreases from top to bottom of the profiles. In the upper pseudo-brecciated horizon,
1 03
Diversity of continental silicifications
(a)
(c)
(b)
Fig. 14. Silicified clayey deposits (near Provins). (a) Columnar horizon. The columns display a clear prismatic division. Their smooth upper part corresponds to the capping structures. (b) Vertical joint filled with siliceous granules and silica cutans. Scale bar is 1 em. (c) Void in the columnar horizon. Void is filled with granules (1), thin particle-size graded laminae (2) and, towards the top, thinly laminated opal (3). Ceiling of the void is lined by titania-rich stalactites ( 4). Scale bar is 250 Jlm.
centimetre-sized nodules are reworked during the silicification process and the fine-grained material is removed; in the columnar horizon sand- and silt-sized grains are incorporated into the illuviation deposits; finally, in the basal granular horizon, the deposits no longer have discernible grains. This granulometric distribution is typical of the soil profiles and is the result of progressive loss of energy with downward percolation. Mechanisms of silicification and their significance
There are two types of silicification within these profiles: a massive type that maintains the primary structures of the sedimentary materials and a more nodular and intermittent type containing dis tinctive structures. The preservation of the primary structures in the massive type may link it to ground water silicification processes similar to those described in Australia by Thiry & Milnes (1991) and Thiry (1997). Cappings and illuviation structures, as well as the high titanium content of the nodular type, relate the latter facies to pedogenic silcretes. The formation geochemistry and mechanisms of these silicifications remain to be clarified. This would provide valuable information in reconstructing the palaeoenvironmental conditions in which they developed.
The massive silcrete
The massive silcrete facies at the base of the silicified breccia profile are remarkable because they preserve the structures of the primary formations. They devel oped through silicification of the clay-rich matrix of the breccia. The silicified matrix consists of micro crystalline quartz, with minor opal in some places, and always has relatively low titania concentrations (<0.5% Ti02). The geochemical nature of their formation is not yet clear and can only be assumed. The major processes include destruction of clay minerals and immobilization of silica, whereas aluminium and iron, which are always abundant in these types of materials, are leached out of the rocks. It seems likely that reactions of acidic solutions with the origi nal sedimentary clay led to AI and Fe mobilization and produced amorphous silica from the residual clay structures. Such a mechanism has been suggested in order to explain the deep bleached profiles devel oped on the Cretaceous shales from inland Australia (Rayot 1994; Rayot et a/. 1992; Thiry et al. 1995). The Australian weathering conditions, however, appear to have been quite different and there are several unanswered questions, particularly concern ing the development of an acidic environment. The weathering materials subject to silicification are
104
M. Thiry
always highly oxidized and contain neither pyrite nor organic matter, which would provide a source of acidity. Acidity may result from 'ferrolysis' by oxida tion of ferrous iron in hydromorphic horizons or in ground water (van Breemen et al. 1984; McArthur et al. 1991; Grimaldi & Pedro 1996). Another major problem is the impermeability of the silicified clay materials. This impermeability does not permit a flow of water to leach out the aluminium of the formation. Moreover, prjmary structures have not collapsed and there is no tvidence that a significant porosity devel oped during the weathering process and was later cemented. Whereas aluminium and iron are leached out during acidic weathering, the highly soluble amor phous silica released during the alteration of clay minerals seems to be retained. This may result from a decrease in the solubility of the amorphous silica in brine (Marshall & Warakomski 1980; Marshall & Chen 1982). No evidence of evaporitic environments has been observed in the silicified formations, but gypsum is known to have been deposited in the Paris Basin and in the Massif Central during the middle and upper Eocene, at the time when the silcrete profiles developed (Munier-Chalmas 1890; Lorenz & Bavouzet 1981; Gross 1984; Blanc-Valleron & Thiry 1997).
The distribution of the silica phases also reflects the chemistry of the solution from which they pre cipitated (Fig. 13). The development of coarser microcrystalline, amoebic and euhedral quartz in the upper horizons may be due to more dilute solutions than those that produced the minute microcrystalline quartz in the deeper horizons (Thiry & Millot 1987). Palaeoenvironmental conditions
Conspicuous fissure and void fillings in the pedogenic silcretes suggest reworking and flushing of surface soil materials down into the subsoil horizons. Upward grading of grain sizes in these deposits, and the recur rence of interspersed, finely laminated zones of silica and titania, indicate that intermittent but repeated infiltration events gradually clogged the profile. In addition, as the structures are so well preserved, with rare evidence of erosion, it seems probable that induration through deposition of interstitial silica occurred between successive infiltrations, when parts of the profile dried out. Intermittent and repetitive infiltration appears to have alternated with evaporation, which caused the precipitation of silica from solution. At the time of formation of the pedogenic silcrete profiles the climate was probably dominated by alternating wet and dry periods.
Silica distribution in the columnar silcretes
In the columnar silcretes the silica derives from the profile itself. There is abundant evidence of silica dis solution within the profile, especially in the upper horizons. Dissolution of the early massive silicifi cation appears to be the main source of silica for the pedogenic silicification. Titania accumulated when silica was dissolved, and thus the higher titania con centrations in the upper horizon account for succes sive dissolution stages (Fig. 13). Part of the silica derived from dissolution in the upper part of the profile has been widely distributed downwards by a succession of precipitation and leaching events. This leached silica precipitated towards the base of the profile, probably as a result of the progressive concentration of silica and other cations in the downward-moving solutions. The driving force of this concentration would have been evaporation (Thiry & Millot 1987), with a simultane ous lowering of silica solubility as a result of increas ing concentrations of salt (Marshall & Warakomski 1980).
Palaeogeographical distribution
The landscape in which the quartzose pedogenic sil cretes developed can be described. In the southern Paris Basin, silcretes are mainly residual and crop out in too scattered a fashion to allow a reconstruction of the palaeolandscape in which they developed. It may have been similar to those described in Australia, where silcretes appear to have developed in contrast ing landforms and not only on plateaux and pene plains. They armour glacis (gently inclined slopes) on scarp foot pediments (Milnes & Thiry 1992; Simon Coinc;:on et al. 1996), and are thicker in the transition zone between the glacis and the plains. This silcrete distribution shows influence of lateral water flow along the glacis. In the southern part of the Paris Basin, silcretes follow a Tertiary palaeosurface extending from the Massif Central basement to the centre of the basin (Fig. 15). Silica accumulated pref erentially in transitional zones between upland and lowland areas and silcretes developed extensively around the basin. There are scattered remnants from
105
Diversity of continental silicifications
s
MASSIF CENTRAL
PARIS BASIN
N
m
300 200 100
EE:Jjbasement I :::::fFll I Mesozoic sediments CJ Lower Eocene
I£53Upper Eocene �silcrete
Geological section of the southern border of the Paris Basin showing the position of the pedogenic silcretes along an even palaeosurface, the Eocene palaeosurface shaped by the siderolithic clastic discharge during early Eocene times. On the edge of the subsiding basin the silcrete armour is interbedded within the Eocene series, which allows the silicification to be dated. Fig. lS.
ssw
Sable de Lozere
D sandstone � quartzite llll!lill claystone D limestone
g Chalk
·-
Seine Riv.
I
water table
NNE m
120
25 km
Fig. 16. Geological section of the Beauce Plateau. The Fontainebleau Sand is preserved beneath the limestone plateau of the Calcaire de Beauce and form the scarp of the dissected plateau. The Fontainebleau Sand forms a main aquifer flowing out to the north.
the upstream glacis in which the low-lying areas, like the Loire graben, acted as traps. The location of the thickest silicified profiles (more than 15 m deep) in the grabens and channels that formed the lowlands may be a contributing factor in this development. QUA R T Z I T E SAND
LENSES
IN
F O R MAT I O N S
Pure white sand containing quartzite lenses exists in several different Tertiary formations of the Paris Basin. The white sand generally has been considered as relatively pure marine or continental sand deposits that were slightly weathered before burial. The quartzite lenses contained in the sand have been related to early silicification, in relation with pedo genic and/or evaporative processes, before burial of the sand units (Alimen 1936; Pomerol 1965). In fact,
these quartzites appear to be related to groundwater processes occurring during the recent geomorpho logical evolution of the basin. Geology
The Fontainebleau Sand, of Stampian age (early Oligocene or Rupelian), is a 50-70 m thick, fine grained, well-sorted marine sand topped with an aeolian dune sequence (Alimen 1936). South of Paris, the Fontainebleau Sand is preserved under lacustrine limestone of late Oligocene to basal Miocene age and forms the scarps of the dissected Beauce Plateau (Fig. 16). East and north of Paris it forms only residual buttes above the Brie Plateau. The Fontainebleau Sand Formation has never been buried deeper than 50 m. There are three main diagenetic stages: from the primary deposition facies to the pure white sand con taining quartzite pans.
106
M. Thiry
Bleaching of the sand
Although the Fontainebleau Sand is white at outcrop, or at least always oxidized, some drill-holes show that dark, glauconitic, organic- and pyrite-rich sand exists beneath the plateaux. These dark sand facies occur only locally below the water table, and the boundary between the dark and the oxidized facies closely follows the shape of the piezometric level (Fig. 17). This alteration appears to have been most active near the water table and has developed in zones of groundwater outflow. It is directly dependent on the incision of the plateaux. Comparison of the mineralogical and chemical compositions of the bleached and dark facies shows that the bleaching is accompanied by a considerable alteration of the sand. The dark facies are devoid of carbonate, but casts of shells show that carbonate was present in the primary facies. Feldspars are almost entirely altered in the oxidized facies and although smectite is the main clay mineral in the dark facies, it becomes subsidiary in the oxidized facies, where kaolinite is the main clay mineral. Furthermore, the amount of clay minerals decreases considerably in the oxidized facies. Distribution ofthe quartzite
Silicification within the Fontainebleau Sand has pro duced flat-lying, very tightly cemented sandstone or sedimentary quartzite layers 2-8 m thick at different levels in the formation. The quartzite is composed mostly of quartz grains with overgrowths; chalcedony and microcrystalline quartz are only subordinate silica phases. Drill-hole data indicate that the quartzite layers that crop out on the edges of the plateaux and in the
w
COIGNIERE
TRAPPES
valleys do not extend more than a few hundred metres beneath the limestone-covered plateaux (Thiry et al. 1988). The quartzite is, thus, restricted to the outcrop zones of the Fontainebleau Sand and linked to the present or recent morphology (Fig. 18). There is a single quartzite layer or two to four superposed layers at a depth interval of 10-20 m under the limestone cover (Fig. 19a). The quartzite layers are about 2-5 m thick and extend laterally for 50-200 m or are composed of several smaller lenses, sometimes arranged in steps. The thickest and most numerous layers are found in the thalwegs formed by secondary drainage incisions. Silicification occurred after bleaching of the sand. The strong link between the quartzite and the present geomorphology suggests that silicification occurred relatively recently in outcrop zones of the Fontaine bleau Sand. The distribution of the quartzite in super imposed horizontal layers may indicate a relationship with water-table levels. Alteration ofthe white sand and the quartzite lenses
Further alteration of the sand can be observed in the quarries supplying sand to the glass industries at the edge of the Beauce Plateau. Here, near the outcrop, where the limestone cover becomes thinner, the sand is shining white and very pure (more than 99.5% Si02) and tends to be totally devoid of clay minerals. Quartzite lenses also show evidence of alteration within the white sand. The upper quartzite lenses are characterized by numerous cavities and vertical structures and pipes. Quartz overgrowths in the outer layers of these lenses are corroded, suggesting disso lution of the quartzite in the upper part of the sec tions (Milnes & Thiry 1992).
SAC LAY
E
m
Geological section of the Tertiary plateau south of Paris showing the distribution of the primary black and secondary oxidized facies in the Sables de Fontainebleau Formation. The boundary between dark and oxidized facies corresponds approximately to the groundwater level. Borehole data define both the extent of the sand facies and the water table.
Fig. 17.
107
Diversity of continental silicifications W-SW Yvette Riv.
E-NE Rhodon Riv. Merantaise Riv. �r-------� ,-------�
m
160 130 100
_ _ _-_-_-_-_-_-_ -- - _-_-_-_-_-_-_-_-_-_-_-_-_- _ -_ _ --- -
0 EJ
Marnes a Huitres quartzite lense
D D
Sable de Fontainebleau water table
D
I
Calcaire d'Etampes et Argiles a Meulieres drill hole
1 km
Fig. IS. Arrangement of the quartzite lenses in the Fontainebleau Sand on the dissected Beauce Plateau south of Paris. The quartzite lenses, which crop out on the edges of the plateaux and in the valleys, do not extend beneath the limestone-covered plateaux. Borehole data define both the extent of the quartzite lenses and the modern groundwater level.
(a)
(b)
(c)
(d) Fig. 19. Silicifi cation in the Fontainebleau Sand. (a) Silicifi ed zones in the Fontainebleu Sand occur as superposed quartzite lenses. The man in the centre of the picture is about 2m tall. (b) Quartzite lens with a typical wing-shaped morphology and features of superposed silicifi ed layers. (c) Botryoidal or ' custard-like' aspects of silicifi ed lenses. These features result from successive silicifi ed layers that point to centrifugal growth of the silicifi cation. Scale bar is 5 em. (d) Silicified pans displaying areas underlined by tightly cemented fringes that indicate a growth of the pans by successive silicified layers. The outer surface of the pan displays 'custard-like' features related to these silicified layers. Scale bar is 20cm.
108
M. Thiry
Flint pebbles give detailed information of the alteration process. They are round and show a translucent fracture surface at the base of the forma tion and become altered and friable in the upper white sand. Where the pebbles are embedded in the quartzite lenses, they keep their original rounded shape and translucent fracture surface. The pebbles in the white sand obviously have been weathered after the formation of the quartzite lenses. This late alteration occurred in the vadose zone and resulted mainly in silica leaching, such that the upper quartzite lenses dissolved at their edges. In places, dissolution of the remaining fossil-bearing sand layers at the base of the profiles, close to the present water table, is accompained by precipitation of large euhedral calcite crystals cementing the sand grains. Dating of these calcite crystals by 14C indicates an age of about 30 000yr, confirming that this alter ation process is probably still active. Growth of the quartzite lenses
The most outstanding feature of these quartzite lenses is the contrast they show between the very hard, tightly cemented quartzite and the loose and permeable embedding sand. Boundaries are always sharp and the transition from the sand towards the quartzite generally occurs with a friable rim less than 1 em thick. This pattern raises the question of the growth mechanism of the lenses. If the cementing silica was provided by solution the lenses cannot have been formed in one step. As soon as the sand is partly cemented, the porosity decreases and the feeding solutions take a different pathway so that a tight cementation is never obtained. A progressive cen trifugal growth of the silicified bodies, such as concre tions, may explain this arrangement. To find the answers detailed morphological and petrographical observations are necessary. Morphology of the quartzite lenses
Except for the upper lenses that show vertically shaped structures linked to silica dissolution after silicification, most of the silicified bodies are fiat lying. There is a great variety of forms and sizes. The quartzite morphologies are always smoothly rounded, metre-long bodies, frequently shaped like spindles or like aircraft fuselages or wings, with one well-rounded and one thinner edge (Fig. 19b ). More complicated forms also exist and may result from a coalescence of neighbouring bodies, but they always
remain streamlined. A striking feature of most of the quartzite lenses is that they predominantly extend towards the outcrop fringe or the neighbour ing valley. These morphologies, although larger, are similar to the elongated and orientated calcite con cretions thought to have precipitated in the saturated zone parallel to the palaeo-groundwater flow direc tion in clastic aquifers (Jacob 1973; Pirrie 1987; Johnson 1989; McBride et al. 1994; Mozley & Davis 1996). The characteristic shapes of the quartzite lenses clearly link the silicification to palaeohydro logical flows. The outer fringe of the quartzite lenses frequently has botryoidal or 'custard-like' aspects. These fea tures are caused by a succession of silicified layers, suggesting a centrifugal growth of the silicification (Fig. 19c). Generally there is no structure inside the sandstone, but sometimes tightly cemented fringes are visible inside the silicified lenses (Fig. 19d). This confirms that the lenses grew by accumulation of suc cessive silicified layers. Silicification progresses around a core by cementation of successive layers that overlap in a discordant way. The spatial arrange ments of these botryoidal features show the flow pattern of the feeding solution. Petrography ofthe quartzite
The quartzite lenses are composed of tightly cemented quartz sandstone with well-developed quartz overgrowths and low residual porosity. The detrital grains are very clean and in a few examples contain impurities that outline the overgrowths. Cathodoluminescence of the quartzite shows detrital grains and three main habits in the quartz over growths (Marechal 1995). 1 Subeuhedral quartz overgrowths with successive dark and clear stripes (Fig. 20a). The overgrowths are sutured along straight lines that converge towards triple junction points, leading to polygonal contacts. These are syntaxial overgrowths around the detrital grain. 2 Isopachous quartz layers surrounding the detrital grains (Fig. 20b ). The coatings are alternately bright and non-luminescent. These isopachous layers cannot be linked to quartz overgrowths, which would lead to euhedral shapes. They suggest amorphous or poorly ordered silica deposits that later recrystallized into quartz from the detrital core. 3 Some overgrowth stripes have flame- or feather like structures consisting of successive bright and dull sections radiating from the core (Fig. 20c). These
109
Diversity of continental silicifications
(c) (a)
(b)
Fig. 20. Cathodoluminescence of the quartz overgrowths in the Fontainebleau Sandstone. (a) Syntaxial overgrowths. (b) Isopachous layers fi lling the pores: 1 , detrital quartz; 2, quartz layers; 3, fi nal quartz cement occl uding porosity. (c) Flame- or feather-like features radiating from the detrital quartz grain. Scale bars are 50 ).l m.
flame-like structures also may relate to centrifugal recrystallization of a primary silica deposit arround the detrital grain. The quartzite lenses may be cemented by one or the other of the above secondary quartz overgrowths. Sometimes all three types are present and arranged in a sequential fashion: syntaxial overgrowths are fol lowed by isopachous layers and then again syntaxial overgrowths. Corrosion of the syntaxial overgrowths and of the isopachous deposits is sometimes visible at the transition from one zone to the other. The thick ness of the quartz rings, and thus the residual poros ity, varies as well. The sequences are on the scale of a centimetre and correspond to the successive silicified layers revealed by the botryoidal features of the quartzite lenses. The different cements and the sporadic silica dissolutions probably reflect changes in the silicification environment, which may occur periodically and/or laterally along the solution pathway.
lenses in thalwegs adjacent to the main valleys, where ground water is discharged. The quartzite lenses develop behind springs or spring lines, near the valley floor and mark the successive levels of falling water tables (Fig. 21 ). When the water table fell as a result of continued downcutting of the landscape, former sil cretes were isolated in the vadose zone, where they were dissolved along the boundary. The recurring silicification episodes lead also to a lateral sequence of the successive quartzite generations along a valley. During one stage of river downcutting, quartzite lenses would have formed at the top of the sand immediately under the limestone cover at the head of the valley as well as in deeper layers of the sand formation downstream. The quartzite lenses situated at the same altitude at the edge of a valley can, thus, be related to several generations of formation.
Model of groundwater silicification
Groundwater chemistry and silica solubility
The strong relationship of the quartzite lenses with the present geomorphology suggests that silicifi cation occurred relatively recently in the outcrop zones of the Fontainebleau Sand. Moreover, the shapes and orientations of these lenses are thought to relate to precipitation governed by groundwater flow. This hypothesis of silicification by groundwater flow explains the location of the thicker quartzite
The composition of the ground water displays a con sistent pattern throughout the Beauce region, with a general S-N variation (Bariteau 1996). This variation is linked directly to the flowpath followed by the ground water (Fig. 16). 1 In the south, all the ground water is contained in limestones of the Calcaire de Beauce Formation. The water is of calco-carbonaceous character, with a pH
Mechanisms of silicification
110
M. Thiry
-
�
=
�
L
-
-
/
...... ..._ -
------------------------- -- ---- -
/
--
�/
""'
-
-
Calcaire d'Etam pes
Sable de Fontai nebleau unsat u rated/sat u rated
M arnes a H u itres
q uartzite lens
500 m
5 m
I
Fig. 21. Schematic model of successive stages of groundwater silcrete formation in the Fontainebleau Sand. Quartzite lenses develop at the water table behind springs near the valley floor. Downcutting of the drainage with accompaning lowering of the water table leads to deeper quartzite levels. Note that quartzite lenses occurring at a certain level in the sand in different valleys are not related to the same episode of silicification.
of around 6.5. The Si02 content is derived from the alteration of clay minerals contained in the Calcaire de Beauce Formation. The ground water averages 14-17p.p.m. of Si02 and is thus oversaturated with quartz and cristobalite. 2 In the north, the ground water is restricted to the Fontainebleau Sand Formation and the limestone cover is thin. Therefore, the infiltration water that feeds the water-table is less mineralized and progres sively dilutes the ground water toward the outflow areas. The dissolved charge of the ground water decreases towards the outcrop areas of the Fontainebleau Sand, and its silica content decreases to 10-12 p.p.m. Thus, paradoxically, near the valley, where the ground water has a lower silica content, the sand stone is cemented. The mechanism of silica deposi tion probably is complex and related to variations in the physico-chemical characteristics of the ground water as it approaches the outflow areas. These varia tions may result from the mixing of the meteoric or soil water with the ground water. Two factors may intervene: the behaviour of silica in solution and crystal growth laws. Variation in the pH or the temperature cannot explain the silicification linked to the outflow of the
ground water because they are too weak to influence silica precipitation significantly. Organic compounds may play a role. Some dissolved organic compounds may form complexes with silica and increase its solu bility, as observed in natural environments and in the laboratory (Bennett et al. 1988; Bennett & Siegel 1989). If such mechanisms were active in the cemen tation of the Fontainebleau Sand, these silico-organic complexes must have been present in the ground water and then destroyed (oxidized?) near the outflow areas, probably by mixing with infiltration water. Infiltration water always contains more dis solved oxygen than ground water, but it is also slightly more acidic and enriched in organic com pounds (Bariteau 1996). Quartz growth and amorphous silica deposits
Contrary to silica dissolution, fewer studies have been made of precipitation of silica phases, especially at low temperatures. Quartz growth is very slow and difficult to achieve experimentally in surficial condi tions. Generally speaking, crystal surface properties are basic points for crystal nucleation and growth. Some elements inhibit or, on the contrary, favour pre cipitation at the grain surface. Some minerals, such as
111
Diversity of continental silicifications
clay minerals, carbonates, or even iron oxides, coating quartz grains can prevent silica cementation (Heald & Larese 1974; McBride 1989; Ehrenberg 1993) . Metallic cations, such as AP+, adsorbed at the grain surface can inhibit quartz growth (Delmas et al. 1982) . The ions adsorbed at the surface of minerals, quartz in this case, form a barrier that prevents the H4Si04 molecules from 'grafting' themselves on to the existing crystal network and building it up. On the other hand, some ions can act as catalysts and favour crystal growth. Amorphous silica forms by polymer ization of H4Si04, then aggregation of the polymers forms colloids, and finally precipitation of the latter forms opal-A (Iler 1979; Williams et al. 1985; Chan et al. 1995 ) . Aggregation of the polymers and floccula tion of the colloids can be favoured by the presence of some ions in the solution. In the case of the Fontainebleau Sand, phenomena related to quartz crystallization and silica flocculation probably explain the distribution of the silica cement. In the Beauce ground water, where neither silici fication nor quartz overgrowth is discernible, the rela tively high silica oversaturation with regard to quartz solubility is probably attributable to surface pro perties of the quartz grains and/or the presence of inhibiting elements that prevent quartz growth on the detrital grains. Near the outflow areas, the arrival of infiltration water with quite different properties (more dilute and acidic, enriched in dissolved organic compounds, etc. ) , may alter the surface properties of the quartz grains, cancel the precipitation inhibitions and thus allow silica deposition. Modification of the silica speciation in solution may act in a similar way. The arrival of water with different geochemical char acteristics may destroy the silica complexes, and if they are quickly broken down will lead to high over saturation and deposits of amorphous silica. In the case of the Fontainebleau Sand, only processes that involve the surface properties of the quartz and the stability of silica complexes seem able to explain silica precipitation near the groundwater outflow areas; even though the water is more dilute. Whatever the mechanism, it clearly appears that silica precipitation occurs at the interface between the ground water and the infiltration of water that has passed through only a thin limestone cover. Silica precipitation along an interface may explain the sharp (less than a centimetre-wide) boundary between cemented quartzite and loose sand. Like wise, the successive silicified layers that have created the quartzite lenses may be a manifestation of this interface, where silica precipitates.
Cementation and flow rate
A well-cemented and impermeable quartzite lens, one or even several metres thick, cannot form as a single event in a homogeneous and isotropic fashion. When the sand cementation begins the permeability decreases and the water flow is diverted around the silicified core. Therefore it becomes difficult for silica to reach these conspicuously silicified cores and com plete the cementation. To create totally impermeable silicified zones, such as those in the Fontainebleau Sand, silicification has to take place centrifugally around a core. The successive silicified layers at the surface of the quartzite pans and silica mineral sequences shown by cathodoluminescence are con sistent with this type of cementation process, i.e. the formation of successive envelopes around a silicified core. The sand cementation was modelled with a coupled flow-transport code (METIS ) ( Man�chal 1996) . The model was constrained with permeability values, the hydraulic gradient and silica content mea sured in the Fontainebleau Sand and average quartz precipitation kinetics were used. Two models of water flow were constructed: a linear one with only parallel flow towards the valley and a radial one with con verging flow lines. The models emphasize the impor tance of the groundwater flow rate. With the linear flow model there is no substantial cementation after 1 Myr of silica precipitation. With the radial flow model, however, quartzite lenses 3 m thick with only 5-10% residual porosity formed after 100000500000yr. Cementation controlled by water flow explains why the thickest and most numerous lenses are found in the thalwegs formed by secondary drainage incision, where the groundwater outflow lines converge. S I L I C I FI E D L A C U S T R I N E LIMESTO N E S Geology
The extensive Tertiary continental limestones o f the Paris Basin form superimposed plateaux that con stitute the main geomorphological features of this basin. Typical examples are the Brie and Beauce plateaux (Fig. 16 ) . Almost all of these limestones contain cherts. Their size varies from millimetre-sized dots to siliceous bodies several tens of metres long. The silicified facies are distributed very irregularly in the limestones. Locally, they can form over 10% of a
112
M. Thiry
quarry cliff, but be absent a short distance away. On a regional scale they average about 5% of the volume of the lacustrine limestone in the southern Paris Basin (Aubert & Lorain 1977; Menillet 1980) . Although these silicification features were known in the last century (Cuvier & Brongniart 1808; Cayeux 1929; Menillet 1988b) , their genesis has not been explained. Two main types can be distinguished from their position within the limestone. 1 Concordant cherts, in regular layers, formed of opal-CT, generally bound with marls containing fibrous clay minerals (sepiolite and palygorskite) that may have developed during sedimentation or early diagenesis. 2 Discordant cherts, of irregular shape, massive or with numerous limestone inclusions, formed almost exclusively of quartz. The silicification followed recrystallization and dissolution phenomena of the limestone and thus appears as a feature of late epi genesis. The period of silicification is still not known: did it occur between two phases of deposition, as assumed by most authors, or much later through recent weathering processes? The section of the Calcaire de Champigny (upper Eocene) studied here is located in the eastern Brie Plateau (Pecy, Seine-et-Marne) . It is typical of the discordant and irregularly shaped silicification fea tures found in the 'lacustrine' limestones of the Paris Basin (Fig. 22) .
lithology
m
karst features
silicified bodies
lacustrine limestone \ �
marl pervasive silicification
Fig. 22. Schematic diagram of the Calcaire de Champigny Formation (upper Eocene) showing the distribution and shapes of the silicifi ed zones (Pecy, Seine-et-Marne ) Note similarity between the distribution and the shapes of the silicifi ed zones and the dissolution features of the limestone. .
In the middle of the section, the silicified zones are of two types: (i ) pervasive silicification, which pre serves the structures and the dull aspect of the lime stone (Fig. 23b ) and forms an irregular layer up to 1 m thick; it includes (ii) irregularly shaped translucent flints with a cross-section of about 0.2-0.5 dm2 , dis posed in a more or less cross-connected network and containing numerous 'pseudo-clasts' of preserved limestone (Fig. 23c ) . In the pervasively silicified area it frequently is impossible to distinguish between silicified and unsilicified zones without appropriate tests. 3 At the top, silicification appears as scattered small nodules of less than 0.1 dm2 ; they are partially dis solved and often microporous.
2
The lacustrine limestone
The limestone is about 20 m thick with only rough stratification and has very low amounts of detrital quartz grains and clay minerals. It shows nodular and 'pseudo-breccia' facies as a result of many successive modification processes during and after deposition (mud cracks, microkarst, reworking of milli- to cen timetre-sized elements, etc. ) . Joints filled with large sparite crystals are obviously more recent. A micro karst network developed at a later stage. The cherts
Silicification mainly affects the base of the section above an irregular clay layer. Generally, silicified facies are clearly discordant to sedimentary struc tures and ch�rts vary in size and shape from the base to the top of the section. 1 At the base, silicified zones form translucent fiat irregular layers up to 1 m long and 2-10 dm2 in cross section (Fig. 23a) .
Micromorphological features
Petrofabrics of the various silica phases, their distri bution, organization and succession are the keys to understanding the mechanism of silica arrival and settling in the system. Two main types of silicified facies can be distinguished: silica deposits in voids
Diversity ofcontinental silicifications
113
(a)
(c) (b) Fig. 23. Silicifi ed lacustrine limestone. (a) Flat-lying chert layer. (b) Irregularly shaped cherts. (c) Amoeboidal chert (dark) containing 'pseudo-clasts' of preserved limestone (clear). Scale bars are 5 em.
and replacement of a former limestone matrix. Voids are the paths along which the solutions flow through the formation. Their distribution and infillings are of special interest for the interpretation of silicification processes. Quartz is almost the only silica phase in the cherts. Chalcedonite is scarce and opal is only a minor phase and cannot be identified with certainty when associ ated with fine microcrystalline quartz. A wide range of quartz crystal sizes exists in the void deposits and in the replaced matrix. Crystal sizes range from tiny microcrystals under 1 f.!m to well-developed euhedral crystals up to 500 f.!m. 1 Microcrystalline quartz is by far the most common silica phase of the replaced matrix and shows no preferential orientation. The crystals are irregular, amoebic, intergrown and 1-2 f.!m in diameter. The microcrystalline zones generally are very pure and devoid of calcite inclusions. 2 Larger amoebic quartz grains 10-50 f.!m in diame ter develop in places. They are of irregular shape and generally have a dark core with fine calcite inclusions and a clear outer fringe. 3 Dark euhedral quartz with numerous calcite inclu sions is a singul?r facies that develops in the micritic matrix. The size of the inclusions frequently reaches 100 f.!m. They indicate the growth zones. Microprobe analyses show that micro-inclusion contents average about 15-20% and may reach 40% calcite.
Clear euhedral quartz crystals develop in a 'pali sadic' fashion in joints and voids. Their sizes vary from 10 to 100 f.!m. 4
Silicification-porosity relationship
Thin-sections show a systematic relationship be tween silicification and high porosity zones, with the latter being preserved or partly infilled with silica. Voids appear as straight joints, cracks around nodules (Fig. 24a) or cylindrical pipes that form a microkarst network in the limestone matrix. The primary voids average between 100 and 200 f.!m in diameter. The preserved voids plus the traces of former ones (now infilled with silica) frequently represent a porosity of about 30% and locally more. In some places, the fre quency of the voids is such that they must have devel oped during the silicification process after the earlier ones were infilled with silica. On the other hand, silicification does not develop where porosity is very low. The size of the silica deposits varies from one silicified zone to another, but also varies within the same zone, and even in the same joint on a millimetre scale. In the core of a silicified zone the joints can be rimmed with thick quartz deposits, whereas away from the core the quartz deposit as well as the replaced border becomes thinner and even reduced to a thin rim of microcrystalline quartz, and finally disappears alto-
114
M. Thiry
(a)
(b) Fig. 24. Silicification-porosity relationship in the silicified limestone. (a) Silicified zone showing a high primary porosity partly infilled with secondary silica. Voids appear in black, silica infillings show as clear rims, replaced matrix is flecked and micrite is also dark. Silica infillings show a vertical gradient, they are thick at the base and progressively thinner toward the top. (b) The zone at the left is highly porous and strongly silicified, to the right the joints are cemented by early sparite crystals (these also appear clear), and voids and silicification are more sparse (see sketch on Fig. 25): 1, microcrystalline quartz; 2, euhedral quartz; v, void; s, sparite; m, micrite. Scale bars are 2 mm. Crossed polars.
gether (Figs 24b & 25). A similar variation can be observed in facies with disconnected voids, where the whole matrix is replaced by silica when the voids are close together and residual limestone zones remain when voids are more scattered (Fig. 26) . 2 mm
Silica infillings
The voids are entirely or partly infilled with silica. The first deposits on the edges of voids are of opal or microquartz followed by small palisadic quartz crys tals and finally by euhedral quartz crystals toward the centre. The deposit can consist of a narrow strip only a few microns wide if silicification is limited. Otherwise quartz fills up the entire void if silicification has progressed sufficiently. In places, the quartz shows ribbon structures that are of two types: 1 regularly spaced crenellated ribbons indicate the successive growth zones of quartz crystals. These quartz crystals probably developed by crystal growth from solution; 2 botryoidal ribbons developed in the flamboyant · quartz, with ·undulating extinction. In this case the quartz did not grow directly from solution but pro bably by recrystallization of primary opal botryoids.
D micrite matrix 0 microquartz replacements � palisadic quartz filling voids LJ void Fig. 25. Silicified zones always show a high primary porosity infilled with thick secondary silica deposits. Away from the silicified core, in the adjacent joints, the silica deposits and the replaced limestone border become thinner, and at last disappear totally.
Limestone replacement
In the silicified zones the limestone matrix has been replaced by quartz but most of the limestone struc tures have been preserved: it is still possible to recog nize former fossils, nodules, pseudo-clasts and sparite joints in the silicified matrix. These earlier structures can be identified because of changes in the silica
1 15
Diversity of continental silicifications
Fig. 26. The highly silicified zones of the lacustrine limestone show numerous former voids, outlined by flame-like palisadic quartz crystals, with shapes that are similar to microkarsts. Zones of micrite remain when the voids become less frequent. Late clear sparite fills the voids remaining after silicification.
D
micrite
D microq uartz � palisadic quartz -
microfabric: clear zones of pure quartz, dark zones with numerous inclusions of residual calcite, thinner or coarser quartz crystals, etc. Microcrystalline quartz is the main silica form in the replaced zones. In places, euhedral quartz with micro-inclusions of calcite have also developed. Voids play a major role in the replacement of the limestone matrix by quartz. In all silicified zones there is evidence of high primary porosity. On the edges of these zones the replacement of the lime stone matrix is restricted to the wall of the voids and becomes progressively thinner, like the deposits observed in voids (Fig. 25) . Matrix replacement has developed over a distance that does not exceed 0.5-1 mm from the border of a void. Sequences ofsilicification
Silicification also shows quantitative and qualitative variations within a single silicified zone. Frequently it varies inside a centimetre-wide vertical sequence that can be repeated several times in a sample (Fig. 27). 1 Silica infillings in the voids progressively decrease from a strongly silicified basal zone, in which initial pores are entirely filled with silica, to only thin silicified rims at the top of the silicified zone before they disappear entirely. 2 Matrix replacement shows similar variations, with a more or less 'massive' replaced basal zone, topped by discontinuous silicified zones restricted to cracks, which gradually diminish in size and disappear. 3 The size of the quartz crystals also varies in a verti-
spa rite
cal direction. At the base of many silicified zones there are relatively large amoebic patches of quartz (over 10 11m in diameter), some with cores containing calcite inclusions. Their sizes progressively decrease upward and they form only tiny microcrystals toward the top of the zone, where replacement is discontinuous. Interpretation and mechanisms
Origin and distribution ofsilica
Because the limestone is pure, without clay or sand layers, the silica must come from other formations (overlying sand, loess or soil). Owing to the low solu bility of silica in surficial waters, strong flow is needed to continuously renew the silica precipitated from solution. This constraint explains the close relationship between voids and silicification features. It also explains the quantitative variation of silica deposi tion and limestone replacements along a joint: silicification is most intense in the pipe and joint sec tions where flow is the strongest and weaker where the flow is slower. Water flow also explains the verti cal sequences of silicification: the base levels that have always been below the local water table are strongly silicified, whereas higher zones that were probably only periodically under- water are less silicified. Silicification must be considered as a groundwater process similar to those proposed by Khalaf (1988),
116
M. Thiry restricted replacement and only thin silica deposit in the voids
irregular replacement around zones of high porosity
extensive silica replacement with thick silica deposit in the voids preferential silicification of some micrite nodules
D silica (matrix and void deposits) D
micrite matrix
Fig. 27. Sketch of the vertical sequence frequently shown by silicification features. Above a tightly silicified basal zone, the replaced limestone fringe along the pores and joints thins out progressively.
Thiry et al. (1988), Arakel et al. (1989) and Thiry & Ben Brahim (1997).
ment front, the lower is the diffusion gradient. The replacement thus slows down and finally stops. The relative magnitude of silica infillings versus matrix replacement is probably the result of variable condi tions of quartz growth and related crystallization pressures. Age of silicification
The hydrodynamic conditions required to ensure the silica supply imply a hydraulic head gradient within the limestone formation. This is not the case during sedimentation because there are no areas of relatively high relief around the basin. Substantial groundwater flow is possible only after the limestone formation had been uplifted and incised by valleys. This limits the possibility of silicification to the Pliocene and Quaternary Periods and suggests that the phenomenon is relatively rapid. Several authors have noticed that the limestone is frequently more silicified, with wide pervasive silicified layers, in areas where it crops out (Turland 1974; Hatrival 1983). This pattern of distribution for the silicified facies cor roborates the interpretation of the silicification as being relatively recent and after the erosion of the sedimentary cover. M E U L I E R E S : W E AT H E R I N G
OF T H E
S I LI C I F I E D L I M E S T O N E
Mechanism of replacement
The replacement of the limestone by silica with perfect conservation of the structures implies that silica precipitation and calcite dissolution occurred simultaneously. Replacement will occur only if quartz crystallization of itself induces calcite dissolution. Crystallization pressure of quartz over calcite is the more likely mechanism inducing this replacement, as argued by Maliva & Siever (1988), Dewers & Ortoleva (1990) and Maliva (1992). The growth of quartz crystals exerts a pressure on the neighbouring calcite crystals which increases their solubility. Moreover, this mechanism implies silica supply and calcite evacuation on the replacement front. Silica probably is supplied from the joints where water circ,ulates at the same time as dissolved calcite is evacuated tow.ards these joints. Movement of these elements would proceed through diffusion along fluid films on crystal j oints. The diffusion gradi ent conditions the extent of the replacement. The longer the distance between the void and the replace-
A particular type o f siliceous and cavernous rocks in the Paris Basin, called meulieres (millstones), is the product of weathering of the silicified lacustrine limestone facies. The intriguing feature of the meulieres is their cellular aspect. Indeed, the charac teristic meulieres consist of a framework of thin (less than 1 mm thick) siliceous plates outlining centime tre-sized hollows, giving the rock a very low density. In spite of the size of the voids (more than 50% porosity), the meulieres have a low permeability, most of the voids are enclosed or communicate only with a few neighbouring voids. The meulieres were quoted for the first time in the geological literature by Guettard ( 1756). They are generally embedded in variegated clay, and the for mation as a whole is called 'Argiles a Meulieres'. Most authors have regarded the Argiles a Meulieres as a weathering product of the silicified lacustrine lime stone. The only matter under discussion was whether the meulieres resulted simply from dissolution of the limestone or whether the dissolution brought further
Diversity of continental silicifications BRIE PLATEAU
HURE POIX
m a .s.l.
w
200 Calcaire de Beauce
Calcaire de Brie
Sable de Fontainebleau
�
117
pedogenic silcrete armour
�
1 00
ground-water qu artzite l ens
•
0
Argiles a Meuli eres weathering complex
Fig. 28. Geological section through the Brie and Hurepoix (Beauce) Plateaux showing the layout of the 'Argiles a Meulieres' weathering formations. The Argiles a Meulieres shows thicker profiles and more 'mature' facies on the Hurepoix Plateaux than on the Brie Plateau and relate to a longer evolution of the profiles on the upper plateaux than on the more recently denuded lower plateau.
alteration and transformation of the primary facies (Cuvier & Brongniart 1808; Dollfus 1885; Gosselet 1896; Cayeux 1929; Alimen 1936; Cholley 1943; Prost 1961). Dollfus (1885) made the assumption that weathering and silicification of the lacustrine lime stones were linked. He thought that the silicified facies of the Calcaire de Brie might be related to a late and continuing alteration of the limestone as a result of infiltration of rain water dissolving the lime stone at the top and precipitating silica at depth. Recently, Menillet (1987) re-examined these ques tions and made a distinction between inheritance and weathering. We refer to much of his work here. The meulieres have been actively quarried during the last century for mill- and building-stones. Most of the sewer network in Paris and the railways around Paris were built with meulieres. Geology
The 'Argiles a Meulieres' Formation delineates ancient plateau surfaces in the southern Paris Basin (Brie and Beauce Plateaux) (Fig. 28). The Beauce Plateau and the fragmented buttes of the Hurepoix region that constitute its northern extension form the highest surface of the Paris Basin and generally date back to the Miocene (Cholley 1960; Klein 1975). The Brie Plateau corresponds to an older marine erosion surface (Fontainebleau Sand, Oligocene) and has only relatively recently (early Quaternary) been stripped of its sedimentary cover (Lutaud 1948). In Hurepoix, the Argiles a Meulieres cover the Oligocene formations (Fontainebleau Sand and Calcaire d'Etampes) and the Late Pliocene fluviatile deposits (Sables de Lozere) are not affected by the
'meulierization' process. On the Brie Plateau, the Argiles a Meulieres Formation seems very discon tinuous and is even absent in areas where the Calcaire de Brie limestone is at its maximum thickness; on the other hand it is well developed when it rests on more clayey facies (Argiles Vertes de Romainville, Calcaire de Champigny, Barthonian Marls). The Argiles a Meulieres generally has thicker profiles and more 'mature' facies on the Hurepoix Plateaux than on the Brie Plateau (Menillet 1987). The reason is the longer evolution of the profiles on the Hurepoix Plateau than on the more recently cleared Brie Plateau. Profile description
The 'Argiles a Meulieres' is a very heterogeneous for mation, 2-10 m thick. lt consists of irregularly shaped, often cellular, siliceous blocks (called meulieres) packed in a clayey or sandy matrix of a variegated, reddish and ochre to greyish colour. A vertical facies arrangement, however, can be distinguished (Fig. 29). 1 Mostly, the Argiles a Meulieres overlie a lacustrine limestone with numerous microkarst dissolution features. In places, however, it can be superimposed on clay formations, such as the Argiles Vertes de Romainville, or even on sandy ones, such as the Fontainebleau Sand. 2 At the base of the Argiles a Meulieres there is gen erally a clay horizon with residual limestone boulders and silicified limestone masses. The clay is brown and red with hydromorphic features emphasized by grey mottles arranged in a reticular pattern. Metre-size collapse structures related to karst dissolution phe nomena in the underlying limestone are frequent. 3 The lower part of the siliceous horizon is made up
118
M. Thiry
of compact siliceous masses embedded in brown clay. The siliceous masses are of irregular shape and their size varies from 0.5 m to several metres in diameter. They are generally dull, but can become slightly translucent in places and are light in colour (white, yellowish or even greenish). A scoria-like rim sur rounds the blocks and develops along joints. The brown clay contains numerous illuviation features. 4 The upper part of the siliceous horizon consists of porous or cellular siliceous masses with illuvial varie gated clays. The siliceous masses form horizontally elongated slabs 0.5-1 m thick at the base of the horizon, which become smaller towards the top, where they generally form irregularly shaped stones from 0.1 to 0.5 m in diameter. They are pale yellow to ochre. 5 Usually the meulieres do not crop out at the surface of the plateaux, but are topped by a blanket of Quaternary loess and a soil profile. The meulieres formation is rather discontinuous. The meulieres form humps from a few metres to
several tens of metres in diameter without any par ticular structure. These siliceous humps are divided by similarly sized 'hollows' infilled by brown clay and sandy clay. The various meuliere facies
The meulieres possess a large variety of facies, from massive to highly porous 'cellular' facies, and are finally weathered out and desilicified into more friable blocks, even breaking down into a rough sand like material. The parent rocks of all the Argiles a Meulieres facies are the silicified lacustrine lime stones of the Paris Basin, as described above. Weath ering and karst development within these limestones provide the original siliceous material. As the lime stones generally have very little insoluble residue, the embedding clay is mainly part supplied by the overly ing loess and sand, or even inherited directly from the underlying clay and marl formations. The inherited silicified limestones
m soil on Quaternary loess 8
cel l u lar meuliere 6 massive meul iere
The inherited silicified limestones retain most of the features found in the limestones. The large silicified masses generally have a dull aspect, preserving the nodular and pseudo-breccia fabrics. It is even pos sible to recognize fossils just as in the silicified lime stones. They retain calcareous nodules and small pseudo-clasts within the siliceous network. This inherited silicified limestones facies forms the massive lower horizon of the meulieres profile. The dissolution of the calcareous remains at the edges of the blocks and along the joints creates a scoria-like nm.
4 brown hydromorphic clay with residual limestone and silicified limestone
Decalcification and desilicification ofthe silicified limestones 2
karst dissolution of the lacustrine limestone
Fig. 29. Sketch of macromorphological organization of the 'Argiles a Meulieres' profi le in weathered limestone materials on plateaux of the Paris Basin. The meulieres clearly develop from the silicifi ed limestone by dissolution of the residual limestone patches to form the highly porous 'cellular' facies and deposition of secondary silica in the voids.
Decalcification of the silicified limestones is the first step in the weathering process that leads to develop ment of the cellular facies (Fig. 30a). Dissolution of the limestone residues produces a more or less porous or cellular facies, depending on the magnitude of the primary silica deposits and replacements. The framework of these decalcified facies is consti tuted by veins and joints coated with silica deposits (palisadic quartz and chalcedony), sometimes sealed by subeuhedral quartz of the silicified limestones. Behind and alongside these silica deposits there lies a more or less thick fringe of various silica petrofabrics, corresponding to the replaced fringe of the limestone
119
Diversity of continental silicifications
along the joints (Fig. 30b & c). There is a great variety of petrofabrics, similar to those known in the silicified limestones: tiny microcrystallline quartz, amoebic and euhedral quartz with dark inclusions, etc. The replacements made of larger quartz crystals give way to nodules and zones of microporous fabrics, where the residual calcite between the crystals has been dissolved. The size and frequency of the resulting hollows or lacuna are very variable and depend directly on the density of the primary silicification. Desilicification of the silicified limestones follows the decalcification higher up in the profile. There, and/or around the siliceous masses, the hollows become progressively larger: in places, the silica framework is almost reduced to the primary silica deposits in the joints, the replaced limestone fringe is only residual, being completely dissolved in places. This silica dissolution affects mainly the tiny micro crystallline quartz petrofabrics, which may contain some opal, and appears to preserve the petrofabrics with larger quartz crystals. It starts by creating small hollows that widen progressively and join together to form larger ones. Finally, when the desilicification process is completed, the silica network breaks down, forming a coarse sand around the meulieres blocks in the higher parts of the profile.
(a)
(b)
The hollows may develop with illuviation of clay cutans. In places, the outer hollows of the meuliere slabs are infilled with ferruginous clay illuviation cutans. The clays are well orientated and formed of a mixture of kaolinite, illite and in some places smectite. Secondary silica deposits
There are several types of secondary silica deposits in the Argiles a Meulieres profile. The most noticeable is euhedral quartz growing in the secondary hollows created by decalcification and desilicification. These crystals grow in opposite direction back to back, in relation to the primary crystals of the silicified lime stone (Fig. 31). They can only grow after the develop ment of secondary voids. These quartz crystals can form secondary walls, which contribute to the development of the cellular meuliere facies when further dissolution of the primary silica matrix will occur. The silica deposits in the voids and joints of the silicified limestone are usually clear and lack any colouring pigment or foreign mineral. On the other hand, some thick silica deposits of the lower massive meuliere facies (mostly chalcedony, seldom opal) are
(c)
Fig. 30. The cellular siliceous facies of the meulieres. (a) Typical cellular facies formed of thin silica walls isolating prismatic voids. Scale bar 2 cm. (b) Residual silica deposits around primary porosity in the former silicified limestone: 1, primary void; 2, silica deposits around the primary void; 3, replacement of the primary limestone; 4, secondary void resulting from the dissolution of the limestone zones. Scale bar is 100 J.lm. (c) The same. Crossed polars.
120
M. Thiry
D microquartz � direction of crystal g rowth rm 2nd sub-euhedral quartz deposit - 2v: secondary void � 1 st sub-euhedral quartz deposit - 1 v: primary void
brown and in places films of iron oxide are inserted between the successive silica concretions. Clearly, these silica deposits have formed in near-surface environments where iron oxides are mobile, even in the massive meuliere horizon itself or in very shallow limestone, near the weathering front. In any case, they seem to be connected with the meuliere profile. Finally, silica also developed within the clay matrix and illuviations of the Argiles a Meulieres Formation. In the scoria-like rims of the massive meulieres, some illuviated clay cutans are bleached and have lost their birefringence. They may be replaced by microcrys talline quartz that pierces the illuviation structures. In other places, complex nodular silicified facies, formed of microcrystalline quartz containing iron oxide pigments, include detrital quartz grains. As the silicified limestone never contains detrital quartz grains, these facies definitely result from silicification of the clay matrix embedding the meuliere masses. These silicified clay materials are always restricted to the lower part of the Argiles a Meulieres profile. Weathering complex and silicification-desilicification
The weathering stages
The Argiles a Meulieres is a weathering complex that rests on limestone formations of various ages con taining silicified masses. These silicified masses form the main part of the insoluble residues released by
Fig. 31. Secondary silica deposits after dissolution of the limestone remains in the meulieres. The growth of the quartz crystals in opposite directions, back to back, indicates quartz deposition after formation of the secondary voids.
the limestone during weathering. The first step of the weathering process is a karstification of the lime stones. The karst infillings are formed by silicified masses released from the limestones and sand and clay reworked from the overlying sediments and soils. The karst infillings are very heterogeneous, sometimes sandy and permeable, sometimes clayey and sealed. The matrix of the Argiles a Meulieres Formation consists of clay and sand, brown and red in colour, with hydromorphic features emphasized by grey mottles arranged in a reticular pattern. Often, the Argiles a Meulieres contain a local groundwater layer. The clay mineral suite of the Argiles a Meulieres partly reflects inheritance from the parent rocks and shows clear development of the kaolinite within the weathering profile (Prost 1961; Menillet 1988a). Illuviated clay materials frequently form well-stratified units that extend over several tens of centimetres. The hydromorphic and illuviation fea tures assign the lower part of the Argiles a Meulieres to the B horizon of a soil profile. The silicified masses show several alterations in the weathering complex of the Argiles a Meulieres. These alterations occur at different scales: along the profile from the base towards the top and from the outer fringe towards the inner core of the silicified blocks and masses. They can be arranged in a double sequence: one leading to progressive dissolution of the primary structures, the second upgrading the silicified facies (Fig. 32).
121
Diversity of continental silicifications
upper part of the profi le silica dissolution
�
limestone
l'i/H:i!i'HI
�
m icrocrystalline q u a rtz matrix
euhedral quartz in secondary voids
�
c=J
palisadic q u a rtz of the primary voids
clay ill uviations
void
"" 2 mm
Fig. 32. Schematic sketch of the development of the meuliere facies. Dissolution of the residual limestone patches ( 1 ) gives a first cellular facies (2). Hence, there are two evolution paths. In the lower part of the profile secondary silica deposits develop with euhedral quartz growing in the voids (3) and even by alteration of the illuviation cutans into microcrystalline quartz ( 4). In the upper part of the profile, the silica framework inherited from the silicified limestone is dissolved progressively and voids are enlarged (5) until the silica network collapses and becomes coarse sand, which mixes with the soil materials (6).
First there is a progressive decalcification of the silicified limestone masses released by the weather ing. This decalcification leaves a very porous silica material with numerous irregularly shaped hollows. Further weathering brings progressive dissolution of the weakest crystallized silica phases (opal and tiny microcrystalline quartz), usually leaving only the silica deposited in the voids of the limestone and leading to a very porous silica material of cellular appearance. Important silica dissolution in the upper part of the profile brings about a collapse of the silica frame and the formation of a coarse sand. Secondary silica may form in the lower horizons of the profile. Euhedral quartz crystals and sometimes chalcedony develop in the secondary voids created by the dissolution of the limestone. Secondary micro crystalline quartz also develops within the illuviated clay minerals. There is obvious leaching of the iron oxides and alteration of the clay minerals, with leach ing of the aluminium and immobilization of the silica released. In the most mature profiles, on the · Hurepoix Plateaux, the silicification of the clay min erals can produce welding together of the primary
silicified masses, thus contributing to the massive aspect of the lower horizon. Mechanisms and environments
The Argiles a Meulieres result from a weathering sequence, starting with a karst filled with residual materials and developing into a true weathering profile where the primary structures are destroyed. Silica redistribution plays a major role in the devel opment of the meulieres facies. The silica phases are progressively dissolved towards the top of the forma tion depending on their degree of crystallinity. In the end, the large quartz crystals filling the voids of silicified limestone are also corroded in the upper most horizons. TI1en, part of the silica released is pre cipitated in the lower horizons. This behaviour shows some similarities with pedogenic silcretes, which also are sinking progressively into the landscape by successive silica redistribution. This process causes a gradual loss of the trace elements contained in the silica material at the top of the profile, and is accom panied by an enrichment of titania in the upper clay
M. Thiry
122
materials and trace elements and iron oxides in the lower ones (Menillet 1987). As in the pedogenic silcretes, alteration of the clay minerals accompanied by silica deposits is the geo chemical mechanism triggering the development of secondary silica deposits. The reason may be an acidic alteration caused by oxidation of ferrous iron in hydromorphic horizons or in ground water, as has been suggested for quite different environments (van Breemen et al. 1984; McArthur et al. 1991; Grimaldi & Pedro 1996) but also observed in modern hydro morphic soils in western Europe (Brinkman et al. 1973).
TI1e lower part of this profile acts as the parent rock and as an argillaceous substratum for its upper part. This impermeable layer keeps the water from infiltrating from above and favours silica accumula tion and redistribution at the base of the profile. This may explain why the Argiles a Meulieres Formation becomes thicker towards the border of the lacustrine limestone formations, where the latter are thinner and more marly. A clay formation at the bottom of the lacustrine limestones (such as the Argiles Vertes de Romainville beneath the Calcaire de Brie) has a dual role: it favours the silicification of the limestone by groundwater processes and contributes to the for mation of a sealed and hydromorphic weathering complex. Where the Argiles a Meulieres rest directly on sandstone, however, such as the Fontainebleau Sand, the profiles contain only a few or no secondary silica deposits and the dissolution of the inherited silicified limestone is enhanced. Warm temperate climates with alternating wet and dry periods, together with particuliar geological and topographical situations, are likely causes of the meulierization process. S U M M A RY: E X T E N T A N D SIGNIFICANCE
OF THE
C O N T I N E NTA L S I L I C I F I C AT I O N F E AT U R E S
I n the Cenozoic continental formations o f the Paris Basin, silicification is widespread. As these forma tions have never been deeply buried, if at all, the various silicifications can only result from surficial or subsurface processes in a continental environment. There are two main types of silicification within a wide range of silicified facies: pedogenic and ground water silicification. These two types show entirely different features and distributions and have totally
different significance. The nature of the silicified facies must be described before they are interpreted in terms of palaeoenvironmental significance. Pedogenic silcretes are linked directly with the surface. They show numerous geopetal structures and well-differentiated features between the top and the base of the sections. There is always only one silicified horizon. Pedogenic silicification relates either to a relative accumulation of silica with leaching of the other elements or to absolute silica accumulation. It is important to be able to identify pedogenic sil cretes when reconstructing palaeoenvironments and palaeogeography in continental settings. Pedogenic silcretes are generally good markers of palaeosur faces. Moreover, they are signs of periods of palaeo landscape stability under alternating wet and dry climates. In contrast, groundwater silcretes are not linked directly with the surface and soils. They develop at depth in relation to groundwater pathways and water-table levels. Groundwater silicification fea tures are linked to absolute silica accumulation processes. They may develop in superimposed layers as a result of water-table fluctuations and/or down cutting of the landscape. From this standpoint they can be valuable guides to the geodynamic evolution of a basin. Quartzite lenses embedded in white sand facies, similar to those described in the Fontainebleau Sand, exist throughout all sand formations in the Paris Basin. In most of these formations, alteration and bleaching of glauconitic marine sands seem to have produced the pure white sand. All these deposits have the same structural layout as the Fontainebleau Sand. They are sandwiched between thick marl and limestone formations, and mark the main aquifers that discharge at the edge of the plateaux (Fig. 33). All of the silicification features were probably formed by groundwater flow and are related to the extensive erosion and stripping of the Tertiary deposits during the Pliocene and Quaternary Periods. Silicified limestone is also widespread in most, if not all, the Tertiary limestones in the Paris Basin. In some areas, as at the north-western edge of the basin (Picardie region), residual silicified limestone lying on the Chalk is the only remaining evidence of the initial extent of the former limestone for mations (Thiry et al. 1983). This silicified limestone may be related to groundwater outflow regimes similar to those of the quartzite lenses in the sand formations. As water pathways are more complicated
Diversity ofcontinental silicifications
and heterogeneous in the limestones owing to karst circulation, successive water-table positions within the limestone formations are more difficult to determine. Groundwater silicification features within the sand and limestone formations may relate to a single episode. For instance, at some points on the Hurepoix Plateaux, the dune ridges at the top of the Fontainebleau Sand have thick quartzite lenses, whereas the neighbouring limestone of the interdunal depressions is extensively silicified. The sand and limestone silicification may have developed simultaneously in relation with a water table that cut across the sand and limestone boundary. Silicified evaporites, especially replacement of gypsum crystals, are also common in the Paris Basin (Arbey 1980; Toulemon 1982). They are irregularly and discordantly distributed in relation to the stratig raphy, which suggests a post-depositional silicifi cation process. The silicification processes may not be different fundamentally from those found in the limestone and described above. The frequency and size of the groundwater silicification features in the Paris Basin appear remarkable. Groundwater silicification features may be as large in other basins but they have often gone unnoticed or even been misinterpreted. Our under standing of the frequency of such groundwater silicification features might be helped by remember ing that ground water on the continents is usually supersaturated in silica in respect to quartz. This means that ground water is always able to feed silicification if the conditions of quartz precipitation become favourable. These conditions depend on geo-
123
chemistry and crystal growth. Furthermore, high flow rates are necessary to provide the amount of silica that precipitates. The case of the secondary silica deposits of the Argiles a Meulieres is somewhat special. There is a great deal of highly soluble silica (poorly crystallized quartz or even opal) within the weathering complex that is inherited directly from the silicified limestone. These soluble silica phases may favour an increase of soluble silica content in the percolating water and cause some secondary silica deposits. One must remember that the interpretation and understanding of the various continental silicification features of the Paris Basin are only possible because the basin is relatively young and because its palaeo geographical framework is well known. The case pro bably would be different if these silicification features were buried beneath a thick sedimentary cover and only identified by boreholes. Most of the various silicification features probably would not be differentiated but lumped together and related to a single broad silicification event connected to the transgressive or unconformable palaeosurface. AC K N O W L E D G E M E N T S
I would like to acknowledge several individuals who have helped me in my research on silicification fea tures. Discussions with Regine Simon-Coin�fon and Jean-Michel Schmitt at the laboratory in Fontaine bleau, Fran�fois Menillet from BRGM at Orleans, Mohamed Ben Brahim from the University of Oujda (Morocco) and Tony Milnes at CSIRO, Adelaide,
ssw
NNE BEAUCE
BRIE
PARISIS
SOISSONNAIS PICARDIE
Loire Riv.
Seine Riv.
m
a.s.l.
1 50
50
kii-ilii'l
Tertiary limestone
0
main Tertiary sandstone formations
lllilll
main claystone fm.
KJ Chalk
Fig. 33. Schematic sketch across the superimposed plateaux of the Paris Basin and the successive sand formations, which all have bleached facies and contain quartzite lenses. All of them form main aquifers that discharge into valleys at the edge of the plateaux.
M. Thiry
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helped the emergence and the maturation of the ideas developed in this study. Each of them will rec ognize his own contribution.
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Spec. Pubis int. Ass. Sediment. (1999) 27, 129-15 1
Authigenic clay minerals i n continental evaporitic environments
J. P. C A LV O * , M . M . B LA N C -VA LLE R O N t , J. P. R O D R I G U E Z - A R A N D I A * , J . M . R O U C H Y t and M . E . S A N Z * *Departamento d e Petrologia y Geoquimica, Facultad Ciencias Geol6gicas, Universidad Complutense, 28040 Madrid, Spain; t Laboratoire de Geologie, Museum National d 'Histoire Naturelle, 43 rue Buffon, 75005 Paris, France
A B S T R AC T
A variety o f clay minerals are commonly found i n modern continental saline environments and clays also form a considerable portion of the sedimentary successions accumulated under evaporitic conditions in continental settings of the past. Research conducted world-wide for more than three decades indicates that specific clay assemblages form by authigenesis in such arid to semi-arid environments. Clay authigen esis, whether through direct precipitation from solution (neoformation) or by transformation of precursor minerals, mainly pyroclastics and detrital clays, takes place both in soil profiles and in various saline lake subenvironments. Examples of cation exchanges leading to authigenic clay minerals have been reported from many modern lakes and soils of North and South America, Africa and Australia. In most cases, significant removal of Mg+2 , enhanced by reaction with silica from the lake waters, results in the formation of magnesium silicates. Magnesium-smectites (stevensite, saponite, hectorite), kerolite and sepiolite have been determined as rather common authigenically formed minerals. In addition to these occurrences, palygorskite, either associated with sepiolite or as the predominant clay mineral, is present in many soils of arid to semi-arid regions. Minor occurrence of other clay minerals, e.g. corrensite, nontronite, Li- and Al-Mg-rich clays together with a variety of mixed-layer minerals, has also been reported from both modern and ancient continental saline settings. Illitization of precursor clay minerals is also thought to be a significant process of clay authigenesis in these environments. A review of clay occurrences from Tertiary evaporite successions in France, Spain, western USA, and Pleistocene formations of east Africa provides evidence that the aforementioned clay mineral assem blages, especially those dominated by Mg-rich clays, are also typical in sediments accumulated in ancient continental evaporitic settings. The comparison between basins developed in active tectonic systems, for instance the grabens related to the western European Rift System during the Paleogene, and basins evolv ing under relatively quiescent conditions, shows that in the latter case the formation of authigenic clays was favoured because of the lower sedimentation rates in the basins. There is usually a greater abundance of authigenic clays in lake-margin environments owing to large variations in salinity, pH and pC02 which can induce their formation. The sensitivity of clays to salinity fluctuations in the lake waters ma kes the stratigraphical progression of the clay mineralogies a useful indicator of lake-level fluctuation, probably related to climatic changes, in continental evaporitic settings.
INTRODUCTION
Continental saline deposits commonly are formed in closed basins of variable size located in arid and semi-arid areas. These may be subtropical high-pres sure belts, intracontinental arid regions, intramoun tainous desertic areas, as well as polar regions. Sedimentation in these settings is largely controlled by climate, tectonism arid physiography. According to
Handford (1991 ), these settings may be as different as tectonic basins, including fault-bounded basins (rifts, foreland, etc.) and intracratonic sags, interdune depressions, wind deflation hollows, floodplain areas, volcanic or meteorite impact craters, and even glacial valleys. These closed basins are characterized by unstable
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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Table 1. Summary of the main authigenic clay minerals occurring in continental evaporitic environments (chemical formula from Nickel & Nichols 1991 ; Mg/Si and Al/Si values for hectorite and palygorskite from data in Newman & Brown 1987) Mineral
Structural formula
Saponite Kerolite Stevensite Hectorite Sepiolite Palygorskite
hydrological conditions, so that small variations in cli matic or tectonic factors can induce large variations of lake water level and in the chemical composition of waters. Hardie (1984) distinguished two main types of brines derived from meteoric waters, i.e. alkaline brines (Na-K-C03-SOcCl) , poor in Ca and Mg, and neutral brines (Na-K-Mg-Ca-SOcCl) , poor in carbonate-bicarbonate, which are not differ ent from sea-water derived brines; hydrothermal inputs usually produce CaC12 -enriched brines, whereas volcanogenic waters are responsible for acid conditions. In addition, some basins can be connected to the sea temporarily, especially during periods of ocean-level rise. Climate changes can be responsible for rapid and high-amplitude variations of water level in small and shallow depressions, resulting in an alternation of flooding and desiccation stages. Evaporite deposition takes place usually during the drier periods characterized by drop in water level, which sometimes results in total desiccation of the lake area. By contrast, wetter periods induce lake highstands, which usually are recorded by biogenic carbonates on the edges of the lake basin and clay deposits in the basin centre ( Rouchy et al. 1993). Ter rigenous input by torrential fioodings is the common mechanism for the accumulation of a great volume of detrital clays in the lake areas. Accordingly, the most common sedimentary suc cession of continental saline environments is com posed of evaporitic deposits interbedded with clastic sediments, especially clays, which correspond mostly to episodes of non-evaporitic conditions. Neverthe less, the high ionic concentrations that prevail in these environments are considered to enable chemi cal reactions between minerals and water so that authigenic clay minerals may form (Jones & Bowser 1978; Jones 1986; Charnley 1989). For instance, some palygorskite deposits and mixed-layer clay corren site, commonly observed in Triassic evaporitic basins of Europe, have been envisaged as authigenic clays
Mg/Si
Al/Si
0.76-0.87 0.75 0.75 0.67 0.67 0.17-0.49
0.07-0.18 0.0 0.0 0.0 0.0 0.02-0.31
formed locally under evaporitic conditions (Jeans 1978) although corrensite is also considered as an indicator of burial diagenesis (Hillier 1993). More over, most of these clay deposits are not intercalated directly with the evaporite-bearing successions but occur lateral to them. This paper is focused on the relationships between clay minerals and evaporite sedimentation in order to summarize current under standing of the detrital and authigenic clays in conti nental evaporitic environments. Table 1 summarizes the composition of the most common authigenic clay minerals that will be referred to in this review of clays found in continental evaporitic environments.
MODERN SETTINGS General considerations
Hypersaline conditions at present are developed in two distinct hydrological and morphotectonic set tings. The first one is related to tectonically controlled basins. Rift valley settings are the most common, but the large orographic basins, such as the saline lakes of the Qaidam Basin in China (Spencer et al. 1990) and the Altiplano of Bolivia ( Risacher 1992), now remnants of larger and deeper lakes, Lake Urmia in Iran ( Kelts & Shahrabi 1986), the salars ofthe Chilean Atacama, etc., could be included in this first type. The second type corresponds to very fiat palaeoland scapes with either very shallow perennial or tempor ary lakes that have a great extent, e.g. Lake MacLeod (Logan 1987), Kara Boghaz Gol ( Strakhov 1970), Lake Chad ( Gac 1980), Great Salt Lake ( Spencer 1982), or are restricted to small depressions. In most of these latter evaporitic basins, runoff is limited, except for torrential discharges, and most of the water comes from subsurface seepage of groundwater, either fresh or saline. This pattern is especially well developed in central and southern parts of Australia,
Authigenic clay minerals
e.g. Lake Hopkins, Lake Neale, Lake Amadeus and Lake Eyre (Logan 1987; Jacobson et al. 1988). Millot (1964) was first in ascertaining the impor tance of clay authigenesis in saline lakes. As he stated, 'saline, alkaline lakes show a gradation from detrital sedimentation, in which the clays deposited are copies of the argillaceous stock of the rocks, the weathering, and the soils of the drainage basin, to chemical sedi mentation, the effect of this chemical treatment giving rise to transformations of the inherited clays'. The author concluded that 'in lakes with calcareous and siliceous alkaline chemical sedimentation indisputable argillaceous neoformations take place' (Millot 1970; p. 174). Detrital clays
Most of the clays occurring in continental environ ments are of detrital origin and supplied to the basins in periods of flooding or episodic discharges, the two main sources being either pre-existing mudstones or soils developed on rocks by weathering. The process of salt flocculation has been pointed out as an effective mechanism for the sedimentation of clay when freshwater flows into a saline lake (Sonnenfeld 1984; Hillier 1995). Thus the composi tion of the clays is not related initially to saline or hypersaline conditions. Millot (1970) and Charnley (1989) noted that clay associations in many modern saline lakes are detrital and simply reflect the mineral suites resulting from river discharge, rain wash, and aeolian transport, quite similar to the clay associa tions of most freshwater lakes. This is the case, among other examples, of many modern North American lakes and playas of the Mojave Desert. Such detrital clay suites can be used to characterize different catchment areas in a Jake basin, as exemplified by recent sedimentation in Lake Turkana, Kenya (Yuretich 1979). Illite, kaolinite, chlorite and diocta hedral aluminous smectite have been described as common detrital clay components of terrigenous mud in many modern lake basins. A problem arises where detrital clays are partially or fully inherited from arid soils, e.g. calcretes, 'paly cretes' (Singer 1979, 1984a,b; Rodas et al. 1994), or dry mudflats in which authigenic clays formed. The input of these reworked clays to the saline lake could lead to a misleading interpretation of the lacustrine clay assemblage. Except for very specific geological situations in which the soils or the authigenic clays of mudflats are the only source for clays, the reworked clays deposited in the lake consist of mixtures of clay
131
minerals and usually show textural features indica tive of degradation related to transport and al teration under the new hydrological conditions of the lake (Millot 1970; Hillier 1995). Authigenic clays in lake sediments
Distinction between detrital and authigenic clay min erals is often difficult. Authigenic clays, i.e. clays formed 'in situ' through direct precipitation from solution ('neoformation'), reaction of amorphous gels, or by transformation (usually aggradative) of some precursor mineral are unusual but they have been found in both arid soils and saline to hyper saline lakes (Singer 1984a; Jones 1986; Charnley 1989). Saline alkaline lakes in volcanic regions com monly show extensive formation of authigenic zeo lites (Surdam & Sheppard 1978) but formation of authigenic clays also has been documented in some of these lakes. In Lake Albert, Oregon, the compari son between the clay associations supplied from weathered pyroclastic rocks of the catchment area and the clay association observed in the lake sedi ments provides evidence that the clays extracted K, Mg and Si to form interstratified illite and a trioctahe dral Mg-rich smectite, the composition of which resembles that of stevensite (Jones & Weir 1983; Banfield et al. 1991). Illitization of smectite also has been observed in sediments of the lakes Albert and Manyara, east Africa, but there the aggradation occurred during a desiccation stage of the lakes, in relation with the formation of analcime (Singer & Stoffers 1980). In Great Salt Lake, Utah, Spencer (1982) and Spencer et al. (1984) found that the chemistry of the lake clays (< 0.25 f.!m) differs markedly from that of the inflow river clays; most notably the lake clays exhibit higher Mg/AI ratios. Variation of these Mg/AI ratios appeared to be concomitant with the trends observed in the Si/Al ratios of the lake clays, which show much higher proportions of both silica and magnesium than the river clays. Spencer (1982) con cluded that significant removal of Mg+Z, enhanced by reaction with H4Si04, resulted in the formation of authigenic magnesium silicate. Badaut & Risacher (1983) offered a case study for the authigenesis of Mg-smectite from replacement of biogenic silica in diatom frustules accumulated in Bolivian saline lakes. According to these authors, the authigenic process is unambiguous and develops under conditions of saturation with respect to amor phous silica, and pH above 8.2.
132
J. P Calvo et a!.
Recent work by Webster & Jones ( 1994) in Quaternary sediments from the Double Lakes Formation, Texas, reveals that sepiolite, interstratified Mg-smectite and palygorskite formed as authigenic clays in recent lake environments. The dominance of each of these minerals in several sediment horizons is thought to reflect saline shifts induced by evapora tion. Thus, sepiolite is assumed most likely to precipi tate in a brackish lake, whereas the dominance of interstratified Mg-smectite could reflect more saline conditions. Palygorskite is presumed to precipitate from pore-waters during relatively arid soil develop ment prevailing under playa conditions. These results are in agreement with previous assumptions by Jones (1986) that the formation of authigenic Mg-smectites, i.e. stevensite, requires higher saline conditions than those required for the formation of sepiolite in arid to semi-arid lake systems. Palygorskite seems to be a more common clay mineral in sediments of areas sur rounding the lakes or in stages of strong desiccation (playa stage) of the lake systems. Authigenic clays in lake-margin sediments
Many processes of clay authigenesis take place pref erentially in the marginal areas of saline lakes. Highly concentrated solutions can become ponded in small restricted depressions, e.g. interdune areas, isolated channels, small playas, and in the capillary fringe directly above emerging water tables. Cheverry (1974), Tardy et al. (1974), Maglione (1976), Carmouze et al. (1977) and Gac (1980) have reported authigenic magnesian-rich smectite (stevensite) in unconsolidated muds lying below the surficial saline sediments of interdune depressions of the northern border of Lake Chad. A comparison between the compositions of inflow waters of the Chari and Logone rivers and the strongly evaporated waters lying on the interdune depressions showed that the latter were depleted in Ca and Mg, which suggested formation of authigenic Mg-silicate, probably 'Mg rich montmorillonite', within the lacustrine mud (Gac et al. 1977). Stevensite was also recognized in unconsolidated muds below the surficial saline sedi ments (Gac 1980) and within small oolites accumu lated on the lake floor (Darragi & Tardy 1987). The thin sheets of stevensite in oolites are associated with aragonite. An Fe-rich smectite, nontronite, also has been found forming peloids in the recent sediments of the lake (Pedro et al. 1978). Besides these authigenic clay occurrences in shallow areas of the Lake Chad region, Hardie (1968)
found sepiolite associated with calcite, gypsum and dolomite in the playa margins of the Saline Valley, California. He considered this association to have been precipitated from evaporative concentration of ground water in the playa, except for dolomite which could be of detrital origin. A similar situation was rec ognized by Kautz & Porada (1976) at the margins of a saline pan in Kalahari, south-west Africa, where sepiolite clays form in relation to a near-surface aquifer. All these occurrences suggest that formation of authigenic smectites, sepiolite, and palygorskite takes place at the boundary between sedimentation and pedogenesis as subaqueous and subaerial environments alternate. Nevertheless, although for mation of interstratified Mg-smectites has been docu mented properly in several modern saline lakes and their occurrence has been supported by both experi mental and thermodynamic calculations, the investi gation of recent lacustrine sediments at Lake Chad and in modern saline lakes of the western USA has failed to positively identify sepiolite (Jones & Galan 1988). Charnley (1989) notes that the present-day lacustrine formation of fibrous clays remains highly questionable, although it was extensive in several geo logical periods (see Callen 1984 for a review). In our opinion this may be a distorted view because sepiolite develops preferentially in environments with low sedimentation rates which are difficult to recognize in modern settings; thus, modern sepiolite deposits may not be seen or may be considered as older. By contrast, the two chain-structure clays sepiolite and palygorskite have been recognized world-wide in recent soils of arid or semi-arid regions. Soils contain ing palygorskite, sepiolite or both minerals have been found in areas of the south-western USA (Bachman & Machette 1977; Hay & Wiggins 1980; Jones 1983; among others), South Africa (Watts 1980), the Middle East (Singer 1984b ) , Australia (Singer & Norrish 1974; Arakel et al. 1990), and the Mediter ranean region (Yaalon & Wieder 1976; Paquet 1983). The presence of these clays is commonly related to the formation of calcretes and/or dolocretes, which may be either phreatic or vadose. Arakel & McConchie (1982) and Arakel et al. (1990) have documented the occurrence of palygorskite, sepiolite and authigenic smectite within calcretes related to groundwater discharge in playa settings of inland Australia. Replacement of the carbonate and as sociated clays by silica in the calcretes usually was observed, which provides a case study for the origin of silcrete in such an arid environment. Dissolution of the host detrital grains and reprecipitation of silica,
Authigenic clay minerals
most pronounced in the transitional zone between phreatic and vadose calcretes, was suggested as a reliable mechanism for the development of silcrete (Arakel et al. 1990). The process whereby sepiolite and palygorskite are accumulated in soils of arid and semi-arid regions has been a matter of considerable debate (see review in Jones & Galan 1988). Jones (1983) concluded that sepiolite in calcic soils of south-western Nevada appears to have been precipitated from percolating waters that experienced strong evaporation. In that case, silica and solute magnesium were readily avail able from the carbonate and pyroclastic lithologies of the region. An alternative source of the magnesium for the formation of fibrous clays, through recrystal lization of high Mg-calcite to low Mg-calcite in cal cretes, has been proposed by Watts (1980). Contrary to the formation pattern suggested for sepiolite, the presence of palygorskite in calcic soils is thought to result from some kind of incongruent dissolution-precipitation process, which could explain the aluminium and small amounts of iron incorporated to the palygorskite (Jones & Galan 1988). The degradation of fibrous clays in soils when climate becomes more humid has been pointed out by Nahon & Ruellan (1975) and Bigham et al. (1980). Paquet & Millot (1972) concluded that the minerals are unstable and they weather to smectite when the mean rainfall exceeds 300mm. This assessment can be used as a solid basis for the palaeoclimatic and palaeoenvironmental inter pretation of palaeosols bearing sepiolite and/or palygorskite. A N C I E N T S ETTIN G S
In past continental saline environments, two types of basins can be distinguished. The first type comprises basins developed in active tectonic systems (rift, sub siding basins), characterized by relatively stable con ditions and high rates of sedimentary accumulation. In this setting, the most soluble saline facies, embrac ing a large variety of sulphate, chloride and evaporite carbonate, are commonly well-preserved because processes of synsedimentary dissolution or mechani cal reworking are not very active and the deposits are buried quickly. In these basins, detrital input derived from fine-grained sedimentation during periods of non-evaporitic conditions is usually important, and clay minerals are interbedded with the evaporitic
133
facies. The second type of basins comprises shallow lakes or lagoons that often become desiccated, lying on flat landscapes with low sedimentation rates. Calcium sulphate, interbedded with clayey laminated sediments, constitutes the predominant evaporite facies in the central part of the depressions, the clay minerals, whether massive and/or displaying subaerial exposure features, being more abundant in the margins of the lakes. Palaeogene of the western European Rift System and related basins
The continental western European Rift System cor responds to a complex set of individual grabens and smaller related basins of variable size and shape, the formation of which took place discontinuously throughout the Palaeogene (Fig. 1). Thick evaporite deposits, including halite and local potash deposits, accumulated in the deeper areas of some grabens, whereas the smaller depressions contain only thin evaporite beds that are composed of gypsum associ ated with dolomitic carbonates. These two types of basins are clearly differentiated by the nature of the clay mineral assemblages. The grabens:Alsace, Bresse, Valence, Camargue
In the 2700 m thick upper Eocene to lower Oligocene evaporite-bearing deposits of the Alsace rift system, clays are almost exclusively detrital (Sittler 1965; Blanc-Valleron 1990-1991). Palygorskite and sepio lite have been reported only in association with detri tal clay minerals in some marly interbeds within the potash deposits in the central part of the evaporitic basin. Unlike these rift deposits, palustrine carbon ates that formed on the pre-rift erosional surface were characterized by the presence of larger amounts of palygorskite, as in the metre-thick lacus trine Planorbis pseudoammonius limestone (Sittler 1965; Valleron et al. 1983). In the southern segments of the rift (Bresse, Valence, Camargue ), which contain thick sulphate and chloride deposits, the clay minerals are inherited and underwent no, or only few, transformations during their deposition (Moretto 1987; Dumas 1988; Fontes et al. 1996). The most soluble facies are pre served in the grabens, where the burial rate was high. In such basins, peripheral areas with low sedimenta tion rates providing enough time for aggradation, or even neoformation, of clay minerals were absent or little developed. This can explain the detrital
134
J. P Calvo et al.
D • 1\ U
Hercynian basement Main Tertiary volcanism Main alpine thrusts
Clays depos�s quoted in the text
P = Paris S = Salinelles M = Mormoiron
,..-.:;. ....,.-
Main Paleogene grabens
G) Upper Rhine ® Bresse CD Limagne @) Valence ® Ales ® Nimes -Camargue <J) Apt-Manosque-Forcalquier ® Gu� of lion ® Narbonne
character of the clay mineral assemblages in such environments. The small related basins
In Alsace, palustrine carbonates are scattered on the pre-rift erosional surface; these deposits are charac terized by the presence of variable amounts of palygorskite, up to 90% of the clay fraction in the metre-thick Lutetian Planorbis pseudoammonius limestone (Sittler 1965; Valleron et al. 1983). Then, the first rift deposits of upper Eocene to lower Oligocene
• Exte::: Paleogene evaporites v
D •
Scattered gypsum
deposits Main hal�e
depos� Main potash depos�s
Fig. l. Schematic map showing the distribution of the main Palaeogene continental evaporite settings in France (adapted from Blanc Valleron 1990-1991; Maerten & Seranne 1995; Rouchy 1997).
age correspond to the 2700 m thick evaporite-bearing deposits where clays are almost exclusively detrital (Sittler 1965; Blanc-Valleron 1990-1991 ). Fibrous clays have been quoted in some marly interbeds within the potash deposits: mainly traces of paly gorskite, except for two samples coming from the centre of the evaporitic basin and containing, respec tively, 10% palygorskite and 20% sepiolite. Paly gorskite can be inherited from Lutetian deposits, but the presence of sepiolite, never quoted in older for mations, argues for an authigenic origin of these fibrous minerals. Nevertheless, the scarcity of their
Authigenic clay minerals
occurrence implies that authigenesis is negligible as compared with inheritance processes. The Mormoiron Basin is one of these small depres sions located east of the Rhone valley, on the margin of the subalpine range, where continental conditions started during the Cenomanian and lasted through the Palaeogene. Triat & Trauth ( 1972), Trauth (1977) and True (1978) provided a detailed knowledge of the sedimentary processes that occurred in this basin. The sedimentary succession starts with continental deposits composed mainly of reworked weathering profiles (laterites), and palustrine deposits associated with attapulgite-rich calcretes (Jocas Limestone, middle Eocene). Then, true lacustrine conditions developed in response to an extensional tectonic phase leading to the deposition of a sequence com posed of sandy marls and argillites grading upward into gypsum and carbonates, both limestone and dolomite. This sequence is characterized by a clay mineral trend ending with the development of Mg-rich clay minerals: Al-Fe smectite (wyoming) ---7 Al-Mg smectite (cheto) ---7 AI saponite ---7 Mg saponite ---7 stevensite and sepiolite (Triat & Trauth 1972;Trauth 1977). The formation of these magnesian clays, which pre-date the deposition of the gypsum and associated dolomites is considered to result from processes of authigenesis in highly concentrated conditions. The Sommieres Basin, located west of the Rhone valley on the south-western margin of the Massif Central, is a complex depression bounded by two large fault bundles (N!mes and Cevennes faults). The Oligocene lacustrine deposits (clastic sediments and carbonates) include a sepiolite-rich interval mainly composed of carbonates (limestones and dolomites) with claystone interbeds (Salinelles Limestone). The clay mineral assemblage comprises Al-Mg smectites ( cheto ), palygorskite, stevensite and sepiolite (Trauth 1977). As inferred from the dolomite facies, similar to those found in evaporite settings, the sepiolite is interpreted as resulting from precipitation at the final stage of evaporative concentration of the brines, which occurred during the desiccation of peripheral depressions surrounding the main lake (Trauth 1977). Such processes may be compared with those taking place in the peripheral areas of Lake Chad or Saline valley. Nevertheless, the authigenic processes cannot explain satisfactorily the formation of sepiolite-rich layers of claystones, which probably resulted from resedimentation in deeper parts of the lake of sepiolite that formed initially on the marginal areas of the lake.
135
Paris Basin
Middle Eocene to Oligocene formations of the Paris Basin provide a good example of evaporitic depo sition in dominantly continental settings, which resulted from the restriction of a shallow epiconti nental sea. The 'Masses du Gypse' Formation, up to 35 m thick, was deposited in the centre of the Paris Basin during the Marinesian and Ludian. During the Marinesian (late middle Eocene), fibrous clay minerals formed in interdunes areas on the borders of a marine seaway. Both the characteris tics of the palaeolandscapes and the sedimentary processes that occurred in the peripheral area have been described by Blanc-Valleron & Thiry (1993) (Fig. 2). The Marinesian deposits, consisting of clayey lenses and more regular clay layers associated with sands, lie within depressions enclosed between dunes comprising the Auversian Sands, which were formed soon after the withdrawal of the sea. The pure clay lenses are composed of sepiolite and smectite. The sandy layers include smectite, interstratified illite-smectite and traces of illite and kaolinite. The more regular clayey layers, which overlap the dunes, contain palygorskite, smectite and illite. The sedimen tary pattern recognized in these deposits can be compared with that observed at present along the northern border of Lake Chad, with the highest solu tion concentration taking place in the interdune clay pans and ponds where ground water is discharging. In these areas, which are sheltered from detrital input and are subject to strong evaporation, authigenic clays, namely sepiolite form. During the Ludian (upper Eocene), thick gypsum layers, for which the origin (marine or continental?) of the brines from which they precipitated is still discussed, were deposited as the 'Masses du Gypse' in the centre of the basin while marls and lacustrine limestones occupied the peripheral areas. Silcretes and calcretes, developed previously around the sedimentary basins, left a deep imprint on the landscapes. These were later preserved from erosion by the siliceous and calcareous armoured surfaces, which protected the underlying sandy and clayey formations. Thus, chemical deposits were not contaminated by detrital input. Landscapes can be compared with the present-day wide lime stone plateaus of the hammadas in northern Africa as well as the silcrete-capped plateaus of inland Australia. The clay minerals form three main assemblages (Trauth 1977) (Fig. 3):
J P Calvo et al.
136
� lacuslrine limestone
0 marine and eolian sand
. KAOLINITE [22d iLLITE � ILLITE-SMECTITE
ferruginous sand
� SMECTITE Q PALYGORSKITE
humus-rich sand with palaeosol structures
Q sEPIOLITE
·�1,;, • • · · � ·
� ·
· · Sm ·
Fig. 2. Sketch regarding the palaeogeography of the Paris Basin during the Marinesian, cross-section from Mont Saint-Martin (after Blanc-Valleron & Thiry 1993).
50 km _,// /
/'� v
/(,--y---9-, v
v
',, v
......... ....
v.· ·
;;- : . � v' ; � v :
· -v·: : · _:::.;,:
: ·:v- ·:· · v:·.
_ _�
v-·
fl:-:..., :::-., · %�""-i'��..-
marl
EYJ G
Fig. 3. Sketch showing the main clay mineral assemblages occurring in the Ludian formations of the Paris Basin (after Trauth 1977).
Authigenic clay minerals 1 palygorskite and Al-Mg smectites, which are the most common clays; 2 the Al-Fe smectites present along the southern margin of the basin; 3 the sepiolite and Mg-smectites that occur along the south-western margin of the gypsum deposit. The dis tribution of these clay assemblages does not coincide with the palaeogeographical zones defined by facies and bulk rock mineralogy. Thus, the magnesian clays are not always associated with dolomite; the sepiolite being related to calcite-rich layers, whereas dolomite layers usually contain Al-Mg clay minerals (smec tites and palygorskite). This can be explained by the fact that clay minerals were in competition with carbonates to incorporate magnesium. Silica con trolled the process: when silica was absent, magne sium was incorporated into carbonates and dolomite was formed. Alternatively, when silica was available, sepiolite and Mg-smectites developed by extracting magnesium and so calcite was precipitated together with magnesian silicates. Distribution of sepiolite and Mg smectite along the southern margin of the basin probably is related to silica-rich water inflows, origi nating from the flint-bearing chalk that formed the local bedrock.
Fig. 4. Geological map of the Iberian Peninsula showing the location of the main Tertiary basins, with indication of the continental basins cited in text.
137
Spanish Tertiary basins
An extensive development of continental evaporite basins occurred in Spain during the Tertiary (see Friend & Dabrio 1996 for a review). The three largest Tertiary basins, Madrid, Ebro and Duero basins, located in the interior of the Iberian Peninsula, contain thick evaporite deposits compris ing both gypsum and other more soluble salts, i.e. halite, glauberite and thenardite, which commonly are associated with mudstone, marlstone and carbonate facies (Corrochano & Armenteros 1989; Ortf & Salvany 1990; Salvany et al. 1994; Calvo et al. 1 996). In addition to these three major basins, thick continental evaporite formations also are present in Calatayud, Teruel and Guadix-Baza (Fig. 4).
Authigenic clay deposits are especially well developed in palaeosols and in the margins of the saline lake systems. A considerable amount of paly gorskite was accumulated in soil profiles in both the northern and southern parts of the Madrid (Tajo) Basin as well as in the southern Duero Basin during the Palaeogene. These palygorskite-rich palaeosols (palygorskite often constitutes 100% of the clay
- Siliceous terrain (Mainly Palaeozoic) CJ Carbonate terrain (Mainly Mesozoic) D Tertiary Basins
J. P. Calvo et al.
138
of the basins. The contribution, through erosion, of the palygorskite facies to the nearby lake environ ment has not yet been estimated. The middle Miocene formations of the Madrid Basin contain the most complete clay minerals asso ciation recognized in fossil evaporite lake systems of Spain. During the middle Miocene, the distribution of continental sediments in the basin shows a concentric facies arrangement (Fig. 5). Fibrous clay minerals, both sepiolite and palygorskite, occur closely related to Mg-smectites (stevensite, saponite, kerolite stevensite mixed-layers) across a transition zone from distal alluvial fan facies into lake-margin envi ronments (Fig. 6). In this setting, the presence of paly gorskite is restricted mainly to palaeosols, whether developed in arkosic distal fans or in the lake margins in periods of net subaerial exposure (Leguey et al.
fraction within the soil) have been referred to as 'palycretes' by Rodas et al. (1994). The palaeosols consist of indurated sediments that are cemented by secondary palygorskite and which display morpho logical features and structures similar to calcrete, i.e. pellets and pisoliths, sepic and striated fabrics (Brewer 1964; Hay & Wiggins 1980), as well as fea tures related to the growth of rhizoliths. Other paly gorskite-rich palaeosols have been recognized in the Miocene of the southern Duero Basin (Bercimuel deposit). Therein, the palygorskite has been inter preted as an authigenic transformation product from detrital mica (Suarez et al. 1994). Both the Palaeo gene (upper Eocene?) and Miocene examples of palygorskite-rich palaeosols formed in distal alluvial fan facies around saline lake systems and record a change towards more arid conditions in the evolution
0
teJI I � :: I I ' I"-""I X :x
30Km
ALLUVIAL FAN !Arkose) MUDFLAT-LACUSTRIN E MARGIN (Clay) LAKE/POND (Carbonate) SALINE LAKE (Gypsum)
Fig. 5. Sketch of depositional systems in the Madrid Basin during the middle Miocene. Bar on the sketch indicates the location of the profile represented in Fig. 6.
139
Authigenic clay minerals Lake margin
Open lake
Alluvial
�
Groundwater e
�
�
Carbonate facies
Clay mineralogy Carbonate mineralogy
Mudflats
t
t
t
t
t
===i :: ----= ========;: :: = :: :====:: ===;�::::� :: G z;::z:;� g;-----;1;:-y; � �;: ----l� <> <>
:::s:::::s::
----=
=--
== �
-
Evaporation
* Oscillation of
lake water level
Massive bedded Calcretes/ Nodular Nodular Stromatolites, Bedded limestones and dolostones, dolostones dolocretes limestones dolostones tepees gypsum molds Mg-smectites . . . Polygorsk1te/ Sep10l1te (stevensite, saponite) Illite sepiolite Kerolite/stevensite Calcite/ Calcite/dolomite Dolomite dolomite Calcite Dolomite
Fig. 6. Idealized sketch of marginal to open lake environments in the Madrid Basin during the middle Miocene. The distribution of clay minerals throughout the profiles corresponds exclusively to authigenic clays, which commonly are associated with carbonate facies (modified after Calvo et al. 1995a).
Fig. 7. Outcrop view of middle Miocene sepiolite deposits from the Vicalvaro mine, near Madrid. TI1e sepiolite beds, shown as white colour, are intercalated light brown clay with thin silicified carbonate interbeds. Quaternary arkosic deposits overlie the sepiolite beds disconformably, the contact being marked by a dark horizon which locally stains the sepiolite. The height of the section is about 6 m.
1985; Bellanca et at. 1992) . In addition to its presence
in palaeosol profiles, sepiolite forms huge economic deposits along the western and northern parts of the basin ( Galan & Castillo 1984; Doval et al. 1986) (Fig. 7) . Some carbonate, mainly dolomite, and trioctahe dral Mg-smectites are intercalated locally within the sepiolite deposits. This association is interpreted as deposition in ponds and marshes that extended on to the mudflat areas of the lakes ( Galan & Castillo 1984; Calvo et al. 1986, 1989; Doval et al. 1986; Ord6iiez et al. 1991). Within the ponds, the sepiolite could be formed by precipitation from silica-bearing waters, derived from granitic uplands, that were mixed with
brackish or saline Mg-enriched waters from the nearby lake. In this setting, the continuous ground water recharge into previously formed authigenic clay minerals and carbonates could also contribute to new diagenetically formed clay minerals, i.e. sepiolite, which are found as cavity-lining and/or fissure-fill clays ( Leguey et al. 1995). As aforementioned, the magnesian smectites of the Madrid Basin display a varied mineralogy, in cluding stevensite, saponite, and kerolite-stevensite mixed-layers. These clay deposits occur in massive, metre-thick beds alternating with carbonates, mainly dolomite ( Fig. 8), which are partially silicified ( Calvo
140
J. P Calvo et al.
Fig. 8. Close-up view of a clay-carbonate sequence containing Mg-smectite and kerolite-stevensite mixed· layers from middle Miocene formations of the Madrid Basin. From bottom to top, the sequence is formed of greenish clays (A), pink clays (B) and nodular carbonate, mainly dolomite, affected by root bioturbation (C). Length of the trowel for scale, 23 em.
et al. 1995a,b ). The clay beds exhibit typical greenish
grey to cream colours except for some pink-coloured interbeds which consist mainly of kerolite-stevensite mixed-layers with minor stevensite, sepiolite, illite, quartz and calcite (Martfn de Vidales et al. 1991; Pozo et al. 1994). The textural characteristics of some of these clays are shown in Fig. 9. The clay mineral asso ciation recognized in the Miocene of the Madrid Basin is interpreted to have been deposited in a mar ginal saline lake environment, which grades basin wards into calcium sulphate and probably sodium sulphate deposits (Calvo et al. 1989; Garcia et al. 1990; Ord6fiez et al. 1991). These lake-margin sediments extend along a narrow fringe, about 1-2km in width, between the cities of Madrid and Toledo, flanking distal alluvial fan facies at the toe of an arkosic allu-
vium. In this setting, the Mg-smectites formed by authigenesis, mainly by transformation of detrital dioctahedral smectites but also by clay neoformation locally (Doval et al. 1986; Pozo et al. 1994). The formation of the pink clays, dominated by kerolite-stevensite mixed-layers, merits some comment because they provide evidence about the changing hydrochemical conditions in the marginal saline lake environment. Thus, the pink clay layers, which display soil features such as root bioturbation and mottling, occur towards the top of the greenish Mg-smectite beds, with a gradual transition between the two lithofacies that suggests a progressive shift from 'normal' marginal lake sedimentation to palustrine conditions with intervening pedogenic processes. According to this pattern, the kerolite stevensite mixed-layers represent periods of in creased salinity and pH rise compared with those in which the underlying Mg-smectites formed. Scat tered sepiolite occurrences within the pink clays could be indicative of episodic drops in water salinity (Pozo et al. 1994). The influence of lake-level changes in the clay min eralogy of middle Miocene sediments of the Madrid Basin has been studied in detail by Bellanca et al. (1992). Lowstand lake periods are marked by a sedi mentary facies association comprising nodular car bonates and marls, massive dolostones, bioclastic limestones, mudstones and chert, which partially replaces these deposits resulting in the formation of silcrete. The mudstones consist of smectite, both di and trioctahedral types, sepiolite, palygorskite (Fig. 9d) and illite, thus characterizing a very shallow lake margin subjected to episodic subaerial exposure. In contrast, the overlying deposits comprise mainly diatomaceous marls, laminated bioclastic limestones and marls, massive and intraclast bearing mudstones, and basal conglomerates. Bellanca et al. (1992) inter preted this facies association as accumulating during a highstand lake period. The abundant sepiolite and palygorskite intraclasts recognized in these deposits provide evidence that reworking of authigenic clays must be envisaged as a very effective mechanism in such a depositional environment. A similar arrangement of authigenic lithofacies has been recognized in Miocene formations of the Duero Basin, north-central Spain (Fig. 10). Armenteros et al. (1995) found that in this basin smectite, palygorskite and sepiolite are widely distributed in the transition zone from distal alluvial fan, mudflat and saline mudflat-ephemeral lake environments, the most intense transformation of the sediments having taken
Authigenic clay minerals
141
(a)
(b)
(d)
(c) Fig. 9. Scanning electron microscopy (SEM) photomicrographs from authigenic clay minerals occurring within lake-margin sequences in middle to upper Miocene formations of the Madrid Basin. (a) Saponite forming the bulk of some clay beds in Yuncos, Toledo province; bar = 1 !-LID. (b) Kerolite-stevensite mixed-layers displaying a honeycomb structure, Esquivias, Toledo province; bar= 1 !-LID. (c) Aggregate of long sepiolite fibres cementing a void within Mg-smectite-rich beds from Esquivias; bar= 1 !-lll1. (d) Long fibres of palygorskite associated with calcite crystals in palaeosols developed on a saline lake margin, Malcovadeso area, Toledo province; bar = 1 !-LID. (Photomicrographs courtesy of Dr M. Pozo.)
place in the mudflat areas. Palygorskite is the pre dominant clay mineral in mudflats and its formation was related mainly to the development of soils of ver tisol type. A variety of palaeosols, embracing calcrete, dolocrete and silcrete, formed in this setting, which together with the occurrence of the clays seems to be a common lithofacies assemblage in ancient arid to semi-arid lake environments (see also Wright & Sandler 1994). The saline facies in the Duero Basin are mainly in the form of gypsum crusts, which cycli cally alternate with sepiolite and Mg-Al smectite that also display pedogenideatures.
Palygorskite is also found as the predominant clay mineral in floodplain, mudflat and playa-lake envi ronments of the continental Palaeogene of the Ebro Basin (Fig. 1 1 ) (Ingles & Anad6n 1991). In the case study, palygorskite possesses various clay morpholo gies and fabrics, i.e. short fibres about 1 11m long which commonly form fibrous borders of platy detrital clays in sediments deposited in floodplain and mudflat environments; in contrast, tree-like palygorskite fibres up to 10 11m long occur mainly in playa-lake deposits. According to these features and the distri bution of clay mineral associations, Ingles & Anad6n
PALAEOZOIC ROCKS (MOSTlY METAMORPHC)
A
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NEOGENE (lACUSTJIINE)
IGNEOUS ROCKS
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-
CLAYS
-
-
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Mg - A I Smectite Pol yoorsk i te S e p i olite C A L C RE T E S DDL O C RETES S I L CRETES G Y P S U IOI C R U S T S
-
-
-
B
A UT H I GE N I C C L A Y S A I - M Q Smectite
-
-
-
-
-
Fig. IO. (A) Geological map of the Duero Basin, northern Spain, with indication (see bar) of the location of the palaeoenvironmental sketch shown in B. (B) Sedimentological model for Neogene lacustrine deposits in the eastern part of the Duero Basin showing the distribution of both detrital and authigenic clays as well as main pedogenic deposits throughout distinct sedimentary environments (1, proximal and middle alluvial fan; 2, distal alluvial fan; 3, dry mudflat; 4, saline mudflat ephemeral lake) (modified after Armenteros et a!. 1995).
Authigenic clay minerals
PLAYA-LAKE
8
MUD F L AT
SHALLOW LAKE
ALLUVIAL
143
Calatayud Basin (Arauzo et al. 1991) and the volcanic region of Campos de Calatrava (Pozo & Martin de Vidales 1989; Galan et al. 1994). Sepiolite deposits in Calatayud extend into the marginal areas of the basin, the central parts being formed by an alternation of mudstone, gypsum and glauberite (Na2Ca(SO4)2) deposits. This location of the sepiolite beds suggests that they formed in shallow lakes developed on shadow zones between alluvial fans; in this setting, strong evaporative concentration of Si and Mg-rich waters resulted in the formation of the authigenic clay minerals. In the Campos de Calatrava volcanic region, Plio Pleistocene lake sediments rich in palygorskite, dioc tahedral smectite, illite and occasional sepiolite occur intercalated with dolomite and pyroclastic materials. According to Pozo & Martin de Vidales (1989), sepio lite formed during periods of relatively high aridity in the basin and its formation probably was related to a hydrated amorphous magnesian silicate precursor rather than direct precipitation from solution. Paly gorskite in these lake sediments of the Campos de Calatrava region has been interpreted to be formed by transformation of precursor clays, namely diocta hedral smectite and illite in a highly Si- and Mg-active environment (Galan et al. 1994). Pozo & Martin de Vidales (1989) suggested that the transformation process could be achieved effectively by partial dedolomitization of the associated carbonates.
DEPOS ITION
Fig. ll. (A) Geological map of the eastern part of the Ebro Basin, with indication (see bar) of the location of the palaeoenvironmental sketch shown in B. (B) Idealized sketch for the environmental distribution of clay minerals in the south-east Ebro Basin during the Early Bartonian (P: palygorskite; I: illite; Ch: chlorite; K: kaolinite; S: smectite) (modified after Ingles & Anad6n 1991).
(1991) suggested that palygorskite may form from
transformation of pre-existing clay minerals (smec tite, chlorite-smectite mixed-layers and illite) and in situ neoformation from Si- and Mg-rich waters of the ephemeral lake. The results obtained from the Eocene deposits of the Ebro Basin are relevant because they provide evidence that the formation of palygorskite in semi-arid saline lake settings may be accomplished according to different mechanisms that ultimately are related to specific environments. Other occurrences of fibrous clay minerals in Tertiary basins of Spain comprise those of sepiolite deposits in middle Miocene formations of the
USA examples
The formation of authigenic clay minerals in ancient saline lake settings has been reported mainly from two geological locations in western USA: the Eocene Green River Formation (Bradley & Fahey 1962; Dyni 1976; Tettenhorst & Moore 1978) and the Plio Pleistocene deposits of the Amargosa Desert (Fig. 12) (Papke 1972; Eberl et al. 1 982; Khoury et al. 1982; Hay et al. 1986). First reported by Bradley & Fahey (1962), steven site occurs in the saline facies of the Wilkins Peak Member of the Eocene Green River Formation in south-west Wyoming. Additional insights by Dyni (1976) indicate that authigenic trioctahedral smec tite, probably hectorite, is widespread within clay stone and oil-shale beds from the Parachute Creek Member of this formation in the south-western part of the Uinta Basin, Utah. The authigenic clays are associated mainly with dolomite in the claystone beds whereas calcite is more dominant in the oil shale, thus suggesting that the interstitial waters of
J. P Calvo et al.
144 0
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Fig. ll. Geological sketch of the Amargosa Desert, southern Nevada, showing the location of the Amargosa Flat and modern playa environment (modified after Hay eta!. 1 986).
the claystones during deposition and (or) lithification were more saline than those of the oil-shales. Accord ing to the facies distribution and clay contents of the various lake facies, Dyni (1976) concluded that authi genic trioctahedral smectite formed in a magnesium rich, nearshore lacustrine environment and the clay
associations clearly reflect a zonation from basin edge to basin-centre. Besides this occurrence of authigenic trioctahedral smectites, Tettenhorst & Moore (1978) have described unusual deposits formed of stevensite oolites in the Green River Formation of central Utah. The oolites were inter preted as having been precipitated in place from the lake water, thus rejecting a detrital or transforma tion origin from tuffs. Darragi & Tardy (1987) have reported a possible analogue for the stevensite oolites of the Green River in present saline lakes of Chad. The Amargosa Desert, located at the California Nevada boundary (Fig. 12), contains economic deposits of sepiolite and Mg-smectite that formed mainly during the Pliocene (Hay et al. 1986). The area is surrounded by Palaeozoic limestone and dolomite relief as well as Tertiary clastic sediments and tuffs. The Pliocene clay deposits are associated with varied carbonate lithofacies that were deposited in playas, marshland, ponds and floodplains, the whole being fed by springs with a water chemistry that was prob ably not much different from that of modern springs (Khoury et al. 1982). Brecciated limestone and dolomite seepage mounds are present along zones of groundwater leakage within the ancient playas (Hay et al. 1986; Calvo et al. 1995b) (Fig. 13). Both clay and carbonate lithofacies exhibit palustrine features such as root marks and desiccation cracks indicative of a very shallow lake environment (Fig. 14). Although no highly evaporitic phases, except for some halite traces (Khoury et al. 1982), have been found in the Amargosa Desert, saline, alkaline water lake envi-
Fig. 13. Outcrop view of a seepage mound deposit developed within Pliocene lacustrine sequences from the Amargosa Desert, Carson Slough locality. The white, massive carbonate, mainly formed of dolomite with variable amount of calcite and silica, deforms the overlying, well-laminated clay deposits consisting mainly of Mg smectite. Mound in the photograph is about 3 m high and 5 m wide.
Authigenic clay minerals
145
Fig. 14. Close-up view of sepiolite and Mg-smectite deposits alternating with nodular carbonates of palustrine origin. Moretti Mine, Amargosa Desert, Pliocene.
ronment with episodic climatic oscillation has been suggested for that area throughout the Pliocene (Hay et al. 1986). Changing evaporative conditions could explain the occurrences of sepiolite and mixed-layer kerolite-stevensite, the latter requiring more saline water and higher pH than that required to precipitate sepiolite (Khoury et al. 1 982; Jones 1986). Pleistocene formations related to East African lakes
Since the pioneering works on the Pleistocene and recent formations of the East African lakes (see Frostick et al. 1986), considerable attention has been devoted to clay-rich deposits occurring in these areas. Many of the clay deposits have been exploited eco nomically for several manufacturing purposes, such as pottery and brick-making, decolorizing, and as components for other industries (paper, rubber, fer tilizers, etc.) (Tiercelin 1991). Large deposits of bentonite and sepiolite are mined in Pleistocene for mations, specially those located in the Amboseli Basin, on the Tanzania-Kenya boundary. The extracted bentonite is used for oil-well drilling and foundry sands whereas pure white sepiolite is manu factured as smoking pipes. In Amboseli, the magne sian clays are associated with carbonates (both calcite and dolomite) and marls that belong to the Sinya Beds, a formation of early to middle Pleis tocene age that was deposited in a semi-arid lake basin (Stoessell & Hay 1978; Hay & Stoessel! 1984; Hay et al. 1995).The clay mineral assemblage of these beds is dominated by sepiolite and mixed-layered
kerolite-stevensite, which form veins and cavities within the carbonate. The Amboseli deposits constitute a good case study of the clay-phase relationships between kero lite and sepiolite. Hay et al. (1995) have carried out a detailed analysis of the influence of salinity on the formation of the kerolite-smectite (Ke-St) mixed layers. Based on the ()180 values obtained from these clays, Hay et al. (1995) point out that high salinities favour a high content of stevensite in the Ke-St whereas the kerolite-rich Ke-St formed under lower salinity conditions. Both kerolite-smectite types were chemically precipitated from Si02-rich and Mg2+-rich lake and ground water. In addition to these common kerolite-smectite types, an Al-rich Ke-St mixed-layer also has been recognized in Amboseli. A probable genetic relation between this clay mineral and detrital clays has been suggested (Hay et al. 1995). DISCUSSION AND
C O N C LU S I O N S
There is basic agreement that most o f the clays found in continental evaporite formations are of detrital origin, thus faithfully reflecting the clay composition of older argillaceous formations of the palaeo drainage areas, the products of pedogenic weather ing, or both. Illite, kaolinite, chlorite, dioctahedral smectite and a number of mixed-layer clays have been recognized as common detrital clay minerals in the evaporite formations. The same situation has
146
J P Calvo et al.
been recognized in many modern settings, leading to the assumption that the clay assemblages in saline lakes do not show significant differences from those of freshwater lakes. The investigation of clay minerals in various recent lakes, however, whether alkaline (water with high pH values and enriched in car bonate and alkaline earths) or saline (lower pH, S04 /Cl - brines), throughout the world has demon strated that the formation of authigenic clays is a rather common process whereby detrital clays or other highly reactive substances, especially volcan oclastic deposits, are altered into new mineral phases. Mass-balance calculations carried out in recent lakes, such as Lake Abert, Great Salt Lake, Lake Chad and others (see references above), indicate that solute loss of K, Mg and Si contributes to the formation of the authigenic clays, whether through transformation of pre-existing clays ('transformation by addition' of Millot 1964) or by direct precipitation from the saline solution. In most of the cases, the resulting clays are Mg-rich clays, such as stevensite, saponite and sepio lite, with fewer occurrences of rectorite and hectorite. In addition, palygorskite, an Al-Mg fibrous clay, and kerolite-smectite mixed-layers have been found widely distributed in both recent and ancient con tinental saline environments. The following para graphs discuss several aspects concerning the formation of authigenic clay minerals in evaporitic continental environments. Clastic sedimentation rates
The magnitude of clastic sedimentation rates, i.e. ter rigenous input, within the saline environments is assumed to be a critical factor for the formation of the authigenic clay minerals, as provided by several case studies, particularly from ancient evaporite for mations. The western European Rift System basins that developed in France during the Palaeogene offer a good example of this situation. The thick evaporite successions accumulated within the more rapidly subsiding basins are characterized by intercalated clay deposits of almost exclusively detrital origin, whereas in the smaller related depressions (Mormoiron, Sommieres ) , which had lower sedimen tation rates, a wide assemblage of Mg-rich clays asso ciated with gypsum and carbonate is recorded. Similarly, the sedimentary stages of reduced deposi tion in the larger basins, e.g. initial rift stage, are also characterized by the development of authigenic clay assemblages. This supports the observation that in
both recent and ancient continental evaporite set tings, the authigenesis of clay minerals is favoured mostly in the marginal areas (interdunal depressions, peripheral marshes, muddy or carbonate fiats) of the saline lakes. In these areas, transformation of precur sor clays fed by episodic discharge into the lake envi ronment is very effective. Highly reactive conditions are reached in this setting because of the large varia tion in salinity and other factors such as pH and pC02 . Figure 15 shows an idealized sketch of the various environments and subenvironments in which the formation of authigenic clays in saline settings has been reported to take place. A comparison of the different patterns of lake basin evolution, including sedimentation rates as a main factor involved in the formation of these clays is also represented. Pedogenic processes
The commonly observed extensive development of soils in the margins of saline lakes provides evidence that, in these areas, sedimentation rates are low but also episodic. Pedogenic processes account also for the formation of new clay phases, particularly paly gorskite and sepiolite, the occurrence of which has been widely reported in relation to calcretes, dolocretes and silcretes (Singer 1979; Jones & Galan 1988; Armenteros et al. 1995). In some cases, paly gorskite is the only clay mineral present in these soils, giving rise to a new pedogenic term ('palycrete') (Rodas et al. 1994). Whether subordinate to paly gorskite or as the predominant mineral, sepiolite usually has been found in palaeosols developed under arid to semi-arid saline conditions. It is com monly accepted that, in contrast with palygorskite, sepiolite accumulates as a direct precipitation product within the soil profile (Watts 1980). Exam ples from Kalahari calcretes described by this author suggest that sepiolite constitutes a late-stage mineral after the early formation of palygorskite, which extracts aluminium from the environment and increases the amount of magnesium available for the precipitation of sepiolite. The occurrence of diagen etic zeolite, associated with both authigenic illite and smectite, in calcretes developed in margins of some saline, alkaline lakes of East Africa has been described recently by Renaut (1993). In this setting, the authigenic clay minerals within the palaeosol could be interpreted as a by-product of diagenesis after complex reaction of detrital silicates with Na-rich interstitial brines.
147
Authigenic clay minerals LOW S U B S I D E N C E I LOW B A S E M E NT+-
S E DI M E NTATION R ATES
ALLUVIAL FAN & LAKE MARGIN ENVIRONMENT
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( stttvttnsiftt, saponiftt "- . .j_ 0_ig!!.tt!_ �I'!!!�ondifions)
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_
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DRAINAGE AIIIEAS {OFTEN VOLCANIC ROCKS) SOILS a PALIEOSOLS COARSE: CLASTIC SEDIMENTS
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Fig. IS. Idealized scheme showing the most commonly observed environmental distribution of clay minerals in continental saline settings. The sketch underlines differentiation between basins undergoing distinct subsidence conditions, which result in contrasted sedimentation rates. Processes of authigenic clay minerals formation take place mainly in saline lake-margin environments; furthermore, a larger clay mineral assemblage is found in this environment. The formation of authigenic clay minerals in saline open lake environments is highly dependent upon the variation between increasing and decreasing saline conditions in lake waters.
Groundwater discharge in lake-margin environments
The transitional zone between areas of major terrige nous accumulation, namely alluvial fan systems, and the lake is also largely influenced by groundwater discharge of contrasting hydrochemical composi tions, which contribute to the destabilization of the inherited clay minerals. The continuous groundwater recharge into the margins of the lake can result in the formation of a varied assemblage of authigenic clay minerals in which several diagenetic phases can be recognized (Fig. 15). Groundwater discharge through springs (seepage, artesian flows) into the lake-margin areas is also thought to play some role in the formation of authi genic clays. The common association of fossil mound spring deposits with marginal lacustrine sedimentary sequences bearing authigenic clays, such as in the Amargosa Desert, Amboseli, Madrid Basin (Calvo
et al. 1995b) and older formations as well (Wright &
Sandler 1994), seems to corroborate this assessment. The occurrence of distinctive clay phases formed at different times and occupying various positions within highly deformed carbonate deposits indicates the complex interplay between ground water and saline, alkaline lake water, resulting in the formation of the authigenic clays (Hay et al. 1995). Resedimentation of authigenic clays
Resedimentation of authigenic clays constitutes a reliable mechanism to explain the presence of pre sumably authigenic clays in lacustrine sequences in which sedimentary features are indicative of more dilute saline conditions. This situation has been high lighted in lacustrine sequences from the Miocene formations of the Madrid Basin (Bellanca et al. 1992). Both palygorskite and sepiolite were found in significant amounts as either mud chips or minute
148
J P Calvo et a!.
clay aggregates in the basal deposits of a lacustrine unit accumulated during a rising lake level, which provides a stratigraphical pattern for resedimenta tion of authigenic clays in a lake undergoing a salinity change. Yet data are lacking about the importance of the reworking of clays formed authigenically within soils and/or lake-margin subenvironments and their further supply into more open lake areas. The aeolian contribution of either clay pellets or dust from adja cent areas into the lakes (Talbot et al. 1994) may also contribute to the accumulation of large volumes of authigenic clays in these saline settings.
A C KN O W L E D G E M E N T S
We are thankful t o E. Sanz Rubio, R. Mas and M . Pozo for their help in furnishing some o f the graphic material included in the paper as well as comments and documentation that have supplemented several aspects of the work. This has benefited from the financial support of the Spanish CICYT (Project AMB94-0994) and the CNRS (UA 723 Physico chimie des processus biosedimentaires). REFERENCES
Authigenic clays as indicators of salinity shifts in open lake areas
The sedimentation and possible transformation and/or precipitation (neoformation) of the clays in open lake areas merit some comments. Recent inves tigation by Webster & Jones (1994) demonstrates that clay assemblages are highly sensitive to salinity shifts in these areas, with characteristic clay assem blages corresponding to brackish, saline (perennial) or ephemeral-lake (playa) conditions. Furthermore, the sequential arrangement of the various clay min eralogies may be used as a reliable indicator of the lake-level fluctuation (cyclic or non-cyclic) in conti nental evaporitic environments. Millot (1964, 1970) stated that the spatial distribu tion of clay minerals in continental saline settings follows a rather well-defined zonation pattern, showing a general trend from the most aluminous clays at the periphery of the lakes and more magne sian clays basinwards (Millot 1964, 1970). Evidence from many reported case studies, however, indicates that this pattern is not always realistic and that Millot's model must be used with caution. For instance, most of the sepiolite accumulated in the Madrid and Calatayud basins during the Miocene was formed in marginal settings of saline lakes instead of the central parts of the basin (Calvo et al. 1989; Galan & Castillo 1 984). Likewise, palygorskite can consti tute the predominant clay mineral in open lake areas even though the bulk of palygorskite deposits is deposited in the associated mudflats (Ingles & Anad6n 1991). These observations make necessary additional systematic research on the presence of clays in continental saline settings. Often a lack of multidisciplinary teams working in both present and ancient evaporitic lake systems has resulted in little if any progress in knowledge of the actual importance of authigenic clays in the se settings.
ARAKEL, A.V. & McCoNCHIE, D. (1982) Classification and genesis of calcrete and gypsite lithofacies in paleo drainage systems of inland Australia and their relation ship to carnotite mineralization. J sediment. Petrol. , 52, 1 1 49-1170. ARAKEL, A.V., JACOBSON, G. & LYONS, W.B. ( 1990) Sedi ment-water interaction as a control on geochemical evo lution of playa lake systems in the Australian arid interior. Hydrobiologia, 197, 1-12. ARAUZO, M., GONZALEZ LOPEZ, J.M. & LOPEZ AGUAYO, F. (1991) Caracterizaci6n mineral6gica, qufmica y evoluci6n geoqufmica de los materiales terciarios del area del rio Perejiles. Bal. Soc. Esp. Miner. , 14, 21 1-221 . ARMENTEROS, 1 . , BuSTILLO, M. A . & BLANCO, J.A . (1995) Pedogenic and groundwater processes in a closed Miocene basin (northern Spain). Sediment. Ceo!. , 99, 17-36. BACHMAN, G.O. & MACHETTE, M.N. (1977) Calcic soils and calcretes in the southwestern United States. U.S. geol. Surv. Open File Rep. , 77-794. BADAUT, D. & RISACHER, F. (1983) Authigenic smectite on diatom frustules in Bolivian saline lakes. Ceochim. Cos mochim. Acta, 47, 363-375. BANFIELD, J.F., JoNES, B.F. & VEBLIN, D.R. (1991) An AEM-TEM study of weathering and diagenesis, Albert Lake, Oregon: II. Diagenetic modification of the sedi mentary assemblage. Ceochim. Cosmochim. Acta, 55, 2795-2810. B ELLANCA, A., CALVO , J.P., CENS!, P., NERI, R. & Pozo, M. (1 992) Recognition of lake-level changes in Miocene lacustrine units, Madrid Basin, Spain. Evidence from facies analysis, isotope geochemistry and clay mineralogy. Sediment. Ceo!. , 76, 135-153. BIGHAM, J.M., JAYNES, W.T. & ALLEN, B.L. (1980) Pedogenic degradation of sepiolite and palygorskite on the Texas High Plains. Soil Sci.Soc. Am. J, 44, 159-167. B LANC-VALLERON, M.M. (1990-1991) Les formations paleogenes evaporitiques du bassin potassique de Mul house et des bassins plus septentrionau.x d'Alsace. These, Doct. Sci., University of Strasbourg and Documents du Bureaus de Recherches Geologiques et Mimieres, 204, Orleans, 350 pp. BLANC-VALLERON, M.M. & THIRY, M. ( 1 993) Mineraux argileux, paleoalterations, paleopaysages et sequence climatique: exemple du Paleogene continental de France. In: Sedimentologie et Ceochimie de Ia Surface
Authigenic clay minerals (Eds Paquet, H. & Clauer, N. ) , pp. 199-216. Colloques Academie Sciences Cadas (a la memoire de Georges Millot), Paris. BRADLEY, W.H. & FAHEY, J.J. (1962) Occurrence of steven site in the Green River Formation of Wyoming. Am. Mineral. , 47, 996-998. BREWER, R. (1 964) Fabric and Mineral Analysis of Soils. Wiley & Sons, New York. CALLEN, R.A. (1984) Clays of the palygorskite-sepiolite group: depositional environment, age and distribution. In: Palygorskite-Sepiolite Occurrences, Genesis and Uses (Eds Singer, A. & Galan, E. ), 1-37. Developments in Sed imentology, 37. Elsevier, Amsterdam. CALVo, J. P., ALONSO, A.M. & GARCIA D EL CURA , M.A. (1986) Depositional sedimentary controls on sepiolite occur rence in Paracuellos de Jarama, Madrid Basin. Geogaceta, 1, 25-28. CALVO, J.P.,ALONSO ZARZA,A.M. & GARCIA D EL CURA, M.A. (1989) Models of Miocene marginal lacustrine sedimen tation in response to varied depositional regimes and source areas in the Madrid Basin (central Spain). Palaeo geogr. Palaeoclimatol. Palaeoecol. , 70, 199-214. CALVO, J.P., JONES, B.F., B USTILLO, M., FORT, R., ALONSO ZARZA , A.M. & KENDALL, C. (1995a) Sedimentology and geochemistry of carbonates from lacustrine sequences in the Madrid Basin, central Spain. Chem. Ceo!. , 123, 173-191. CALVO, J.P., Pozo, M. & JONES, B.F. (1995b) Preliminary report of seepage mound occurrences in Spain. Compari son with carbonate mounds from the Amargosa Desert, western USA. Geogaceta , 18, 67-70. CALVO, J.P., ALONSO ZARZA, A.M., GARCIA DEL CURA, M.A., ORDONEZ, S., RODRIGUEZ-ARANDA, J.P. & SANZ MONTERO, E. (1996) Sedimentary evolution of lake systems through the Miocene of the Madrid Basin: paleoclimatic and pal eohydrological constraints. In: Tertiary Basins of Spain (Eds Friend, P.F. & Dabrio, C.), pp. 264-269. Cambridge University Press, Cambridge. CARMOUZE, J.P., PEDRO, G. & BERRIER, J. (1977) Sur la nature des smectites de neoformation du Lac Tchad et leur distribution spatiale en fonction des conditions hydrogeochimiques. C. R. Acad. Sci. Paris, 284, 615-618. CHAMLEY, H. (1989) Clay Sedimentology. Springer-Verlag, Berlin. CHEVERRY, C. (1974) Contribution d !'etude pedologique des polders du lac Tchad. Dynamique des sels en milieu continental sub-w·ide dans des sediments argileux et organiques. These, Sciences Naturelles, University of Strasbourg. CORROCHANO, A. & ARMENTEROS, I. (1989) Los sistemas lacustres de la Cuenca del Duero. Acta geol. Hisp. , 23, 259-279. DARRAGI, F. & TARDY, Y. (1987) Authigenic trioctahedral smectites controlling pH, alkalinity, silica and magnesium concentrations in alkaline lakes. Chem. Ceo!. , 63, 5972. DOYAL, M., CALVO, J.P., BRELL, J.M. & JoNES, B.F. (1986) Clay mineralogy of the Madrid Basin: comparison with other lacustrine closed basins (abstract). Symposium on Geo chemistry of Earth Surface Processes and Mineral Forma tion, Granada March 47-20, 1986, pp. 188-189. DuMAS, D. (1988) Le Paleogene salifere du bassin de Valence (Sud-Est de La France): geometrie et sedimentologie des
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Spec. Pubis int. Ass. Sediment. (1999) 27, 153-188
Saprolite-bauxite facies of ferralitic duricrusts on palaeosurfaces of former Pangaea
I . VAL E T O N Am Hohen Tore 4a, 381 18 Braunschweig, Germany
A B S T R AC T
The tectonic and morphogenetic evolution o f Pangaea with special respect t o the late Mesozoic t o early Tertiary history of the landscape and the early Tertiary weathering cover are described. Within the fer ralitic duricrust of this time span a saprolite-bauxite facies pattern on hilly landscapes and on downwarp ing platforms is developed in extended newly formed coastal areas after the break-up of Pangaea. The early to middle Eocene was still a time of world-wide flat relief, of world-wide warm current systems in the oceans and therefore of a humid warm climate. The relief of the pre-, syn- and post-bauxitic landscapes indicates tectonic lability and short times for bauxite formation. The facies distribution of the vertically and laterally well-developed saprolite-bauxite facies pattern depends on parent-rock variables, morphol ogy and drainage patterns. The mineralogy and chemistry of saprolite-bauxite and the quality pattern in bauxite deposits are dis cussed with respect to the supergene processes. In contrast to 'normal' laterites, a strict separation of A! and Si by an effective extraction of silica has prevented the formation of A! silicates in parts of the Box horizon, leading mainly to neomineralization of gibbsite, boehmite and diaspore. Post-bauxitic tectonic activities have transformed the very flat near-sea-level landscape by subsidence or uplift. Changes of relief and of climate since the Eocene have led to a differentiation of soils dependent on altitude and on climatic zones. Results are either truncated ferralitic profiles and erosional landforms or polygenetic overprinting of saprolites and bauxites by younger soils, forming a complex 'solum'. Alu minization by ferralitic weathering destroyed the main geochemical parent-rock characteristics, resulting in supergene geochemical environments dominated mainly by Al, Zr, Ti, Ga and Fe, but still marked by some trace element associations indicative of the original parent-rock composition. These specific super gene geochemical domains in the ferralitic duricrusts are very useful as lithostratigraphical marker hori zons in terrestrial environments. INTRODUCTION
The aim o f this article i s not t o present another detailed description of bauxitic occurrences or deposits, because a wealth of information is already available about the geological situation, mineralogi cal data and mining patterns of bauxite on a world wide scale ( see e.g., Bardossy & Aleva 1990; Rouillier 1990; Patterson et al. 1994). The goal rather is to evolve a general concept of the special morpho tectonic and climatic environments leading to eco nomically important aluminium concentration in duricrusts by supergene alteration, and to reconstruct these time- and space-related weathering processes. They are connected with the evolution of the Alpine orogen and the reorganization of Pangaea during · late Mesozoic to Tertiary times, the development
of terrestrial topography on tectonic platforms along passive margins, and the growth of immense river systems on the continents, creating special climatic and hydrographic conditions for ferralitic weath ering. Time- and space-related aluminization by supergene alteration has led to the separation of aluminium and silica, and to the formation of clay and bauxite deposits. OCCURRENCE
O F BAUXITE
D E P O SITS
Definition and properties of lateritic bauxites are described by, among others, Millot (1964), Tardy
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
153
154
I. Valeton
( 1969) , Valeton ( 1972, 1983a,b, 1994) , McFarlane ( 1976) , Grandin ( 1976) and Boulange ( 1984) . Tardy ( 1993 ) presents a very detailed description and
genetic interpretation of the different types of lat erites and of tropical soils. In this context the terms 'hard lateritic soils' or 'ferralitic duricrust' are used for deep and hard in situ chemical weathering crusts on various types of non-carbonate rocks, with preser vation of relict textures of the parent rock mainly in the lower saprolitic part and additional neoformed textures in the upper oxic and hardened parts. They consist of layer silicates in the 'saprolite' (sialite ) , and of oxides and oxyhydroxides of iron and aluminium in the 'oxic zones' (ferralite, alite ) . The term 'ferri crete' is very confusing and should not be used in this context. Mechanical or chemical displacement of these materials leads to allochthonous products of 'laterite-derived facies' (LDF) . The thickness of these duricrusts ranges between several tens and more than hundreds of metres.
Bauxites represent the only raw material from which aluminium can be produced economically; this is mainly as a result of the threefold increase of A�03 from parent rocks ( 15-19% ) to bauxites (45-60% ) during weathering. In 1996 bauxite world production amounted to 111 Mt, and world reserves are estimated to be around 23 Gt. Deposits of lateri tic bauxite are located mainly in Australia, South America, Africa and India (U.S. Bureau of Mines 1995 ) , as shown in Fig. l. As lateritic bauxite can change vertically and later ally into kaolinitic or smectitic saprolites, diverse qualities of bauxites and clays and their by-products are developed. The production of special-grade bauxite for abrasives, cement, chemicals, refractories and other non-metallic purposes ranges between 2 and 3 Mt yr-l. Saprolitic high-A! clays, high-quality kaolinites and/or smectites can be important by economic products in connection with bauxite mining. In some bauxites, elements such as Ti, V, Ga,
Bauxite ores and reserves on: Igneous and metamorphic rocks "' Clastic sedimentary rocks •
Fig. 1. Distribution of lateritic bauxites on Pangaea (based on Patterson et a/. 1994).
Pangaean saprolite-bauxite facies
Ge and Au can be enriched, which eventually can be separated during processing. Bauxites on top of metamorphic rocks, especially in greenstone belts, can be enriched considerably in Au. In weathering crusts on alkaline massifs, high concentrations of phosphates and elements such as Sr, Zr, Ti,Th, U, Ga, Nb,Ta and REE can be of economic interest. In addition to their value as mineral deposits, saprolites and bauxites in tropical areas can be important aquifers as a result of their high porosity and permeability. Ferralitic duricrusts are also used for road construction. SEQUENCES
OF P A L A E O S U R FAC E S
A N D T H E S P E C I F I C W E AT H E R I N G P R O FI L E S T H AT D E V E L O P
ON
THEM
King ( 1953, 1962) first directed attention t o palaeo surfaces and their interpretation in 'canons of land scape evolution'. Only a short summary of the world-wide literature available on this topic can be presented here: Palaeosurfaces have developed during different times in Earths history in response to specific tectonic activities under various climatic, biological and pedogenic conditions. Different surface-covering layers such as saprolites, laterites, bauxites, silcretes, calcretes and other duricrusts (e.g. iron caps), are the essential testimony of the prevailing palaeoenviron ment. Palaeosurfaces of different ages are character ized by specific properties corresponding to their evolution, and can be used as lithostratigraphical marker horizons. Saprolites connected with lateritic bauxites formed during the time intervals of early Cambrian to early Ordovician, Late Devonian to early Carboniferous and Cretaceous to Tertiary. A first period of bauxite formation started during the Middle Devonian and reached a peak during the early Carboniferous (Bardossy 1993). The global climate, but also the tec tonic situation, were generally unfavourable for bauxite formation during Triassic, Early and Middle Jurassic times. During Late Jurassic to Cretaceous times the first bauxites appeared on carbonate plat forms of the Tethyan geosyncline in the European, Asian and African Mediterranean area (Bardossy & Combes, this volume, pp. 189-206). Most of the dated bauxites on Pangaea are early Tertiary in age; the bauxites in the Caribbean region belong to the late Tertiary. On all continents of Pangaea the oldest planation
155
belongs to the Gondwana surface, which is dated to pre-Jurassic and Jurassic in time. It is followed by the Cretaceous post-Gondwana surface. The Gondwana palaeosurfaces on the Australian craton and the history of their evolution is well described by Twidale (1994). These relict surfaces frequently are charac terized by deep erosion, and with regard to their age they may correspond to much older 'retaken' land surfaces in parts, belonging to the 'genetically complex' type of plains. During the Late Cretaceous to early Tertiary pla nation, extended land surfaces (the 'les grandes surfaces' of Grandin & Thiry (1983)) were formed on most of the Gondwana continents. The so-called 'Sulamericana', 'African' and 'Indian' land surfaces (Table 1 ) are covered by early Tertiary ferralitic duri crusts, which contain saprolite-bauxite facies in areas of optimum drainage. Platform situations characterized by sequences of sediment accumulation, intercalation of eroded land scapes covered by soils or other weathering products, permit dating and reconstruction of the terrestrial weathering history. Stratigraphical dating or absolute age determination of lateritic bauxites on Pangaea outside of the Tethyan area is possible only when they occur in sedimentary or volcanic sequences. In these cases a post-Upper Cretaceous and pre-Oligocene, i.e. a Palaeocene to Eocene age is indicated (Gordon et al. 1958; Valeton 1967, 1985; Aleva 1979, 1983, 1984; Grandin & Thiry 1983; Bardossy & Aleva 1990; Valeton et al. 1991; Tardy 1993; Valeton & Wilke 1993). The thickness of these weathering profiles ranges between several tens to a lOOm and they represent a unique geochemical marker horizon containing eco nomically important supergene deposits of 'lateritic ores' mined for AI (bauxite), Fe/Mn, Ni/Co, Cu, Au, PGE, REE and phosphate (see Valeton 1994; and ref erences therein). Fault activities during late Tertiary and Quaternary times caused the dislocation of palaeosurfaces, fre quently with very high uplift rates. In uplifted areas the oldest surfaces are in the highest position, fol lowed by younger surfaces in lower positions. On downwarped shelves or graben structures the oldest plains are deepest followed upward by younger accu mulation surfaces (Fig. 2). Lateral drainage of silica saturated water out of the elevated and intersected plateaus into lower plains led to silcrete formation (e.g. the thick groundwater silcretes in internal depressions in Australia (Simon-Coin\on et al. 1996)). Grandin & Thiry (1983) mention two periods of intensive silcrete formation following alitization
156
I. Valeton
and ferralization, one during the late Eocene to Oligocene and a later one during the Late Miocene to Early Pleistocene. The late Eocene/Oligocene to Early Miocene times were characterized by start of relief accentua tion and a cooler and drier climate. During the Middle to Late Miocene, which was marked by a warm humid climate, the development of deep lat eritic weathering profiles was restricted to basaltic rocks and extreme tropical monsoon climates (Oregon, USA) . During the Quaternary, world-wide tectonic activities led to the evolution of the present-day morphology of the landscape and to the prevailing differentiation in climate, vegetation and soil types. The early Tertiary surface in places thus became uplifted to altitudes of more than 2000 m. During the Quaternary, ferralitic weathering is restricted to volcanic rocks of few local tropical areas only (Hawaii, South-east Asia, South Vietnam; Bardossy & Aleva 1990). The north-eastern part of the Guyana Shield
in South America represents a particularly good example of complex history and the succession of planation plains; along the uplifted Guiana block on one side and the north-eastern subsiding platform on the other side (Aleva 1983; Fig. 2). LATE
M E S O Z O I C TO
E A R LY
T E R T I A RY P A L A E O S U R FAC E S A N D EXTENSION O F THEIR FERRA LITIC D UR I C RU S T S WITH A S A P R O L I T E - B AU X I T E FAC I E S
PATT E R N
To understand the relationship between: 1 the world-wide tectonic evolution, 2 the characteristics and genesis of pre-, syn- or post bauxitic landscape, 3 the trends of climate and evolution of subsurface drainage patterns, 4 the palaeogeographical position of saprolite-
Table 1. Sequence of planation plains in Guiana and Brazil (Modified from Aleva, 1981) Sequence
Approximate absolute age (Ma)
v
IV III
1 5 25
II
50 65 and more
Relative age Pleistocene, recent Late Tertiary II, Pliocene Late Tertiary I, Oligocene -Early Miocene Early Tertiary, Palaeocene -Eocene Jurassic-Cretaceous
Planation plain in Brazil, Guyana Paragacu Late Velhas Early Velhas Sulamericana Gondwana
Fig. 2. (Opposite) Coastal plain bauxite deposits in the Paranam-Onverdacht-Lelydorp area showing the relationship between morpho tectonic genesis and spatial extension of the bauxite (redrawn from Aleva 1973, in Bardossy & Aleva 1990). (a) During the Late Cretaceous, first uplift activities of the Guiana block, followed by subsidence and deposition of sand and clay on top of the Gondwana-post-Gondwana plain (G-Pl-P); during the Palaeocene, continous slow subsidence and transgression on the coastal platform resulted in the deposition of thick layers of sand and clay in a landscape of tidal flats, sand bars, mud flats and marshes; tropical climate with palms, mangrove and ferns. (b) During the Eocene, bauxitization started on top of the Sulamericana plain (S-Pl), accompanied by sea-level and groundwater oscillation; a weak regression favoured slight dissection of the platform. TI1e following period of warm humid climate promoted weathering and lateritization and the formation of bauxite over most of the area. Preservation of a high groundwater level on the platform supported the removal of alkalis, calcium and silica in solution through groundwater and creeks. Area too high for mangrove vegetation. (c) During the Oligocene, a distinct marine caused renewed erosion and the formation of wide valleys; the laterite-bauxite caps prevented underlying soft sediments from erosion, flat-topped hills were considerably reduced in size. The area is too high for mangrove vegetation. (d) Since the late Oligocene, renewed transgressions and renewed deposition of thick layers of sand and clay prevailed in a hot and humid climate with abundant mangrove vegetation. Sedimentation first infilled the valleys between the laterite-bauxite capped hills, later it covered most of these hills. At least five periods of drier climate with extensive occurrences of grass in the landscape are discernible. In the top layers the remnants of Amerindian cultures have been found.
Pangaean saprolite-bauxite facies S D u ring and after subsidence deposition of thick layers of sand and clay
During long stable period leaching and laterization resulting in formation of bauxite caps on surface
157 N
ca.
60 Myr
ca.
55-40 Myr
Leaching and laterization
Sulamericana � Piain (S-PI)
....
"+- (G-P-PI) (b) ca.
35 Myr
During regression, incision of wide valleys and formation of table mountains Continued humid climate Bauxite
Remnants of older sedimentary cover
Val ley
_
S-PI
....
G-P-PI
(c)
During and after subsidence deposition of sediments
(d) Distance
�------ ca.
ca.
30 Myr to present time
- G-P-PI 1 50 km --------�
158
I. Valeton
Eocene 45 Myr
\ J
-
-
-
-
---
-
1 80° -----. marine current system - large river system lateritic bauxites
---
� marine Lower Tertiary
1 20°
1 80°
� marine Cretaceous � (only for Australia)
Fig. 3. New coast lines after reorganization of Pangaea; direction of river systems and bauxite distribution. Circulation pattern of warm surface waters in the oceans at 45 Ma (middle Eocene) with a still intact warm equatorial current system.
bauxite belts within ferralitic duricrusts and their spatial mineralogical-geochemical pattern, it is nec essary to take a brief look at Pangaea's history: During the late Mesozoic to early Tertiary-after the reorganization of the cratonic shield areas of Pangaea-the properties of the new continents changed drastically. Alkaline ring structures and large complexes of flood basalts (Cameroon, Parana Basin in Brazil, Deccan Traps in India) are related to structural lineaments. Graben structures and triple junctions cause deep marine embayments (e.g. Mississippi, Amazon, Benue Trough). The world-wide formation of new river systems led to a maximum of erosion of the deeply weathered terrestrial hinter land and to the accumulation of fluvio-deltaic and lit toral clay-silt-sand associations in the coastal areas. The creation of new coast lines is the world-wide precondition for the formation of extended deep fer ralitic duricrusts on the Cretaceous-Palaeogene pla nation plains (Fig. 3). Peneplation took place during times of marine regression, whereas bauxitization
started with the onset of marine transgressions and rising groundwater levels. Tectonic evolution and exposure of parent rock areas on Pangaea
The tectonic evolution since the break-up of Pangaea led to the dislocation of continental blocks, and to the exposure of four main groups of parent rocks. Baux itization took place on: 1 pre-Cretaceous basement rocks and on Palaeozoic to early Mesozoic sediments; 2 Precambrian and Jurassic-Cretaceous-Tertiary alkaline ring structures; 3 Cretaceous to early Tertiary flood basalt sequences; 4 Cretaceous to early Tertiary fluvio-deltaic and lit toral clay-silt-sand associations. The phenomenon of world-wide evolution of flexure zones or horst and graben systems along the new borders of the continents has influenced land-
Pangaean saprolite-bauxite facies
159
\ .J
NW (m a.s.l.) Sulamericana-Piain , , Serra do Relo � 10 1 200 ', Serra dos Paulos Rio ', Pombe 800
SE (m a.s.l.) 1 200
•1
400
Rio
400
h I
(a)
0 � 0 1 2 3 km
0
NW p 1 837 p 1 838
landslide
p 1 665
I
Fig. 4. (a) NW-SE section of the Cataguasis region, Serra da MantiqueiTa, south-east Brazil indicating neotectonic displacement of the early Tertiary Sulamericana plain on the Charnockite belt. The crests of the relict plateaus, covered by bauxite, occur today at different levels between 400 and 1300 m (Beissner 1989). (b) Vertical zonation of thick in situ bauxite profiles on crests and higher slopes; colluvial transport and reworki11g dominates on lower slopes and valleys of Serra dos Paulos, Cataguasis (Beissner 1989).
800
mnniiJ) r-:;::3
�
w��
I
I
p 1779
I
SSE ( m a.s.l.)
Recent soil Reworked clayey material Reworked bauxite
Bauxite, in situ � � residual breccia Bauxite, in situ
(b)
SERRA DOS PAULOS
scape evolution since late Mesozoic times (Table 2, Fig. 4). The formation of ferralitic duricrusts with a saprolite-bauxite facies pattern took place on those parts of Late Cretaceous to early Tertiary planation plains that are related to new coastal lines along extended platforms on passive and subsiding shelves. Those duricrusts are developed in the topog raphy near-sea-level altitudes and their formation is closely connected with sea-level oscillations (Figs 2, 5 & 6). Pre-Cretaceous basement rocks and Palaeozoic to early Mesozoic sediments
Bauxite deposits on shield areas develop on anorthosites, charnockites, granulites, gneisses,
lrl§ l.; .�:::.j Saprolite 1•'< ,."111 Parent rock
greenstone-belt lithologies or Proterozoic to Palaeo zoic slates, phyllites, shales and, rarely, on Karoo sediments. In South America bauxites occur around the north-eastern, eastern and south-eastern part of the Guiana Shield and the eastern part of the Central Brazil and Atlantic shield. The early Tertiary Orinoco-Sulamericana plain from Venezuela to south-east Brazil is covered by a ferralitic alteration crust several tens of metres deep (Aleva 1984; Bardossy & Aleva 1990). In the uplifted basement areas, the bauxite displays a regional distribution on hilly landscapes with an accentuated relief (crests on 'half-orange' topography). The economic importance of these bauxites is based on the wide distribution of mineable deposits and the chemical quality of the
I. Valeton
160
Table 2. Evolution of weathering cycles since late Mesozoic time (Modified from Yaleton, 1994) Soil sequence Cycle
Sequence
Age
Plate tectonic
Evolution of
within the
situation
planation plain
weathering cycle
Supergene Climate
chemistry
Precambrian II
Early Palaeozoic
Ill
Pre- and early
Break-up of
Mesozoic
Pangaea
IV
Jurassic and Cretaceous
Reorganization of new Pangaea continents Main time of
Gondwana surface Post-Gondwana surface Reorganization
Eocene
Beginning of
Warm to temperate
from
of large river
'equatorial zones' far to the north
intrusions Palaeocene-
weathering
systems
anorogenic 2
Sialitic
and the south Flat
Formation of
Warm-
collision-
morphology
deep and
wet
subduction-
on extended
hardened
greenhouse
orogeny
platforms
saprolite-
effect
along passive
bauxite-
C02
margins
ferralite
of the new
duricrusts
continents
along passive
Sulamericana surface
Alitization, ferralitization
margins; local silcretes
African surface 3
Indian surface Oligocene
Activation of
Morphological
Formation of
vertical
differentiation
groundwater
movements
of the relief
silcrete in
on Pangaea
lower plains
Cooling down
Increasing reworking, sedimentation of LDF Increasing geochemical differentiation owing to climatic, morpho- and pedogenetic differentiation
4
Late TertiaryMiocene
Red earth
Increasing
Warm-wet optimum
h01·st and
(kaolinite,
graben
gibbsite,
Middle
tectonics
haematite);
Miocene
Over bauxites: in situ
brecciation of bauxites (residual breccia) and red earth formation
5
Quaternary
Strong vertical movements
Evolution of
World-wide
Cooling
younger
increasing
down,
plantation
differentiation
glaciation
plains in
of soils: warm
uplifted areas
-wet -yellow soils topped by black soils; arid -calcrete -saltcretesilcrete; Temperatewetpodzolic soils
161
Pangaean saprolite-bauxite facies 70
75 IIlJ!]] Laterite without Bauxite D Tra p Basalt (Cretaceous to Palaeocene) ffi Precambria ?::i
�
N I
; .......
� ; �<
Ahma "bad 1
I
i I
30
I
: '
! {
20
20
15
15 miD Area of younger
erosion 12Z3 Tertiary, marine Eo=Eocene M io= Miocene • • • Bauxite Deposit c::I D eccan Trap 0 1 00 500 km
10
70
75
80
10
85
Fig. 5 . Early Tertiary bauxite belt o n the Deccan Peninsula and in Gujerat showing its relationship t o the early Tertiary coast line (Valeton 1983a,b).
ores. TI1e bauxitic duricrusts in crest areas of the Serra da Mantiqueira, central Brazil, were submitted to optimal vertical drainage conditions and thus present direct transformation from parent rock into bauxite without an intermediate saprolite (Figs 4 & 15). Neotectonic uplift has moved the early Tertiary surface to altitudes between 400 and 2000 m. In West and equatorial Africa (Fig. 3) large bauxite deposits occur on Ordovician sandstones, Silurian slates, Devonian shales and sandstones (penetrated by Mesozoic alkaline intrusions). Bardossy & Aleva (1990) distinguish between the Guinean subprovince on the Precambrian and Palaeozoic basement area
(de Weisse 1954; Saposhnikov et al. 1976; Mamedov et al. 1985; Boski 1987) and the Cameroon subprovince
on Late Cretaceous and early Tertiary lava flows (Belinga 1972; Hieronymus 1985). The bauxites and bauxitic laterites on the Jos Plateau in Nigeria are situated between the two provinces on the western border of the Benue Trough, covering Precambrian basement rocks, Mesozoic alkaline ring complexes and 'fluvio-volcanic series' (the equivalent of the 'flood basalts' in Cameroon). Early Tertiary bauxites are overlain in this area by lake sediments containing an Oligo-Miocene flora (Takahashi & Jux 1989). In East Africa bauxites appear on the eastern granulite
+----- Kutch-Mainland -------
S
Palaeocene/ U. Cretaceous
v
. ., ., Trap-Basalt ., v .... . : .
v
v
Palaeocene - Upper Cretaceous
v
.:t::: :J "' u._ ...!.
g"'
Jurassic (a)
N
approx. 50 km
><:
, • o ' , ,
.
.
Lower Eocene ca.
Laki. � ----� � � L. Eocene/ Palaeocene ...
.
•Sandstones, claystones, limestones _.:..--:--- , . an "ba sement" . Preca mbn .
55-40
,·
m i l l i o n years a g o
�����5&�������������am����ms����������������--
�
s
(b)
Middle Eocene
(c) Pliocene Miocene .... Oligocene+-
Pliocene
1� ���i������=�� s
\1
N
N
N
Quaternary to present day
Northern
(d)
(e)
D . La �
Quaternary Pliocene Miocene
�. Ill ...
Oligocene M. Eocene Nummulitic Limestone Trans-Regression
8 Ill
E. Eocene, Laki-Formation Laterite-Bauxite
u � D.
N
Basalt Early Cretaceous Jurassic
Fig. 6. N-S section in the Kutch area of Gujerat showing the late Mesozoic-Cenozoic evolution of the platform and steps of subsidence during Jurassic transgression, Early Cretaceous regression; cover by flood basalts (a); formation of a first ferralitic duricrust during the Palaeocene to early Eocene (b); reworking and seaward resedimentation of laterite-derived material (LDF) and its bauxitization during the early Eocene (several smaller regressions and transgressions) (c); subsidence and marine regression during the late Eocene and Miocene (d); the actual situation of the deeply eroded landscape (Wilke 1987).
Pangaean saprolite-bauxite facies
complexes in Tanzania (Mutakyahwa & Valeton 1995) and on Mesozoic Karoo sediments and Jurassic dolorite sills in the Natal district, South Africa (Fitzpatrick 1983; Horn 1983). In India bauxites cover granulites and gneisses in the southern part of the Deccan Peninsula (Valeton 1967, 1972). In Australia the Darling Range bauxite district is situated over Archaean granites and gneisses of the Yilgarn block.The North Kimberley bauxite deposits are underlain by platform rocks (Bardossy & Aleva 1990). Precambrian and Jurassic-Tertiary alkaline ring structures
Precambrian and Cretaceous to early Tertiary alka line intrusions that are related to tectonic lineaments occur in all parts of Pangaea. Bauxites over alkaline rocks are of high economic importance because of their special mineralogy and chemistry. Intermediate rocks such as syenite are low in iron content, but they can be enriched in phosphorus, zirconium, niobium and REE. The rapid weathering of foids such as nepheline produced high porosity and permeability, and therefore optimal drainage conditions. The Precambrian P-bearing alkaline rocks of Para and Maranhao, north-east Brazil, are good examples of bauxites enriched in phosphorus (Schwab & Oliviera 1981 in Schwab et at. 1983). Meso-Cenozoic complexes are situated in south-east Brazil, e.g. Pot;os de Caldas, Lages, Araxa and others (Melfi & Carvalho 1983; Valeton et at. 1988, 1997; Formoso et al. 1989; Schumann 1994). Extremely well-developed vertical and lateral facies patterns of saprolite-bauxite are characterized by saprolite in depressions or lower slopes and high quality bauxite on slopes, altered to bauxites and enriched in P, Zr, REE and/or Nb. Cretaceous to early Tertiary flood basalt sequences
Economic bauxite deposits on flood basalts occur in Cameroon and India, and younger subeconomic bauxite of this type is described from Oregon and South Vietnam. In India, plateau basalts of the Deccan Peninsula and the Kathiawar block are covered by a bauxite belt that follows the early Tertiary border of marine sedimentation (Figs 5 & 6). The basalt layers of Kathiawar and Kutch are inclined from NE to SW. They were strongly eroded during a Palaeocene ter-
163
restrial phase directly before the evolution of the lat eritic-bauxitic duricrust. Saprolites and bauxites are developed over basement rocks, Early Cretaceous sandstones, Late Cretaceous to Palaeocene trap basalt and Palaeocene sediments of the early to middle Eocene Indian land surface. A 'bauxite belt' with a maximum lateral extension of 30-60 km was formed. Three bauxite horizons that developed sub sequently over basalt and the near-coastal terrestrial sand-clay-lignite association of the early Eocene Laki formation are the result of subsidence, sea-level oscillations and synchronous groundwater move ments (Fig. 7). Bauxitization is succeeded by an extensive marine transgression of the middle Eocene. Nummulitic limestone overlies those parts of the bauxites reached by the transgression. A lateral facies differentiation from marine shales into the saprolite-bauxite belt and finally into a Fe-rich laterite catena is well developed in the Kutch area, Gujerat, India (Fig. 8; Valeton 1983a,b; Wilke 1987). The facies member next to the shore line is a very thick and irregular smectitic saprolite with perfect preservation of relict basaltic textures. In the continental direction this smectitic saprolite is gradu ally substituted by a kaolinitic saprolite. Finally, a 'typical' vertically zoned bauxite profile has devel oped (Fig. 9). The indurated bauxite is of light colour and poor in iron near to the shore line. It becomes more and more replaced by a hard iron-rich bauxite or ferralitic crust in the landward direction. This iron rich bauxite corresponds to bauxite deposits of the uplifted plateaus of the Deccan Peninsula. Cretaceous to early Tertiary fluvio-deltaic littoral clay-silt-sand associations
Bauxite belts that formed over fluvio-deltaic, estu arine and littoral sediments following the Palaeogene coastal lines comprise the largest ore deposits of the world. The fluvio-deltaic to marine sand-clay facies with lignite intercalations and lagoonal sediments indicate fluctuation of sea-level by regressions and transgression. The oscillation of sea-level and the related mobility of the groundwater table are essen tial factors for the genesis of this type of bauxite. It is characterized by laterally differentiated bauxite laterite facies and conspicuous vertically zoned profiles. These bauxites are always accompanied by synchronous and/or subsequent reworking and resedimentation of laterite-derived material. Erosion of the deeply weathered hinterland has created a par ticular sedimentary facies association (Valeton 1971;
164
I. Valeton
grey claystones
bright claystones, root horizons
iron crust
f-U-UIJU..U.J.J.J..I..Ll.u.J.u.J.LL.LJ.J..I.Jj
bauxite kaolinitic saprolite
- - - - - -- ----·
mea �������ijWJJ
�
- - -
... : >
+--'--'--'...;...;."'-'--.;._-.-:..j
1 m
clayey conglomerate bauxite
<::
' section 2
lignite
iron crust transition zone: bauxite/iron oxides saprolite, kaolinitic, partly with corestones
smectitic saprolite
•
v
v
generalized section 1
v
v
Trap basalt
Ratardia river section
Wilke 1987; Valeton & Wilke 1993; Morgan 1995), which is characterized by the following features (Figs 7 & 10). 1 Very uneven facies changes of alternating fan glomeratic fans, with very irregular lenticular to cross-bedded stratification, and also horizontal stratification. 2 Permanent contrast in colours between red, purple, white, grey or dark grey to black. 3 Continuous alternation of poorly sorted, low-
Fig. 7. Standard profile of the three subsequent ferralitic weathering pofiles (I, II, III) at the early to Middle Eocene Indian land surface in Kutch (Iwanoff 1985; in Vale ton & Wilke 1993).
maturity sediments and high-maturity sediments with rounded clasts. The low-maturity rocks can be composed of a large variety of angular or subrounded rock components, quartz, feldspars, angular clay pebbles, angular particles of iron crusts embedded in a clayey matrix rich in kaolinite. The high-maturity rocks consist of well-sorted and well-rounded monomict components of, e.g., pure quartz sand, kaolinitic clay or black layers of heavy minerals. 4 Cementation of selected layers by precipitation of
(Opposite. ) Lateral facies evolution of ferralitic profiles on flood basalts in India from the coastal situation (left) to the landward direction (right). Profile (a), Goniasar, presents a very thick smectitic and kaolinitic saprolite which thins out towards the hinterland. The cover, a ferralitic horizon, consists in the seaward direction of pure white bauxite with a nodular, 'boulder-like' or pisolitic texture. Toward the hinterland the bauxite becomes richer in iron, which occurs in layers or concretions, and finally is replaced by a ferralitic crust (after Wilke 1987). Similar sections in Gujerat are presented in (b-d). Profile (e) is situated in the uplifted Western Ghats of the Deccan Traps area (after Valeton 1967; 1983a, 1984; Valeton & Wilke 1993).
Fig. 8.
GONIASSAR
N 50 E
N 20 E
S 20 W
S 50 W
N 60 E
S 60 W ',,
---- 10 m---(a) Kutch: HST
- hinterland E truncated section
�IIIWIIIIIIIIHIUU@Illth.. 1 5-20..;;- : ---- � --- : :..-
�t
·
.
·
.
�
Miocene,..__
seaward - @)
- - - --
.
·
.
·
. � '
: �--:�
t
..
-�
Pliocene
�
" · < : ·. .: ·: · · " .. � � y
:� ==:=:-::.: kaolinitic clay
with quartz spheroidal retic texture :.... +--- 60 m (? Laki. F.) +----- 40 m +- 3 m __..
Udagiri Plateau/Western Ghats
;;p ;:s
c;§ "'
$::) ;:s
�
(b)
® Kutch: LST
2.
�·
<:)-' $::)
B.
�
'i:j> ("")
(c)
�-
0 Mewasa/Kathiawar: LST
NE
sw
. �=....:.� . -�==-�..:_:::_...:.-:::-==-=-.::=:�==--==--:..-'P -5 m �::::_:...= .
.
. ·
"' =--�-�-=-=:;:--����--=-::=--:_� -=--:=� -=v-:�--=.., ..., .., "
v
v
v
v
v
v
v
v
(d)
v
y
v
v
-300 m
v
v
v
v
v
v
v v •
=
@ Udagiri Plateau/Western Ghats: LST
!a,m/ ::J�[ SW
.::::: :_�·
.
·
v-; . ' . .
.... ----: :- ....-... -:-....-... � "'-" � -:, ..... � -v � ;;-
(e)
-
...
-300 m
>.
NE ����� om
..,-.. ---;;
vv v v"
Miocene Limestone Nummulitic Limestone (Mid-Up. Eocene) Laki F ormation iron-rich bauxite Boehmite, diaspore (bo, di) bauxite alucrete gibbsite (gi) bauxitic talus (ka) LST saprolite kaolinitic bentonic (bt) HST Trap Basalt altered fresh
...
v -
>--' 0\ Vl
166
I. Valeton M I N E RALS
HORIZON
'local synonyms'
Box
ferric rete
and other characteristics
-· · ·cc..oOIIN..-
'iron - crust'
fe
TEXTURE
newly-formed t.
highly porous, cavernous, red to dark red
newly-formed t.
pisolitic t. 'fluidal t.'
alucrete
Box81
in-situ
-brecciated relict t i n single
'bauxite'
grains
porous, hard, reddish brown, yellow, cream
vesicular t.
Br Box
relict of roots
white - red, soft - hard
newly-formed t.
Kaolinitic saprolite Br
k
i n upper parts
white, soft, dense, partly red
'lithomarge'
relic t.
c
relic t.
tongue
green-g rey, sticks with
Br5
B/C
smectitic saprolite 'bentonite'
in lower parts
V y
v /,. v
y
I
- - - - - - - - ·
I V II � --""
v
......,\
y
y
v
altered _
_
_
_
parent rock
('trap-basalt')
fresh
y y
y
feldspar v pyroxene
crumbly, dark grey
v
olivine, glass, chlorophaerte
v
v
v
silica, iron or colloidal clayey material, in the form of lenses, concretions or tracing burrows of animals or roots. Depending on the sedimentary environment, iron occurs in oxide minerals or as siderite. 5 The more quiet sedimentary environments are characterized by the presence of plant remains, roots and trace fossils of animals and finally of peat or lignite. The high content of clayey clasts, clay balls, col loidal clay material and of immature compounds, favours the transformation of this type of laterite derived sediments (LDF) into high-quality bauxites. Examples of this type are the bauxites surrounding the Arkansas nepheline syenite and found on Paleo gene sediments in the Mississippi embayment, USA (Fig. 1 1 ) (Gordon et al. 1958), bauxites of the Amazon area and the north-eastern part surrounding the
Fig. 9. Vertical section of saprolite-bauxite bearing laterite over basalt, Kutch/India (after Wilke 1987).
Guiana Shield (Valeton 1971; Aleva 1984; Lucas 1989; Truckenbrodt et al. 1995), as well as bauxites from the Indus and Gujerat area in India (Valeton & Wilke 1993), and the Carpentaria Gulf area with the deposits of Weipa and Gove in Queensland and Northern Territory, Australia (Loughnan & Baylis 1961; White 1976; Loughnan & Sadleir 1984; Schaap 1984, 1985; Morgan 1995). Aleva (1 965, 1979, 1981) developed a very instruc tive model of subsidence in the eastern part of the Guyana Shield and of sedimentation, subsequent bauxitization and alitization of the sediments, the syn- and post-bauxitic dissection of the near-coastal planation plain by valleys and their coverage by younger sediments. He also discussed the importance of lateral groundwater flow in explaining the lateral facies variation in saprolite-bauxite. The extraction
Fig. lO. (Opposite. ) Laterite-derived facies (LDF) on clastic sediments (Valeton 1971 ). (a)Irregular kaolinitic clay beds with alternation of coarse-grained angular kaolinitic clasts, fine kaolinitic lenses and heavy mineral layers, Sura leo haul road, St Helena, Surinam (Valeton 1971 ) . (b )Pre-bauxitic red and white sediments with alternating boulder-clay layers, rich in haematite, and heavy mineral layers; post-sedimentary gibbsitization; Sura leo haul road, St. Helena, Surinam (Valeton 1971 ). (c) Reworked and stratified laterite overlying bauxite deposits,Jos plateau, Nigeria (Valeton 1991 ) . (d) Reworked and stratified pisolitic laterite, Maktesh Rahman, Israel. (e) Siderite layers in Laki formation (LDF), Kutch, India. (f) Alternating layers of fine kaolinite, siderite, and sandy layers with haematite, penetrated by fossile root horizons, Kutch, India (Valeton, unpublished).
Pangaean saprolite-bauxite facies
1 67
(a)
(b)
(c)
(d)
(e)
168
I. Valeton
N ... INTERIOR LOW PLATEAUS PROVINCE
1 00
(a) TYPE 2 Colluvial deposits at the base of the
TYPE 3
TYPE 4 Conglomeratic deposits at the base of the Saline
TYPE
200
300 km
1
Fig. ll. (a) Palaeogeographical situation of the Late Cretaceous to early Tertiary Gulf Coast area, USA and bauxite formation during Palaeocene-Eocene time (after Overstreet 1964): 1. TWN - Eocene nonmarine; 2, TWM -Eocene marine; 3, TM - Palaeocene marine; 4, bauxites; 5, actual coastal line; 6, Eocene coast line; 7, inner line of the bauxite belt. (b) Vertical section across the Arkansas nepheline syenite complex, presenting its in situ ferralitic weathering, reworking of saprolites and bauxites and resedimentation of laterite-derived material (LDF) together with lignites during Palaeocene-Eocene time (after Gordon et a!. 1958).
of silica, but also of iron and aluminium and their transportation, are indicated by arrows in Fig. 2b.This bauxite belt extends to the immense, economically very important, bauxite deposits in the Amazon
region (Truckenbrodt et al. 1995; Lucas 1989): Creta ceous and Palaeocene sediments are the parent rocks for early Tertiary lateritization-bauxitization (Fig. 12 & Plate 1, facing p. 158). The southward extension of
169
Pangaean saprolite-bauxite facies Loose layer
0 --
Yellow clay facies
Nodular layers Cemented layers
White and purple clay layers 20
White and purple clay facies m
Fig. 12. Evolution of a saprolite-bauxite profile on Cretaceous-Palaeocene sediments in the Amazon region, which is overprinted by a polyphase and polygenetic 'sol urn' development on top (ferruginous facies and yellow clay) (Lucas 1989).
this bauxite belt is found in the phosphate-bearing bauxite over alkaline massifs and phyllites in the states of Para and Maranhao, Brazil. The bauxite belts of Weipa and Gove in the Carpentaria Gulf area, Australia are underlain by Jurassic-Cretaceous sediments in clay-sand conglomerate facies, with an upper sequence of Late Cretaceous marine transgressive glauconitic sand stones in the Weipa region (Weipa Cycle of Rolling Downs Group; Grubb 1971, D.J. Burke in Bardossy & Aleva 1990). In the Gove district, the crystalline base ment is overlain by the Early Cretaceous Mullaman Beds, intercalated with lignites containing an Albian microflora. At Weipa the bauxite deposits still repre sent the most elongated continuous bauxite blankets on a peneplain that has been uplifted only slightly (10-50 m above sea-level). The bauxites have been exposed at the surface since their formation and sub jected to reworking since early Tertiary time. In large areas of their distribution they are covered by irregu lar sedimentary layers enriched in bauxitic pisolites (Plates 1 & 2, facing p. 158). Characteristics and genesis of pre-, syn- and post-bauxitic landscapes
Time-equivalent surfaces are the Sulamericana plain in South America, the African surface in Africa, the Indian planation plain in India and equiva lent plains in the high pediments in Australia. The general characterization of these land surfaces in relation to their parent-rock properties, tectonic structures and the evolution of their various weather ing crusts has still to be finalized. Two main morpho logical types of landscapes covered by bauxite can be distinguished:
1 Surfaces with a hilly relief and bauxite formation on plateaus and slopes mainly on top of basement rocks with a 'half-orange' topography (Serra da Mantiqueira, Brazil), or on partly steep slopes of alkaline intrusions (Pot;os de Caldas, Minas Gerais, Brazil). Bauxite deposits of this situation are often local and small and of limited economic quality. 2 Platforms with extreme peneplanation on which large, economically very important bauxite belts are developed (saprolite-bauxite covered platforms around the Guiana Shield, in Equatorial Africa, in South-east Asia and around the Gulf of Carpentaria, Australia). Processes of pre-, syn- and post-bauxitic tectonic dislocation and contemporaneous change in drainage have influenced, in addition to the parent-rock properties, the specific forms of the early Tertiary landscape, as can be observed in platform-bauxite belts in India, North and South America, Africa and Australia. Good examples for tectonic activity causing con temporaneous land-forming processes during the early Tertiary are represented by bauxites over flood basalts and over Late Cretaceous to early Tertiary sediments in Gujerat, India (Wilke 1987; Valeton & Wilke 1993). In this area, some hundred metres of basalt were eroded before and during the formation of ferralites and bauxites (Fig. 6b-d). Three horizons of bauxites are developed in the same platform situa tion, occurring as a sequence of bauxites in a tectoni cally active zone (Fig. 7). This indicates that the time of duricrust formation was not necessarily a long quiet period, but that bauxite formation could· occur in relatively short time intervals under favourable environmental conditions. The landscape morphology during the time of
170
I. Valeton
saprolite-bauxite formation on the early Tertiary Sulamericana plain was first described in detail from Surinam by Aleva (1965). It is characterized on base ment rocks as well as on early Tertiary sediments by a more-or-less flat relief dissected only by small river channels, not deeper than 10-20 m (Fig. 2b). Bauxit ization here took place at a near-sea-level altitude. Similar successions of peneplained surfaces with extended bauxite belts on Late Cretaceous to early Tertiary landscapes are well-known from the Arkansas and Gulf Coast area in the USA, equatorial Africa, along the Western and the Eastern Ghats of the Deccan Peninsula and the Gulf of Carpentaria, Australia (see Figs 3, 5 , 6 & l l a & b).According to the age of under- and overlying volcanics and sediments, these large bauxite occurrences were formed in an early Tertiary time interval on subsiding platforms in a topographic near-sea-level position and have been partly dislocated by younger post-bauxitic tectonic activities. The vertical displacement during the late Tertiary and Quaternary resulted in an increased relief (Table 2), and the deposits are actually situated at very different altitudes ranging between 400 m (Indus valley in the western border area of India) to more than 1600 m above sea-level in the Western Ghats, India. Similar situations are characteristic of the bauxite occurrences in the near-Atlantic areas of South America or equatorial Africa (Bardossy & Aleva 1990). The post-bauxitic evolution of the landscape depended on: 1 the type and rate of younger tectonic uplift; 2 the properties of the ferralitic duricrust; saprolites became selectively eroded, whereas alucretes and sil cretes formed resistant crests; 3 the intensity of younger polygenetic destruction of the duricru.st. The relict areas of early Tertiary plains and plateaus with duricrusts are marked by typical fea tures such as extended swamps, meandering rivers, lakes and dambos filled by reworked products of sil cretes, pebbles or pisolites, rich in iron or gibbsitic concretions. A good example for these post-bauxitic phenomena is provided by the Songea area in Tanzania (Mutakyahwa & Valeton 1995). The slopes of intersected valleys and along inselbergs, too, repre sent specific features of the early Tertiary morphol ogy. The reconstruction of post-bauxitic plateau borders in the Cataguasis area, Brazil was described by Beissner (1989) and Valeton et al. ( 1991) (Fig. 13). The well-developed saprolite below the bauxite func tioned as an aquifer and caused slope erosion, mass
movements, landslides, talus deposits, boulder streams by reworking of core stones, exposure of fresh parent rocks and the filling of valleys by laterite-derived material. Quaternary growth of peat and lignite in local depressions is typical of those environments. Climatic trends and evolution of subsurface drainage patterns
The plate tectonic and oceanographic situation during the early Tertiary is characterized by a still intact, warm, circum-equatorial marine current system (Frakes et al. 1994) (Fig. 3). The circum Antarctic cold current system did not exist before the Oligocene. The world-wide relief was very flat previous to the appearance of the alpine mountain belts. These morphotectonic conditions gave rise to a world-wide well-balanced warm and humid climate. The extended platforms and large river systems favoured extremely wet climate conditions. The lateral facies differentiation of the saprolite bauxite belts is connected with these climatic environments. With regard to the speed and direction of the groundwater flow -both in the vertical and the lateral direction -two end-members of the sapro lite-bauxite-laterite association can be distinguished (Valeton 1983a) (Fig. 14): 1 bauxite formed above the groundwater table 'bauxite in uplifted areas'; 2 bauxite formed below the groundwater table 'bauxite on subsiding platforms'. An interaction between sea-level oscillations and groundwater movements promotes an extreme chemical alteration in rocks, which is related to the vertical and lateral migration of solutions and repre cipitation during feralitisation (Gordon et al. 1958; AIeva 1983; Valeton & Wilke 1993). A greenhouse resulting from an elevated C02 content in the atmosphere (Beck et al. 1995; Fawcett et al. 1995; Nesbitt et al. 1995; see fig. 1 of Miller et al. 1987 in Flower & Kennett 1994) additionally caused the strong chemical weathering during that time span. The regional extension of lateritic-bauxitic duricrusts therefore could extend far beyond the actual tropical climate zones in the northern and southern hemispheres. The Palaeogene bauxite boundary to the north is located on basalts in the Antrim massif of Northern Ireland (Smith & McAllister 1987), and the southernmost bauxite deposits are found in Lages, South Brazil (Melfi &
171
Pangaean saprolite-bauxite facies
Ll
N
0
500
1 000 m
� roads landslides, reworked plains (with direction of slide) !'), sharp incised val leys • exposed fresh parent rock l i m it of valley floor -c�·:, boulder streams, talus deposits pits • •
llll!ll in situ bauxites on planation plains !=:: : ::::) val ley fill (white silt and clay) + 2 1 o02'S G',;i·;.;_o� Boulder fields c:J landslide masses
Fig. 13. Morphotectonic units of the 'half-orange' topography at the Sulamericana plain with early Tertiary in situ bauxites over the charnockite belt. Late Tertiary to Quaternary destruction of the surface resulted from neotectonic uplift. The actual groundwater table lies at the base of the weathering profiles. Valley incisions in the deeply weathered landscape are followed by mass movements and landslides. Downslope displacement of lateritic material causes exposure of fresh bedrock in the upper steep slopes and swampy block fields in the basal parts, which locally are covered by peat. Block-filled streams consist of fresh rocks, bauxitic cortex with shelly textures, bauxites and saprolites. Fine kaolinitic material covers the valley floors (Beissner 1989).
Carvalho 1983), and in the southern mainland of Aus tralia and Tasmania ( Loughnan & Sadleir 1984).
R E L AT I O N S H I P B E T W E E N LANDSCAPE
E VO L U T I O N ,
G R O U N D WAT E R R E G I M E A N D Palaeogeography of saprolite-bauxite belts within ferralitic duricrusts
Ferralitic duricrusts extend across the early Tertiary landscape as widespread specific weathering prod ucts. Supergene alteration over parent-rock precon centrations of iron, manganese, nickel, copper, gold and phosphates form economically important ore deposits ( Valeton 1 994). In contrast to many of those deposits that are found far from coastal regions, the supergene concen tration of aluminium, caused by separation of silica and aluminium, is related to near-coastal zones with optimal humidity and drainage. Thus, the morpho tectonic and climatic evolution of the Pangaea continents limit the distribution pattern of the (late Mesozoic) early Tertiary bauxite belts.
M I N E R A L O G I C A L - C H E M I C A L FAC I E S P AT T E R N I N S A P R O L I T E - B A U X I T E FAC I E S
( Q UA L I T Y PATT E R N I N BAUX I T E S )
Landscape evolution and the groundwater regime within the ferralitic duricrusts have caused a well developed vertical and lateral zonation of the weath ering products. It is not by accident that saprolite or bauxite or their absence occur within the duricrust. Exploration and mining of bauxite are based on these vertical and lateral variations of bauxite distribution. Bauxite quality does not depend only on parent-rock petrology, rather, quality variations are determined mainly by the pedogenic environment. Mineralogy, chemistry, texture and porosity of the bauxite deposits differ with respect to their occurrence on
172
I. Valeton Relationship between groundwater and
CD
main elements vertical and lateral removal impregnation
vertical differentiiltion of alteration profile
direction ofdrain;:�ge
®
0
Fe + � � s· Fe, AI G .W.-
I
1,
weak
extremely good
weak
downward
mainly upward
weak
mainly relic
mineralogy
@
type of alteration profile
gi,go (he)
- mainly neo formatio n - meinly ne9 formation - mainly relic -he, ka - g i lbo, di) - ka - s m
hilly highlands or flat platforms and in accordance with the tectonic and morphogenetic history of the plain on which they formed (Valeton 1972; Aleva 1984;Wilke 1987; Bardossy & Aleva 1990;Tardy 1993; Valeton & Wilke 1993). Determing the chemical reorganization that has occurred during the processes of saprolite and bauxite formation is based on various methods: Millot & Bonifas (1955) developed the isovolumetric method, which stems from the fact that a certain volume of the parent rock is replaced by the same volume of saprolite or bauxite. Other methods use the assumption that certain minerals (zircon) or ele ments (Zr, Ti) remain immobile during weathering. No single mineral or element, however, is completely stable in an open system; thus only approximations of the chemical balance of the system are possible. With the help of multivariate statistical methods, for example, cluster analysis, trends of similar behaviour of the chemical constituents during weathering can be derived. The comparison of mean values of parent rock chemistry with that of the alteration products could indicate at least tendencies of the enrichment or depletion behaviour of elements during weath ering. Only the combination of several of these methods make it possible to form a good approxima tion of the natural systems. Initial and diagenetic formation of bauxite above the groundwater table
These bauxites develop on hilly landscapes as a result of vertical drainage (Fig. 14). They occur in the
mainly neo formation
-ka {he) - ka ldi, bo)
Fig. 14. Relationship between groundwater table and type of alteration (after Valeton 1983a): (1) formation of bauxite at various levels above the water table without separation of AI and Fe; (2a) low silica bauxite and (2b) high-silica bauxite at the top of the section near the surface of the groundwater level, with strong separation of AI and Fe; (3) formation of flint-clay below groundwater level by total removal of Fe.
more continental parts of the Pangaea-derived land masses and originate from parent rocks with low iron contents, such as phyllites, charnockites, anorthosites and syenites. Direct transition from parent rock into bauxite may occur without inter mediate saprolite in optimally drained crest areas and on upper slopes. This pure bauxite facies merges, with retarded water circulation, vertically and laterally into a bauxite-saprolite facies or into a pure saprolite. Good examples of this type are the bauxites over charnockites in the Cataguasis area of the Serra da Mantiqueira (Fig. 15) and over nepheline syenites at Poc;os de Caldas, Brazil (Valeton et al. 1997). Vertical evolution
The transition zone from fresh parent rock into the duricrust is characterized by the formation of core stones grading into saprolite on lower slopes or less well-drained areas (Fig. 16, locality 7) and directly into bauxite on well-drained upper slopes (Fig. 16, locality 6). The transition from fresh parent rock into bauxite is achieved under conditions extremely con ducive to the extraction of silica, thus preventing the formation of aluminium silicates. The saprolite is distinguished by excellent relict textures; kaolinite and halloysite are the only Al silicates; high-A! goethite and high-A! haematite replace mafic minerals; three-layer silicates and maghemite are absent in these profiles. Residual min erals can be quartz, illite and heavy minerals (Fig. 17). Diagenetically generated kaolinite and iron minerals
173
Pangaean saprolite-bauxite facies recent soil solum ,-'-.,
I I
�� 1�1 w
�
m
_
� � 1 .Q 1 � 'i3 � 1 � 1 t >Q
I �
Fig. lS. Model of the polyphase and polygenic history of a ferralitic weathering profile without saprolite for the Cataguasis region, south-east Brazil (unpublished report; Vale ton 1985) .
't
�
l c
't .c
:� I
]t
j
residual layer (quartz, heavy minerals)
eluvial horizon
I� I� g I ·� I� I 1 ·� I I I ··S� I u
l
yellow soil
::::� :::
red soil and residual breccia of bauxite, rich in gibbsite
.9 � � Ui =-=
:g
E
�
c�
actual 2 ground water level
(Iow-A! goethite and low-A! haematite) fill the pore spaces. The bauxite is composed of gibbsite and high-A! goethite, plus a small percentage of kaolinite, as the only neoformed minerals, and a relict texture is typi cally well-preserved. In these profiles iron is nearly immobile. Aluminium and Fe and their related ele ments stay together and are not separated. Quartz, garnet, pyroxenes and phosphates can survive as corroded residual grains of parent rocks. The foids (nepheline) of the magmatic rocks of Po«os de Caldas were dissolved at a very early stage, producing extremely high porosity and permeability. During diagenesis, a sequence of generations of gibbsite and goethite was formed. The first generation replaces parent rock minerals, the later ones appear as cutinae or as coarser grained crystals in pore spaces (Lemke 1986; Valeton et al. 1991 ; Schumann 1994). The AI substitution in iron minerals decreases with each younger generation. The mineral composition of the profiles depends on the parent-rock petrology (Fig. 16) and their chemistry reflects the parent-rock chemical composi tion. During ferralitic weathering on 'half-orange' topography over charnockites, the AI : Fe ratio has remained constant; whereas Si : Fe has decreased (see Fig. 22a). Parent-rock chemistry can be estimated from the ratio Zr/Ti02 : Si02 of the bauxites (Fig. 18). The trace element content in bauxites depends partly on preconcentrations in parent rocks: amphibolites possess an elevated preconcentration of Cr, Ni and V and low values of Zr and REE (Ce, La; Nd); gneisses are marked by elevated values of Ba, Sr, Ce, La and Nd and low values of Cr and Ni. Positive relationships exist between Fe203 and Ti, P, V and Cr. The correla tion of Ga with Al203 indicates an isomorphous replacement of AI by Ga in gibbsite. The elements Zr,
Ti and Nb are related positively, as are Ce, La and Nd. In weathering products on alkaline rocks, which gen erally possess higher contents of REE, Zr and Nb (owing to parent-rock chemistry), similar trends can be observed. These elements can be mobile, produc ing !ate AI-Ti-Zr-Nb-Ce-bearing precipitates (Melfi et al. 1992; Schumann 1994). Lateral evolution
Lateral differentiation depending on variable drainage activity is well expressed in the sections on charnockite (Fig. 4) and on alkaline rocks. On the same parent rock, bauxites with well-preserved relict textures pass laterally into kaolinitic saprolites, with conservation of relict textures. Neoformed textures are restricted to the infilling of pore spaces, joints and root channels. No indication of longer transport or of precipitation of AI and Fe in neoformed textures can be found in these bauxite types formed above the groundwater table. Initial and diagenetic formation of bauxite below the groundwater table
These bauxites are characterized by good vertical zonation of the profiles (Fig. 9) and by more or less well-pronounced lateral facies differentiation (catena) caused by the high mobility of subsurface drainage of solutions below the groundwater table (Fig. 8). On platforms, groundwater mobility is con trolled by sea-level oscillations. The lateral facies pattern of saprolite bauxites passes into bauxitic duricrusts high in iron. These are widespread over extended areas of former platforms and are by far the most frequent type of bauxite, occurring over plutonic and metamorphic basement
..... --..) .j:>.
soil
Locality 6. Section (Pit) 1 660 Sample No - 1 1 03 1T 0 - 1 1 02 2
1 1 01
3
- 1 1 00
4
- 1 099 [ 1 098] - 1 097 �
bauxite 5
B�x
8
- 1 096 � [1 095] - 1 094 [ 1 093] 1 092 -
9
1 091
6 7
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1 00 0 20 40 60 0 50 o joint filling Si02 •AI203 •Fe203 M i neral Composition (wt. %) •
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Locality 7. Section Sample No 519 520 - 521 522 523 - 524 - 525 - 526 527 - 528 1 082 - 1 08 1 _
Gi
.. � � I:
soil
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: [ 1 086]
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1
I�
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0 20 40 60 o joint fi lling • Si02 •AI203 •Fe203 /secondary 0 1 2 bet/rea I + Ti02 pipe filling
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100 50 0 Mineral Composition (wt. %)
Fig. 16. Mineral composition and chemistry of a purely bauxitic profile on amphibolite and of a saprolite-bauxite profile on amphibole-biotite gneisses, Cataguasis region, south-east Brazil (Gi, gibbsite; K, kaolinite; Qtz, quartz; An, anatase; Fe, iron minerals) (Beissner 1989).
Pangaean saprolite-bauxite facies
175
rocks and over basalts. Bauxite belts over sandy-clayey sediments on the initially low-lying platforms along the border of passive shelves cover parts of the Guyana platform, the Amazon Paranaiba Basin, the equatorial African zone and the border parts of the Deccan Traps basalt in India, and they occur on the Early Cretaceous sandstone around the Gulf of Carpentaria, Australia (Plate 2). In many parts of Pangaea these bauxite-covered plat forms were uplifted during late Tertiary and Quater nary times. Their relics are preserved on high plateaus of the inselberg landscapes. (a)
Vertical evolution
(b)
(c) Fig. 17. (a) Relict texture of charnockite in a kaolinitic saprolite, showing the outlines (goethite) of mafic minerals, Cataguasis, Brazil. (b) Relict texture of goethite, pseudomorphous after garnet; bauxite on charnockite, Cataguasis, Brazil (Lemke 1 986). (c) Relict staurolite grain replaced by kaolinite in saprolite, Ricanao plateau, Suriname (Valeton 1971).
Well-examined examples are the bauxites over Deccan Traps basalts (Valeton 1967, 1972; Valeton & Wilke 1993). The saprolite may be several metres or tens of metres thick, always with well-preserved relict textures. It can be purely kaolinitic or underlain by smectitic saprolite (Fig. 9). Apart from subordinate occurrences of goethite, haematite is the dominant iron mineral, although secondary maghemite does occur as the typical iron mineral, but mainly in smec titic saprolites. Ilmenite skeletons are frequent relict minerals. The diagenetic transformation into sapro lite is expressed by replacement textures and by occasional infilling of pore spaces by kaolinite and haematite. The bauxite horizon is several metres or more in thickness and is composed initially of primary gibb site and kaolinite. Iron frequently possesses a higher mobility than aluminium, which can lead to a pure yellow or whitish bauxitic layer. This bauxite horizon can be capped or replaced laterally by lenses or hori zons rich in iron. The main minerals are gibbsite, often with kaolinite, haematite and subordinate goethite. The diagenetic transformation of alitic or ferralitic horizons is characterized by a texture change from relict to neoformed gel-like, spongy, vesicular, nodular and pisolitic textures (Fig. 19), indi cating the relatively high mobility of aluminium and iron, which become redistributed vertically and later ally. The diagenetic change in texture is accompanied by changes in mineralogy, which cause a separation of Si, AI and Fe accompanied by crystallization of coarse-grained gibbsite. Lateral evolution
The lateral differentiation of saprolites and bauxites that evolves below the groundwater table is discussed
176
I. Valeton
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AMPHIBOLITE (POINT 5+6) @ Parent rock "' Bauxite (·parent rock 'group I') 1' Saprolite x Joint filling .. Soil AMPH.-BT. - GNEISSES ( POINT 7+8)
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GARNET - G N EISSES ( POINT 1 4)
@ Parent rock o
+
0.01 Zr/Ti02
Bauxite (•parent rock 'group I I ' ) Saprolite Joint fil ling Soil
Bauxite (•parent rock 'group Ill') Soil
GARNET - GNEISSES (POINT 1 4) 0·1 ® Parent rock " Bauxite (•parent rock 'group IV') o Saprolite • Soil Z x
Red clay (point 1 2) Yellow clay (point 1 2 )
Fig. IS. The Si02 : Zrffi02 ratio permits classification of ferralitic weathering products with respect to the parent rock chemistry, Cataguasis region, south-east Brazil (Beissner 1989).
with reference to occurrences on flood basalts in India (Fig. 8). The transformation of basalts into saprolites, fer ralites and bauxites is a process of almost complete neomineralization, mainly via solution. All elements pass into solution, the sequence and rate of their reprecipitation depend on environmental factors, such as chemistry of the primary and secondary mineral phases, on the concentration of elements in aqueous solution and on the intensity and rate of drainage. The lateral mineralogical evolution of the Box horizon is characterized by a transition from a facies rich in aluminium into a facies rich in iron. The dominant iron mineral is haematite in addition to traces of goethite. These two minerals occur in several subsequent generations, with a high rate of AI substitution in the beginning but with a decreasing trend of AI substitution in later generations. The lateral textural evolution in the Box horizon is
very impressive. A massive, hard bauxite -partly red and rich in iron -changes in the seaward direction into a yellow to white bauxite-poor in iron-with a boulder-like texture. Indurated bauxite concretions or nodules are embedded in a soft gibbsitic matrix. In near-coastal areas with a subsiding tendency during bauxitization, a very special white pisolitic and boehmite-diaspore facies occurs (Fig. 20). The nuclei of the pisoliths are built up of gibbsite or kaolinite, indicating well-preserved relict textures of basalt. The zonal formed cortex subsequently around pisoliths is built up by very fine-grained boehmite and can be replaced by a small percentage of coarsely crystallized diaspore. The pisoliths are intersected by desiccation cracks. This pisolitic gibbsite-boehmite and diaspore facies has developed below the ground water table, in depressions, indicating hydromorphic low pH and low Eh conditions. The environment was rich in humic acids and sulphur.
Pangaean saprolite-bauxite facies
177
(b)
(a)
(c)
(d) Fig. l9. Textural structures in bauxites. (a) Large shrinkage crack in high-quality bauxite on Deccan Traps basalt, B agru Hill, India. (b ) 'Brecciated' bauxite in uppermost part of the profile on charnockite, Shevaroy Hills, India. (c) Vesicular texture. (d) Nodular texture. (e) Relict texture of quartz and feldspar in bauxite on charnockite, Shevaroy Hills, India. (f) Pisolitic texture: nuclei with basaltic relict texture, kaolinite and gibbsite; zonal cortex built up of gibbsite and boehmite (rarely diaspore); matrix mainly gibbsitic; on basalt, Bagru Hill, India (all photographs: I. Vale ton).
The geochemistry of bauxites on basalts is delin eated by the presence of: 1 main elements -Si02, Al203, Fe203, Ti02 and H20;
minor elements - CaO, Nap, K20, MgO, MnO and P205; 3 trace elements -Ba, Ce, Co, Cr, Cu, Ga, La, Nb, Nd, Ni, Pb, Rb, Sr, Th, U, V, Y, Zn and Zr.
2
178
I. Valeton
(e)
(f) Fig. 19. Continued.
�
'?_.,_,.
bo = boehmite
bo
di = diaspere
bo 5 - 1 0%
less than 5%
bo more than 1 0% 12
bo,di
NALlYA •
locality
• alu-crete (bauxite) D kaolinitic & smectitic
�
saprolite
\
v
v
26
di �
di �
di
BHUJ •
J u rassic to B h uj Formation
N
t
Eocene to recent
Fig. 20. Distribution of diagenetic boehmite and diaspore in the white bauxitic Box horizon on the weathering profiles (striped) over basalt (hooks) in central Kutch, covered by Eocene to Recent sediments (white). The signature in the circles indicates the relationship between saprolite and alucrete. 'bo' and 'di' represent the amount of boehmite and diaspore in addition to gibbsite in the Box horizon. Areas of predominance of boehmite and diaspore indicate syn-pedogenic depression zones with high groundwater table, hydromorphic low-pH and low-Eh conditions (Wilke 1987).
Pangaean saprolite-bauxite facies
The main elements (calculated as oxides), i.e. Si02, Al203, Fe203 and Ti02, form 70-80 wt. % of the bauxite, the rest are minor and trace elements. The H20+ contents range from 10 to 30%. The decrease in Si02 and Fe203, and the increase in Al203 and Ti02 from basalt via saprolite to the Box horizon, demon strate the initial desilification and deferrication and the trend of relative concentration of Al203 (Figs 21 & 22). The mobilized iron is partly reprecipitated under absolute enrichment conditions in various horizons and partly extracted from the profiles and reprecipitated as siderite or pyrite in lacustrine and lagoonal sediments. Cluster analysis shows clearly the reorganization from sandstone or basalt geo chemistry into a new element association of the diverse saprolites and bauxite members with or without iron (Fig. 21a-c). Minor elements in the basalt are K, Na, Ca, Mg, Mn and P. Magnesium, which originates from pyroxenes and olivines, reappears in the smectite serpentine mineral group of the lower saprolite and is practically absent in the bauxites. Calcium occurs as impregnations of calcite in the pore spaces of various soil horizons, dominantly in the kaolinitic saprolite. The timing of calcite formation is still under discussion. The major calcitization took place during Quaternary calcrete formation, which overprinted the bauxite. Potassium is, to some extent, adsorbed on the interlayers of smectite in the lower saprolite, but is lost in the kaolinitic saprolite and bauxite. It seems to be the most mobile minor element in the system. Sodium is adsorbed by the smectites in the lower saprolite, forming a sodium-smectite. The Na20 content of 1.3-1.5% probably originates from seawater influence on the groundwater chemistry during bauxitization. Additionally, sodium occurs in the alunite and halite of Quaternary overprinted kaolinitic saprolite and bauxite. Manganese is a very mobile element that has been mobilized several times. As an initial step, it is bound in the smectite of the smectitic saprolite, but during younger polyphase overprints it becomes displaced down ward and locally reprecipitated, together with Fe, REE and other mobile trace elements. Phosphorous presents a particularly interesting pattern. It is adsorbed in the interlayers or on the surface of not only layer silicates but also iron and aluminium oxides. Its abundance therefore shows no upward decrease. The behaviour of trace elements can be subdivided into groups, depending on the mineral association of the various soil horizons. To define the relative mobil-
179
ity of these elements, titanium is regarded as the most stable element. The concentration or depletion of some trace elements in relation to titanium in the different soil horizons over basalts as parent rocks illustrates the geochemical trends during surficial alteration. The high adsorption capacity of smectite is responsible for high contents of elements such as Cu, Zn, Co, Ce, Nd, La and Y in the lower saprolite, whereas Ga, Zr and Pb are mainly concentrated, together with AI, in the bauxite. Vertical and lateral facies patterns related to landscape evolution and groundwater regime
The saprolite-bauxite facies within the ferralitic duricrusts is limited to optimal climatic zones and to zones of special groundwater regimes, forming two main facies types (Fig. 14). Above the ground water table on well-drained hilly landscapes bauxite could form directly from the parent rock without intermediate saprolite. This type of ferralite, poor in silica, is characterized by gibbsite and goethite as the only initial minerals. Bauxites formed below the groundwater table are underlain or laterally replaced by kaolinitic (Si : AI 1 : 1) or even smec titic (Si : AI 2 : 1) saprolites. Main initial minerals are gibbsite and haematite. Boehmite, diaspore and magnetite have grown diagenetically. The environ mental conditions of their formation are still under discussion. Bauxites with a pisolitic texture and boehmite con tents are known from many localities. Koster (1961) also describes diaspore in the bauxites of the Western Ghats, India. White bauxite with boehmite contents from equatorial Africa are mentioned by Bourdeau (1991) and Tardy & Roquin (1992). Tardy (1993) describes a very similar facies of white pisolitic bauxites, poor in iron but with elevated contents of boehmite from the bauxite blanket on the Famansa plateau in South Mali, Africa. He points out that the facies of white boehmitic bauxite is situated mainly in the central part and not at the borders of the plateau. The dehydration connected with the transformation from red into white bauxite seems to be controlled by groundwater conditions. The regional increase of boehmitic bauxite towards the north of West Africa is interpreted by Tardy (1993) as the result of palaeocli matic zonation since early Tertiary time. Boehmite bearing pisolitic kaolins occur in Late Cretaceous sediments of Upper Egypt as a distal member of a flint-clay facies (Germann et al. 1994). Boulange (1984) points out the close genetic relationship =
=
I. Valeton
180 -0.08
-0. 1 1
0.27 0.44 0.62 0.79 0.97
0.10
0.07
0.25 0.43 0.61 0.79 0.97
Si02 Na20 MgO CaO Fe203 MnO Ni K20 Rb Ba P203 Al203 Ga Zr Ti02 Nb Cr Th
Si02 Na20 MgO CaO Ni MnO K 20 Rb Ba Sr P203 Ce La Nd y
Co Zn Cu Al203 Ga Zr Ti02 Nb Cr Th
u
Pb Sr Ce La Nd
y
u
Co Zn Cu
Pb Fe203
v
v
BSL - .... - BXT
BSL - .... - BXT (a)
-0.08
samples poor in iron
0.10
(b)
including samples rich in iron
0.27 0.44 0.62 0.79 0.97 Si02 MgO Na20 K20 CaO Th Zr y
Co Cu Ni Zn Al203 Sr u
Cr Ga Rb Fe203 v
P203 MnO ,-----i Ti02 Nb Ba Ce La Nd Pb MAMUARA ( 1 0) (c)
bauxite on sandstone
Fig. 21. Cluster analysis (dendrograms) of selected lateritic duricrusts on basalt and sandstone in Gujerat, India: (a) bauxites on basalt poor in iron; (b) bauxites on basalt rich in iron; (c) bauxites on sandstones. Note the different element associations of the three types of bauxites. T11e correlation coefficient (0.0-0.99) increases from the left-hand to the right-hand side. It indicates a similar behaviour of elements during bauxitization, which leads to specific supergene element associations, as a result of the chemistry of residual or neoformed minerals. The relationship of Al2 03-Ga-Zr TiOrNb-Cr-Th-U-Pb of bauxites on basalts, poor in iron, is still present in bauxites rich in iron, but Fe and V show another relationship. (Wilke 1987).
181
Pangaean saprolite-bauxite facies
60 50 40 30 20 10
10
Al203
I
10 20 30 40 50 60 70 80 90
Fe203*x5
(a)
Fig. 22. Triangular diagram of Al203-Fe203-Si0 2 showing the fields of chemical variation during weathering. (a) Fresh alkaline parent rocks, hydrothermally affected 'potassic rocks', saprolites, bauxites and solum in the Po'
between pisolitization and boehmite formation and explains the occurrence of boehmite as a process of dehydration resulting from exposure of bauxite at the surface of plateaus. This idea is also expressed by Tardy (1993). Certainly, more observations of the regional facies distribution of gibbsitic and boehmitic bauxite and of processes of deferrification under reducing conditions would lead to a better under standing of boehmite or diaspore genesis. Experi mental precipitation of aluminium hydroxides and oxyhydrates in the presence of organic ligands, sulphates or chlorides, leads to pseudoboehmite (Violante & Violante 1980; Violante & Jackson 1981; Violante & Huang 1985; Prodromou & Pavlantou Ve 1995). A model for diaspore genesis under low temperature and low-pressure conditions is still requested.
Al203
10 20 30 40 50 60 70 80 90
Fe203 *
(b)
Al203
10 20 30 40 50 60 70 80 90
Fe203
(c)
F O R M ATI O N
OF S O L U M
ATT R I B U TA B L E T O P O LY P H A S E
AND
YO U N G E R
P O LY G E N E T I C
P E D O GENESIS
The soils developed b y post-bauxitic pedogenic overprinting reveal a polyphase and polygenetic history. Soils formed over saprolites and bauxites are described as 'solum', a term also used by Felix Henningsen (1994). Mechanical and chemical alter ation by 'degradation of bauxites' takes place as a result of climatic change, uplift, valley intersection and dislocation of the groundwater table. Late Ter tiary and Quaternary times are distinguished by: 1 truncation of bauxites and erosional landforms; 2 formation of solum in different climatic zones under modified groundwater conditions.
182
I. Valeton
Truncation of bauxites and erosional landforms
From their near-sea-level altitude of formation, some bauxites must have been uplifted to various topo graphical levels as blocks by younger tectonic activ ity. These bauxites, which normally are not overlain by sediments, have been exposed to mechanical and chemical alteration. The type and intensity of this alteration depended on variations in climate, vegeta tion and modification of the post-bauxitic drainage pattern. Intact planation plains on larger plateaux in tropical areas were characterized by a dense drainage pattern, meandering rivers, swamps and dambos, which resulted from the stagnation of the water table above 'intact' saprolites. After dissection of the land surface by rivers, the groundwater and drainage conditions changed dras tically. Softer and less hardened parts of the weath ering crust especially were eroded and removed. Weathering profiles have been capped to various depths, leaving behind a truncated landscape. Its surface is often littered with scattered blocks left behind from the former iron crust, residual bauxite breccia or core stones from the B/C horizon. In upper slopes, the hard duricrust or even fresh parent rocks are exposed in steep rocky scarps (Fig. 13 ) . Headwater regions and source horizons on slopes mark the previous and often the present position of the groundwater table in the weathering blanket. Landslides, mass movements and subsequently inter sected rivers are the result of the periodic excess of runoff during rainy periods, creating a very specific slope morphology. Boulder streams and coarse grained talus material of core stones, saprolites, fer ralites and bauxites form block fields along the lower slopes and provide evidence for younger, intensive wash-out of the clay-silt matrix, which accumulates in lower flat valleys ( Kehlenbeck 1986; Beissner 1989) . Formation of solum under different climates
Solum formation by polyphase alteration of bauxites during the late Tertiary and Quaternary is described by Valeton (1985), Wilke (1987), Beissner (1989), Lucas (1989), Valeton et al. (1991, 1997), Valeton & Wilke (1993), Sohumann (1994), Truckenbrodt et al. (1995). Tardy (1993) discusses the preservation and exposure of bauxites in various present climatic zones. Good examples of the variation of solum with changes in climate from north to south occur in Brazil. Successive steps of alteration with time are: 1 brecciation of the topward part of the in situ bauxite;
2
red soil formation from bauxite as parent rock
( possibly during the late Tertiary) ; 3 yellow soil formation ( possibly during the Quater nary) ; and 4 black soil formation ( present-day soil formation ) .
Amazon region
In northern equatorial Brazil Lucas (1989) and Truckenbrodt et al. (1995) mention the polyphase alteration of bauxites in the Amazon Basin under tropical conditions. They describe the vertical sequence from indurated bauxite with well preserved sedimentary relict textures upward into various zones of hard but strongly reorganized bauxite towards the top of the profile. The compact bauxite transforms into a brecciated bauxite and finally into a 'nodular' facies, which passes upward into a very thick yellow in situ clay material ( known as Belterra Clay) (Fig. 12). Lucas (1989) determined the mineralogical and chemical alterations in the dif ferent zones, which clearly show the changes in mineral stability during the initial phase of bauxitiza tion and during the Neogene and Quaternary phase of solum formation. South-east Brazil
The intensity of solum-formation decreases from the Amazon Basin via Serra da Mantiqueira in the north to Lages in the south. The surface-exposed bauxites over the Precambrian charnockite-amphibolite belt on the Sulamericana plain in the Serra da Mantiqueira were uplifted during Miocene to Qua ternary times. Subsequently, they were submitted to polyphase soil formation (Fig. 15; Table 2) ( Beissner 1989; Valeton et al. 1991). The bauxite profiles grade upward into a zone of: red in situ brecciated bauxite, a red earth; and finally in a yellow soft earthy zone which is overlain by a black humic soil. On the top of these plateaus, these profiles seem to be continuous without any break. On smooth slopes, stone lines in the yellow zone and river dissection indicate an inter val of erosion and landsliding during Quaternary times. 1 Mineralogically, the brecciation of the indurated bauxite started from the surface by penetration of roots. During bauxite brecciation ( residual bauxite breccia ) , a soft red soil has been formed by the remo bilization of iron and aluminium and their reprecipi tation as haematite and of a second generation of gibbsite. 2 Downward crystallization of haematite and
183
Pangaean saprolite-bauxite facies 1825 0
Hm/Gt
0.2
0.6
C02org.wt.%
pHH20
5.02
5.06
AI subst. in 15
25
mole %
Kaoi/Gi
0.5
1. 5
'recent' soil yellow soil
2 red earth
Fig. 23. Change of mineralogical and chemical parameters during transformation from bauxite into yellow and black soils, Cataguasis region, Brazil. (a) Organic carbon increases toward the top of the recent soil profile producing lower values in pH. Towards the top, haematite is replaced by goethite, AI substitution in goethite increases and ratio of kaolinite : gibbsite contents oscillates; (b) concentration of AI, Fe and Mn soluble in oxalic acid (0) and dithionite (D) (Lemke 1986).
bauxite
3 4 m
5
(a)
0.01
FejFe,
AIJAI, MnJMn, 0.1
1
'recent' soil yellow soil
2 red earth bauxite (b)
3 4
m
5
coarse-grained secondary gibbsite II has taken place in the pore space of the indurated bauxite. Kaolinite seems to have remained stable and quartz has become relatively enriched. 3 Strong chemical leaching during the younger Quaternary weathering has led to a yellow soft clayey soil (Fig. 23), resulting in the relative enrichment of quartz, stable heavy minerals and of kaolinite. Gibb site broke down in the Box horizon and eventually has been secondarily enriched in the saprolite. During yellow soil formation, from the underlying red earth or bauxite, gibbsite and haematite were the least stable, causing their strong depletion. Amorphous or weakly crystallized kaolinite and halloysite are neo formed precipitates in the yellow soil. The growth of fine-grained goethite with an extremely high substi tution by aluminium (up to 32%) causes the yellow colour. The content of aluminium and iron, dissoluble in dithionite, increases towards the top of the yellow soil. In the profiles, a segregation and downward dis placement of manganese and of the highly soluble REE took place. The lowermost part of the Box horizon and the saprolite form the actual aquifer. They are character ized by a very high porosity, leading to a 'skeletal bauxite' or a 'collapsed residual breccia'. Reprecipi tation of gibbsite and poorly crystallized layer sili cates (halloysite) takes place in the increasing pore space near the recent groundwater level. Manganese minerals, halloysite and REB-bearing minerals occur
in veins and layers in the basal part of the profiles. Elements such as Ba, Cu, Ni, Zn, Ce, La and Nd are concentrated together with Mn. The mechanical downward displacement of fine grained, colloidal material from the upper parts of the profiles during the rainy seasons provokes eluvia tion from the surface and a downward illuvation. The surface of the former bauxite can be covered by a residual sand layer, composed of angular quartz and dark heavy-mineral grains (Fig. 15). 4 The humic soil (20-80 cm thick), high in cor•' and the upper yellow zone of the solum are strongly related to the recent relief of the land surface. They are certainly the youngest members, which, on the hill tops or plateaus, form an in situ overprint on the fer ralitic profiles. On the slopes or on valley floors, however, stone lines of reworked bauxitic products occur at the base of the yellow soils or intercalated in it. Similar polyphase soils are developed on the baux ites of Poc;os de Caldas (Schumann 1994; Valeton et al. 1997) and on the bauxites of Lages, north of Porto Alegre in south-east Brazil (Valeton et al. 1988; Formoso et al. 1989). In the Lages area the bauxite is covered by only a very thin layer of yellow soil in the solum. Indian
Bauxites submitted to subtropical or arid climates, such as in India, present variable polyphase histories.
I. Valeton
184
ha = halite cc less than 5% = calcite CC 5 - 1 0% g y = gypsum CC more than 1 0% a/ = a l u n ite 1 2 loca l ity ap = apatite • alu-crete (bauxite) D kaolinitic saprolite D smectitic saprolite cc
NALlYA •
Fig. 24. Distribution pattern of Quaternary calcite, sulphates and halite in the Box horizon (striped) of the same locality with similar groundwater conditions as in Fig. 20(a) (Wilke 1987).
In the Bihar Mountains, Eastern Ghats, the bauxite duricrusts are overprinted by residual brecciation of laterites and bauxites and are now overlain by so called 'grey cotton soils' (Valeton 1984). In Kutch, Gujerat, the bauxites were overlain by marine sedi ments during middle to late Eocene and Miocene time. After erosion of the overburden, Quaternary pedogenic alteration under arid conditions has led to saltcretes and calcretes (Fig. 24). In general, forma tions of solum over lateritic bauxites in arid regions are characterized by calcretes, saltcretes and analo gous soils. The distribution patterns of younger polyphase mineral associations in the Quaternary solum show a certain similarity to the distribution pattern of boehmite and diaspore (Fig. 20). Possibly, analogous environmental conditions during diagenesis and during Quaternary alteration are responsible for the concentration of boehmite and diaspore as well as of alunite, gypsum, halite and calcite in the same struc tural depression (Fig. 24). O B L I T E R AT I O N
OF P A R E N T R O C K
G E O C H E MI S T RY A N D C R E AT I O N O F S U P E R G E N E A L - D O M I N AT E D G E O C H E M I C A L P R OV I N C E S
To compare the element associations o f parent rocks, saprolites, bauxites and solum, parameters such as
parent-rock texture, mineralogy and chemistry, and vertical and lateral facies changes during the differ ent steps of weathering have to be defined (see Table 2).
Aluminization took place from late Mesozoic to early Tertiary times. It was followed by 'solum' forma tion, which overprinted and degraded the 'lateritic bauxites' during late Tertiary and Quaternary times. The preservation of the resulting 'lateritic bauxite' blanket depended on younger morphotectonic activ ity and erosion. The formation of 'lateritic bauxites' is a supergene alteration process that takes place in an open system by the input and output of water and inorganic and organic solutions. The reactions between the solid phases of the parent rocks and the solutions led to the neoformation of H20-bearing layer silicates, oxyhydrates, oxides, phosphates and traces of other minerals. The formation of 'solum' during younger phases of the climatic and hydromorphic evolution led to additional adaption of the chemical and mineralogi cal system in the bauxite belts. In tropical regions red earth probably was produced during the late Tertiary, whereas the yellow earth and black soils are believed to have been formed during Quaternary times. In arid regions, calcretes and saltcretes are typical 'solum' products of arid climates on lateritic bauxites. Altogether, lateritization and solum formation
185
Pangaean saprolite-bauxite facies
have resulted in a nearly complete obliteration of the original 'parent-rock element association' and the development of the new 'supergene element associa tion'. For each element the amount of its depletion from the system depends on the stability of the primary (parent rock) minerals and of the mineral phases neoformed during the successive steps of alteration. Even stable elements such as aluminium or titanium may have been mobilized, e.g. as organa complexes and removed from the system. The loss of material during lateritic weathering is very high. Porosity in saprolites may rise up to 30 or even 60 vol. % , whereas in bauxites porosity varies between 20 and 50 vol.% . I n Fig. 22b & c the fields o f main element associa tion are shown for parent rocks, saprolite-bauxite and the solum over charnockites and alkaline rocks. Weathering products over all types of parent rocks reveal a similar trend of relative or absolute enrich ment or depletion of elements (Fig. 25a & b). In lateritic bauxites, the main elements AI and Ti, frequently also Fe and P, become enriched. Concen trations of trace elements such as V, Pb, Ga, Nb, Zr, Th and U are increased. Depending on their primary contents in parent rocks, these secondary element concentrations may achieve considerable economic interest. The gailium content for instance, which is used in the semiconductor industry, can be greatly increased in bauxites. This element, which is not enriched in residual minerals, seems to replace Al in the lattice of gibbsite and Fe in the lattice of iron min erals. In bauxites on basement rocks the concentra tions of Ga vary between 35 and 40 p.p.m., and on alkaline rocks of Poc;:os de Caldas the Ga content rises to very high concentrations of around 150 p.p.m. Even in a totally altered weathering mantle, however, some parent-rock parameters such as texture, the structure of residual minerals and the ratio of stable elements (Fig. 18) still permit the determination of parent-rock petrology. After detailed petrographical studies, cluster analysis (see Fig. 21a-c) can be useful to recognize the common heritage of certain element associations. In the den drograms the bauxites from Mamuara (Fig. 21c), for instance, derived from sandstones with layers of stable heavy minerals, display a very close relation ship between Zr, Th and Y (zircon), Fe, V, P, Mn, Ti and Nb (opaque minerals) and Ce, La and N d (mon azite). The geochemistry of bauxite on basalts, which contain only a small percentage of opaque minerals (ilmenite, magnetite) and which have been submitted to strong chemical alteration during bauxitization
D Phonol ite ( n = 1 8)
0 Bauxite ( n= 27) + Solum ( n = 39)
10
0.1
0.01
Si02 1102 Al203 Fe203Mn0 MgO CaO Na,O K20 P205 Lol (a) 1 oo :r�������_L_L�
10
0.1
0.01
Cr N i Co
V
Cu P b Z n Rb B a Sr Ga N b Z r Y T h U L a C e N d
(b)
Fig. 25. (a) Major element pattern and LOI of phonolites, nepheline syenites, sa pro lites, bauxites and solum, normalized to average alkaline rocks (nepheline syenites and phonolites) of Po�os de Caldas (Schumann 1994). (b) Trace element pattern of phonolites, nepheline syenites, saprolites, bauxites and solum, normalized to average alkaline rocks (nepheline syenites and phonolites) of Po�os de Caldas (Schumann 1994).
(Fig. 21a), is nevertheless characterized by the domi nant association ofAl with Ti, Zr, Ga and Nb, and also the elements Cr, Th and U are very closely associated with Al. In bauxite facies, members rich in iron (Fig. 21 b), Fe and V may be reiated. The cementation of land surfaces by duricrusts is not limited only to ferralitic weathering processes. After the tectonic reorganization of Pangaea and the appearance of the alpine mountain chains, the evolu tion of a weathering sequence was established in the time interval from the end of the Mesozoic to the Holocene (Table 2). This weathering sequence starts with ferralitic duricrust formation and is followed by a differentiation of landscapes and of soils. The super-
186
I. Valeton
gene geochemical reorganization during this late Mesozoic to Holocene weathering sequence is one of the most important geochemical processes during Earth's history, having created Al-dominated geo chemical provinces on a global scale. R E FE R E N C E S
ALEVA, G.J.J. (1965) The buried bauxite deposits of Onver dacht, Suriname, South America. Geol. Mijnbouw, 44, 45-58. ALEVA, G.J.J. (1979) Bauxitic and other duricrusts in Suri name: a review. Geol. Mijnbouw, 58, 321-336. ALEVA, G.J.J. (1981) Bauxitic and other duricrusts on the Guiana Shield, South America. In: Lateritisation Processes, Proceedings of International Seminar on Lat eritisation Processes, December 1979, Trivandrum India (Ed. Banerji, P.K. ), pp. 261-269. Oxford and IBH Publish ers, New Delhi (also Balkema, Rotterdam). ALEVA, G.J.J. (1983) On weathering and denudation of humid tropical interftuves and their triple planation sur faces. Geol. Mijnbouw, 62, 383-388. ALEVA, G.J.J. (1984) Lateritization, bauxitization and cyclic landscape development in the Guiana Shield. In: Bauxite Production 1984 (Ed. Jacob, L., Jr.), pp. 297-318. B auxite Symposium, Los Angeles. American Institute of Mining Engineers, New York. BA.RDOSSY, G. (1993) Carboniferous to Jurassic bauxite deposits as paleoclimatic and paleogeographic indicators. PANGEA: Global Environment and Resources, Mem. Can. Soc. petrol. Geol., Calgary, 17, 283-293. BA.RDOSSY, G. & ALEVA, G.J.J. (1990) Lateritic Bauxites. Developments in Economic Geology, Vol. 27 Elsevier, Amsterdam. BECK, R.A., BURBANK, D.W., SERCOMBE, W.J., OLSEN, T.L. & KHAN, A.M. (1995) Organic carbon exhumation and global warming during the early Himalayan collision. Geology, 23(5), 387- 390. BEISSNER, H. (1989) Geologie, Mineralogie und Geochemie der Bauxite aufpriikambrischem Basement im Gebiet von Astolfo Dutra, SW-Catguases, Brasilien. Dr. thesis, Uni versitat Hamburg. BELINGA, S.M. (1972) L'alteration des roches basaltiques et le processus de bauxitization dans l'Adamaoua (Camer oun). These de doctorat d'etat, Universite Pierre et Marie Curie, Paris; no. A. 0. 6562. BosKI, T. (1987) Lateritic bauxites from SE Guinea Bissau. Dr. thesis, University of Bruxelles. BoULANGE, B. ( 1984) Les formations bauxitique lateriques de Cote d'Ivoire. Trav. Doc. ORSTOM, 175,341 pp. BouRDEAU, A. ( 1991) Les bauxites du Mali. Geochimie et mineralogie. These, de l'Universite Louis Pasteur, Strasbourg. DE WEISSE, G. (1954) Notes sur quelques types de laterite de Ia Guinee Portugaise. Proceedings of the 19th Interna tional Geological Congress, Alger, 1 992, XXI, 171-179. FAWCEIT, PJ., GIBBS, M.T. & BARRON, E.J. (1995) Changes in continental hydrology through geologic time: the role of paleogeography and atmospheric carbon dioxide. The 1st SEPM Congress on Sedimentary Geology, 13-16 August, St. Pete Beach, Florida, Abstracts, p. 52.
FELIX-HENNINGSEN (1994) Mesozoic-Tertiary weathering and soil formation on slates of the Rhenish Massif/ Germany. Catena, 21, 229-242. FITZPATRICK, R.W. (1983) Occurrence and mineralogy of fer ruginous bauxites along the Eastern seaboard of South Africa. Geol. Soc. SAfr. Spec. Pub!. , 7, 87-96. FLOWER, B.P & KENNEIT, J.P. ( 1994) The Middle Miocene climatic transition: East Antarctic ice sheet development, deep ocean circulation and global carbon cycling. Palaeogeogr. Palaeoclimatol. Palaeoecol. , 108, 537-555. FORMOSO, M., RETZMANN, K. & VALETON, I. (1989) Fraction ation of rare elements in weathering profiles on phono lites in the area of Lages, Santa Catarina, Brazil. Geochim. Brasiliensis, 3(1 ) , 5 1-61. FRAKES, L.A., PROBST, J.L. & LUDWIG, W. (1994) Latitudinal distribution of paleotemperature on land and sea from early Cretaceous to middle Miocene. C.R.Acad. Sci. Paris, 318(serie II), 1209-1218. GERMANN, K., SCHWARZ, T. & WIPKI, M. (1994) Mineral deposit formation in Phanerozoic sedimentary basins of north-east Africa: the contribution of weathering. Geol. Rundsch. , 83, 787-798. GoRDON,M.,JR, TRACY, J.I.,JR & ELLis, M.V. (1958) Geology of Arkansas bauxite region. US. geol. Surv. Prof Pap. , 229, 268 pp. GRANDIN, G. (1976) Aplanissements cuirasses et enrichisse ment des gisement de manganese dans quelques regions d'Afrique de !'Ouest. Mem. ORSTOM, 82, 275 pp. GRANDIN, G. & THIRY, M. (1983) Les grandes surfaces conti nentales tertiaires des regions chaudes succession des types d'alteration. Cah. ORSTOM Ser. Geol, XIII(1 ) , 3-18. GRUBB, P.L.C. (1971) Genesis of the Weipa Bauxite Deposits, N E Australia. Mineral Deposits, 6, 265-274. HIERONYMUS, B. (1985) Etude de !'alteration des roches erup tives de !'ouest de Cameroun. Laterisation, bauxitization et evolution bauxitique. These de doctorat d'etat, Universite Pierre et Marie Curie, no. 85-20, Paris. HORN, G.F.J. (1983) Comments on the genesis of bauxite in Natal. Trans. Geol. Soc. SAfr. , 86, 99-104. IwANOFF, A. (1985) Sedimentary petrography of Lower Eocene rocks from the Kutch peninsula, State of Gujerat, India. Dr. thesis, University of Hamburg. KEHLENBECK, U. (1986) Beziehung zwischen der Bauxitbil dung auf der Charnockitserie und Reliefentwicklung im Gebiet von Mirai, nordlich von Catagueses/Brasilien. Thesis, University of Hamburg. KING, L.C. (1953) Canons of landscape evolution. Geol. Soc. London Bull. , 64, 721-752. KING, L.C. (1962) Morphology of the Earth. Oliver and Boyd, Edinburgh. KoSTER, H.M. (1961) Vergleich einiger Methoden zur Untersuchung von geochemischen Vorgangen bei der Verwitterung. Beitr. Mineral. Petrog. , 8, 69-83. LEMKE, P. ( 1986) GefUge, Mineralogie und Geochemie der Bauxite, Hill 812, nord!. von Mirai, M.G., Brasilien. Thesis, University of Hamburg. LOUGHNAN, F. C. & BAYLISS, P. (1961) The mineralogy of the bauxite deposits near Weipa, Queensland. Am. Mineral. , 46, 209-217. LouGHNAN, F.C. & SADLEIR, S.B. (1984) Geology of estab lished bauxite-producing areas in Australia. In: Bauxite Production 1 984 (Ed. JACOB, L., Jr), pp. 436-450. Bauxite
Pangaean saprolite-bauxite facies Symposium, Los Angeles. American Institute of Mining Engineers, New York. LUCAS, Y. (1989) Systemes pedologiques en Amazonie Bresilienne. Equilibres, clesequilibres et transformations. These, University of Poitiers. MAMEDOV, V.L., ANUFRIEV, I.K. & JAKUBOVITS, & SUMA, N.L. (1985) Particularities of the Sangaredi bauxite deposit, Guinea (in Russian). Izv. Vyssh. Uchebn. Zaved. geol. Razeved. , 4, 38-47. McFARLANE, M.J. (1976) Laterite and Landscape. Academic Press, London. MELFI, A.J. & CARVALHO, A. (1983) Bauxitization of alkaline rocks in southern Brazil. Sci. Geol. Mem. , 73, 161-172. MELFI, A.J., SOUBIES, F. , NAHON, D. & FORMOSO, M.L.L. (1996) Behaviour of Zirconium mobility in bauxites of the Southern Brazil. South America Geol. 7, 9(3/4). MILLOT, G. (1964) Geologie des argiles. Masson, Paris. (English edition Geology of Clay, 1970.) MILLOT, G. & BONIFAS, M. (1955) Transformation iso volumetriques dans les phenomenes de lateritisation et de bauxitization. Bull. Serv Carte geol. Alsace Lorraine, 8(1) , 3-19. MoRGAN, C.M. (1995) Geology of the spheres at Weipa. Trav. JCSOBA, 22(24), 61-74. MUTAKYAHWA, M.K.D. & VALETON, I. (1995) Late Creta ceous-Lower Tertiary weathering event and its laterite bauxite formation in Tanzania. Mitt. Geol. paliiont. Inst. Univ. Hamburg, 78, 1-66. NESBITT, B. E., MENDOZA, C. A. & KERRICK, D.M. (1995) Sur face fluid convection during Cordilleran extension and the generation of metamorphic C0 2 contributions to Cenozoic atmospheres. Geology, 23(2), 99- 101. OvERSTREET, E.C. (1964) Geology of the southeastern bauxite deposits. US. geol. Surv. Bull. , 1199A, 19 pp. PATTERSON, S.H., KURTZ, H.F., OLSON, J.C. & NEELEY, C.L. (1994) World bauxite resources. Geology and resources of aluminium. US. geol. Surv. Prof Pap. , 1076-B, 151 pp. PRODROMOU, K.P. & PAVLATOU-VE, A.S. (1995) Formation of aluminium hydroxides as influenced by aluminium salts and bases. Clays Clay Minera/. , 43(1), 1 1 1-115. RouiLLIER, J.P. (1990) Bauxite Deposits, Mining Operations and Producers of the World. Aluminium-Verlag GmbH, Dusseldorf. SAPOSHNIKOV, D.G., BOGATYREV, B.A. & BARKOV, V.V. (1976) The bauxites and weathering crusts of Guinea (in Russian). Kora Vyvetrivanya, Moscow, 15, 3-50. SCHAAP, A.D. (1984) Summary of Mineralogical Investiga tion of the Weipa Kaolin Deposit. Comalco Aluminium Ltd, Weipa. ScHAAP, A.D. (1985) Environmental considerations in developing a kaolin industry at Weipa. Proceedings ofthe North Australian Mine Rehabilitation Workshop (Ed. Lawrie, J.W.), 231-239. ScHUMANN, A. ( 1994) Chemische Verwitterungsprozesse auf Alkaligesteinen und Hinweise auf hydrothermale Ein wirkungen seit dem ausgehenden Mesozoikum, Poc;os de Caldas, Minas Gerais, Brasilien. Dr. thesis, University of Hamburg. ScHWAB, R.G., LIMA D A Co sTA , M. & D E OuviERA, N.P. (1983) About the formation of bauxites and phosphate laterites in the Gurupi region (Northern Brazil). Zen tralbl. Geol. Paliiontol. Part I, 3/4, 563-580. SIMON-CO!NyON, R., M I LNES, A.R., THIRY, M. & WRIGHT, M.J.
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(1996) Evolution of landscapes in northern South Aus tralia in relation to the distribution and formation of sil cretes. J geol. Soc London, 153, 467-480. SMITH, B.J. & McALISTER, A. (1987) Tertiary Weathering Environments and Products in Northeast Ireland (ed. Gardiner, V.) International Geomorphology 1986 Part II, pp. 1007-1031. John Wiley & Sons, Chichester. TAKAHASHI, K. & Jux, U. (1989) Palynology of Middle Ter tiary lacustrine deposits from the Jos Plateaux, Nigeria. Bull. Fac. Libr. Arts, Nakasaki Univ. Nat. Sci. , 29(2), 181367. TARDY, Y. (1969) Geochimie des alterations. Etude des arenes et des eaux de quelques massifs cristallins d'Eu rope et de !'Afrique. Mem. Serv. Carte geol. Alsace Lorraine, 31, 199 pp. TARDY, Y. (1993) ?etrologie des Laterites et des Sols Tropi caux. Masson, Paris. TARDY, Y. & RoQUIN, C. (1992) Geochemistry and evolution of lateritic landscapes. In: Weathering, Soils and Paleosols (Eds Martini, I.P. & Chesworth, W.), pp. 407-443. Else vier,Amsterdam. TRUCKENBRODT, W., KOTSCHOUBEY, B. & HIERONYMUS, B. (1995) Aluminization: an important process in the evolu tion of Amazonian bauxites. Trav. ICSOBA, 22(24), 2742. TwiDALE, C.R. (1994) Gondwanan (Late Jurassic and Creta ceous) Paleosurfaces of the Australian craton. Palaeo geogr. Palaeoclimatol. Palaeoecol. , 112, 157-186. U.S. BuREAU OF MINES (1995) Mineral Commodity. Sum. Jan. 1995. VALETON, I. (1967) Laterite und ihre Lagerstatten. Fortschr. Mineral. , 44(1), 67-130. VALETON, I. (1971) Tubular fossils in the bauxites and the underlying sediments of Surinam and Guyana. Geol. Mijnbouw, 50(6), 733-741. VALETON, I. (1972) Bauxites. Developments in Soil Sciences I. Elsevier, Amsterdam. VALETON, I. (1983a) Paleoenvironment of lateritic bauxites with vertical and lateral differentiation. In: Residual Deposits (Ed. Wilson, C.). Spec. Pub!. geol. Soc. London, No. 11, pp. 77-90. Blackwell Scientific Publications, Oxford. VALETON, I. (1983b) Paleogeographic interpretation of the world-wide distribution of oxisols (laterite, bauxite) in the Lower Tertiary. 5th International Congress of ICSOBA, Zagreb, Yugoslavia, 26-28 Septembe1; Vol. 13(18), pp. 1 1-13. VALETON, I. (1984) The Cretaceous-Tertiary history of the B agru Hill area, Bihar/India and the paleoenvironment of their bauxite formation. In: Products and Processes of Rock Weathering, Recent Researches in Geology (Eds Sinha-RoY, S. & Gosh, S.K.), Vo1. 1 1 , pp. 74-81 . Hindustan Publishers F, India. VALETON, I. (1985) Time-space model in relation to tecto and morphogeneses of supergene ore formation. Monogr. Ser. Mineral. Deposita, 25, 197-212. VALETON, I. (1987) Bauxite und Kaolinlagerstatten in Aus tralien. Geowissenschaften Unserer Zeit, 5, 149-156. VALETON, I. (1991) Bauxites and associated terrestrial sedi ments in Nigeria and their position in the bauxite belt of Africa. I. Afr. Earth Sci. , 12(1/2), 297-310. VALETON, I. (1994) Element concentration and formation of ore deposits by weathering. Catena, 21, 99-129. VALETON, I. & WILKE, H. (1993) Tertiary bauxites in subsi-
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Spec. Pubis int. Ass. Sediment. (1 999) 27, 1 89-206
Karst bauxites: interfingering of deposition and palaeoweathering
G. B A R D O S S Y * and P. -J. C O M B E S t *Hungarian Academy of Sciences, Budapest, Hungary H-1387; t University ofMontpellier II, case 66, Geology of Ore deposits, Place Eugene Bataillon, 34095 Montpellier Cedex 05, France and CNRS URA n °] 405
A B S T R AC T
This article presents the main data for the characterization o f karst bauxites, starting with the different types of classification and followed by the distribution of the deposits in space and time. A short review is presented about the main geochemical and mineralogical features of the karst bauxites. The particular importance of the lithological and sedimentological features of karst bauxites is emphasized. The position of bauxite deposits within sedimentary sequences is analysed and compared with palaeogeographical reconstructions. These new approaches allow the reconstruction of the main scenarios leading to the for mation of bauxite deposits: autochthonous, parautochthonous, allochthonous and parallochthonous. In the following a simple typology is presented, based on palaeogeography and the tectonic position of the deposits, and demonstrated by some of the most representative examples. The attention is focused on karst bauxites as geodynamic indicator rocks, with respect to palaeoclimate, eustatic movements and tectonic instability. Thus the bauxites, particularly the karst bauxite deposits, must be considered as essential indica tors in the reconstruction of the external geodynamic processes. Finally, the new approaches and methods that elaborate the origin of karst bauxite deposits also can be applied successfully to the exploration of new deposits.
INTR O D U CTION
Since the discovery o f bauxite b y Berthier i n 1821 at Les Baux, Provence, south-eastern France, many investigations have been carried out on karst bauxite deposits. Starting in the second half of the 19th century, the exploration for new deposits reached a peak during the 20th century and encompassed the whole world. The results of these explorations demonstrated that the majority of the karst bauxite deposits are related to shallow, marine carbonate platform sediments. The filling of karst depressions and cavities by bauxite has been observed at many places, leading to the general acceptance of the term 'karst bauxites' for this particular type of deposit. Concerning the origin of the deposits, progress was much slower, because from the very beginning of the investigations a striking contradiction has been apparent: the two rock types, karst bauxites and car bonate rocks, so closely related in the deposits, are essentially different in their geochemical composi-
tion.The bauxites are extremely enriched in hydrates of aluminium, whereas carbonate rocks contain only traces of this element. In most cases it was impossible to prove the carbonate basement as being the exclu sive parent rock of the bauxite (note that the lateritic bauxites are accepted to be derived from their base ment of alumosilicate composition). Numerous and contradictory genetic theories have been elaborated in order to explain the origin of the karst bauxite deposits. Often local features have been generalized, without good reason, and attributed to the totality of this type of deposit throughout the world. During the last 30 years, however, considerable progress has been achieved by the application of new multidisciplinary approaches. At present we have a much better understanding than before of the geological situation and the lithological, geochemical and mineralogical composition of the karst bauxite deposits. The most important advances
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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were achieved when studies embraced the bauxite deposits and their basement and cover rocks. The results enabled palaeogeographical reconstructions to be made for the area of the deposits. Furthermore, new concepts of global tectonics, sequence stratigra phy and palaeoclimatic modelling have been applied to the study of the deposits. Palaeotectonic investi gations, particularly for the early phases of devel opment, have helped to explain the complicated relationships between rock fracturing and kars tification and the trapping of bauxite material. All these new findings have contributed to a much better understanding of the origin of karst bauxite. The deposits can be incorporated into a model of the 'sed imentary landscape', the development of which is controlled by climate, tectonic stability of the region, eustatic movements and by the sedimentary environ ment serving as a trap for the bauxitic sedimentary material. The limited extent of this article does not allow a synthesis of all the new results on karst bauxite deposits to be presented. Instead we highlight only those points that seem to us as essential for a better understanding of the new approaches, and which may show the new directions and perspectives for future research.
The karst bauxite deposits can be classified in dif ferent ways. 1 According to the internal structure of the deposits (Bardossy 1982): (a) Mediterranean-Jamaican type; (b) Kazakhstan type; (c) Ariege type; (d) Timan type; (e) Salento type. 2 According to sedimentary processes and type of bauxitization (Valeton 1976; Bardossy 1982; Combes 1990) associated with the deposits: (a) autochthonous; (b) parautochthonous; (c) parallochthonous; (d) allochthonous. 3 According to palaeogeography and tectonic posi tion of the deposits (Combes & Bardossy 1994): (a) intracontinental, medium to relatively high level; (b) pericontinental, medium to low-level; (c) low, insular platforms. Other classifications have been developed accord ing to mineralogical composition, geomorphology, hydrogeological position (Mindszenty et a!. 1987; D'Argenio & Mindszenty 1995) and the form of the deposits.
C L A S S I F I C AT I O N O F K A R S T BAUXITE D E P O SITS
Both bauxite and laterite are products of intense con tinental subaerial weathering. Laterite is composed of nearly equal parts of clay minerals (mainly kaolin ite), alumina, and ferric iron hydrate minerals and titanium oxides. Bauxite is distinguished from laterite by its higher content in alumina and iron min erals. Three main types of bauxite deposits can be proposed: 1 lateritic bauxite deposits, formed by in situ chemical weathering of underlying aluminosilicate rocks; 2 karst bauxite deposits, covering the more or less karstified surface of limestones, dolomites and, rarely, of marls; 3 Tikhvin-type deposits, overlying the eroded surface of aluminosilicate rocks, being the product of the erosion of lateritic bauxite deposits. According to our calculations, approximately 88% of the global quantity of bauxite produced by nature belong to the lateritic bauxite deposits; 1 1 % of them belong to the karst bauxites and only less than 1 % to the Tikhvin-type deposits (Bardossy & Aleva 1990).
D I S T R I B UT I O N
OF B A U X I T E
D E P O SITS IN S PACE A N D TIME
The majority of karst bauxite deposits are concen trated in seven regions, listed below in order of their overall tonnage: 1 the Caribbean area (mainly Jamaica, followed by Haiti and the Dominican Republic); 2 the northern part of the Mediterranean area (mainly France, Italy, Hungary, Croatia, Bosnia, Greece and Turkey); 3 the Ural Mountains and Central Asia (Russia and Kazakhstan); 4 eastern Asia (China and Vietnam); 5 the Irano-Himalayan area (Iran, Pakistan, Afganistan); 6 the south-western part of the Pacific (mainly Solomon Islands, Philippines); 7 the south-eastern part of the USA. These deposits are situated mostly on shallow water carbonate platforms of the tectonically mobile Tethys belt (Fig. 1). In contrast, the lateritic bauxite deposits are connected genetically to the more stable,
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large cratonic areas situated to the south of the Tethys. The oldest karst bauxite deposits are found in the Eastern Sayab Mountains, Siberia, Russia and are of Early Cambrian age. Large karst bauxite deposits were formed during the Middle and Late Devonian in the Ural Mountains, and during the Carboniferous in China. Very few karst bauxite deposits are known from the Permian and the Triassic periods. A large expansion of karst bauxite formation started with the Middle Jurassic in Croatia, Montenegro and Greece and its peak was reached during the Middle Cretaceous (France, Italy, Croatia, Bosnia, Greece, Turkey) . Large deposits of Late Cretaceous and Palaeocene to middle Eocene age occur in Hungary, Croatia and Bosnia. There was a second minimum of karst bauxite formation during the late Eocene and Oligocene, but again large-scale bauxite forma tion occurred during the Late Miocene and Early Pliocene in Jamaica. Karst bauxite deposits of Pleis tocene age occur on Samar island (Philippines) and on the Solomon islands. It also appears that karst
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Fig. 2. Relative intensity of bauxite formation during the Phanerozoic.
bauxite is being formed at present on some islands of the south-western Pacific. The relative amount of bauxite formation, expressed in Mt Myr-1, was calculated by Bardossy (1994) as shown in Fig. 2. The amount of eroded bauxite cannot be determined, but it must be related to the total amount of bauxite formed in the given time interval. Furthermore, it is likely that the bauxite of some allochthonous deposits was formed earlier and was redeposited several times before arriving at its present location. Thus the original time of bauxite formation for several Late Cretaceous to Palaeocene deposits was presumably the Middle Cretaceous. C H E MI C A L A N D M I N E R A L O G I C A L C O M P O S IT I O N
Bauxites are characterized b y high Al203, medium Fe203 and +H20, and low Ti02 and Si02 contents. This basic feature is common for lateritic, karst
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Karst bauxites
and Tikhvin-type bauxites. Major differences are found in the mineralogy: lateritic bauxites are pre dominantly gibbsitic, whereas karst bauxites are equally divided into gibbsitic, boehmitic and dias poric types, with transition between the gibbsitic and boehmitic, and between the boehmitic and diasporic types. Bauxites of mixed gibbsitic and diasporic composition so far have not been found as a primary paragenesis. The gibbsitic bauxites contain 4550% Al2 03, and the boehmitic and diasporic ones 50-60% . The Fe203 content of the karst bauxites varies generally from 15 to 25% . Reduced ferrous bauxites are srrb'ordinate relative to the ferric ones; the main iron minerals of ferric bauxites are haematite and goethite. Some Palaeozoic karst bauxites are enriched in berthierine (chamosite), and locally karst bauxites also are enriched in siderite. Bauxites that were covered by marshy sediments suffered dia genetic reduction in many places: the original iron minerals were partly leached and partly replaced by pyrite and marcasite. The true karst bauxites contain less than 10% Si02 , almost exclusively in the form of clay minerals, mainly kaolinite. This is in major contrast to lateritic bauxites, which in many places contains a relatively large percentage of residual quartz grains, e.g. Darling Ranges, Australia. Anatase is the main tita nium mineral, accompanied by subordinate rutile and ilmenite. The Ti02 content generally varies from 2 to 4%. Let us stress that bauxites are the most hydrated rocks on the Earth's surface: the chemically bound +H20 content may reach 30% in the gibbsitic baux ites. On the other hand, boehmitic and diasporic bauxites contain only 1 1-14% +H20. In contrast to lateritic bauxites, karst bauxites often contain carbonate minerals, mainly calcite and locally also siderite, in the form of secondary infiltrations that fill fissures, pores and small cavities. Intercalations consisting of clastic limestones and dolomite pebbles occur in some allochthonous bauxite deposits of Hungary, e.g. Halimba. Most karst bauxites are enriched in some minor and trace elements, i.e. highly enriched in Mn, Ag, As, Bi, Cr, Hf and Th, and moderately enriched in Ga, Nb, Ni, Pb, Sc, U, V, Zr and the rare earths, whereas they are depleted in Co, Cu, F, Ge, K, Mg, Na, P, Sr and Zn. The processes of their enrichment and depletion are complicated and depend on several factors, such as composition of the parent rock, Eh and pH of the ground water, presence of organic complexes, rede-
position, burial and secondary processes, and last but not least on the geochemical affinities of the given elements. LITH O L O G I C A L A N D S E D I M E N T O L O G I C A L FE ATU R E S
Lateritic and karst bauxite deposits differ from each other also by their main lithological features. Most lateritic bauxites are characterized by residual and aphanitic textures; some of them are nodular, pisoidal and concretionary. Residual structures prevail and locally blocky, bouldery and columnar structures are observed. All these features are the result of in situ autochthonous development of the deposits plus, at some locations, secondary, mainly diagenetic processes (Bardossy & Aleva 1990). The majority of karst bauxites are allochthonous, or at least parallochthonous and parautochthonous. As a consequence, clastic textures are very frequent, such as microclastic, arenitic and conglomeratic textures (Nicolas 1968; Combes 1969, 1973; Bardossy 1982; Mindszenty 1985). The clastic bauxites exhibit more or less distinct stratification. Graded bedding of bauxite is observed in several deposits of France, Hungary and Jamaica; subordinately cross stratification also occurs. Aphanitic texture is also very frequent, e.g. in Jamaica. Ooidal and pisoidal textures have been observed mainly in European karst bauxites and are attributed to early diagenetic processes. Where conditions of drainage remained fav ourable, the dissolution of the limestone and dolomite underlying the deposits continued after the deposition of the bauxite. Deep karstic depressions and funnel-shaped sink-holes were formed this way. The bauxite was affected by this development through repeated collapse and subsidence into depressions, and collapse structures observed at such places are evidence of this. Very heterogeneous bauxitic material fills such depressions: blocks of hard, massive bauxite are embedded into a softer and more clayey bauxitic groundmass, e.g. at Bedarieux, France (Combes 1969, 1973, 1984). At many places, pebbles and blocks of dolomite and limestone were washed into the bauxite deposit from adjacent, higher surfaces. These materials gen erally are located along the rims of the deposits and they are 1-20 cm in diameter. Locally, clastic interca lations consisting of limestone and dolomite pebbles also occur, e.g. at Halimba, Hungary. Even in these
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places the majority of the deposits still consist of bauxitic material, this being the major lithological feature of the Mediterranean-Jamaican type deposit (see classification in previous section). On the other hand, the Kazakhstan-type deposits consist mainly of non-bauxitic, clastic terrigenous material, and the bauxite forms only lenses or layers within them. These bauxites are generally stratified and have clastic textures, indicating their allochthonous origin (Bardossy 1982). The Ariege-type deposits, which occur mainly in southern France, show particular lithological fea tures: sedimentary clays and marls directly cover the weakly karstified surface of marine limestones. There is a gradual transition from these rocks into the overlying bauxite, interpreted as the result of an autochthonous bauxitization process. The texture of the bauxite is mainly pisolitic, oolitic and aphanitic (Combes 1969, 1990). Secondary diagenetic and epigenetic processes tend to eliminate the original lithological features of bauxite and to replace them by new ones. The original red to ochre colour of bauxite is replaced by yellow, white, pink, purple, grey and even green; mottled bauxites also are frequent. These colours are the result mainly of the partial dissolution and reduction of ferric iron and of the formation of new clay miner als such as berthierine and chlorite. The original clastic textures and structures are increasingly obscured by these processes and they are replaced mainly by concretionary and nodular textures. Cavi ties and fissures of bauxite are filled by secondary calcite and clay minerals. These transformations are limited in most cases to the top 0.5-3 m parts of the deposits. It is quite rare that the entire deposit is transformed, as, for example, the complete resilification (kaolinization) of the Ollieres deposit, France (Bardossy 1982). Deposits that remained on the surface, e.g. in Jamaica, suffered Jess secondary transformation than those that were buried, e.g. the European deposits. In conclusion, textures and structures are the best indicators for interpreting the genetic processes of bauxite deposits. Their genetic interpretation will be discussed in the next section. R E L AT I O N S H I P D E P O S I T S TO
OF K A R S T B AU X I T E
P A L A E O W E AT H E R I N G
Practically all types o f bauxite deposits are geneti cally related to lateritic (= ferrallitic) palaeoweather-
ing processes. The ongm of the lateritic bauxite deposits is discussed in other chapters of this volume (Schmitt, pp. 21--41; Valeton, pp. 153-188). We have concentrated our attention to the circumstances of karst bauxite deposits. We discuss their development according to the palaeogeographical and tectonic classification of Combes & Bardossy (1994), as schematically represented in Fig. 3. Similar consider ations have been presented by Carannante et at. (1994). High-level intracontinental deposits
(Fig. 3)
Over the internal parts of the continents lateritic bauxite deposits were formed mainly at relatively higher altitude. Only few karst bauxite deposits are found in this intracontinental position, where ancient carbonate platforms were incorporated into the con tinent. Kazakhstan-type deposits are characteristic of this position, with karstic depressions reaching 50-300m depths. The depressions were filled by several layers of continental sediments: clay, sand and conglomerate, and the clastic bauxite lenses occur at several horizons above each other. The bauxite horizons are covered in many places by sediments representing marsh environments, mainly by lignitic clay and lignite. Kazakhstan-type deposits represent a particular mode of allochthonous bauxite accumu lation in a continental environment (Bardossy 1982). Some of these deposits are situated along the contact of carbonate and aluminosilicate rocks of the crys talline basement. Here, the transport distance of the eroded lateritic weathering products was probably the shortest. Most of the Kazakhstan-type deposits occur in Kazakhstan and southern Siberia, but they have been found also in Missouri and in the south eastern part of USA (Clarke 1984). It should be stressed that all these bauxites are of gibbsitic composition. Low-level pericontinental deposits
(Fig. 3)
The majority of karst bauxite deposits are situated along the rims of ancient continents, either on the seaward slopes of continental platforms or on coastal plains, and were formed by the emergence of marine carbonate platforms. According to our research, the majority of these pericontinental deposits originated, more or less directly, from the partial, sometimes total erosion of the higher, neighbouring lateritic bauxite deposits. The transport of the eroded bauxitic material occurred in fluvial form and in several sedi-
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Karst bauxites
LOW-LEVEL PERICONTINENTAL DEPOSITS
HIG H-LEVEL INTRACONTINENTAL DEPOSITS
KARST BAUXITES
LATERITIC BAUXITES
P l a t f o r m
A
DEPOSITS ON INSULAR PLATFORMS KARST BAUXITES
�--�-�--�--�-�--�-�-C:����==��=r-�-�-- O m
B
Fig. 3. Schematic diagrams of the three main types of deposits according to the palaeogeographical and tectonic classification of Combes & B ardossy (1994). The most frequent substrata: (A) Intracontinental deposits: Palaeozoic basement, in some cases abducted ophiolites on internal areas (e.g. in Greece) ; ancient carbonate platforms incorporated into the continent (on external areas). Pericontinental deposits: carbonate platforms. (B) Deposits on insular platforms: carbonate platforms and carbonate shoal.
mentological steps, as demonstrated by the compo site bauxite pebbles found in many deposits. Some of the deposits had a relatively high position above the ground-water level; thus karstic depressions 50-150 m deep could develop. These deposits generally occupy longer stratigraphical intervals (tens of mil lions of years maximum). The bauxite is partly gibb sitic, partly boehmitic. In some places bauxitization resumed after the arrival of the bauxitic material in the depressions, e.g. in France and Hungary. These deposits are considered by us to be partly parautochthonous, partly parallochthonous. Closer to the coast, with diminishing elevation, the depth of the underlying karstic depressions becomes less pronounced (5-40 m). In some places entire alluvial fans of allochthonous, fluvially transported bauxite were detected, e.g. Szbc area, Hungary. Unconformities separate the main arrival of clastic bauxitic material. The younger material is incised into the underlying older bauxite in the form of river beds, as illustrated in Fig. 4 (Bardossy & Kovacs 1995). Intercalations of limestone, dolomite pebbles or detrital dolomitic sand occur in several bauxite deposits of France, Hungary and Turkey (Combes 1973, 1990; Ozlti 1977; Bardossy 1982). These interca-
lations indicate short periods of stronger erosion that separate the periods of much lower energy transport represented by the finer bauxitic material. Under particularly favourable conditions large karstic caves, situated below the main palaeokarst surface, were filled by bauxitic material. This was revealed by bauxite exploration and subsequent mining at Knezopolje, near Mostar, Bosnia (Kormos 1943). On the low-lying coastal plains the position of the groundwater level was close to the surface, thus only shallow karst depressions could develop (<5 m). The periods of emergence of the carbonate platform were short, as indicated by the short stratigraphical gaps occupied by the deposits (1-3 Myr minimum). In Croatia, Bosnia and Montenegro several bauxite horizons occur above each other, separated by marine carbonate platforms (Sakac & Sinkovec 1991). In such low-level deposits the geochemical conditions of bauxite accumulation varied from slightly oxidizing to slightly reducing, as the ground water level fluctuated periodically (Mindszenty et al. 1987). Boehmite and diaspore are the main alumina minerals of these deposits, accompanied by goethite, haematite and locally by berthierine, siderite and chlorite. In the Devonian deposits of the North Ural
G. Bardossy and P. -J Combes
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w
E 7
Fig. 4. Geological-palaeomorphological profile of the Kislod bauxite deposit, Hungary. l, bauxitic clay; 2, clayey bauxite; 3, clayey bauxite conglomerate; 4, bauxite; 5, high-grade bauxite; 6, bauxite conglomerate; 7, early Eocene lignitic clay (directly overlying the deposit).
Mountains, Russia, the coast was so close that part of the bauxite was washed into lagoons, quiet bays and sea coves (Bushinsky 1975). Green and grey bauxite accumulated under the geochemical influence of the sea water. Marine fossils found in the bauxite and thin marine limestone intercalations demonstrate the marine environment of these bauxite accumulations. It should be stressed, however, that the bauxitic ma terial originated from the continent, i.e. from lateritic weathering, as in all other bauxite areas. Sedimentological studies have revealed that in the Parnasse Mountains, Greece, allochthonous bauxite accumulated in lagoons and shallow marine environ ments over internal carbonate platform limestones. Intercalations of limestones and the presence of mol luscs in the bauxite prove the validity of this assump tion (Combes 1979; Combes & Mangin 1983). A particular type of pericontinental low-level bauxite formation is represented by the Ariege-type deposits, mentioned briefly in a previous section. Their genetic interpretation is illustrated by the synthetic profiles of Fig. 5 (Combes & Peybernes 1 99 1 ) . As a first step, lateritization and bauxitization occurred on the continent. Different weathering products were eroded and transported to the coast, where they were deposited as littoral clayey and marly sediments during a period of transgression and marine highstand. These ferruginous clays and marls emerged during the following regression and under favourable climatic conditions their bauxitization started immediately. In some places the entire emerged sediment was bauxitized down to its lime stone basement; in other places bauxitization was limited to the upper part of the profile. Gradual tran sition of the marine clays and marls into bauxite can be observed at such places. Oscillations of sea-level resulted in repetition of the process, as expressed by the superposition of several bauxite profiles. The
main control of these oscillations was eustatic change of sea-level. Analysis by sequence stratigraphy indi cates that basin pelites, platform limestones and bauxites are depositional units, which can be interpreted in terms of systems tracts. Detailed min eralogical and geochemical studies of the deposits in the Ariege area proved the lateritic nature of this palaeoweathering process, e.g. upward enrichment of AI, Fe and Ti, balanced by depletion of Si. Local erosion and redeposition also was observed (Combes 1 969, 1 990) . Deposits on low, insular platforms (Fig. 3)
These types of deposit developed on islands that have never been connected with the continent. The most important examples are in Jamaica, the Dominican Republic, Haiti and Rennell Island; much smaller deposits are found in the Central Appennines, Italy. The main genetic problem for a long time was the lack of suitable parent material on these islands. The pure limestone basement could not furnish the large tonnage of bauxite represented by the deposits, as presumed by the so called 'terra rossa theory' (de Weisse 1 964). In the case of Jamaica, fluvial transport from the Central Inlier must be excluded because the bauxite deposits are at higher altitude. Waterman ( 1 962) suggested that wind-blown volcanic ash falling on to the karst surface was the parent material for the Jamaican bauxite. This idea was supported by Comer ( 1 974), who found bentonite intercalations in the contemporaneous Miocene Montpelier limestone. According to Comer, volcanic ash that fell on the emerged part of the island was bauxitized, whereas ash that fell into the shallow sea surrounding the island was turned into bentonite. Field observations by Bardossy ( 1 982) revealed mound- and hummock-shaped bauxite accumula-
197
Karst bauxites PYRENEAN AXIAL BASIN
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.
Karstic paleosurface · . •_ ,· ....___ a Collapse breccia Tecto-sedimentary breccia
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BAUXITIZATION
TST Transgressive systems tract HST1 Early highstand systems tract HST2 Late highstand systems tract Lowstand systems tract CS Condensed section {mfs) Maximum flooding surface Sequence boundary LST
SB
Fig. 5. Reconstruction of the Ariege-type bauxite sequence and its correlation with systems tracts according to the terminology of Vai l et al. (1987). (Modified after Combes & Peybernes 1991.)
tions on some plateau areas, as shown in Fig. 6, which cannot be explained in terms of erosion and fluvial accumulation and favours the wind-blown pyroclas tic origin. The bauxitization of the loose and highly permeable material must have been very rapid. Subsequent erosion and short-distance transport into neighbouring valleys and fiat basins ( e.g. Essex Valley) produced large, fiat stratiform deposits cover ing areas of several square kilometres. The weak stratification and local graded bedding observed in these deposits is a proof of the above-mentioned events. Similar palaeogeographical conditions, on smaller islands, have been demonstrated in the Central Appennines, Italy, where remains of the initial pyro-
clastic material have been found at the base of the Campo Felice deposit (Bardossy et al. 1977 ) . Later studies carried out by D'Argenio et a/. ( 1987 ) supported the pyroclastic origin, corroborated by the occurrence of intense contemporaneous volcanic activity to the north and east of the bauxite area. A highly instructive assemblage of bauxite deposits, representing three types of origin, have been studied and described from north-west Sardinia by Combes et al. ( 1993 ) . Their main features are as follows. 1 Continuous stratiform deposit with uniform thick ness of 2-3 m, extended over several square kilo metres. Illitic marls of Berriasian age weathered into bauxite of boehmitic composition. A gradual
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G. Bardossy and P. -J Combes A
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3
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0
10
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transition can be observed in many places from the marl to the bauxite. Consequently the deposit is interpreted as autochthonous in origin. 2 At higher elevation the weathering of the Berri asian marl was more complete and even continued into the underlying Tithonian and Kimmeridgian limestone. A karstified surface developed this way with moderately deep karstic depressions. The thickness of the bauxite deposit varies from 5 to 1 0 m. Repeated collapse and subsidence of the bauxite material into the depressions produced highly het erogeneous structures with bauxite debris, nodules, pisoliths and concretions, embedded in a pelitomor phous matrix. Locally bauxitization continued during and after this development, as demonstrated by geo chemical profiles. This type of deposit is regarded as parallochthonous. 3 In the low-lying areas the bauxite beds cover Valanginian to Aptian limestone. The footwall surface is weakly karstified; the thickness of the deposit is 0.5-4m. The bauxite has a clastic texture, and stratification, graded bedding and bauxite con glomerate occur locally. As a consequence these deposits are considered to be of allochthonous origin. They developed from the erosion and transport of bauxite of higher-level deposits of types 1 and 2. Locally bauxitization continued after the accumula tion of the bauxitic material (parautochthonous origin) . The main features of the three types of deposit and their genetic interrelation are presented in Fig. 7. Berriasian marl was the common parent rock for all
50
60m
Fig. 6. Geological profiles of lenticular bauxite deposits on the Manchester Plateau, Jamaica. (A) road cut; (B) bauxite hill explored by drilling.
these deposits. Their development was determined by the varied movements of tectonic blocks induced by the Austrian tectonic phase. This study proved that even in a relatively small area different types of bauxite deposits could originate if the tectonic processes furnished significant differences in mor phology, elevation, rate of weathering, karst develop ment, erosion and subsequent sedimentation of the bauxitic material. Summary
The relationship of karst bauxite development to palaeoweathering can be summarized as follows: lateritic bauxite deposits developed mainly on the interior of tectonically stable cratonic regions. The formation of the deposits depended on favourable climatic conditions, the development of planation surfaces, good drainage conditions with a high infiltration rate and suitable parent rocks. Karst bauxite deposits, on the other hand, follow tectoni cally unstable belts, characterized by widespread distribution of carbonate platforms that served as basement and traps for bauxite accumulation. We have found no difference between the climatic conditions of lateritic and karst bauxite develop ment. The process of lateritization and bauxitization also is essentially the same. The difference is in the geochemical environment, i.e. aluminosilicate over the cratons and determined by the carbonate plat form during karst bauxite accumulation. Further differences arise from the autochthonous character
199
Karst bauxites WEATHERING ZONE OF MARLS
DEPOSITIONAL ZONE
WEATHERING
WEATHERING EROSION
(AND EXTENSION
EROSION AND TRANSIT
AND SUPERIMPOSITION
OF WEATHERING)
I
I
to KIMMERIDGIAN
BERRIASIAN (base level of the ground-water)
UPLIFT RATE
UPLIFT RATE IMPORTANT
UPLIFT RATE MODERATE
WEAK
ALLOCHTHONY PARAUTOCHTHONY
PARALLOCHTHONY
AUTOCHTONY
TYPE 3
TYPE 2
POST-BERRIASIAN
ANTE-BERRIASIAN
FOOT-WALL
FOOT-WALL Important W
Fig. 7. Geodynamic model, types of deposits and main characteristics of the bauxites in north-west Sardinia, Italy. (Modified after Combes et al. 1993.)
[]0 � M O s;o,
At2o.,
Without extension of weathering
of the lateritic bauxite deposits and the overwhelm ingly allochthonous, parallochthonous and parautochthonous karst bauxite accumulations. Intensive research in several countries has tried to clarify the still unresolved problems of karst bauxite genesis, such as the nature of the parent rock, the sed imentology of bauxite transport and accumulation, and the nature of geodynamic aspects. In our opinion essential progress has been achieved, and these results enable us to proceed to applications in the field of general palaeoclimatology and palaeogeogra phy, to be outlined in the following section.
PALAE O CLIMATIC AND PALAE O G E O G RAPHICAL SIG NIFICAN C E O F KA R S T BAUXITE D EP O SITS
Palaeoclimatology was founded on palaeontological evidence, by accepting the concept of uniformitarian ism. Its development has been promoted con siderably during recent decades by new, inorganic methods, such as: 1 palaeomagnetic measurements, which allow deter mination of the former latitude of a given area, a
Si02
AI203
Extension of weathering
Partial weathering
Partial weathering
Important weathering
crucial indicator of palaeoclimate -continental drift is the basic concept for such reconstructions; 2 stable-isotope measurements (e.g. ()lSQ and ()BC) for temperature determination of the palaeoatmos phere and sea water, and for estimates of the C02 and oxygen contents of the palaeoatmosphere; 3 computer-based construction of palaeoclimatic models, simulation of palaeo-ocean currents and rainfall patterns for different periods of the Phanerozoic; 4 computer-based sensitivity analyses with the aim of establishing the relative importance of factors influencing the palaeoclimate; 5 calculations of palaeoclimatic cyclicity, as deter mined by cyclic changes of the Earth's orbital insola tion rates, e.g. Milankovic cycles. These measurements and calculations have resulted in a rapid development of palaeoclimatol ogy. Nevertheless, all these results are of an indirect nature. To be definitively accepted they must be verified by direct geological evidence. There are two ways of doing so: first by the above-mentioned palaeontological methods and second by the applica tion of palaeoclimatic indicator rocks. All sedimentary rocks contain some form of palaeoclimatological information. The indicator
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rocks are distinguished by distinctive, unequivocal palaeoclimatic information, referring to well-defined, narrow climatic conditions. Systematic studies have revealed that continental sediments and residual rocks are the most informative indicator rocks, such as palaeosols, laterites, bauxites, lignites and coals, aeolian sediments, red beds, continental evap orites, and some glacial and periglacial sediments. For a long time laterites and bauxites were neglected in this respect, because of disputes over their genesis and partly because of their uncertain stratigraphical positions (Ziegler et al. 1 987). The cli matic conditions of laterite and bauxite formation have been studied intensely during the last few decades and general agreement has been achieved among most researchers (Bardossy & Aleva 1 990). The stratigraphical position of the majority of deposits also has been defined more accurately. In this section we discuss only the karst bauxite deposits. Their position is favourable in this respect because they are limited in most cases to relatively short stratigraphical intervals. Thus, at our present level of knowledge, karst bauxites are among the best palaeo climatic indicator rocks, representing tropical, wet cli matic conditions. A crucial question is the validity of palaeoclimatic information furnished by indicator rocks, when going back in geological time. In other words, how far back in time are the limits of our uniformitarian approach? It seems to be accepted that the fundamental geochemical and mineralogical rules were the same throughout Phanerozoic and even Proterozoic time. Changes of global temperature and of the atmos phere's composition certainly influenced the inten sity of bauxitization, but not its general nature. Control of the vegetation cover
Bauxite and laterite are strongly dependent on the type of vegetation cover, which at present varies from savannah-type woodland to tropical forest. We presume a similar vegetation cover for the past, as far back as the Middle Devonian, when the surface of the continents was invaded by land plants. In our opinion, this is the time limit for the uniformitarian application of bauxites and laterites as palaeoclimatic indicator rocks. Small karst bauxite deposits are known from the Ordovician, Silurian and Cambrian as well, but the climatic and other conditions of their formation are uncertain. Another significant change occurred during the Cretaceous, in a time interval 80-120 Ma,
when angiosperm-deciduous ecosystems gradually replaced much of the former conifer-evergreen ecosystems. Knoll & James ( 1987) and Yolk (1989) showed the angiosperm-deciduous ecosystems produced higher weathering rates, as a result of their annual leaf loss, as compared with the conifer-ever greens. They did not study the particular case of baux ites, but it seems reasonable to expect that this change affected the rates of bauxitization as well. The rapid increase in bauxite formation during the Cretaceous, illustrated in Fig. 2, was the result of several factors, such as favourable climatic conditions, palaeogeogra phy, and presumably also the change of ecosystems outlined above. In our opinion this is not a limitation of the uniformitarian approach, but it certainly represents a change in bauxitization rates, and the completeness of the bauxitization process. Rainfall distribution
Parrish et al. (1982) published computer based palaeorainfall maps for several periods of the Phanerozoic. Bardossy & Aleva ( 1990) and Bardossy (1994) discussed the correspondance of these maps with the location of the bauxite deposits. There is a general agreement in most places, which we interpret as verification of the palaeorainfall maps. The Early Triassic conditions, with Pangaea still in existence, were characterized by a continental interior with an arid and semiarid climate (Fig. 8). The known bauxite deposits, all karst bauxites, are situated along the eastern coast, in the tropical and subtropical belt. The area is characterized by high to medium-high rainfall, and thus the correspondence is perfect. Note the absence of bauxite deposits along the west coast, despite high rainfall at low latitudes. It can be explained by the presence of cold ocean currents that certainly strongly influenced the local climate. The Cenomanian conditions are characterized by the break-up of Pangaea and extension of the Tethys ocean to the west (Fig. 9). Distribution of the numer ous karst bauxite deposits in the Mediterranean area correspond well with this palaeogeography and with the rainfall map. In our opinion, even more rainfall could be expected, considering the convergent coast lines of the sea, the warm ocean current, and the west erly wind direction. There is a striking coincidence between the Cenomanian rainfall pattern for the Arabian Peninsula and the presence of a large lat eritic bauxite deposit at Az Zabirah, which is now located in the very centre of the desert. This appears to verify the palaeorainfall map.
201
Karst bauxites
D Dry SO Semi-dry
•
I d
Wet
D
Very wet
Bauxite deposits
___. Warm ocean currents -- + Cold ocean currents
Fig. 8. Palaeogeographical and palaeorainfall map of the Early Triassic after Parrish et a/. (1982), complemented by the bauxite deposits. Pluviometry (em yr-I): < 50 = low rainfall; 50-lOO = moderately high rainfall;> 200= high rainfall.
D Dry SO Semi-dry
•
I
/I
Wet
CJ
Very wet __,.
Warm ocean currents
--
+
Bauxite deposits Cold ocean currents
Fig. 9. Palaeogeographical and palaeorainfall map of the Cenomanian, after Parrish et a/. (1982), complemented by the bauxite deposits. Pluviometry (em yr--1 ): < 50= low rainfall; 50- 100= moderately high rainfall;> 200 = high rainfall.
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G. Bardossy and P-J Combes
Palaeoatmosphere composition
According to palaeoclimatologists, the composition ofthe atmosphere was crucial for the long-term devel opment of the Earth's climate. The C02 content was the most important factor in this respect, followed by oxygen, methane and sulphur containing aerosol par ticles. High C02 concentrations gave rise to a warm and wet climate ('greenhouse effect'); low C02 con centrations on the other hand created cooler climates with glaciations in the polar regions ('icehouse effect'). The relative C02 and temperature curves calculated by Berner (1994) demonstrate this effect quite clearly (Fig. 10). According to Budyko et al. (1987) the oxygen content had three peaks during the Phanerozoic; during the Ordovician, the Early Carboniferous, and the Cretaceous, the latest being the largest. The explanation for these changes is rather compli cated. Palaeoclimatologists presume that increases in C02 content result mainly from global volcanic activ ity. Decay of plants and animals contribute also to the C02 increase. In contrast, rapid burial of large forests could seal organic material from oxidation and from its conversion to atmospheric C02. The C02 content of the atmosphere is continually lowered by the photosynthesis of plants. Continental weathering of magmatic and metamorphic rocks also consumes atmospheric carbon dioxide. This fact has been accepted by palaeoclimatologists, and its conse quences have been included in the computer-based palaeoclimatic models (Crowley 1993; Berner 1994). However, their calculations assumed clay minerals as the main weathering products. Huge amounts of lat erite and bauxite, however, are produced in tropical and subtropical wet climatic conditions, and this type of weathering causes further significant changes in the chemical composition of the parent rocks. Combes & Bardossy (1995) calculated the average oxygen content of bauxite and laterite (55.7 % ) and that of the corresponding parent rocks (45.7 % ), finding a difference of round 10%. They carried out model calculations for two time intervals (114-91 Ma and 91-65 Ma) of the Cretaceous with four different rates of weathering (2, 3 , 4 and 5 m Myr-1). The results show 2-1 1 % of atmospheric oxygen fixed by bauxite and laterite weathering during these time intervals. We suppose that the high oxygen content of the atmosphere during this time promoted a higher rate of weathering, but that the higher consumption out lined above gradually lowered the oxygen content of the atmosphere, slowing down the process.
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In conclusion, bauxites and laterites are among the best palaeoclimatic indicator rocks for warm and wet palaeoclimates, and they can be used to verify palaeoclimatic maps and models. The highly variable intensity of bauxite and laterite formation during the Phanerozoic is an important factor in explaining the long-term development of the atmosphere's composition.
(b)
203
Karst bauxites Palaeogeographical and geodynamic significance
Karst bauxite deposits also have a particular palaeo geographical significance. Their presence indicates continental conditions, or at least the emergence of one or more islands. Bardossy & Dercourt ( 1990) detected a number of such emergences by the presence of bauxite deposits in the Tethys realm, at a place where only a shallow sea was indicated on former palaeogeographical maps. Many karst bauxite deposits are situated along ancient coast lines. Their presence is a powerful tool in defining more precisely the border of the continent. Karst bauxite deposits also can be used to estimate the relative elevation of the land surface, by measure ment of the depth of the karst depressions that form the basal surface of the deposits. The depth of the depressions is related directly to the position of the karst water-level below the surface. The deeper the depressions, the higher was the relative elevation of the deposits above the karst water-level. On the other hand, there are other deposits underlain only by very shallow karst depressions. They are inter preted to have formed in the fluctuation zone of the karst water-level, or even within it. Such deposits are characterized by particular bauxite textures, colour, geochemistry and mineralogy. Mindszenty et al. (1987) and D' Argenio & Mindszenty ( 1995) distin guished, on this basis, 'vadose' and 'saturated' envi ronments for karst bauxite deposits. Karst bauxite deposits are suitable for detailed sedimentological evaluations as well. As pointed out in the foregoing sections, the majority of karst bauxite deposits are of allochthonous or at least parautochthonous or parallochthonous origin. Trans port of the bauxitic material by rivers or by sheet flow could be demonstrated in several areas. Unconformi ties have been detected within allochthonous bauxite deposits in France and Hungary. Intercalations of dolomite and limestone pebbles indicate abrupt changes in erosion rates, e.g. at Halimba (Hungary) and Bedarieux (France). It is even possible to determine the direction of bauxite transport and the extent of the main transport routes by evaluation of sedimentological and geochemical features, e.g. in the Szoc bauxite district, Hungary. Some karst bauxites accumulated in a marshy environment, resulting in partial leaching and reduction of their iron con tent, e.g. in China (Liao Shifan 1 991). Combes ( 1984) demonstrated the accumulation of bauxite in coastal lagoons in the Parnassos zone of central Greece (Fig. 1 1 ) . The grey and green bauxites of the North Ural
Mountains represent littoral, shallow-marine condi tions of bauxite accumulation, proven by the pres ence of marine fossils and by thin intercalations of marine limestone (Bushinsky 1975). As a conse quence, well-studied karst bauxite deposits can furnish essential contributions to the sedimentologi cal interpretation of a selected area. Bauxite deposits also reveal important infor mation on geodynamic aspects. It has been demon strated by Combes et al. (1991) that eustatic sea level oscillations controlled the formation and the preservation of karst bauxite deposits in many areas, particularly in coastal lowlands. The presence of karst bauxite deposits is indicative of tectonic instability. This is demonstrated by local uplifts and subsidences and by changing rates of erosion. Lateritic bauxite deposits were eroded and their material was transported to emerged carbonate platforms. On the other hand, long periods of tectonic stabil ity are not favourable for the erosion of lateritic weathering profiles, even in the interior of the conti nents. Large-scale planation surfaces may survive under such conditions for millions of years. As a consequence, much less bauxitic material will arrive at the coastal zones, and less bauxite will be accumu lated in karstic depressions. The Caribbean territory is an exception in this respect, as wind-blown pyro clastic material presumably was the source of these deposits.
E C O N O MIC S I G NIFICAN C E O F KA R S T BAUXIT E D E P O SIT S
Bauxites have a high economic importance because they are the raw materials for the production of met allurgical alumina and as a final product for metallic aluminium. The world production of bauxite for metallurgical purposes was llO Mt in 1994 (Mineral Commodity Summaries 1995, US Bureau of Mines). Of this production, 73.3% came from lateritic bauxite deposits, 26.2% from karst bauxite deposits and only 0.5 % from the Tikhvin-type deposits. The reason for the large proportion of lateritic bauxite deposits lies in their large dimensions and their position on the surface, allowing lower mining costs. The production of karst bauxite came from 10 countries, as listed in Table 1. Jamaica and Kazakhstan produce bauxite from open pit mines, whereas in the other countries part of the production comes from the underground mines.
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Table 1. World production of karst bauxites in 1994
Country Jamaica Russia People's Republic of China Kaz akhstan Greece Hungary Montenegro Turkey Italy Romania
Production (Mt) 12.0 4.6 4.5 2.8 2.0 0.9 0.9 0.6 0.3 0.2
An additional 7 Mt of bauxite were produced in 1994 for non-metallurgical purposes, such as for alu minous cements, slag adjusters, proppants and for the chemical industry, e.g. alum. Special low iron (max. 2% Fe2 03), low titania (max. 2.5 % Ti02) bauxites are the raw materials for the production of refractories and abrasives (Sehnke 1992). Specialty aluminas, such as activated or calcined aluminas, are used in the chemical and pharmacological industries. In 1994 more than 1 Mt of special white bauxites were mined for refractory purposes. The main producer is the People's Republic of China, with about 750 000 t. The ore comes from karst bauxite deposits in Guizhou, Henan, and Shanxi provinces. The remaining tonnage
Karst bauxites Table 2. Tonnages of identified karst bauxite reserves
Country Jamaica Greece People's Republic of China Ru ssia Turkey Kazakhstan Montenegro and Serbia Italy France Hungary Croatia Solomon Islands Dominican Republic Haiti Iran Bosnia Romania Philippines Pakistan Cuba
Reserves ( Mt) 2000 650 400 250 110 100 90 80 80 70 60 60 40 40 40 30 25 20 10 5
Much time and expense can be saved by the scientific examination and evaluation of the deposits. This is why the information and problems outlined in the foregoing sections are not only of purely scientific interest, but represent a high economic interest as well.
AC K N O W L E D G E M E N T S
Th e authors express their sincere gratitude t o A Palmer and M. Thiry for their important and con structive comments and suggestions that helped us to improve the manuscript. The authors are very indebted to P. Eichene for the presentation of the figures.
R E F E R E NC E S
G. (1982) Karst Bauxites. Bauxite Deposits on Carbonate Rocks. Elsevier, Amsterdam. BARDOSSY, G. (1994) Carboniferous to Jurassic bau xite deposits as paleoclimatic and paleogeographic indicators. In: Pangea (Eds Embry, A.F., Beauchamp, B. & Glass, D.J. ) , Mem. Can. Soc. Petrol. Geol., Calgary, 18, 283-293 .. BARDOSSY, G. & ALEVA, G.J.J. ( 1990) Lateritic Bauxites. Else vier, Amsterdam. BARDOSSY, G. & BOURKE, D.J. ( 1993) An assessment of world bauxite deposits as source of greenfield alumina plant developments. Aluminium (D ilsseldorf), 69, 888-894. BARDOSSY, G. & DERCOURT, J. (1990) Les gisements de bauxite tethysiens (Mediterranee, Proche et Moyen Orient ) ; cadre paleogeographiqu e et contr6les gene tiques. Bull. Soc. Geol. Fr. , S (VI) , 869-888. BARDOSSY, G. & KOVACS, L.O. ( 1995) A multivariate statis tical and geostatistical stu dy on the geochemistry of allochthonous karst bauxite deposits in Hungary. Non;e newable Resour. , 4(2), 138-153. BARDOSSY, G., BON!, M., DALL'AGLIO, M., D'ARGENIO, B. & PANTO, G. (1977) Bauxites of Peninsular Italy: Composi tion, origin and geotectonic significance. Monogr. Ser. Mineral Deposits (Berlin-Stuttgart), 15,1-61. BERNER, R.A. (1994) 3 GEOCARB II: a revised model of atmospheric C02 over Phanerozoic time. Am. I. Sci. , 294, 56-91 . BERTHIER,P. (1821) Analyse d e l'alumine hydratee des Baux, departement des Bou ches-du-Rh6ne. Ann. Mines, 6, 531-534. BunYKo, M.J., RoNov, A.B. & YANSHIN, A.L. (1987) History of the Earth's Atmosphere. Springer-Verlag, Berlin. BusHINSKY, G.L. (1975) Geology of Bauxites ( in Russian ) . Izdatelstvo, Moscow. CARANNANTE, G., D' ARGENIO, B., MINDSZENTY, A., RUBERT!, D. & SIMONE, L. ( 1994) Cretaceous-Miocene shallow water carbonate sequences. Regional u nconformities and facies patterns. In: 15th International Association of Sedi mentologists Regional Meeting, April, Ischia, Italy, pp. 2560. BARDOSSY,
is mined in Guyana, Brazil and the USA, all from lat erite bauxite deposits. The identified bauxite reserves are estimated to be 43 billion tons (Bardossy & Bourke 1993). Addi tional 15-20 billion tons of undiscovered bauxite resources-both hypothetical and speculative - are presumed to exist. Of the identified global reserves, 89.9% belong to lateritic bauxite deposits, 9.7% to karst bauxite deposits and only 0.4% to the Tikhvin type deposits. The tonnages of karst bauxite reserves are listed on Table 2. By far the largest reserves are in Jamaica and Greece. Part of the reserves in France and Hungary cannot be mined - at least by the present day mining technology -being situated below the karst water-table. This circumstance poses high technical dangers and would require excessive mining costs. The reserves of low-iron, white, refractory grade bauxite are much more limited. Guyana and Brazil possess the largest identified reserves, both in lateritic bauxite deposits. The People's Republic of China also has large reserves of refractory grade bauxite in karst bauxite deposits. Small reserves are known from Greece and Montenegro. Exploration of buried, 'hidden' karst bauxite deposits is a difficult and costly activity, because of the relatively small dimensions of the deposits and the variability of bauxite quality. The proper under standing of the origin of the deposits has proved to be very useful in planning the exploration strategy.
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CLARKE, O.M. ( 1984) Bauxite deposits of the United States. In: Bauxite (Ed. Jacob, L., JR ) , pp. 245-261 . American Institute of Mining, Metallurgical and Petroleum Engi neers, New York. CoMBES, P.J. (1969) Recherches sur Ia genese des bauxites dans le Nord-Est de l'Espagne, le Languedoc et I' Ariege (France). These, Doct. Etat University of Montpellier; Mem. C.E.R. G.H.,J-4, 1-375. CoMBES, P.J. (1973) Etude geologique sur les conditions de mise en place d'une bauxite allochtone a substratum carbonate: le gisement de B edarieux, Herault, France. In: ICSOBA, 3e Congres International Nice (Ed. Nicolas, J.), pp. 89-108. COMBES, P.J. (1979) Observations sedimentologiques, paleogeographiques, mineralogiques et geochimiques sur les bauxites du deuxieme horizon dans Ia zone du Par nasse, Grece. Bull. Soc. geol. Fr. , 7(XXI), 485-494. COMBES, P.J. ( 1984) Regards sur La geologie des bauxites; aspects recents sur La genese de quelques gisements a sub stratum carbonate. Bull. Centres Rech. Explor. Prod. Elf Aquitaine, 8(1 ), 251-274. CoMBES, P.J. ( 1990) Typologie, cadre geodynamique et genese des bauxites fran<;:aises. Geodyn. Acta (Paris), 4, 91-109. CoMBES, P.J. & BARDOSSY, G. (1994) Typologie et controle geodynamique des bauxites tethysiennes. C. R. Acad. Sci. Paris, 318(serie II), 359-366. CoMBES, P.J. & BARDOSSY, G. (1995) Influence des bauxites et laterites sur La composition de !'atmosphere. C. R. Acad. Sci. Paris, 320(serie IIa), 109-116. CoMBES, P.J. & MoNGIN, D. (1983) Decouverte de Gasteropodes dans le gisement de bauxite d' Anthimos (deuxieme horizon), zone du Parnasse (Grece). Geol. Mediterraneenne, X(1), 15-24. COMBES, P.J. & PEYBERNES, B. (1991) Role de l'eustatisme dans La genese des bauxites de type Ariege,Pyrenees Cen trales. C. R. Acad. Sci. Paris, 313(serie II), 669-676. CoMBEs, P.J., OGGIANo, G. & TEMussr, I. (1993) Geody namique des bauxites sardes, typologie, genese et con trole paleotectonique. C. R. Acad. Sci. Paris, 316(serie II), 403-409. CoMER, J.B. (1974) Genesis of Jamaican Bauxite. Econ. Geol. , 69 (8), 1251-1264. CROWLEY,T.J. (1993) Climate change on tectonic time scales. Tectonophysics, 222, 277-294. D'ARGENIO, B. & MINDSZENTY, A. (1995) B auxites and related paleokarst: tectonic and climatic event markers at regional unconformities. Eclogae Geol. Helv. , 88(3), 453-499. D'ARGENIO, B., MINDSZENTY,A., BARDossY, G., JuHASz , E. & BoNr, M. (1987) Bauxites of southern Italy revisited. Rend. Soc. Geol. Italy , 9, 253-268. DE WEISSE, G. (1964) B auxite lateritique et bauxite kars tique. In Symposium sur les Bauxites, Oxydes et hydroxy-
des d'Aluminium (Ed. Karsulin, M.), pp. 7-29. Academie Yougoslave des Sciences et des Arts, Zagreb. (October 1-3, 1963.) KNOLL, M.A. & JAMES, W.C. (1987) Effect of the advent and diversification of vascular land plants on mineral weathering through geologic time. Geology, 15, 10991 102. KoRMOS, T. (1943) Bauxitablagerung in Hohlen. Foldt. Kozl. (Budapest) , 73, 296-305. LIAO SHIFAN (1989) Bauxite Geology of China. Science and Technology Publishing House of Ghizou, China. MINDSZENTY, A. (1985) The lithology of some Hungarian bauxites. A contribution to the Cretaceous paleogeo graphy of the Transdanubian Central Range. Acta Geol. Hung. , 27(3-4), 445-457. MINDSZENTY, A., D'ARGENIO, B. & BoGNAR, L. (1987) Creta ceous bauxites of Austria and Hungary: lithology and paleotectonic implications. Trav. ICSOBA (Zagreb), 20, 13-31. NICOLAS, J. (1968) Nouvelles donnees sur Ia genese des bauxites a mur karstique du Sud-Est de Ia France. Mineral. Deposita, 1(3), 18-33. OZL, N. (1977) Differents modes de formation de breches calcaro-bauxitiques dans les bauxites des Taurides occidentales, Turquie meridionale. C. R. Acad. Sci. Paris, 284(serie II), 1021-1023. PARRISH, J.T., ZIEGLER, A.M. & ScoTESE, C.R. (1982) Rainfall patterns and the distribution of coals and evaporites in the Mesozoic and Cenozoic. Palaeogeogr. Palaeoclimatol. Palaeoecol. , 40, 67-101. SAKAC, C. & SrNKOVEC, B. (1991) The bauxites of the Dinar ides. Trav. ICSOBA (Zagreb), 23, 1-12. SEHNKE, E.D. (1992) Refractory grade bauxite: an overview. Proceedings SME Annual Meeting, Phoenix, Arizona, pp. 1-1 1 . VAIL. P.R., COLIN, J.P., JAN D U CHENE, R., KUCHLY, J., MEDIAVILLA, F. & TRIFILIEFF, V. (1987) La stratigraphie sequentielle et son application aux con-elations chrono stratigraphiques dans le Jurassique du bassin de Paris. Bull. Soc. Geol. Fr. , 8(III), 1301-1321. VALETON, I. (1976) Criteria for autochthonous and allochthonous source material of bauxitic ores on car bonaceous rocks. Trav. ICSOBA, 13, 21-36. VoLK,T. (1989) Rise of Angiosperms as a factor in long-term climatic cooling. Geology, 17, 107-110. WATERMAN, G.C. (1962) Some chemical aspects of bauxite genesis in Jamaica. Econ. Geol. , 57, 829-830. ZIEGLER, A.M., RAYMOND, A.L., GIERLOWSKI, T.C. & RowLEY, D.B. (1987) Coal, climate and terrestrial produc tivity: the present and early Cretaceous compared. In: Coal and Coal-Bearing Strata: Recent Advances (Ed. Scott, A. C.), Spec. Pub!. geol. Soc. London, No. 32, pp. 25-49. Blackwell Scientific Publications, Oxford.
Spec. Pubis int. Ass. Sediment. (1999) 27, 207-221
Precambrian palaeosols: a view from the Canadian shield
Q. GALL 1-1190 Richmond Road,
Ottawa, Ontario, Canada K2B 813
A B S T RACT
Diligent research over the last two decades has found that records of palaeoweathering (palaeosols) and palaeosurfaces can be identified from rocks of Precambrian age in many shield terranes around the world. In the Canadian Shield, for example, approximately 40 palaeosols have been identified despite the camouflage effects of diagenesis, metamorphism and structural overprinting. Both physical and chemical criteria have proven useful in identifying Precambrian palaeosols; and include: 1 stratigraphical position along unconformities; 2 ascending protolith disruption; 3 the presence of macrostructures (e.g. karstification, cores tones); 4 the presence of microstructures (e.g. peds and cutans); 5 mineralogy and colour contrast with adjacent lithologies; and 6 ascending element depletion with concomitant change in chemical weathering indices. There appear to be difficulties in applying modern soil classifications to Precambrian palaeosols owing to the general paucity of horizonation; suggesting that drab-coloured saproliths may be the normal end result of weathering during the Precambrian. In Precambrian terranes, physical and chemical overprinting by burial diagenesis or metamorphism also appear to be common, and tend to mask original pedogenic features. Difficulties in accurately determining the age and duration of palaeoweathering are also common place in Precambrian terranes. Despite these difficulties in studying Precambrian palaeosols, once identified, they may be useful in: (i) reinforcing old, or suggesting new, regional stratigraphical correlations and, hence, in identifying regional to world-wide episodes of palaeoweathering, and (ii) in helping to track unconformity-related mineralization.
I N T R O D UC TI O N
fore more open to interpretation and debate. Palaeosols, which are easily recognized on physical appearance alone in Phanerozoic sequences, are not as obvious in Precambrian terranes and can be identified as such only after careful consideration of a set of criteria, which often includes mineralogy and geochemical profiles as well as physical features. Despite the difficulties of identifying palaeosols of Precambrian age, it is accepted by most geologists that they do exist, and have been found in Archaean to Neoproterozoic-age rocks within shield areas. Following a brief summary of criteria that have been used to identify Precambrian palaeosols, this paper attempts to update the reader on what is
Although palaeosols and palaeoweathering surfaces have not received as much attention as other geologi cal phenomena, they are recognized as bona fide features from which past physical and chemical envi ronments on, or just beneath, Earth's surface can be determined. Unfortunately palaeosols, as with other primary features, become obscured with time, and consequently their identification and usefulness as palaeoenvironmental indicators diminishes. This is particularly true in Precambrian terranes which, commonly, have undergone protracted and/or multi ple episodes of deformation and metamorphism. This obscuring of palaeoweathering features with time makes their identification more difficult, and there-
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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known about Precambrian palaeosols in the various shield areas around the world, and concludes by identifying problem areas, and areas that warrant further investigation, in Precambrian palaeosol research. Although there is an unavoidable bias towards the Canadian data base used by the author, it should be recognized that the Canadian Shield is one of the best studied shield areas in the world (e.g. Hoffman, 1989, with extended references), and that it contains numerous phenomena that are interpreted to be Precambrian palaeosols (Gall, 1992a).
I D E N TIFYI N G A P R ECAM B RIAN PALAE O S O L
The criteria that are most useful in identifying palaeosols of Precambrian age are essentially those that are used to identify younger Phanerozoic palaeosols (e.g. Birkland, 1984; Retallack, 1988; Allen & Wright, 1 989), and characterize modern soils. One must bear in mind that, however, compared with the present, there appears to have been a reduced level of organic activity within Precambrian palaeosols, and a relatively unoxygenated Archaean atmosphere. Fur thermore, as noted for most Precambrian terranes, the extent of deformation and metamorphism must be considered before deciding whether a particular geological phenomenon is a palaeosol. Both physical and chemical criteria can be used to identify a Precambrian palaeosol; the former being immediately applicable in the field, the latter gener ally taking more time and expense to assess. The main physical features to consider are locality, destratification-horizonation, macrostructures and microstructures, and mineralogy. Colour, reflecting mineral chemistry, also can be considered in con junction with geochemical data from normalized geochemical profiles and chemical weathering indices. Palaeosols, by their very nature, should be strati form features that are conformable to unconform ities. Those located along angular unconformities, and obvious disconformities, are easier to see than those developed along subtle paraconformities, espe cially if a structural deformation fabric, or porphy roblastic metamorphic mineral growth, has overprinted the unconformity. Destratification, or the structural disruption of the protolith, is more prevalent in Precambrian palaeosols than. soil stratification (horizonation). In part this may be the result of erosion of the upper parts of the palaeosol
profiles. In addition, a diminished organic activity may have contributed to there generally being less physical and chemical stratification in Precambrian palaeosols. Typically, a sharp-topped saprolitic C horizon changes downward to uneven, incipiently weathered or unweathered protolith (R). Less typical are clay-rich horizons (At) above the saprolith. The South African Waterval Onder palaeosol (Retallack, 1986), Scottish Sheigra palaeosol (Allison et al., 1 992; Retallack & Mindszenty, 1994), and Canadian Steep Rock, Thelon and Elliot Lake palaeosols (references in Gall, 1992a) are examples of Precambrian palae osols with uppermost clay-rich zones. Macrostructures that have been useful in identify ing Precambrian palaeosols include karst features, duricrusts and corestones. Irregular topography, dolines, grikes, cavernous porosity, collapse struc tures and coated grains have been identified in Pre cambrian palaeosols developed on Precambrian carbonate sequences. The correlative Kanuyak Island (Pelechaty & James, 1 991) and September Lakes (Kerans et al. , 1 981) palaeosols in Canada, and the Malmani Dolomite in South Africa (Button & Tyler, 1979), for example, contain some of these karst features. Nodular to massive mineral accumulations along unconformities have been interpreted as duri crusts associated with Precambrian palaeosols and palaeosurfaces. The variety of such mineral accumu lations include calcic Bk horizons or caliche (Sochava et al. , 1975; Kalliokoski, 1986), dolocrete (Gall, 1992b ), silcrete (Mustard & Donaldson, 1 990; Martini, 1 994) and manganese-iron (Button & Tyler, 1981). One must be cautious that these accumula tions may in fact be post-burial, subsurface features. For example, the author (Gall, 1 994) has re-examined the Thelon silcrete (Ross & Chiarenzelli, 1985) and Fault River dolocrete (Gall, 1992b) occurrences and found that early diagenetic quartz and dolomite cementation in sandstones overlying, respectively, the Thelon palaeosol and Fault River palaeosol, were indistinguishable from the duricrust cements. This suggests that some of these mineral accumulations are post-burial features, spatially related to the palaeosurface, and may have formed by similar processes as those described for groundwater silcrete (Thiry & Milnes, 1991) . Corestones, some times exhibiting concentric exfoliation shells, can be easily recognized above a protolith. The protolith itself commonly displays an upward structural/tex tural disruption from weathering. The Flin Flon palaeosol in Canada, and the Sheigra palaeosol in Scotland (Retallack & Mindszenty, 1994) are two
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Precambrian palaeosols examples that display a disrupted protolith overlain by cores tones. Microstructural cutans (ferrans and argillans) and peds are important features in palaeosol identi fication. Precambrian examples of iron- and/or clay rich cutans surrounding aggregates or grains (peds) of less altered material have been described from the Thelon (Gall & Donaldson, 1990) and Elliot Lake (Gay & Grandstaff, 1980) palaeosols in Canada, and pre-Pongola (Grandstaff et al. , 1986) and Water val Onder (Retallack, 1986; Retallack & Krinsley, 1 993) palaeosols in South Africa. Mineral assemblages alone are not indicative of palaeoweathering. A stratiform rock with a mineral ogy that is in contrast with adjacent rock units, however, may be used in conjunction with other crite ria as evidence for palaeoweathering. A wide variety of original protolith and alteration minerals may exist in a palaeoweathering profile. A survey of reported Precambrian palaeosols in Canada (Gall, 1992a) indicates that protolith mineral alteration typ ically decreases with stratigraphical depth, and that even the theoretically most resistant minerals (e.g. quartz, alkali feldspar, zircon, muscovite) can be altered. Neomorphic pedogenic minerals include illite, kaolinite, chlorite, calcite, dolomite, haematite, magnetite, leucoxene, anatase, quartz, tourmaline, goethite, mixed-layer phyllosilicates, vermiculite and gibbsite. Some metamorphosed palaeosols from the same survey (Gall, 1992a) had the follow ing metamorphic mineral assemblages: white mica-biotite, quartz-sillimanite-feldspar, mixed layer paragonite-illite, hornblende-biotite-garnet carbonate mineral, and paragonite-muscovite chlorite. Colour, reflecting mineralogy, is another useful identifying criterion. Typically, palaeosols developed after c. 2.24 Ga under a more oxygenated atmosphere are reddened from iron oxidation. Older palaeosols are typically a drab colour; yet, even subtle colour contrasts with adjacent protolith colour may help to identify a palaeosol. Geochemical data can complement the aforemen tioned physical criteria towards identifying Precam brian palaeosols. Mineral weathering is accompanied by a change in relative element concentrations. For Precambrian palaeosols in Canada, it has been the author's experience that a characteristic of increased weathering within the saprolitic zone is an increase in element depletion as the unconformity is ap proached. Most palaeosol geochemical profiles are normalized to relatively immobile (insoluble) ele ments (e.g. AI, Ti, Zr); yet there is evidence (Gall,
1992a) that even these elements may have been mo bile during palaeoweathering. Caution must be used when interpreting such data profiles, as there is growing evidence of element addition to palaeosols, particularly potassium, following palaeosol burial. Another useful geochemical tool are palaeoweather ing indices (e.g. chemical index of alteration, Nesbitt & Young (1982); weathering potential index, Reiche (1943); weathering ratio, Chittleborough (1991)). Similar to element profiles within the palaeoweath ered zone, these indices reflect an increased loss of mobile elements relative to less-mobile elements with progressive weathering. Such weathering indices have been helpful in defining a few Precambrian palaeosols (Marmo, 1992; Sutton & Maynard, 1992; Gall, 1994).
PALAE O S O L S --t N P R ECAM BRIAN SHIE L D A R E A S North America
_,,.,...··
A long history of mapping of the Canadian Shield by geologists from the Geological Survey of Canada, provincial surveys and universities, has led to the identification of approximately 40 palaeosols, or features interpreted to be palaeosols, at 36 localities (Fig. 1). Unfortunately, Precambrian palaeosols in the Canadian Shield are generally poorly described, if at all, because the palaeosols are seldom the main focus of the field activity, and often difficult to discern. The apparent paucity of documented Precambrian palaeosols in other shield areas maybe for similar fundamental reasons. The majority of the approxi mately 40 Precambrian palaeosols in Canada are described in a paragraph or less. The most extensively documented Precambrian palaeosols in Canada are those associated with the sub-Huronian Supergroup unconformity (numbers 22 and 24, Fig. 1 ; e.g. Roscoe, 1969; Gay & Grandstaff, 1980; G.-Farrow & Mossman, 1988; Rainbird et al., 1990) and Maton abbee unconformity (numbers 2, 6, 10 and 12, Fig. 1 ; e.g. Macdonald, 1980; Tremblay, 1983; Gall, 1992a, 1 994) primarily as a result of their respective associations with world-class uraniferous quartz pebble conglomerate and unconformity-type uranium deposits. Precambrian palaeosols have been found across Canada, including Precambrian terranes within the western Cordilleran and eastern Appalachian orogens. They range in age from Archaean to
210
Q. Gall USA and South America
Canada Ga
1,34
29
1 .0
30
2.0
�.,10, 12 16
9 35
24 22
3.0
1
8 17,23 18 19 20
1 .0
15 14 27 6 13
jl,.
?
Ga
Ga
33
1,2 5,6
12 11
6 10
1 .0
Australia a n d I n d ia
Europe and Asia
Africa
Ga
Ga 1 .0
1,3,5 1 .0
2.0
9
4,7 2 .0
2 .0
31,32 36 10
2.0 31,32
2,6,7
36
5,6,7,8 9
2,8 12
25
3.0
5 1 13
3.0
11 4
3.0
3.0
Fig. I. Temporal distribution of Precambrian palaeosols in Canada, USA and South America, Africa, Europe and Asia and Australia and India. Numbered palaeosols are placed depending on whether the age of development is well known (beside the time line), or poorly constrained (under?). References for the five columns are as follows: Canada 1 , Jefferson & Ruelle (1986); 2, Gall (1992a); 3, Kerans et a/. (1981); 4, Schau & Henderson (1983); 5, Pelechaty & James (1991); 6, Gall (1992a); 7, Stanworth & Badham ( 1984) ; 8 , Falck et a/. (1991); 9, Aspler & Donaldson (1986); 10, Macdonald (1980);Tremblay (1983); 11, Tremblay ( 1972); 12, Gall ( 1994a,b ); 13, LeCheminant et a/. (1981); 14, Frisch (1982); 15, Jackson & Iannelli (1981); 16, Holland et al. (1989); 17, Corkery (1983); 18, Herd et at. (1976); 19, De Kemp (1987); 20, Wilks & Nisbet (1988); 2 1 , J.M. Franklin (personal communication 1990); 22, G.-Farrow & Mossman (1988); 23, Kimberley & Grandstaff (1986); 24, Rain bird (1980); Rain bird et al. (1990); 25, Vogel (1975); 26, Chandler (1988); 27, Eade (1966); 28, Chown & Caty (1983); 29, Bostock (1983); 30, Ryan (1981); 3 1 , Smyth (1976); 32, Knight & Morgan (1981); 33, Patel (1977); 34, Mustard & Donaldson (1990); 35, Campbell & Cecile (1981); Grotzinger et al. (1989); 36, Ray (1975). USA and South America 1 , Kalliokoski (1975); 2, Elston & Scott (1976); 3, Sharp (1940);4, Southwick et at. (1986); 5, Shride (1967); 6, Zempolich et al. (1988); 7, Grambling & Williams (1985); 8, Gibson (1987); 9, Kroonenberg (1978); 10, Cox (1967); 11, Zalba et al. (1992); 12, B arrio et al. (1990). Africa-1, Retallack (1986); 2, Kimberley & Grandstaff (1986); 3, Edelman et at. (1983); 4, Kimberley & Grandstaff (1986); 5, Wiggering & Beukes (1990); 6, Bertrand-Sarfati & Moussine-Pouchkine (1983); 7, Coetzee (1940); Martini (1994); 9, Button & Tyler (1981); 10, Mendelsohn (1961); 11, Hunter (1962); 12, De Jager ( 1964); 13, Button & Tyler (1981). Europe and Asia 1 , Williams (1968); 2, Marmo ( 1992); 3, Kabata-Pendias (1984);4, Chayka & Zaviyaka (1968); 5, Sklyarov & Khromtsov (1972); 6, Sochava et al. (1975); 7, Koryakin (1971); 8, Dodatko et al. (1973); 9, Rankama (1955). Australia and India 1, Miller et a/. (1992);2, Macfarlane et al. ( 1 994); 3 , Lowe (1983) ; 4, Muir (1983 ) ; 5, Dash et al. (1987) ; 6, Kamineni & Rao (1988); 7, Sengupta et at. (1990); 8, Sharma (1979); 9, Golani (1989).
Neoproterozoic. The question-marked columns in Fig. 1 (above), however, indicate that the age of palaeosol development is often poorly constrained. The age uncertainties are not unique to the Canadian Shield. One of the problems with determining the time of palaeosol development in any shield area is the lack of fossils or minerals suitable for dating, hence it is difficult to actually date evolutionary trends in palaeosol development and, thus, any inferred environmental or atmospheric changes. Although its development is poorly constrained in time, the palaeosol at Point Lake (no. 4; Fig. 1) in the
Slave structural province, described by Schau & Henderson (1983) and Easton (1985), is potentially the oldest palaeosol in Canada. It is developed on 3.155 Ga granite and is overlain by coarse siliciclastic rocks and minor volcanic rocks of the Yellowknife Supergroup Keskarrah Formation. Elsewhere, vol canic rocks of the Yellowknife Supergroup were extruded between 2.65 Ga and 2.70 Ga. The palaeosol, which has been subjected to low-grade metamorphism, has not been studied in detail. Chem ical analyses of this palaeosol, which contains white mica, quartz, feldspar and biotite, indicate that it has
Precambrian palaeosols been depleted i n Fe2+, Ca, Mg, S r and N a , and enriched in AI, Ti, Zr, K, Fe3+ and H20 during palaeo weathering (Schau & Henderson, 1983). Subsequent diagenetic dolomite cementation appears to have overprinted some parts of the palaeosol. Better constrained in time are two Archaean palaeosols near Steep Rock Lake, Ontario (no. 20; Fig. 1). One palaeosol, developed on 3.0 Ga grano diorite and tonalite of the Marmion batholith and covered by the 2.9 Ga(?) to 2.7 Ga Steep Rock Group, has been described by Schau & Henderson (1983) and Wilks & Nisbet (1988). The 1-2 m thick white- and brown-mottled palaeosol developed on the granodiorite, has a metamorphic mineral assem blage of paragonite, muscovite, chlorite, quartz and opaques. Chemical analyses display an ascending decrease in Si, Ca, Ba, Mg, V and Fe2+, and an ascend ing increase inTi, Li, B, Zr and H2 0 within the profile. Schau & Henderson's study also offers a rare exam ination of REE mobility in Precambrian palaeosols. They demonstrate that La, Ce, Pr and Nd are concen trated (perhaps in clay minerals) whereas Tb, Dy, Ho, Er, Tm and Yb are depleted with respect to the fresh granodiorite. A zone of chlorite, 'clays', muscovite and calcite developed on the tonalite is considered to be a potential remnant pedogenic C horizon by Wilks & Nisbet ( 1988). The second palaeosol, within the Steep Rock Group, has been described by Wilks & Nisbet (1988), although recognized originally by Wegnast ( 1954) and Jolliffe ( 1955). The Buckshot Ore within the Manganese Paint Rock Member consists of dark, haematitic pisolites and lithic fragments set in a light coloured matrix of haematite, goethite, kaolinite and gibbsite. The Manganese Paint Rock Member discon formably overlies the stromatolitic Mosher Carbon ate. Overlying the Manganese Paint Rock Member is the Goethite Member consisting of brecciated to stratified rocks composed of haematite, geothite, quartz and kaolinite, similar to the Buckshot Ore matrix composition. The Buckshot Ore may repre sent an iron bauxite developed during deposition of a mudstone rich in AI, Fe and Mn following, or, during karstification of the Mosher Carbonate. The youngest (best time-constrained) Precam brian palaeosols occur in the Canadian Cordillera. In the Redstone Copper Belt of the Mackenzie Moun tains (no. 1; Fig. 1 ) , Jefferson (1983) and Jefferson & Ruelle ( 1986) reported a palaeosol formed on the Little Dal Group basalts and overlain by the Coates Lake Group sedimentary rocks. It is characterized by a reddening and saprolitic disruption of the basalt,
211
which has been depleted in sodium and enriched in titanium during palaeoweathering. The palaeoweath ering event is bracketed between a Rb-Sr date of 766 Ma for sills intruding the Tsezotene Formation, which are considered to be similar to sills within the Little Dal Group, and an age of 751 Ma for the informally named Mount Harper Group, which can be corre lated with the Coates Lake Group (Roots & Parrish, 1988). In the Ogilvie Mountains of the Yukon Terri tory (no. 34; Fig. 1), Mustard & Donaldson ( 1990) have documented various features along the dis conformity between the Mount Harper Group and underlying Fifteenmile Group, which they attribute to subaerial exposure and palaeoweathering of the Fifteenmile Group. These palaeoweathering .features include in situ brecciated fragments of the Fifteen mile Group dolostone cemented by dolomite, chert, chalcedony and megaquartz, vugs and geopetal sedi ment fill, all possibly related to karst weathering, cal cretization and subsequent silcretization. The' rest of the Precambrian palaeosols in the Canadian Shield are, temporally, fairly evenly spread between the oldest and youngest occurrences (Fig. l ) , the exception being four palaeosols associated with the Matonabbee unconformity clustered at about 1 .74 Ga in Fig. 1 (numbers 2, 10, 12 and 6). As men tioned previously, because of the large, high-grade uranium deposits associated with the Matonabbee unconformity, the palaeosols developed beneath the Proterozoic Athabasca and Thelon basins are two of the more thoroughly documented Precambrian examples in Canada. The sub-Athabasca Basin palaeosol has been doc umented at many localities in Saskatchewan and Alberta (e.g. Sproule, 1938; Macdonald, 1980; Trem blay, 1983; Wilson, 1 985). The time of formation of this palaeosol is constrained by a c. 1 .89 Ga cooling age for Hudsonian metamorphism of palaeosol protolith lithologies, and a 1 .70-1 .65 Ga age for dia genetic phosphates within the Athabasca Group. Perhaps the most striking feature of this palaeosol is an ascending horizontal colour (mineral) zonation from fresh protolith to grey-green (chloritic) basal palaeosol, to white (illitic), to a red (haematitic) upper palaeosol. This is particularly evident in metasedimentary protoliths (Macdonald, 1 980), where the profiles are up to 70m thick including an uppermost bleached zone, attributed to reduction and/or additional illite and kaolinite formation during diagenesis in the overlying Athabasca Group. Macdonald (1980) has likened the palaeosol pro files to modern laterites. Both Macdonald (1980) and
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Tremblay (1983) noted consistent ascending decreases in Si, Fe2+, Ca and Na, and an increase in Fe3+, within the palaeosol compared to fresh protolith. The undeformed and unmetamorphosed sub Thelon Basin palaeosol (no. 12, Fig. 1) has also been described from widely separated localities beneath the basin (e.g. Chiarenzelli, 1983; Davidson & Gandhi, 1989; Gall, 1994a,b ). The time of palaeosol formation is constrained between the youngest pro tolith, c. l .76 Ga Pitz Formation, and c. l .72 Ga diagenetic apatite in the overlying Thelon Formation. The Thelon palaeosol, which attains thicknesses of 60 m, is generally characterized by the following changes from fresh protolith, through palaeosol, to the base of the overlying Thelon Formation: 1 saprolitic disruption of the proto lith; 2 reddening of the palaeosol as a consequence of iron oxidation; 3 increase in dissolution of feldspar and ferromagne sian minerals accompanied by an increase in · authi genic kaolinite, illite and haematite cement; 4 a loss of Ti, total Fe, Fe2+, Mn, Ca, Na, K, P, U, Zr, Rb, Sr and Ba; 5 depletion in REE, but no fractionation between LREE and HREE; 6 a gain in Fe3+; 7 an increase in chemical index of alteration (CIA of Nesbit & Young, 1982) values. Within the USA, the North American Shield is not as well exposed, and there are fewer documented palaeosols in the, dominantly, Proterozoic age rocks (Fig. 1 ) . The oldest occurrence of palaeoweathering, known to the author, is that des.crib�d by Sharp (1940) within the Grand Canyon sequence, along the Ep-Archaean unconformity beneath the Hotauta Conglomerate. Up to 3 m of red-brown grus is developed where plagioclase and ferro-magnesium minerals in the Archaean granite and schist have been altered. Hi'gher in the Grand Canyon Super group, Elston & Scott (1976) have noted a ferrugi nous zone 1 5 m thick on the c. l . 1 Ga Cardenas Lavas (basaltic) beneath sandstone of the Nankoweap Formation. Two palaeosols of interest because of their tempo ral correlation are those beneath the Palaeoprotero zoic Sioux Quartzite in Minnesota (no. 4; Fig. 1) and Palaeoproterozoic Ortega Group in New Mexico (no. 7; Fig. 1). Dott (1983) and Soegaard & Eriksson (1986) have correlated these and other quartzites, based on broad stratigraphical and geochronological similarities. In the first area, Southwick & Mossier
(1984) and Southwick et al. (1986) have described a palaeosol 9.8 m to > 23 m thick developed on Archaean lithologies, replete with mineral and tex tural degradation, including corestone development. In New Mexico, the potential protolith for a meta morphosed manganese- and aluminium-rich layer along the contact between the 1 .72 Ga Vadito Group schist and the Ortega Group quartzite (pre-1 .69 Ga), is a weathered felsic volcanic rock (Grambling & Williams, 1985). Studies by Gable & Sims (1969) and Gibson (1987) have also identified Mesoproterozoic metamor phosed palaeosols in Colorado (no. 8; Fig. 1 ) . Based on mineralogy, Gable & Sims (1969) have suggested that cordierite-gedrite-bearing gneisses may have originated as weathered rock. Possibly correlative with the cordierite-gedrite rocks, is a phyllite unit in the Needle Mountains, which has been interpreted as a metamorphosed and deformed palaeosol by Gibson (1987). Originally formed sometime between 1690 Ma and 1430 Ma, the phyllite beneath the sedi mentary Uncompahgre Group displays vertical element depletions and relict gneissic and granitic textures within the phyllite zone, which are reminis cent of a palaeosol. Examples of subaerial exposure and weathering of carbonate rocks have been documented in the Neoproterozoic Mescal Limestone (Shride, 1967; Beeunas & Knauth, 1985) and Beck Spring Dolomite (Zempolich et al. , 1988) in the south-western USA. Apart from the Mescal Limestone displaying typical karst features, Beeunas & Knauth (1985) and Zem polich et al. (1988) have found that the carbonate palaeosurfaces have a meteoric stable isotope signature. South America
A few Proterozoic-age palaeosols have been identified in shield regions of the South American continent. The oldest palaeosol identified (no. 10; Fig. 1) appears to be that described by Cox (1967) and Hendrickson (1984), for example, found beneath the gold-bearing Jacobina Series conglomerate within the Atlantic Shield, Brazil. A quartz, muscovite and kyanite layer along the unconformity between the granitic gneiss and overlying auriferous conglomer ate is interpreted to be a metamorphosed saprolite. Kroonenberg (1978) described a palaeosol beneath the Roraima Formation within the Guiana Shield. The palaeoweathered, reddened granite protolith contains a mica and carbonate alteration
Precambrian palaeosols mineralogy in structures reminiscent of a skelsepic plasma fabric. This palaeosol is of further interest considering the role of the palaeoweathering event on local gold and placer diamond occurrences, and, as will be reiterated, because the Roraima Formation and underlying palaeosol can be correlated strati graphically with the Sioux Quartzite in the USA (Rogers et al. , 1984), and in turn, chronostratigraphi cally correlated with the Athabasca, Thelon, Elu Inlet and Hornby Bay quartz arenites and palaeosols in Canada (numbers 10, 12, 6 and 2, respectively, Fig. 1 ). In the Platian Shield, Argentina, Barrio et al. (1990) and Zalba et al. (1992) , respectively, described Neo proterozoic karst and saprolitic unconformities below and within the Tandilia System. Europe-Asia
Bridging the gap between North America and Europe as part of the North Atlantic Craton, the exposed Precambrian rocks in Scotland contain one of the better documented palaeosols (e.g. Williams, 1968; Russell & Allison, 1985; Allison et al. , 1992; Retallack & Mindszenty, 1994). Noted earlier for its clay-rich zone, the 1-3 m thick palaeosol is developed on gneiss and amphibolite and is overlain by Torri donian sandstone. Formed sometime between 1 .7 Ga and 0.8 Ga (Fig. 1 ) , the palaeosol is interesting not only because it contains an upper At horizon, but also because it contains well developed corestones, cutans and peds, pedogenic smectite (still preserved ! ) , and evidence of physical and chemical diagenetic over printing. Perhaps the oldest, and best documented, palaeosol in the Baltic Shield is the Palaeoprotero zoic Hokkalampi palaeosol in Finland and its equiva lents in central Karelia, Russia. Marmo ( 1992) has thoroughly described this metamorphosed palaeosol, developed on granitic rocks and Sariola Group sedi mentary rocks, which formed between 2.5 Ga and 2.2 Ga. Up to 80 m thick, the Hokkalampi palaeosol is a quartz-sericite ± kyanite and andalusite schist which, despite metamorphism, appears to have been depleted in Fe2+, Na, Ca and Mg during palaeo weathering. In central Karelia, Koryakin (1971) and Sochava et al. (1975) documented palaeoweathered granite, and caliche, at the base of the Jatulian Series. Perhaps also correlatable with the Karelian palaeosols, is an in situ diorite breccia overlain by Bothnian schist, which, owing to the dominance of ferrous iron, has been interpreted by Sederholm (193 1 ) and Rankama (1955) as a c. 2.0 Ga palaeo-
213
weathering zone formed under anoxic atmospheric conditions. In north-eastern Poland, an outlier of the Baltic Shield contains one of the youngest Precambrian palaeosols in the East European Craton. Kabata Pendias (1984) and Kabata-Pendias & Ryka (1989) have described palaeosols, up to 30 m thick, devel oped on several pre-Vendian metamorphic and igneous protoliths. Perhaps the oldest noted weather ing (Fig. 1) is located in the East European Craton within the central Ukranian Shield. Dodatko et al. (1973), for example, have interpreted a quartz sericite schist developed on plagiogranite at the base of the Krivoy Rog Series, as a metamorphosed Pre cambrian weathering crust. Examples of Precambrian palaeosols also can be found in exposed parts of the Siberian Craton. Within the Anabar Shield, for example, Chayka & Zaviyaka (1968) and Chayka (1970) described oxidized weath ering crusts on flow tops of the c. 1485 Ma Mukum Series basalt. The Aldan Shield to the east, also con tains Precambrian palaeosols. In the Uchur-Maya region, Sklyarov & Khromtsov (1972) describe meta morphosed Neoproterozoic rocks containing corun dum, sillimanite and andalusite, which they interpret as metamorphosed bauxites. Africa
A long history of geological exploration in Africa, and the recognition of the economic importance of Precambrian unconformities (e.g. Button & Tyler, 1981), has resulted in the identification of at least 13 Precambrian palaeosols (Fig. 1 ) , primarily in South Africa. The oldest palaeosol in Africa appears to be developed on a c. 3.0 Ga to 3 . 1 Ga granitic to gneissic Archaean protolith beneath the Inuzi Subgroup of the Pongola Supergroup (Matthews & Scharrer, 1 968; Edelman et al. , 1983; Kimberley & Grandstaff, 1986). In Swaziland, an aluminous phyllite unit within the Inuzi Subgroup, developed on basalt, has been interpreted by Button & Tyler (1981) as a palaeo weathered zone; although Hunter (1962) suggested previously that the phyllite unit is a metamorphosed aluminous sediment. Palaeoweathered Archaean granite also underlies basal conglomerate of the c. 2.8 Ga to 2.9 Ga Dominion Reef Group (no. 3; Fig. 1 ), Witwatersrand Supergroup (no. 9; Fig. 1), and quartzite of the Black Reef (no. 2; Fig. 1). A sericite, pyrophyllite, chloritoid and leucoxene mineral assemblage in the Jeppestown Shale within the basal Witwatersrand Supergroup has been interpreted by
214
Q. Gall
De Jager ( 1964) as a metamorphosed palaeosol (no. 12; Fig. 1 ) . A metamorphosed bauxite is considered to be the precursor to sillimanite- and corundum bearing rocks in Namaqualand by Coetzee (1940), De Jager (1963) and Frick & Coetzee (1974) . A rare example of palaeoweathered ultramafic rocks has been thoroughly documented by Martini ( 1994). The c. 2.6 Ga palaeosol (no. 8; Fig. 1) beneath the Black Reef Quartzite (Chuniespoort Group), consists of a saprolite developed on serpentinized dunite, and altered to varying degrees by silcrete and dolocrete. Geochemical profiles through the sapro lite show major and trace element mobility, including minor nickel mineralization. A few younger, Palaeo proterozoic palaeosols have been described from southern Africa (Fig. 1 ) . Notable amongst these are the Watervaal Onder palaeosol (Retallack, 1986; Retallack & Krinsley, 1993), which contains microstructures similar to those found in modern vertisols; and, perhaps, the world's oldest karst surface formed on the Malmani Dolomite (Button & Tyler, 1981). The youngest evidence of Precambrian palaeoweathering in Africa consists of pedogenically altered dolomite of the Neoproterozoic Sarnyere Formation, Mali (Bertrand-Sarfati & Moussine Pouchkine, 1983). India
Few Precambrian palaeosols have been described from India, despite there being numerous exposures of Precambrian rocks. Certainly, Precambrian meta morphism and Phanerozoic weathering episodes have made identification of the Precambrian palaeo sols more difficult. One of these rare, and metamor phosed, Precambrian palaeosols has been described by Sharma (1 979) from the Bundelkhand complex in north-central India (no. 8, Fig. 1). Here, a phyllite from within the Archaean Palar Formation contains nodular diaspore in pyrophyllite, and has been inter preted as a metamorphosed laterite. A number of studies have suggested that the granulite-grade gneiss-granite terrane of the eastern Ghats contain metamorphosed Archaean bauxite deposits (e.g. Dash et al. , 1987; Kamineni & Rao, 1988; Golani, 1989; Sengupta et al. , 1990). Thin layers and lenses containing quartz-sillimanite-garnet, quartz-sillimanite, sapphirine-quartz-orthopyrox ene-cordierite-sillimanite and massive sillimanite corundum assemblages, have been interpreted as metamorphosed palaeosols, more because of their bulk rock and mineral chemistry (e.g. high Al2 03
content and Fe20iFeO ratio) than their physical attributes. Further investigation, however, of these lithologies in India (Nanda & Pati, 1991) and in China (Condie et al. , 1992) point out that the thickness and extent of such lithologies, a lack of 'unweathered' protolith, and inconclusive geochemical data, suggest that these aluminous lithologies in granulite terranes may have a sedimentary origin, rather than an in situ palaeoweathering origin. Indeed, the author's re examination of physical and geochemical data on the sapphirine-bearing granulites in Labrador, Canada, suggests that these too cannot satisfactorily be inter preted as metamorphosed laterites, as proposed by Meng & Moore (1972). Australia
Some of the oldest examples of subaerial exposure and palaeoweathering in Australian Precambrian sequences are described from sedimentary litho logies interpreted as representing near-shore, shal low water environments. In western Australia, for example, dissolution-collapse microstructures in the Archaean Strelley Pool Chert, have been interpreted as exposure features by Lowe (1983). In the Northern Territory, Muir (1983) describes calcrete in the Meso proterozoic Amos Formation, McArthur Group, as forming during emergence and meteoric weathering of shallow-water carbonate sediments. More complete palaeosol profiles also have been described from Australia (Fig. 1). Between basalt flows in the Archaean Fortescue Group, Macfarlane et al. (1994) describe 15 m thick sericite-rich zones grading into chlorite-rich zones, and then unweath ered basalt. Major and trace element profiles suggest that the sericite zone and underlying chlorite zone reflect eluvial and illuvial horizons, respectively. In addition, by assuming certain degrees of element mobility (including REE), Macfarlane et al. (1994) suggest that weathering took place under an oxygen poor atmosphere, and that subsequent metamor phism enriched the upper part of the palaeosol in alkali and alkaline earth elements. A younger palaeosol (no. 1 , Fig. 1 ) underlies the Kombolgie For mation sandstone, and is developed on igneous and metamorphic protoliths of the Pine Creek Geo syncline (Miller et al. , 1992). Consistent between numerous saprolitic profiles, is a descending vertical colour/mineral gradation, up to 47.5 m thick, from red/haematitic, to pink-green/illite and chlorite, to a green-grey/chlorite zone above fresh protolith. The palaeoweathering resulted in the ascending deple-
Precambrian palaeosols tion of many elements. However, diagenetic over printing also has been assumed as a result of iron redistribution, the presence of phosphate minerals similar to those found in the overlying Kombolgie Formation, and illite recrystallization immediately below the contact with the Kombolgie Formation. As will be discussed in the section on correlations, the sub-Kombolgie Formation palaeosol is physically, chemically and chronostratigraphically similar to the Matonabbee unconformity palaeosols in Canada.
P R O B L E M S I N R E C O G NITI O N A N D I N T E RP R E TATI O N
There are a number o f problems, o r hurdles, one must overcome in order to recognize Precambrian palaeosols in the first place, and then in interpreting the record of palaeoweathering in a modern context. The basic physical and geochemical criteria that can be used to recognize a Precambrian palaeosol have been discussed at the beginning of this paper. Two major problems remain, however, once a Precam brian palaeosol has been identified. The first, is that generally the record of palaeoweathering appears truncated compared with modern profiles. The majority of Precambrian palaeosols appear to represent only the lower saprolitic, or C horizon, part of the original palaeoweathering profile. This leaves the researcher with an incomplete palaeosol that is hard to classify using modern classification schemes, and therefore to place within a palaeoenvironmental framework. The second problem, is that post-weath ering physical and chemical overprinting of the palaeosol, to some degree, is common in Precambrian palaeosols. Thus, one must try to assess the textural and geochemical changes in the palaeosol profile that can be attributed to burial diagenesis or metamor phism of the palaeosol. Despite the apparently incomplete and truncated nature of most Precambrian palaeosols, it is still pos sible to determine that (i) the phenomenon does represent a palaeosol, and (ii) regardless of which contemporary soil classification scheme is used, a C horizon, regolith or saprolith can be recognized. As noted previously, clay-rich horizons above a saprolith have been described (e.g. Retallack, 1 986; Allison et al. , 1992), and loosely equated to modern At hori zons. Some carbonate-rich, or Bk, horizons also have been recognized. However, there are only a few Pre cambrian palaeosols that contain enough features that allow them to be classified using contemporary
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schemes and therefore the palaeoenvironment in which they developed can be better defined. The apparently incomplete, truncated nature of Precambrian palaeosols, where fully developed palaeosols, equivalent to modern profiles, are the exception rather than the rule, may be indicating something fundamentally different to our present line of thinking. That is, perhaps, under a less oxygenated Precambrian atmosphere where a terrestrial organic carapace was sparse, or not present at all, a saprolite may be all that formed during weathering. The incipient weathering observed in most Precambrian palaeosols and, where developed, the drab appearance of clays in the palaeosols, prompted Retallack (1986) to refer to the palaeosols as 'green clays' that may represent an extinct order of soils. The problem of identifying original features once the palaeosol has been physically or chemically over printed by burial diagenesis or metamorphism, is also of concern; especially in Precambrian terranes, which typically have been subjected to protracted and repeated episodes of deformation and metamor phism (Gall, 1 992a). Generally, it has been found that basinal (diagenetic) fluids alter underlying 'base ment' lithologies along basal unconformities (e.g. Al Gailani, 1 981; Duffin et al. , 1 989; Bethke & Marshak, 1 990). Where palaeosols, including those of Precam brian age, exist along the unconformity, they too are overprinted. Potassium overprinting appears to be common in Precambrian palaeosols (e.g. Matthews & Scharrer, 1 968; Eriksson & Soegaard, 1 985), often resulting in illitized kaolinite, or other phyllosilicate precursors; but examples of overprinting by carbon ate, quartz and iron oxide minerals also have been described. Further details on how to recognize physi cal and chemical overprinting of palaeosols can be found in Nesbitt & Young (1989), Retallack (1991) and Gall (1992a). The task of identifying original pedogenic features is just as daunting when metamorphic textures have not developed, as incipient diagenetic or low-grade metamorphic changes can be detected only by detailed petrography, including mineral analyses, and geochemical profile analysis. An example of subtle post-weathering alteration is found in the Palaeopro terozoic Thelon palaeosol in Canada (Gall, 1994b) . The Thelon palaeosol has been structurally and chemically altered during burial diagenesis of the overlying Thelon Formation to produce: 1 stylolites; 2 local quartz and haematite veins;
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3
a microcrystalline aluminium phosphate-sulphate mineral; and 4 the pervasive illitization of kaolinite. Phyllosili cates in the Thelon palaeosol are commonly indistin guishable, physically and chemically, from those formed diagenetically within the overlying Thelon Formation. Furthermore, it was found that 8D and 8180 compositions of kaolinite in the Thelon palaeosol are almost identical to kaolinite in the overlying Thelon Formation, and is considered to have been re-equilibriated with burial diagenetic fluids emanating from the Thelon Formation. Perhaps the key message here is that overprinting of a buried palaeosol will affect adjacent lithologies as well. Hence, it is critical to recognize the physical, mineral and geochemical trends in adjacent litho logies before original pedogenic features can be identified in the palaeosol. Once a Precambrian palaeosol has been recog nized, another problem of particular concern is deter mining its age. Difficulties arise in determining when, and for how long, a particular palaeosol (palaeoenvi ronment) existed and therefore in accurately deter mining palaeoenvironmental trends through time. Owing to the generallly poor preservation of the palaeosols, and paucity of organic material, Precam brian palaeosols do not contain datable material. As demonstrated in Fig. 1 , the ages of many Precambrian palaeosols are poorly constrained in time. As a result of diagenetic or metamorphic overprinting, phyllosil icate age determinations typically yield ages younger than the episode of weathering. The Precambrian palaeosols for which the time of formation is best constrained, come from areas in which the geo chronology of immediately adjacent lithologies has been determined (i.e. ages of youngest protolith and oldest overlying stratigraphical unit), or from areas where stratigraphy can be correlated with areas where the geological history is better known.
N E W H O RI Z O N S
Although much work remains to b e done i n basic recognition and analyses of Precambrian palaeosols, once recognized, they may be useful for more than reconstruction of the immediate palaeoenvironment. Two areas in which the identification of palaeosols in Precambrian terranes may be of further help are: 1 in aiding regional stratigraphical correlation and, hence, identifying regional to world-wide episodes of palaeoweathering;
2
in helping to track unconformity-related mineralization. With respect to stratigraphical correlation, once a palaeosol has been placed within a stratigraphical context, the recognition of a weathered unconfor mity, especially in deformed and metamorphosed ter ranes, may help in establishing new correlations with other regions where weathered unconformities exist, or it may help reinforce previously suggested correla tions. Examples of such correlations exist within the Canadian Shield, and between the Canadian Shield and other shield areas. In eastern Canada, the pres ence of Precambrian palaeosols beneath the Ramah Group and Mugford Group (numbers 3 1 and 32; Fig. 1) have strengthened the correlation proposed by Smyth & Knight (1978) and Ermanovics et at. (1989) based on stratigraphical and structural similarities between the two overlying groups. Unfortunately, poor time constraints on the formation of these two cOl-relatable palaeosols does not lead to a good esti mation for the age of this palaeo weathering event. In the north-west part of the Canadian Shield, the four palaeosols (numbers 2, 6, 10 and 12, Fig. 1 ) associated with the Matonabbee unconformity also may be cor related (Gall, 1992a). Similarities between these palaeosols, their stratigraphical position and con strained time of formation, suggest that a c. 1 .72 Ga weathering event affected thousands of square kilometres of the Canadian Shield. In a tectono stratigraphical context, development of a palaeo weathered surface over such a large area would require relative tectonic quiescence apd emergence over that same area. This is exactly thy scenario that would have existed in the north-western part of the Canadian Shield, following amalgamation of Archaean and Palaeoproterozoic crust and the for mation of the Laurentian supercontinent between 1 .9 Ga and 1 .7 Ga, as described by Hoffman (1 989). As previously mentioned, in the USA there are Palaeoproterozoic palaeosols beneath the Sioux Quartzite and Ortega Group that can be temporally and, in general, stratigraphically correlated. These sedimentary units and palaeosols can, in turn, be cor related with the quartz arenite units and underlying Matonabbee unconformity palaeosols in the north western Canadian Shield (Dott, 1983; Soegaard & Eriksson, 1986; Hoffman, 1989). It is interesting to note that these Palaeoproterozoic quartzose sedi mentary units and underlying palaeosols in North America are, generally, stratigraphically and tem porally, correlatable with the Palaeoproterozoic Roraima Formation and underlying palaeosol in
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Precambrian palaeosols Venezuela and Guyana (Gibbs & Barron, 1983; Rogers et al. , 1984), and beneath the Kombolgie For mation (Miller et al. , 1992) in Australia. If the tectonic reconstruction for Laurentia (e.g. Dalziel, 1992) is considered, these presently widely separated areas are brought much closer together. It is feasible therefore that Precambrian palaeosol records can be used to support old, and suggest new, regional corre lations, and delimit areas and times of widespread weathering. Similarly, in the Fennoscandian Shield, Marmo (1992) has noted that the c. 2.2 Ga Hokkalampi palaeosol can be correlated with palaeosols else where in Finland and Russia to define a period of continent-wide Palaeoproterozoic weathering. Ojakangas (1988) and Marmo (1992) also have sug gested that the Hokkalampi palaeosol may be corre lated temporally and stratigraphically with the Elliot Lake palaeosol in Canada. This last correlation has important economic implications, because the Elliot Lake palaeosol is overlain by uraniferous quartz pebble conglomerate ore deposits. Further investiga tion of this correlation is necessary to see if the Hokkalampi palaeosol correlates with the Elliot Lake palaeosol, or the slightly younger and strati graphically higher Lake Timiskaming palaeosol (Fig. 1), and therefore to evaluate the economic potential of the Hokkalampi palaeosol and adjacent conglomerates. With respect to unconformity-related mineraliza tion, weathered unconformities and their attendant palaeosols may be linked to ore-forming processes at various stages in their development. They may represent: 1 weathered source regions for placer deposits; 2 sites for both residual and supergene mineralization; 3 following burial, they may act as physical or chemical 'barriers' to mineralized hydrothermal fluids. A variety of Precambrian mineral deposits appear to be linked to weathered unconformities. For example, deeply weathered igneous rocks that have been found beneath, and peripheral to, placer uranium and gold deposits in Canada and South Africa, may have been the source for the detrital ores (Button & Tyler, 1981; Pretorius, 1981). Similarly, Pagel (1991) has suggested that lateritic weathering during the development of the Athabasca palaeosol helped liberate and pre-concentrate uranium prior to formation of the world-class unconformity-type uranium deposits in northern Saskatchewan. Some
models for the formation of the unconformity-type uranium deposits in northern Saskatchewan (e.g. Boeve & Sibbald, 1 978), view the same weathered Palaeoproterozoic unconformity as being a 'barrier' and site for precipitation of uranium ore from migrat ing hydrothermal fluids. Examples of Precambrian palaeosols as residual ore deposits also have been described. In Finland and South Africa, kyanite and sillimanite have, respectively, been mined from rocks interpreted to be metamorphosed aluminous Pre cambrian palaeosols (Coetzee, 1 940; Marmo, 1 992) . The manganese oxide and fluorite-lead-zinc deposits associated with the karst unconformity at the top of the Palaeoproterozoic Malmani Dolomite, South Africa, also can be linked to the weathering and post-burial history of the unconformity (Button & Tyler, 1981 ). Through this brief summary the author is suggesting that once recognized, weathered Pre cambrian unconformities and their attendant palaeosols have the potential for hosting, or being linked to, a number of different mineral deposits for a variety of reasons. For example, as with younger soils, there is also potential for Precambrian palaeosols to host, or lie adjacent to, supergene or lateritic Cu, Ni, and Co-Mo-Ag mineralization. Especially in areas where unconformity-related min eralization is known, recognizing palaeoweathering is important not only in model-driven exploration, but also in assessing the potential for certain types of mineralization to be associated with the unconformity.
AC K N O W L E D G E M E N T S
The author would like to acknowledge his colleagues and their institutions who are part of this IGCP Project 317, for their unreserved cooperation during the course of this project, and for their helpful sugges tions and careful review of this manuscript. Any inac curacies or misrepresentations in this overview of Precambrian palaeosols rest solely on the shoulders of the author.
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Precambrian palaeosols ERIKSSON, K .A. & SOEGAARD, K. (1985) The petrography and geochemistry of Archean and Early Proterozoic sedi ments: implications for crustal compositions and surface processes. Geol. Surv. Fin!. Bull. , 331, 8-32. ERMANOVJCS, J.F., VAN KRANENDONK, M., CORRIVEAU, L., MENGEL, F. , BRIDGWATER, D. & SHERLOCK, R. (1989) The boundary zone of the Nain-Churchill provinces in the North River-Nutak map areas, Labrador. Geol. Surv. Can. Pap. , 89-1C, 385-394. FALCK, H., DoNALDSON, J.A. & HALL, L. (1991) Regolith beneath the Archean Jackson Lake Formation: its impli cations for Yellowknife volcanic belt stratigraphy, Slave province, NWT. Geol. Assoc. Can. Mineral. Assoc. Can. Program Abstr. , 16, A35. FRICK, C. & COETZEE, C.B. (1974) The mineralogy and the petrology of the sillimanite deposits west of Pofadder, Namaqualand. Trans. Geol. Soc. S. Afr. , 77, 169-183. FRISCH, T. (1982) Precambrian geology of the Prince Albert Hills, western Melville Peninsula, Northwest Territories. Geol. Surv. Can. Bull. , 346, 70 pp. G.-FARROW, C.E. & MossMAN, D.J. (1988) Geology of Pre cambrian paleosols at the base of the Huronian Super group, Elliot Lake, Ontario, Canada. Precam. Res. , 42, 107-139. GABLE, D.J. & SIMS, P.K. (1969) Geology and regional meta morphism of some high-grade cordierite gneisses, Front Range, Colorado. The Geological Society of America, Special Paper 128. GALL, Q. (1992a) Precambrian paleosols in Canada. Can. J Earth Sci. , 29,2530-2536. GALL, Q. (1992b) The early Proterozoic Thelon paleosol as part of the Matonabbee unconformity, northwestern Canadian Shield. In: Mineralogical and Geochemical Records ofPaleoweathering, IGCP 317. (Eds Schmitt,J.M. & Gall, Q.). E.N. S. M. P Mem. Sci. de La Terre, 18, 163-174. GALL, Q. (1994a) The Proterozoic Thelon paleosol, North west Territories, Canada. Precam. Res. , 68, 1 1 5-137. GALL, Q. (1994b) Proterozoic paleoweathering and diagene sis, Thelon Basin, Northwest Territories, Canada. PhD thesis, Carleton University, Ottawa, Ontario. GALL, Q. & DONALDSON J.A. (1990) The sub-Thelon Forma tion paleosol, Northwest Territories. Geol. Surv. Can. Pap. , 90-1C, 271-277. GAY, A. & GRANDSTAFF, D.E. (1980) Chemistry and mineral ogy of Precambrian paleosols at Elliot Lake, Ontario. Precam. Res. , 12, 349-373. GIBBs,A.K. & BARRON, C.N. (1983) The Guiana Shield revis ited. Episodes, 1983, 7-14. GIBSON, R.G. (1987) Structural studies in a Proterozoic gneiss complex and adjacent cover rocks, West Needle Mountains, Colorado. PhD thesis, Virginia Polytechnic Institute and State University, Blacksburg, Virginia. GoLAN I, P.R. (1989) Sillimanite-corundum deposits of Son apahar, Meghalaya, India: a metamorphosed Precam brian paleosol. Precam. Res. , 43, 175-189. GRAMBLING, J.A. & WILLIAMS, M.L. (1985) The effects of Fe3+ and Mn3+ on aluminum silicate phase relations in north-central New Mexico, U. S. A. J. Petrol. , 26, 324-354. GRANDSTAFF, D. E., EDELMAN, M.J., FOSTER, R.W., ZBINDEN, E. & KIMBERLEY, M.M. (1986) Chemistry and mineralogy of Precambrian paleosols at the base of the Dominion and Pongola Groups. Precam. Res. , 32, 97-131. GROTZINGER, J.P., ADAMS, R.D., McCORMICK, D.S. & MYROW, P. (1989) Sequence stratigraphy, correlations between ,
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Wopmay Orogen and Kilohigok Basin, and further inves tigations of the Bear Creek Group (Goulburn Super group), District of Mackenzie, N.W.T. Geol. Surv. Can. Pap. , 89-1C, 107-119. HENDRICKSON, B.R. (1984) Stratigraphic position, mineral ogy, depositional environment, and gold distribution ofthe main reef at Morro do Cusnz and Morro do Vento near Jacobina, Bahia, Brazil. Report, South Dakota School of Mines and Technology, 157 pp. HERD, R.K., CHANDLER, F.W. & ERMANOVICS, J.F. (1976) Weathering of Archean granitoid rocks. Island Lake, Manitoba. Geol. Assoc. Can. Mineral. Assoc. Can. Program Abstr. , 1, 72. HoEvE,J. & SIBBALD,T. (1978) On the genesis of Rabbit Lake and other unconformity-type uranium deposits in north ern Saskatchewan, Canada. Econ. Geol. , 73, 1450-1473. HoFFMAN, P.F. ( 1989) Precambrian geology and tectonic history of North America. In: The Geology of North America, Volume a, the Geology of North America -an Overview (Eds Bally, A.W. & Palmer, A.R.), pp. 447-512. Geological Society of America, Boulder, USA. HOLLAND, H.D., FEAKES, C.R. & ZBINDEN, E.A. ( 1989) The Flin Flon paleosol and the composition of the atmosphere 1.8 BYBP. Am. J Sci. , 289, 262-289. HuNTER, D.R. (1962) The mineral resources of Swaziland. Swaziland Geol. Surv. Mines Dept. Bull. , 2, 111 pp. JACKSON, G.D. & IANNELLI, T.R. (1981) Rift-related cyclic sedimentation in the Neohelikian Borden Basin, north ern Baffin Island. Geol. Surv. Can. Pap. , 81-10, 269-302. JEFFERSON, C. W. (1983) The Upper Proterozoic Redstone Copper Belt, Mackenzie Mountains, Northwest Territories. PhD thesis, University of Western Ontario, London, Ontario. JEFFERSON, C.W. & RUELLE, J.C.L. (1986) The Late Protero zoic Redstone Copper Belt, Mackenzie Mountains, Northwest Territories. In: Mineral Deposits of Northern Cordillera. (Ed. Morin, J.A.). Can. Inst. Mining Metal!. 36 154-168. JOLLIFFE, A.W. (1955) Geology and iron ores of Steep Rock Lake. Econ. Geol. , 50,373-398. KABATA-PENDJAS, A. ( 1 984) Chemical weathering of meta morphic rocks of the crystalline basement in north eastern Poland. Bull. lnst. Geol. (Warsaw), 347, 27-80. KABATA-PENDIAS,A. & RYKA, W. (1989) Precambrian weath ering crust in north-eastern Poland. In: Weathering, its Products and Deposits, (Eds B alasubramanian, K.S. et al.) pp. 241-255. Theophrastus Publications, S.A,Athens. KALLIOKOSKJ, J. (1975) Chemistry and mineralogy of Pre cambrian paleosols in northern Michigan. Geol. Soc. Am. Bull. , 86,371-376. KALLIOKOSKJ, J. (1986) Calcium carbonate cement (caliche) in Keweenawan sedimentary rocks (-1.1 Ga), Upper Peninsula, Michigan. Precam. Res. , 32, 243-259. KAMIN EN!, D.C. & RAo, A.T. (1988) Sapphirine granulites from the Kakanuru area, eastern Ghats, India. Am. Mineral. , 73, 692-700. KERANS, C., Ross, G.M., DONALDSON, J.A. & GELDSETZER, H.J. (1981) Tectonism and depositional history of the Helikian Hornby Bay and Dismal Lakes groups, District of Mackenzie. Geol. Surv. Can. Pap. , 81-10, 157-182. KIMBERLEY, M.M. & GRANDSTAFF, D.E. ( 1 986) Profiles of elemental concentrations in Precambrian paleosols on basaltic and granitic parent materials. Precam. Res. , 32, 133-154. ,
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R e gional palaeosurface and p alae owe athering reconstructions
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
Spec. Pubis int. Ass. Sediment. (1999) 27, 225-243
Palaeolandscape reconstruction of the south-western Massif Central (France)
R . S I M O N - C O IN <; O N E C. I. G., cole Nationale Superieure des Mines de Paris, 35 m e St-Honore, 77305 Fontainebleau Cedex, France and UMR SISYPHE C 7619, Universite Pierre et Marie Curie, Paris, France
A B S T RACT
In subsiding areas, successive sequences of deposits progressively bury and protect each other. The sedi mentary record in such areas is influenced fundamentally by the nature of the surrounding source regions, which control the lithological characteristics of the sedimentary infill. Continental uplands, on the other hand, enduring the operation of long and complex subaerial processes throughout changing tectonic and climatic environments, may keep only faint vestiges of these former imprints, making accurate geomor phological interpretation problematical. Clues for unravelling the geomorphological evolution of con tinents were sought by using usual tools of geomorphology (geometrical relationship between the landforms, weathering mantles, erosional processes), complemented by the results of mineralogical and geochemical studies. The south-western Massif Central and its margins provide good case studies for palaeolandscape reconstruction. Its evolutionary history begins with the last uplift of the Hercynian range, at about 290 Ma and ends with the general inundation of the region 20 Ma. Landscape change beyond this time is not con sidered in this paper. Several periods have left distinctive traces in the modern landscape. For example, albitization and dolomitization mark facets of 'the post-Hercynian surface' and are the characteristic products of peculiar geochemical environments. In the same way, calcretes mark an aridification of climates and seal the contemporary Tertiary palaeolandscapes. The typical weathered mantles of crystalline basement appear ubiquitous and are not so reliable for palaeolandscape reconstruc tion because they may have evolved through various palaeoenvironments. Palaeolandforms can survive through a variety of morphogenetic systems and may be partly or totally inherited within younger landscapes. It is obvious that the major part of a landscape is polygenetic, having been submitted to successive degradational and rejuvenation events. Successive marine or lacustrine transgressions may seal and preserve the earlier landforms and palaeoweathered profiles. When exhumed, these palaeolandforms may be integrated in more recent landscapes and will profoundly affect their evolution.
I N T R O D UC TI O N
It is known that on platforms, major planation sur faces often relate to stratigraphical unconformities resulting from erosional reworking during a trans gression with a slow rise of sea-level. By contrast, when sedimentary cover is lacking, recognition of planation surfaces is more ambiguous, especially in the case of extensive, almost flat or hilly surfaces. The term palaeosurface will be used in this paper, in a similar fashion to that suggested by Widdowson (1997a), to mean extensive and even surfaces of a regional scale, and which mark major changes in
This study is concerned with landforms and events taking place on continental areas and, more precisely, on old cratons (shields and platforms). In general, the geomorphology of these cratonic regions (i.e. includ ing shields and platforms) is often characterized by extensive fiat surfaces, broad domes, inselbergs, fiat topped hills and prominent scarps. As noted by Widdowson (1997a), it is difficult to define the origin and evolution of constituent palaeolandforms, and to identify exogenic or endogenic agents that have been responsible for their development.
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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R. Simon-Coinr;on
the geological, geomorphological and sedimentary history of an area. Such surfaces may be of different evolutionary origins and be preserved over long periods of geological time. A useful definition describing palaeolandforms, palaeolandscapes or palaeotopographies is more ambiguous and perhaps more difficult to achieve but, in general, these terms will be reserved here for those features or combination of features occurring at the local scale. In particular the term palaeolandscape will be used to describe a combination of geomor phological features or palaeolandforms existing together at a specific point in time. Reconstruction and assessment of palaeoland scapes form the central theme to this work. The aims are to document the factors controlling (weathering, differential rock resistance, erosion, sedimentation, tectonic changes), and the landforms resulting from, the action of landforming processes over geologically shorter time-scales within long periods of evolution. The evolution of the south-western Massif Central during the Mesozoic and Cenozoic periods is taken as the time-frame to illustrate several landform devel opment processes and the role of inherited landforms in these processes.
P R O B L E M S O F LAN D S CAP E EVO LUTI O N
Continental areas are rather quiet o r stable tectoni cally, and they are composed of roots of Proterozoic and/or Palaeozoic foldbelts. They often include so called 'shield area', where old crystalline basements have long been exposed, and associated marginal platform areas, with thick younger fiat-lying sedi mentary covers. In a very broad view, given that high grade metamorphic rocks are commonly exposed, one can consider that cratonic regions have under gone a slow but sustained rise relative to sea-level; whereas continental platforms have been subject to tectonic changes involving relative uplift and/or sub sidence over time and are associated with deposition of sediments resulting from these processes. Variations in the palaeogeography of basins can be determined readily from their stratigraphical data, which provide detailed information regarding the sedimentological history of the basins and regarding the erosional history of the surrounding emergent regions. As any continental deposits formed above sea-level were subject to erosion and weathering, palaeolandscape reconstructions are much more
difficult. On continental uplands, evidence of land scape change is often scarce and typically azoic. A detailed geomorphological study should provide information about the relief of an area, the mutual relations between the landforms themselves, their origins and age. The investigation devices used in landscape reconstruction include the geometrical study of landforms and their correlatives; the sedi ments of the marginal and intracratonic basins; the residual formations and palaeoweathering profiles. In this fashion it should be possible to follow, through space and time, the development of successive land scapes that have contributed and ultimately led to the morphological characteristics observed today. Geometrical relationships
Studies evaluating the geometrical relationships between the different elements of landscape can be useful in helping understand the evolutionary histo ries of areas (e.g. Lacika 1997; Ringrose & Migon 1997). Using such an approach, it is often relatively easy to map an extremely fiat palaeosurface; for instance, palaeosurface reconstructions are often based on the accordance of summit levels (e.g. Widdowson 1997b ). A palaeosurface could represent the earliest stage of landsurface evolution that can be perceived and be the result of a long history of events. In interpreting such a landscape, the range of affected areas and vertical movements must be taken into account, because they will influence the mor phology and extent of planation and the degree of preservation of the surfaces. Surfaces of different ages are bounded and separated from each other by long, continuous scarps. In such stepped topogra phies, relative dating techniques typically yield a chronological sequence, with the oldest surface in the most elevated position and the youngest ones at the bottom. Correlative deposits
On the border of cratons, where passive continental margins and foreland basins are situated, basement has subsided because of tectonic downwarping and/or sedimentary loading (McKenzie 1984; Gilchrist & Summerfield 1990; Gilchrist & Summerfield 1994). The development of thick sedi ments in these areas may be related to the operation of erosive processes on the continental margins, which often generates planation surfaces. In the ter restrial environment, dating can be achieved, in some
227
Palaeolandscape reconstruction cases, from sediments or volcanic rocks that lie directly upon the surfaces. Alternatively, dating can be traced down into the sedimentary record of adjacent basinal areas and there, recognized from changes in sedimentary record and/or from stratigraphical gaps. Correlations can be established between an erosional cycle on land and deposition in basins at continental margins (e.g. Partridge & Maud 1987). Intracratonic basins also preserve thick deposits, which provide evidence of the evolution of the surrounding land mass. Deposits in basins, however, although they are dependent on continental areas, may not exactly reflect contemporaneous continental evolution. Problems can arise where older successions become reworked by erosion or where other changes in the sedimentary environment cause the record to be degraded or erased. Inheritance
Old peneplains are enduring relief forms and can survive through a variety of different morphogenetic systems without major alteration. Their long-term persistence means that they can often act as reposi tories for sediments in transit, and they can support the development of new weathering covers and soils in response to changing climatic and morphotectonic conditions. Deposits and weathering in this case have little to do with the initial shaping of the surface, and therefore can provide only a minimum age. Never theless, they provide a crucial component of the palaeolandscape, and they constitute a direct record of palaeoclimates through their mineralogical com position, geochemistry and the spatial distribution of weathering products. Palaeoweathering records
The study of landforms, sedimentary investigations of successions preserved at the margins or within intracratonic basins, and the study of residual deposits developed on palaeosurfaces provide the classic methods of investigating the geomorphologi cal evolution of a region (Andre et al. 1994). Of these, the study of palaeoweathering provides a powerful indicator for past climates and environments. Many studies, however, demonstrate that exposed weather ing profiles are often truncated by subsequent erosion and so, in an effort to further decipher the palaeoenvironmental record, attention has now been turned towards detailed study of the mineralogical
and geochemical characteristics of the deeper parts of these horizons, and even down to the apparently fresh rock beneath.
THE S O U TH- W E S T E R N MAS SIF C E N T RA L
The study area covers the north-eastern part of the Aquitaine Basin to the west and the south-western part of the Massif Central to the east (Fig. 1). Although the transition between continental and marine environments migrated widely across the study area during Mesozoic and Cenozoic history, three main morphodynamic domains can be defined from west to east: 1 the subsiding basin area mainly occupied by Tertiary formations; 2 the limestone borderland corresponding to the ancient calcareous platforms of the Cretaceous and Jurassic; 3 the Rouergue marginal hinge zone, which is an interlink area between the Aquitaine Basin and the heart of the crystalline massif. The area has undergone a long and complex geo logical history, characterized by successive periods of continental evolution with erosion and weathering, and marine oscillations with burial of the former landscapes. Three main geodynamic stages can be recognized (Fig. 2): 1 the Hercynian and post-Hercynian stage, cor responding to the uplift and dismantling of the Her cynian Range and the development of wide palaeosurfaces with deep geochemical imprints; 2 Mesozoic marine oscillations during Jurassic and Upper Cretaceous, leading to the burial of the palaeosurfaces by thick calcareous deposits; 3 Tertiary continental evolution dominated by ero sional and weathering processes leading to the exhumation and rejuvenation of the inherited palaeolandscapes. The 'post-Hercynian' palaeosurfaces
In the south-west of the Massif Central (France) (Fig. 1), particularly in the Rouergue and Margeride plateaux, and on the margins of the sedimentary basin of the Grands Causses, almost planar or gently undulating surfaces have been related to the 'post Hercynian' peneplain. These surfaces cut a crystalline basement complex consisting of igneous and meta morphic rocks, comprising granite, gneiss and schist
228
D r�l
R. Simon-Coinr;on
Quaternary alluvium Eocene formations
U'·i!'·i muj: -
Plio-Quaternary sediments Cretaceous sediments
Q l!lllll
l+ +l
Oligocene-Miocene
Trias and Jurassic
Tertiary volcanism
Hercynian basement
Fig. I. Geological map of the south-western Massif Central and the adjoining sedimentary basins.
of Palaeozoic age. This basement region is sur rounded by sedimentary basins of Carboniferous, Permian, Mesosoic and Tertiary age. Different stages in the geomorphological evolution of this area have been revealed, through the preservation of two main palaeoweathering phenomena, characterized by the development of albite and dolomite in the profile. The Rouergue and the Margeride plateaux display various landforms, such as plateaux and ridges, and preserve remnants of pre-Tertiary palaeosurfaces and palaeoweathering horizons. As each plateau has undergone a slightly differing tectonic evolution, however, subtle differences in their respective geo morphological development are preserved. Uplift ofthe mountain range during Carboniferous times Carboniferous continental sediments now preserved in grabens provide the earliest indications of land scape development. Successive torrential, fluviatile or lacustrine formations with coal seams and volcanic material (rhyolitic dykes and ashes) infill deep and narrow grabens. The presence of palaeosols has been noted (Wetter 1968; Fuchs 1969; Delsahut 1981) and
they are thought to indicate alternating cycles of wet and hot climates with vegetation and soil develop ment and periods of erosion influenced by tectonic activity. At this time upland areas and slopes under went active erosion processes and the resulting sedi ments fed the lowland basins with coarse material, whereas on the lower slopes, bauxitic soils developed periodically. When eroded these later also became incorporated into the sediments deposited in the low lands and basins (Fig. 3). All the rock fragments are of eruptive or metamor phic origin. Indeed, as early as the Stephanian, the deep crystalline core of the Hercynian range was . exposed and its sedimentary envelope had been eroded. According to Pin (1979), the area was uplifted by as much as 5000 m before Permian times, which led to rapid erosion of the uplifted area, as a result of vertical movements in a post-orogenic com pressional system. Such rapid uplift, followed by long phases of erosion, is characteristic of cordillera mountain systems. Recognition of periglacial effects and deposits in the Stephanian-Permian deposits of .the French Massif Central (Becq-Giraudon et al. 1996) supports conclusions of large-scale Carbonifer ous uplift. According to the latitude of the Hercynian
Palaeolandscape reconstruction TIME
ECTONIC
w �
uplift
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of the Masssif Pliocene
Miocene
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Central
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fluviatile deposits
-
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DEPOSITS
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Oligocene
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EROSION
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ho sts
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fluviatile
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and
lacustrine deposits
r
kaolinic soils and ba uxite
I
grabens, indicating a reduction in erosion related to a tectonic quiescence. At that time, the landscape was formed of ridges of crystalline basement that sepa rated long narrow floodplains flowing into lakes (Fig. 4). Desposits are composed of fluvial conglom erates and/or sandstones in palaeochannels, mud rocks in palaeofloodplains and lakes. Coal seams are interbedded with fluvial sediments rather than with lacustrine deposits.
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WEATHERING
229
� (!) u.. Z o <�: a: Z z
Q q; (1) -
o z a: >-
w� w
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Fig. 2. Major geological and morphological events.
range at that time (0°-15°N), they postulate altitudes of order of 4000-SOOO m. This elevation may apply to the higher part of the range, but the Rouergue area may have been at a lower altitude than the central part of the ridge. By the end of Carboniferous times, deposition of fine-grained sediments dominated, with the wide spread development of lacustrine phases in the
Dismantling ofthe Hercynian range during Permian times During the Permian, large tectonic basins formed on the borders of the high-relief areas. The geological setting was not very different to that during Car boniferous times. The major changes were an exten sion of the basins, together with erosion of the areas of high relief and a progressive landscape planation. Climate became progressively drier and even arid in sheltered areas. The sediment record shows that several environ ments coexisted: fans of coarse clastic deposits existed near source-areas proximal to the elevated zones, these passed downstream to sandstones into braided-river channels, while silt mudrocks con taining evaporites (dolomite and sulphates) were laid down in playas developed in the distal zones. The uppermost part of Permian sedimentation is fine grained, with widespread evaporites indicating a slowing in the rate of erosion and the develop ment of endoreic conditions. There are numerous stratigraphical gaps and numerous minor unconfor mities corresponding to breaks in sedimentation or to ravinement. Each formation overlapped the former ones, owing to an extension in size of the basins. By the end of the Permian, the region resembled a subdued landscape with very wide and deep subsiding endoreic basins, amply filled with sediments. The correlative erosion during Carboniferous and Permian times destroyed the Hercynian range. Up to 4000 m of Permian molasse is known to lie below the Mesozoic limestone plateau on the border of the Aquitaine and Grands Causses basins. The apparent slowing down of the rate of uplift during Late Permian times may be related to a general lowering of the relief or to gravitational collapse of the Hercynian range (Becq-Giraudon et al. 1996). This hypothesis may explain the apparent correspondence between erosion and sedimentation rates (Simon Coin�on 1989). Basins may never have appeared as
230
R. Simon-Coinr;on
[ERosiON and TRANSPORT
0 �
basement breccia
� -
bauxitic paleosols coarse-grained sediments
EJ EJ
lacustrine deposits fault
Fig. 3. Distribution of erosion, weathering and sedimentation along the slopes of carboniferous gTabens. Note active erosion on uplands and bauxitic soil development on the lower slopes and lowlands.
N
GRANDS CAUSSES
ROUERGUE
ALBIGEOIS
-
DETROIT DE RODEZ
1 000 m - 40 km
�
basement
-
Carbon iferous sed i m ents
�
volcano
[SJ 1------:J
fault lake
Fig. 4. Landscape at the end of Carboniferous times. Ridges of crystalline basement separated long-narrow floodplains flowing into lakes. Permanence of the main structures is remarkable - the future Grands Causses and Detroit de Rodez basins are noticeable.
deep sinks in the landscape, despite the thickness of the deposits. The drying of the climate during Late Permian times led to the stripping of former weathering profiles, and etching processes predominated and produced an 'inselberg surface'. Albitization of the regolith during Triassic times Triassic sedimentation was restricted to a few subsi-
dent basins, where fiuvio-deltaic conglomerates and sandstones were deposited. Often the clastic sedi ments show evaporitic characteristics, with dolomite cements and nodules as well as salt-crystal casts (gypsum and halite) indicating confined environ ments. Climate is thought to be arid. Continental Triassic palaeosurfaces underwent weathering characterized by development of neogenic albite (Schmitt 1 986, 1 992). This peculiar weathering profile may reach over 50 m depth in
Palaeolandscape reconstruction
23 1
upland basin
Fig. 5. Triassic regolith in southern France: diagrammatic section showing the arrangement of the weathering zones. Towards the Triassic basins, deeper regolith profiles with heavily albitized zones are observed.
+ 100 m
+
+
1 to 1 0 km - - _
--
�
B
+
+
Fe M n oxihydroxides basement
places. The albite development relates to deep palaeoweathering profiles, which show three main horizons. 1 An upper reddish and clayey zone marked by the presence of kaolinite and iron oxides. 2 A middle bleached zone, where primary rock struc ture is destroyed and smectite and mixed-layered clay minerals form. 3 A lower pink-coloured zone, where rock structure is preserved, and which is characterized by the devel opment of a fine-grained low temperature neogenic albite. Albitization starts with transformation of plagioclases, then K-felspars are recrystallized into albite and finally micas (biotite and muscovite) are also replaced by albite. Drill-cores show clearly that the albite horizon is linked to the Triassic palaeotopographies, it always disappears at depth but can reach 20-50m in thickness. Geochemical modelling has been conducted to test the conditions that may be responsible for albite development in weathering profiles under surficial conditions (Schmitt 1994). Simulation shows that NaCl-rich solutions percolating through a mineral assemblage similar to a granite, lead to the progres sive alteration of all the primary aluminosilicates into albite. Further during simulation (corresponding to more dilute solutions of the upper horizons of the profiles), albite itself is altered into smectite and finally kaolinite develops. Development of this low-temperature feldspar needs significant enrichment in sodium. Albitization was produced by NaCl-rich solutions percolating through basement rocks and older Upper Triassic sediment cover, even of continental nature (fluviatile or lacustrine deposits), exhibits numerous halite crystal casts. This suggests that saline waters or salts of marine origin have played a role in weathering. Sea water might have frequently invaded the lower parts of the landscape, or marine brines might have been
[SJ
+
+
CJ
+ smectitic clays
Permo-Carboniferous
f: : �
ffi]
neogenic albite
Triassic deposits
carried far inland in the broad coastal plains by winds and sprays during storms. Albitized profiles are located around the low-lying areas of the Triassic landscape. Their location on the marginal hinge-line separating the subsiding basin and the elevated land areas is linked to the relative stability of this zone, allowing the long lasting development of weathering profiles and their preservation. Lateral vanatwns existed along regional catenas (Fig. 5). Albitized profiles developed as well as on the crystalline basement rather than on the Carboniferous and Permian deposits. Their thick ness increased downwards in the subdued areas around the basins. The upper reddish zone is present in all profiles, the middle-bleached zone becomes dis continuous in the lower parts of the landscape and the albitized zone thins out in elevated areas. Albiti zation probably occurred in low coastal areas during dry periods, whereas kaolinized and haematized hori zons formed by meteoric waters during wetter periods (Fig. 6). Dolomitization of the Early Jurassic transgression surface During the Early Jurassic, tectonic rejuvenation occurred in response to the opening of the Tethys Sea. The Triassic palaeosurface was faulted and albitized profiles were preserved only in grabens or on the low margins around basins, where they were buried beneath sediments. Fault-bounded grabens appeared and a drainage network developed in the upper parts of the landscape. Fluviatile conglomer ates, sandstones and claystones were deposited in channels or in deltas. This was followed by a marine transgression moving from the south-west, accom panied by the development of a broad coastal plain with shallow lagoons. The marine palaeoshore is marked by the occur rence of dolomitic duricrusts ( dolocretes ), which
232
R. Simon-Coinqon GRANDS CAUSSES ROUERGUE
DETROIT DE ROOEZ
/( N
�-1000m basement
-
Pennian sediments
D
Triassic sediments
dolomitization
Carbon iferous deposits
�
lake and/or sea
marine transgression
+ 5 km
+
basement al bitized Triassic regolith Lower Jurassic fluviatile deposits
Fig. 6. Landscape from the end of Triassic to early infra-Lias times: the grabens are extensively infilled and the 'gulf' of the Grands Causses is open to marine influences.
CJ
r&'jJ
-
Lower-Jurassic marine deposits columnar dolocrete massive dolomitization
reach an average of 5 m thick (Schmitt & Simon Coin<;on 1985). Landwards, dolocretes typically display a large columnar structure with vertical and planar joints; however, the base of the columns is irregular and weakly cemented by carbonates. Near the shore, columnar structures disappear and give way to massive dolocretes. Dolomitization affected the crystalline basement and the Upper Triassic and/or Lower Jurassic sandstones (Fig. 7). Dolomite progressively replaced the host material and erased all the primary lithological structures. Only scattered quartz grains remain in a brownish dolomite matrix. Dolomitized profiles are fossilized and partly reworked by further marine deposits. Dolomite relates most probably to the combina tion of ground water and pedogenic processes. Dolomitization is linked to refiuxing of marine brines through the underlying freshwater aquifers in the mixing zone of the coastal environment. Periods of relative sea-level change (rise or fall) lead to migra tion seaward or landward of fresh or saline water (Purser et at. 1994). Much of the dolomite appears to have formed in the subsurface on the border of the continent. Several similar ancient palaeoenviron-
Fig. 7. Distribution of dolocretes on the Lower Jurassic marine palaeoshore: dolomitization prograded on albitized basement and sediments. Note the columnar structures landwards and the massive dolocretes near the shore.
ments (e.g. Upper Permian of Poland; Magaritz & Peryt 1994) are well known, as are analogues in some modern aquifers (Randazzo & Bloom 1985). Dolomitization was diachronous, preceding the Early Jurassic transgression that progressively invaded the Triassic palaeolandscape. The coastal plain with lagoons, shaped during the early stages of the trans gression, was replaced by a marine environment when the sea finally invaded the whole lowland areas and submerged the plateaux. Thus, the relief became effectively buried beneath a thick Jurassic cover of marls and limestones (Fig. 8). The 'post-Hercynian ' polygenetic palaeosurface The 'post-Hercynian palaeosurface' represents a complex patchwork of landforms sculpted by differ ent palaeoenvironments, and it integrates elements of disparate origins. The mineralogical and geochemi cal characteristics of the different palaeoweathering remnants provide a very useful aid to understand ing the evolution and development of the post Hercynian palaeolandscapes and to identify the different features.
233
Palaeolandscape reconstruction
CEVENNES
Fig. 8. Liassic transgression invading
the landscape: the region was slowly covered by the sea and fossilized by sediments. The higher parts stood above sea-level and remained as shoals until late in the Jurassic.
GRANDS CAUSSES DETROIT DE RODEZ
ROUERGUE
1000m
W �
basenient
Carboniferous deposits
Triassic sediments
1 Accretionary surfaces correspond to the latest phase of Permian sedimentation and represent the uppermost sedimentary successions of 35004000 m deep basins, which accumulated throughout Carboniferous and Permian times. The outcropping sediments and crystalline basement of the low areas became deeply altered and albitized throughout the Trias, whereas adjacent upland regions were eroded and weathered. A series of accretionary surfaces testifies to continued subsidence in the basinal areas. 2 Broad coastal plains of marine erosion marked by dolomitization were created by the slow rise of sea-level at the beginning of Early Jurassic. Despite minor fluctuations of the base level, there was an overall rise of sea-level, leading to general flooding of the region. Submergence and sedimentation appears to have played a major role in preservation of these elements of the post-Hercynian landscapes. 3 Remaining uplands were eroded further through out the Jurassic, until they too were buried with the continuing sedimentation. This peculiar evolution generated stepped topographies, that usually are interpreted as a denudation chronology in response to uplift, each step developing in respect to a distinct base level below the former ones. In fact, the geo morphological features of the uplands are younger than the topographically lower features, which were buried earlier. For instance, the eastern margin of the Rouergue plateau did not become covered by marine deposits until the Middle Jurassic (Ciszak et al. 1996), evolving as a shallow shoal that was bevelled by marine erosion over several tens of mil lions of years. The features forming the 'post-Hercynian palaeo surface' are diachronous because they were buried at
-
Lower J u rassic sediments
Permian sediments
B
lake and/or sea
different times during the Jurassic transgression, which lasted until the end of Jurassic. In fact the 'post-Hercynian' palaeosurface buried beneath the Jurassic deposits appears as a juxtaposition of palaeolandforms that originated under different environments related to different scales of time and space (Schmitt J.M. & Simon-Coin�on R. 1993). The Permo-Carboniferous accretionary palaeosurfaces, the Triassic albitized palaeosurface and the Early Jurassic dolomitized palaeosurface form wide areas beneath the Jurassic deposits, and even parts of the present-day landscape where the sedimentary cover has been cleared (Fig. 9). Successive exhumations and rejuvenations during Tertiary times
At the end of the Jurassic Period (Portlandian), the sea retreated to leave behind a wide calcareous plat form. A weathering and erosion phase resumed again during the Early Cretaceous, removing a significant thickness of sediments and uncovering outcrops of the highest parts of the crystalline basement. Evidence of this period is better preserved in the Aquitaine Basin than in the crystalline areas of the Rouergue. A second major marine transgression took place by the Middle Cretaceous Period (Cenoman ian) and reached the edge of the Massif Central (Astruc 1988) . A wide marine platform was formed, especially on the western Jurassic limestones, which were buried beneath thick sediments, whereas the uplands evolved under continental environments. The sea withdrew definitively by the end of Creta ceous Period (Senonian). The region underwent a continental evolution
234
R. Simon-Coinr;on platforms and/or exhumed parts of the 'post Hercynian' palaeosurface. Kaolinization of the pre-Eocene landscapes The so-called Lower Tertiary 'Siderolithic' deposits are composed mainly of gravels, sands and clays and are rich in kaolinite. Their thickness varies from a few tens of metres near the Massif Central, up to 500 m basinwards (Kulbicki 1956; Capdeville 1989; Dubreuilh 1989) . The onset of the most important detrital discharge of the Tertiary corresponds to the scheme of biostasy-rhexystasy developed by Erhart (1956), and was induced by instability of relief and cli matic changes. This 'Siderolithic' clastic discharge results from the reworking of widespread deep weathering mantles developed since the withdrawal of the Cretaceous Sea, both on the crystalline base ment and on the carbonate platforms (Blanc Valleron & Thiry 1993). Deep kaolinitic weathering profiles are widely dis tributed in the western part of the area studied. The profiles may reach depths exceeding 10 m and show successive horizons, from fresh granite to red palaeosols (Fig. 10) . Despite such pervasive alter ation, weathering seems to have left some granitic areas largely unaffected. Juxtaposition of weathered and fresh rock is commonly thought to be the result of differential resistance to weathering of the various rock lithologies. Here, inspection reveals no apparent link with the mineralogical or chemical composition of the rock. However, the distribution of the palaeo weathering features is closely related to Tertiary sediments and topographies. The Villefranche-de-Rouergue plateaux show zones where fresh granite and deep red kaolinitic
MONTPELLIER
D Permo-carboniferous � Lower-J u rassic palaeotopogra phies
0 Triassic [II] J u rassic
palaeotopographies
D You nger palaeotopogra phies Fig. 9. The patchwork of the 'post-Hercynian surface': the map represents the basement of Jurassic deposits. The 'post-Hercynian surface' is formed by the j uxtaposition of palaeolandforms developed under different environments and differing scales of time and space.
under various climatic conditions (humid to arid) similar to those of the Carboniferous and Permian period; but the effects on weathering and landforms were quite different (Simon-Coinyon et al. 1996). During Carboniferous and Permian times, weather ing and erosion processes operated on a crystalline mountain range and destroyed it, but during the Tertiary, these processes operated on an almost fiat landscape, thereby creating relief from limestone
silcrete and ferricrete remnants red kaolinic 1 m sandy clays
mineralogy 50
100•J.
grain-size 50
100'.4
red
kaolinic saprolite
5m
argiliceous saprolite gravelly saprolite
10m
weathering front
Fig. lO. Grain-size and mineralogical composition of the day fraction in a red kaolinitic weathered mantle: the data are from samples of a drill-core from the granitic plateau ofVillefranchede-Rouergue.
235
Palaeolandscape reconstruction weathered mantle outcrop at the same topographical level. The red kaolinitic weathering profiles are partly covered by Tertiary sediments of middle Eocene age (Muratet 1983). A section south of Villefranche de-Rouergue shows deep kaolinitic weathering profiles developed on granite, buried by Tertiary deposits and unweathered granite near the Mesozoic cover (Fig. 1 1 ) . The Tertiary deposits include debris of Permian, Triassic and Jurassic rocks outcropping on the opposite flank of the present-day Aveyron valley. Their source was a palaeocuesta that developed on the basement during the early Tertiary. After this depositional phase the scarp carried on retreating and uncovered unweathered granite. Therefore, the deep palaeoweathering profiles appear to be of pre middle Eocene age. The current westerly limit of the weathered granite therefore marks the original posi tion of this cuesta. Basinwards, the early Tertiary relief has been buried and fossilized by Eocene deposits and can be reconstructed from boreholes (Fig. 12). Pre-middle Eocene erosion shaped a well-differentiated land scape, marked by a large granitic dome (Montauban Moissac dome) surrounded by successive cuestas developed in resistant layers of Jurassic and Creta ceous rocks. The granitic dome displays deep weath ering profiles similar to those of the Rouergue crystalline basement. During emergence the Jurassic and Cretaceous limestones were subject to erosional and dissolution processes. In some areas they were removed entirely from the more elevated zones and the old basement beneath was exposed. During the early Tertiary and
w
E
even Late Cretaceous, these outcropping crystalline rocks were deeply weathered under warm and humid climates, and lateritic alteration prevailed. Where sedimentary cover remained, the basement was pro tected from leaching. Consequently, the distribution, thickness and degree of alteration probably reflect where and when basement exhumation occurred. To put it another way, the distribution of old weathering mantles reflects the pattern of Mesozoic cover that existed at the end of the Cretaceous Period and during the Palaeocene Epoch. Eocene rejuvenation oflandscapes During the middle Eocene, on the western part of the Rouergue, the kaolinitic weathering mantles and remnants of the Mesozoic cover were eroded. Valleys dissected the palaeolandscape of domes and plateaux resulting from the degradation of the 'post-Hercynian surface'. These valleys were further infilled with middle and upper Eocene and Oligocene sediments. The northwards compressive thrust of Pyrenean orogen during the Eocene led to the formation of an E-W tilted fault block of exposed crystalline base ment, the 'dorsale de Rieupeyroux' (Simon-Coin9on 1989). This ridge is the major component of the Rouergue structure and divides the crystalline base ment into two domains, which differ in the type of sediments and palaeogeomorphological features that have been preserved (Fig. 13). 1 North of the ridge, Jurassic formations have been partly eroded and the Mesozoic surfaces are
W
Bassin Aquitain
Montauban-Moissac Dome
E
Albigeois
•
0 EZJ [Ij
basement Carboniferous - Permian deposits Triassic sediments Tertiary deposits
[J
1?.:'8 D
Jurassic limestone kaolinic weathering mantle
paleo - cuesta
[2]
fault
Fig. ll. Red kaolinitic saprolite in the early Tertiary landscape: present-day topography of the west side of Villefranche-de-Rouergue plateau, and palaeocuesta, showing the close relationships between the outcropping basement during the early Tertiary period and the weathered zones.
[21J basement c::::::J Trias-Lower Jurassic lff-�{iJ Middle Lias � U pper Jurassic
� Upper Cretaceous - weathered granite c::::::J Tertiary deposits / fault
Fig. 12. Buried early Tertiary landscape: the Montauban Moissac dome. The section has been established from boreholes drilled by the Bureau de Recherches Geologiques et Minieres. Note that during early Tertiary times relief was well differentiated. It seems that no 'planation surface' existed.
236
R. Simon-Coinqon
0 � �
•
D
�
� ...;;;;;:
exhumed in places. Tertiary sediments have been pre served in grabens. 2 South of the ridge all the Mesozoic sediments have been removed and Tertiary sediments lie directly upon the crystalline basement. The Tertiary landscapes are recognizable, with domes and insel bergs upstanding above wide glacis. The uplift of the ridge led to the removal of the Mesozoic sedimentary remnants and the red kaolinitic saprolite. This erosional stage did not give rise to new stepped surfaces but to extensive land scape dissection characterized by hills, domes and valleys (Fig. 14a ) . Some domes and inselbergs are partly exhumation features, which had been devel oped in a thick weathered cover by differential in situ weathering during the Cretaceous and the early Ter tiary. They appeared when the soft weathered rocks were eroded. Other landforms are clearly of a tec tonic origin. The valleys begin on the ridge and are incised over 50 m, becoming even larger and deeper downstream (i.e. > l OO m ) . Numerous landslides and catastrophic flow deposits can be observed in the lower part of the infilling deposits upstream in the valleys. They are interpreted to be the product of sudden and rapid high-density debris flows that were able to carry large boulders. Superimposed sediments are formed of rock fragments packed in a sandy and clayey matrix and reflect mass transport during storms. They pass downslope to fluvial deposits in braided channels draining the lowest parts of the valleys. On the rim of the Aquitaine Basin, bordering the upland regions,
crystalline basement
=���:
;boniferous
e
Mesozoic sediments Tertiary sediments Rieupeyroux ridge upland fault and thrust fault lnselberg glacis
Fig. 13. Schematic geomorphological map of Rieupeyroux ridge. Note the inselbergs towering above wide glacis cut in the Tertiary deposits on the southern side of the faulted block.
coalescent alluvial fans, fed by the dismantling of the relief to the north-east, developed at the mouth of the valleys, forming an extensive sedimentary apron. Further basinwards, paludal deposits compris ing fine-grained sands, silts, clays and carbonates occurred in the lowest parts of the basin. This dissected landscape was then progressively buried during the end of Eocene to Oligocene times. Burial ofthe landscape during the Oligocene Epoch Burial of the lowlands began as early as the end of Eocene Epoch, and as a result a new type of land scape developed that was characterized by residual hills (domes and inselbergs ) with radiating pediments covered with debris (Fig. 14a & b) . Accumulation of detrital material prevailed progressively over removal, and channel drainage began clogged. Evolu tion of this was aided significantly by the widespread climatic changes at the end of Eocene times. The warm humid conditions that had prevailed during the Palaeocene and early Eocene began to deteriorate toward the end of Eocene times, with a further drying and cooling during the Oligocene Epoch. On the north-eastern Aquitaine Basin, a range of weathering processes occurred relative to the location in the landscape: silicified armour formed in the high piedmont; calcrete and gypcrete characterize the lower slopes and glacis; and smectite-palygorskite-bearing horizons are common on the floodplains or on lake margins (Fig. 15 ) ( Simon-Coin<;on et al. 1996) . These different alter-
Palaeolandscape reconstruction
237
Dorsale de Rieupeyroux
UJ �
� -
crystalline basement quartz vein molassic deposits with limestones conglomerates and sand stones (Paleocene ?)
1.!111 / �
fault
EJ
lake
red sandy clays with gravels
glacis with detrital cover
Fig. 14. South-western Rouergue palaeo landscape during the Eocene and Oligocene Epochs: (a) Plateaux, probably inherited [rom the 'post-Hercynian surface' were deeply eroded during the early Eocene. This landscape is modelled on the crystalline basement and corresponds to that of the Montauban-Moissac dome and its surrounding cuestas in the sedimentary basin. Conglomerates and sandstones mark a northward palaeodrainage, the age of which is unknown (Late Cretaceous?, Palaeocene?). (b) During the arid phase at end of the Eocene and during the Oligocene Epochs, subsidence of the region led to the formation of an inselberg-glacis landscape. There was regrading of the slopes and burial of the valleys by alluvial cover sediments. The molasse lakes inundated this landscape, their higher stand took place during Miocene (Aquitanian) times.
lacustrine basin
Fig. lS. Weathering and sedimentation on the north-eastern rim of the Aquitaine Basin. With the flooding of the landscape, red kaolinitic sa prolites gave place to various weathering horizons organized as a genuine catena prograding headwards.
glacis
glacis
low piedmont
1±:::±1 crystalline basement [[[!] paleosol features
ations formed catenas, moving upwards from basins to uplands and pervaded progressively the red kaolinitic palaeosols (Archanjo 1982; Astruc 1988; Simon-Coin<;on 1989). A return to more humid con ditions during the Miocene Epoch resulted in the dis solution of the more soluble elements. The original extent of the gypcrete can be determined only from the preservation of boxworks of gypsum crystals in silicified nodules.
� silicification
� red sands and clays f2SJ pedogenic calcrete � lacustrine limestone
coarse deposits
From the late Eocene times, the Massif Central supplied less and less sediment into surrounding basinal areas. Sedimentary input in the north-eastern part of the Aquitaine Basin was replaced by molasse deposits (i.e. calcareous fine-grained sandstones), resulting from the erosion of the rising Pyrenean range. These deposits were transported by large and sluggish rivers, which formed swamps and lakes on the arid edge of the old basement rocks. From late
238
R. Simon-Coinr;on
Eocene to Miocene times, as the Pyrenean fore trough prograded northwards, relief was slowly buried under molasse deposits related to lacustrine flooding. To explain the extent of the submergence, it is necessary to invoke tectonic subsidence of the area. Lakes reached their maximum level during the Aquitanian Age; in addition, sea-level was at its highest stand at the same time. Downcutting ofstreams during late Tertiary and Quaternary times At the end of the Miocene Epoch (Tortonian), a fall of sea-level combined with a general upwarping of the Massif Central, together with more humid cli mates, favoured incision by drainage networks. Suc cessive erosional cycles occurring from the end of Miocene to Quaternary times, removed the soft molasse deposits and exposed the Jurassic limestone plateau of Quercy. The present landscape appears as a juxtaposition of different elements, which give the illusion of a succession of stepped forms and sur-faces. Several authors (Baulig 1928; Enjalbert 1952) thought that the ridges in crystalline basement were remnants of the pre-Triassic surface, whereas plateaux were derived from Palaeogene peneplains and the incised valleys belonged to a Pliocene or early Quaternary cycle. It is possible, however, to see the manner in which the different landscape elements are really linked together. The flattened plateaux are the result of the general marine transgressions of the region or of the re-exhumation of early Tertiary reliefs, which, in turn, had often been inherited from infra-Cenomanian or Early Jurassic landscapes.
I N T E R P R E T I N G S UCC E S SI V E P A L A E O L A N D S C AP E S
Old cratons are usually thought to b e rather stable and affected by slow movements of episodic uplift and subsidence, favouring the persistence of large scale topographies over geological time. Such per sistence landforms are described in terms of a 'cratonic regime' (Fairbridge & Finck! 1980; Finck! 1982). Large peneplains are attributed to fluvial action generated by a fall of base level or an uplift of the continent (Davis 1899); or considered as the result of scarp-retreat (King 1962). Stepped benches are also thought to be generated by a relative uplift of continental areas (Penck 1953). Large erosional sur faces or etchplains characterized by inselbergs are
related to removal of weathering material by sheet floods that control the geomorphological evolution and chemical breakdown, which is a powerful agent for long-term surface lowering (Btidel 1982; Thomas 1989). The Massif Central of France was not stable enough to correspond exactly to the concept of an old craton. It provides a special example of evolution. The main surface, the 'post-Hercynian peneplain' is not the result of a single erosion process. The major tectonic uplift or changes in base level generated neither pediplains by scarp retreat, nor formation of convex stepped benches. Etching processes corre spond better with the observations and data. Even if the 'post-Hercynian surface' is a major component of the present-day landscape, it was often extensively degraded. Intricacy of successive stages of landscape evolution
Detailed study of some outcrops illustrates the intri cacy of successive stages of landscape evolution. Once exhumed, landforms evolve in environments different from those in which they had developed originally. This is certainly the case in places in the Rouergue, where the plateaux are capped by sedi mentary deposits (i.e. fluvial or paludal successions), weathered blankets of red kaolinitic claystones and silcretes of Tertiary age. They also preserve remnants of albitization and/or dolomitization, attesting to the fact that they pre-date the surficial deposits and saprolites. In this case there can be no genetic link between them and the Tertiary formations (Simon Coin<;on 1989). There is a superimposition of differ ent phenomena on the exhumed surfaces. Good examples of this superimposition are pro vided by weathering profiles of the Rouergue crys talline plateaux, where in a few metres one can observe geomorphological history back to more than 250 Ma (Fig. 16): 1 albitization of basement during the Triassic Period; 2 stripping of the upper part of the weathered profile; 3 deposition of Lower Jurassic fluviatile sandstones; 4 dolomitization of sandstones and basement in rela tion to the Early Jurassic marine transgression; 5 burial under a few hundreds of metres of Jurassic sediments; 6 regression of the sea, followed by erosion and dis solution of the whole limestone cover (from the Early Cretaceous to the late Eocene),
Palaeolandscape reconstruction Tertiary deposits Liassic limestones
l----'--'.·��r---,---'='Tf
dolocrete
kaolinic weathering
dolomitization
fluviatile sandstones Lower-J urassic unco nformity a l b itization a l b itized gneiss
Fig. 16. 'Digest' of several hundreds of millions years on an old surface. Imprints of successive weathering, erosion, burial and exhumation.
7 Tertiary weathering, forming pisolitic iron duri crusts and red kaolinitic claystones. Erosion often ceased against the Early Jurassic basement unconformity, which was again buried beneath Eocene and Oligocene fluviatile and lacus trine deposits. Present-day erosion is removing this soft and friable sequence, exhuming the Early Jurassic surface that has been armoured by Tertiary silcretes. The limestone platform of Quercy offers similar examples: different surfaces attributed to Eocene, Oligocene or Miocene ages are partly inherited from the Cretaceous marine transgression (Cenomanian). During the early Tertiary erosion, weathering and dissolution removed the softer Cretaceous deposits. Today, only remnants of sandy material, often silicified and trapped in palaeosinks or palaeokarst voids, attest to the former extent of the Cretaceous cover (Virol 1987; Astruc 1 988; Simon-Coin'
Some surfaces carry distinctive palaeoweathered profiles, which provide means of identification. Geo chemical imprints left in landscapes may be useful tools to understand landscape evolution, by inferring palaeoenvironmental conditions from present or recent global climatic conditions. This information has to be used with caution, because palaeocondi tions, even those of the Permian and Triassic periods, and particularly atmosphere composition, may have
239
been different in the past and produced specific alter ation patterns in rocks. For example, some palaeoen vironments, such as a sulphuric acid atmosphere, no longer exist in the present world (see Schmitt, this volume,pp. 21-41). For instance, processes of albitization and dolomi tization are not well understood, nevertheless these alterations provide effective markers to follow particular units of the polygenetic post-Hercynian surface. On an other hand, the interpretation of saprolites formed on crystalline basement may be more ambiguous. In high and middle latitudes they are often scarce and not very evolved and can be interpreted either as representing recent (post glacial) weathering, or as the roots of old weathered mantles (Godard 1989). The example of the south west Massif Central leads us to think that the in heritance of palaeoweathering cover is of a major importance. Some geochemical imprints can be younger and superimposed and they thus hide the true origin of landscapes. For instance, pre middle Eocene ferruginous red palaeosols of the margins of the Aquitaine Basin were totally destroyed and replaced by calcrete at the time of aridification of climates, without degradation of the landforms. Rates of evolution
Evolution of palaeosurfaces requires long time inter vals for its full development. Even if most of the land forming processes are slow and progressive, sudden and intense events can occur. Any change may diffuse irregularly and in a complex manner through the landscape, which develops in three dimensions and in a space and time framework. Its evolution is diachronic, as noted by Ollier (1981). Various scales therefore can be distinguished: from millions of years to a few hours, from thousands of square kilometres to a few square metres and even to thin-section scale. The post-Hercynian surface of the south-west Massif Central has evolved since more than 50 Ma. In places it was destroyed by erosion during the Palaeocene and Eocene Epochs and gave way to a landscape of inselbergs and glacis, which has evolved since about 20 Ma. Other areas have been submitted to erosion from the very beginning and it is difficult to tell how old they are. The time limits observed for formation of palaeosurfaces in the world are varied. For example, the period of denudation leading to the sub-Jotnian peneplain in Sweden was estimated to
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R. Simon-Coinqon
400 Ma (Lindmar-Bergstrom 1996, 1997), similarly the main surface of Sogn-Jotun in Greenland required 350 Ma for its development (Peulvast 1988), and the west France peneplain formed over 60 Ma (Sellier 1985). Furthermore, 'the ancestors' of the contemporary inselbergs and domes in Central Aus tralia, are thought to be at least of early Tertiary age (Twidale & Bourne 1978). These long periods explain the existence of gaps in their chronology. Long-duration weathering on basement, even without noticeable physical erosion, contributes to dissolution and to an important loss of material. On the Rouergue basement, weathering was particularly important during the formation of the albitized (Triassic) and the red kaolinitic (Cretaceous and lower Tertiary) saprolites. Such surface lowering rates have been calculated for various bedrock types and climatic conditions. Geochemical calculations conducted on weathering profiles in Buganda (Uganda, Africa) allow us to determine that l m thickness of laterites corresponds to the weathering of 4 m thickness of the granitic parent-material (Tren dall 1962). A stratigraphical approach of the flint bearing clays in the Paris Basin shows that a 40-m-thick profile of flint-bearing clay corresponds to the weathering of 160 m of flint-bearing chalk over a period of 60 Myr. In this case, the rate of landscape lowering is about 2 m Myr-t (Quesnel 1997). Etching processes
Etch surfaces develop by erosion of the soft weath ered mantle down to the fresh rock, thus those areas underlain by more resistant rocks are brought into relief as the less resistant rocks are lowered (Rich 1951). They develop directly from the weathering front. Weathering acts in two ways; it directly lowers the landscape by chemical loss, and leads to subsur face softening of the bedrock, which subsequently is stripped away. As the thickness of the weathering mantle depends partly on the water-table position, the weathering front tends to be horizontal, with more or less residual resistant cores. Etch surface development is influenced by the rate of uplift, which controls the lowering of the weathering front and erosion potential. They can be two-age landforms or can be lowered with a dynamic equilibrium between weathering and erosion (Thomas 1989) . In deeply weathered terrain, some harder cores of rock may be preserved to be exposed as fresh bedrock outcrops following stripping of the friable regolith. Etching is capable of application to both tropical and non-
tropical landscapes, but it is most effective under moist tropical climates. In the south-west Massif Central, two periods were favourable to the development of such etch surfaces; during Permo-Triassic and Tertiary times. Residual inselbergs and tors, shaped in albitized crystalline basement, can be observed in relation to Triassic deposits. A second major etch surface developed on the border of the Aquitaine Basin during the Tertiary. In this area, even if inselbergs and tors are partly buried beneath Tertiary deposits, they are features of the present-day landscapes (Fig. 13). Successive burial and exhumation of l andscapes
Interpretation and hypothesis may be influenced by the burial of a landsurface beneath a sedimentary cover followed by later re-exposure. Landscapes often developed through the burial of palaeoto pographies and their subsequent exhumation. Burial and exhumation may have been polycyclic, diachronic and complex. Even if palaeoenvironments had been relatively stable for long periods, minor changes in base-level, climates or tectonic regimes may have occurred and the resulting palaeosurface will have recorded these changes. There are many examples of exhumed landscapes in different parts of the world (Thomas 1989;Twidale 1994). The excellent preservation of some exhumed forms is remarkable. For example, the sub-Mesozoic etch surfaces in southern Sweden were protected under a Cretaceous cover during the greatest part of the Tertiary (Lindmar-Bergstrom 1996, 1997), Permian glacial landforms, roches moutonnees and valleys are now being exposed by erosion in Hallet Cove and Inman Valley of South Australia (Bourman & Alley 1990; Alley & Bourman 1995). In the south-western Massif Central, sediments covered different facets of former landscapes and preserved heterogeneous forms and weathered mantles. Some zones of Rouergue evolved as shoals, during the Early Jurassic transgression and were bev elled by marine erosion. Limestones were deposited around these shoals. When the Mesozoic cover was stripped during the Cretaceous or the early Tertiary, these palaeoshoals stood higher than the earlier Jurassic surfaces. These demonstrated relationships reveal that the classic scheme of inset surfaces, with the older standing above the younger, does not apply in this location. South of the area studied, in the Montagne-du Suquet (Gard), remnants of Triassic arkosic sand-
Palaeolandscape reconstruction stones are preserved between granite boulders on the summit surfaces, which implies that these tors and inselbergs may be the result of exhumation of a pre Jurassic etch-surface from beneath a Lower Jurassic cover. When softer overlying sediments have been removed, it is difficult to determine the origin of the exhumed palaeosurface, although clues may be derived from geochemical imprints. For example, on the northern border of the Massif Central (Les Pierres-Jaumatres), large basement plateaux exhibit residual tors and inselbergs, which can be related to an Early Jurassic landscape because of the albitiza tion of the granite bedrock.
C O N C L U SIO N S
The south-western Massif Central constitutes a 'hinge' zone that recorded and preserved evidence of several successive palaeolandscapes. Recognition of the superimposition of processes of different envi ronments, related to alternating erosion, weathering and sedimentation, brings a new approach to palaeo landscape reconstruction. It appears that revision of the methods used and of the underlying philosophies for hypothesis formulation in geomorphological investigations is needed. Studies of historical land scape evolution based simply on morphology are difficult to interpret, but they can be revitalized through multidisciplinary investigation. Landscape studies of old orogens and cratons have to be placed in a space and time perspective. When the features have been studied closely and identified in the field, and their relationships to weathering imprints and sediments established, their mode of formation and their ages can be investigated. Many factors operate in the formation and preservation of palaeoland scapes and no single model provides a satisfactory explanation for landform evolution. There is not a simple recurrence of evolutionary processes, even if features carry imprints of climatic, tectonic or erosional cycles. In an evolving landscape, similar external conditions will have different effects because they will be operating on a different land scape, with feedback effects originating from inher ited relief. On the other hand, some landscapes, particularly broad inland surfaces developed on basement rocks, preserve their essential features for long periods. This study has shown that the palaeolandscape system is composite and complex, depending not only on the internal or external influences of rock mineral-
241
ogy, climate, tectonism and sea-level, but also on the characteristics of the pre-existing environments, which influence the course of further landscape evo lution and our present-day and future environment. Palaeolandscape reconstructions over long time spans and large areas are one of the aspects of applied geomorphology. Some mineral deposits occur in specific sites of the landscape, for instance, uranium ore concentration is closely related with the albitized Triassic weathering mantles, in the same way, phosphate deposits are associated with the Eocene and Oligocene infillings of palaeokarst voids of the Quercy. The examples given in this paper show that study of palaeoweathering and computer simula tions can be fruitful, producing detailed data, valuable for landscape reconstructions. Landscape reconstruction allows an evaluation of the degree of stability of a region, which is of a great interest for the dumping of waste products, including radioactive waste.
AC K N O W L E D G E M E N T S
I would like t o thank A. Godard, R.P. Bourman and M. Widdowson for their suggestions, critical revision, constructive comments and language corrections of this paper.
R E F E R E NC E S
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Palaeolandscape reconstruction ScHMITI, J.M. (1994) Geochemical modeling and origin of the Triassic albitized regolith in Southern France. 14th International Sedimentology Congress, Recife, Abstracts, SS, 19-21. SCHMITI, J.M. & SIMON-COIN<;:ON, R. (1985) La paleosurface infraliasique en Rouergue; depots sedimentaires et alterations associees. Geol. Fr. , 2, 125-135. ScHMITI, J.M. & SIMON-COIN<;:ON, R. (1993) Encha!nement des paysages et genese de Ia 'peneplaine' post-hercyni enne dans le Sud de Ia France. 3rd International Geo morphology Conference, 26 Aug-3 Sept Hamilton, Programme with Abstracts, p. 239. SELLIER, D. (1985) Les versants du Pays Nantais. Etude geo morphologique. These 3 eme cycle, University of Nantes. SIMON-COIN<;:ON, R. (1989) Le role des paleoalterations et des paleoformes dans les socles: l'exemple du Rouergue, Massif central fran<;ais. E.NS.M.P Mem. Sci. de la Terre, 9, 290 pp. S!MON-COIN<;:ON, R., THIRY, M. & SCHMITI, J.M. (1996) Variety and relationships of weathering features along the early Tertiary palaeosurface in the south-western French Massif Central and the nearby Aquitaine Basin. Palaeo-3, 129, 51-79. THOMAS, M.F. (1989) The role of etch processes in landform development II. Etching and the formation of relief. Z. Geomorphol. NF, 33(3), 257-274. TRENDALL, A.F. (1962) The formation of apparent pene-
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plains by a process of combined laterisation and surface wash. Z. Geomorphol. NF, 6(2), 183-197. TwiDALE, C.R. ( 1994) Gondwanan (Later Jurassic and Cre taceous) paleosurfaces of the Australian craton. Palaeo-3, 112, 157-186. TWIDALE, C.R. & BouRNE, J.A. (1978) B ornhardts. Z. Geo morphol. NF (Suppl.) Band, 31, 1 1-137. VIROL, F. (1987) Le contact Massif central I Bassin aquitain au niveau du Lot moyen et du Cele: enseignements fournis par les formations superficielles d'iige secondaire et terti aire en matiere d'evolution morphologique. These Doct., Univ. Paris I. WETIER, P. (1968) Geologie du bassin houiller de Decazeville, du Detroit de Rodez et du bassin de Figeac. These Houilleres d' Aquitaine,Albi, 2 vols, 445 pp. WIDDOWSON, M. (1997a) The geomorphological and geolog ical importance of paleosurfaces. In: Paleosurfaces: Recognition, Reconstruction and Paleoenvironmental Interpretation (Ed. Widdowson, M.), Spec. Pub!. geol. Soc. London, No. 120, pp. 1-12. Geological Society of London, Bath. WIDDOWSON, M. (1997b) Tertiary paleosurfaces of the SW Deccan, Western India: Implications for passsive margin uplift. In: Paleosurfaces: Recognition, Reconstruction and Paleoenvironmental Interpretation (Ed. Widdowson, M.), Spec. Pub!. geol. Soc. London, No. 120, pp. 221-248. Geo logical Society of London, Bath.
Spec. Pubis int. Ass. Sediment. (1999) 27, 245-274
Lateritization, geomorphology and geodynamics of a passive continental margin: the Konkan and Kanara coastal lowlands of western peninsular India
M . W I D D O W S O N * and Y. G U N N E L L t *Department of Earth Sciences, The Open University, Walton Hall, Milton Keynes MK7 6AA, UK; t Laboratoire de Geographie Physique, CNRS- UPRES-A 6042, Universite Paris 7, case 7001, 2 Place Jussieu, 75251 Paris Cedex 05, France
A B S T RACT
The existence of extensive belts of lateritization along the coastal lowlands of tropical continental margins is a well-documented phenomenon and this research presents a detailed study of the western coast of peninsular India (12°-18°N). The western coast of India is a passive continental margin and its coastal low lands evolved geomorphologically during middle to late Tertiary times. The associated extensive laterites testify to an important and widespread lateritization that affected western India during this time. These laterites developed upon a pedimented surface resulting from the recession of the Western Ghats escarp ment when climatic and tectonic conditions favourable to deep weathering reached an acme; at least two generations of laterite formations may be recognized. The most extensive phase of lateritization was con sequent to the development of the pedimented surface, because this provided an ideal topographical site for protracted deep weathering. Importantly, this pediment surface has been cut into two distinct litho logical terranes, which differ fundamentally in both their geological age and structure. To the north, the !at erite is developed upon the Deccan basalts of Late Cretaceous to early Tertiary age (c. 65-67 Ma) which have provided a laterite protolith of remarkably restricted and predictable composition. South of 16°30', the coastal laterites have evolved upon the heterogeneous pre-Deccan basement, which comprises high grade (i.e. Peninsular gneisses) and low-grade sedimentary lithologies of Archaean-Proterozoic age (i.e. Dharwar metasediments), and irregular shaped granitic intrusions of Early Proterozoic age. The objectives of the current work are threefold: first, to explore similarities of laterite development across these different geological terranes, thereby demonstrating the ubiquity of both lateritization and controlling geomorphological processes. Such ubiquity clearly points to a fundamental morphotectonic control during evolution of the margin. Second, to outline the chemical and physical differences of laterite types across the Konkan-Kanara lowlands, which inevitably exist as a result of fundamental differences in the underlying geology. Thirdly, the morphology of the lateritized pediment surface is investigated with respect to the morphotectonic development of the Indian margin. The elevated position of the Indian coastal laterite belt (c. 60-200 m), together with associated development of an entrenched drainage confirms its geological antiquity, and indicates that widespread and permanent uplift has affected the margin during late Tertiary times. Moreover, cambering of this lateritized coastal palaeosurface is consis tent with current models advocating seaward flexuring of rifted continental margins in response to onshore denudational unloading and concomitant offshore sedimentary loading.
INT R O D UCTI O N
Fermor 1909; Fox 1923). Their importance in the lit erature, however, is not derived simply from this historical precedent. During the nineteenth and twentieth centuries researchers sought to refine characterization of laterite and to understand its rela tionship to both geomorphological and geological
India has long been a n important area for study of laterites. It was in western Peninsular India that Buchanan (1807) made the first historic description that suggested the term laterite, and in the following decades India remained a focus for laterite descrip tion (e.g. Newbold 1 844; Foote 1876; Merril 1897;
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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M. Widdowson and Y Gunnell
environments. Subsequently, many other important areas of lateritization have been recognized in South America, Africa and Australia, yet remarkably, a detailed and systematic interpretation of the Indian examples has, with only a few notable exceptions (e.g. Sahasrabudhe & Deshmukh 1981; Oilier & Rajaguru 1989), received comparatively little attention in recent decades. More recently, study of the distribu tion of Indian examples has been used to demon strate the link between the lateritization process and the development of large-scale denudational fea tures, such as those described in the models formu lated by King (1967) . For example, the recognition of lateritized palaeosurfaces have been considered fun damental to understanding the long-term geomor phological evolution of stable cratonic areas (e.g. Thomas 1994). The antiquity and geological setting of the Indian laterites means they can supply information regarding long-term morphotectonic development (Widdowson 1997a) and, as will be demonstrated here, provide important evidence relating to the evolution of the western Indian conti nental margin.
G E O L O GY A N D LAND SCAP E O F THE W E S T E R N I N DIAN MAR G I N
The region studied comprises the coastal margin of western peninsular India between 12°-18°N (Fig. 1). Geologically the terrane consists of the rifted flank of the Western Dharwar Craton, comprising an Archaean granite-greenstone terrane that is covered in its northern part by Late Cretaceous continental flood basalts (CFB) known as the Deccan Traps. At the tectonic scale, western India is a particularly interesting continental margin to study because it has been subject to two major rifting events (i.e. Mada gascar and Seychelles at c. 88 Ma and 65 Ma, respec tively) since the break-up of Gondwana in early Cretaceous times (Norton & Sclater 1979; Storey et al. 1995). The most widely documented rifting event is that associated with the Deccan CFB but this represents only the latest, albeit perhaps the largest, rift-associated magmatic event affecting western peninsular India. This multiphase rift history clearly has important implications for the morphotectonic evolution and associated erosional effects which have
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Fig. I. Geological sketch map of the western margin of peninsular India. Coast parallel rectangle shows the current study area, which consists of the coastal plain (i.e. Kanara and Konkan regions between 11 and l9°N ) and lies west of the Western Ghats escarpment. Note the varied lithologies that comprise the pre Deccan basement south of l6°N. o
Passive continental margins occurred as a consequence of the widespread epeirogeny that has affected this margin. The CFB margin uplifts are currently interpreted in terms of mechanisms such as magmatic underplating (Cox 1980; 1993) and denudational isostasy (e.g. Summerfield 1988; Gilchrist & Summerfield 1 990), both of which are fundamentally dependant upon the timing of continental rifting events. The Deccan CFB lavas were erupted as the Sey chelles microcontinent separated from India at c. 65 Ma (end Cretaceous to early Tertiary times), and the associated rifting event was induced by the north ward movement of the Indian subcontinent over the Late Cretaceous manifestation of the Reunion plume (Cox 1978; Duncan 1 990; Mitchell & Widdowson 1 99 1 ; Klootwijk et al. 1 992). The Traps currently cover an area of at least 500000km2 reaching a maximum thickness of c. 2 km, and are an excellent example of one of a number of similar continental flood basalt (CFB) provinces (e.g. the Karoo of southern Africa, and Parana of South America), which are intimately associated with continental break-up and incipient ocean development (White & McKenzie 1989; Campbell & Griffiths 1990). The lavas are typical CFB tholeiites and consequently display a very limited compositional range. Initially they were erupted onto, and rapidly drowned the pre-Deccan (Gondwanan?) land surface, which had over very large areas lain exposed to continental weathering for a considerable period of time. However, the thin infra-trappean Lameta and Bagh beds do indicate the existence of fluvial and shallow marine conditions at some localities during the Late Cretaceous (Prasad & Khajuria 1995). Toward the close of the volcanic episode, hiatuses in the eruption allowed the devel opment of thick weathering profiles (i.e. boles) now preserved within the uppermost lava sequences (Widdowson et al 1997). By contrast, the pre-Deccan basement exposed south of the province is lithologically diverse and spans at least the past 3 Gyr of Earth history. Thus, in the south the Deccan Traps overlie this heteroge neous basement which, in the west, may be broadly grouped into the varied lithologies comprising the Archaean Dharwar craton (Rogers & Mauldin 1 994), and the Proterozoic Upper Cuddapah sediments including the Kaladgis east of Belgaum, together with the Bhima Group sediments in the region south-east of Bijapur. Lithologically, this sub-Deccan basement varies from weakly metamorphosed Proterozoic sediments of the Bhima and Kaladgi groups (e.g. limestones, shales, conglomerates, etc.), down .
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through the weakly metamorphosed Dharwar sedi ments (broadly characterized by immature sand stones, greywackes, Banded Iron Formation and associated lithologies, together with tuffs, agglomer ates and mafic lavas), and into the high-grade Archaean Peninsular gneisses and charnockites. The combination of these latter being typical of a granite-gneiss-greenstone terrane. In Goa these Dharwar supracrustals are apparently intruded by a series of younger, Late Archaean-Early Proterozoic (Dhoundial et al. 1987) granitoid and migmatite bodies (i.e. Londa migmatite-gneiss, Chandranath, Dudhsagar and Canacona granites). These granites form topographically elevated and typically un lateritized masses upon the coastal plain between 15° and 1 6°N. The southern limit of the ancient Dharwar craton is defined by the major Palghat-Cauvery shear zone running east-west (c. l l 0N), and south of which lie the high-grade lithologies of the southern Indian Proterozoic mobile belt (Harris et al. 1994). Given such lithological diversity it is clear that the pre-Deccan basement provides a heterogeneous complex of different geological domains, which, potentially, can affect the physical and chemical nature of later residual deposits such as laterites. Despite these major lithological contrasts between the lava and basement terranes, and within the base ment itself, denudation has affected both the lavas and the pre-Deccan basement in such a way as to yield a fairly homogeneous macrogeomorphological landscape along much of the western peninsular margin. This comprises, from west to east: 1 a low-lying coastal plateau with short westerly flowing rivers; 2 a coast-parallel continental-scale escarpment of between 600 and 1900 m elevation known as the Western Ghats; 3 the crest zone of the Western Ghats forming a narrow zone of discrete, elevated mesas in the basalt region, or of ridges and domes in the cratonic basement; 4 an elevated inland plateau (Karnataka and Maha rashtra uplands) with a gentle eastward slope. In general, the Western Ghats form the major drainage divide of peninsular India because many rivers rising on the east of the main escarpment dis charge into the Bay of Bengal some 800-1000km away. Although rivers rarely cross this divide in the Deccan region, some have breached it between 14° and 15°N. These, and several other minor but significant occurrences further south, descend the Ghats as a series of spectacular waterfalls (e.g. Jog
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Falls: 74°49'E - 1 4° 13'N) and thence supplement the otherwise short courses of the westerly flowing Konkan drainage. In the following sections it will be shown that the landscape of the coastal region can be divided further into subregions that contrast by their geological sub stratum, the relative abundance of laterite landforms, and the presence or absence of a broad topographical depression between the Western Ghats escarpment and the coastal belt of laterite plateaux.
LAT E RI T E IN W E S T E R N I N D IA
Overview
Laterite is common in western peninsular India and its origins and distribution are discussed by a number of eminent early workers (e.g. Foote 1876; Medlicott & Blandford 1879; Fermor 1909), and more recently in numerous localized studies (see relevant papers in IGCP 129 Lateritization Processes symposia volumes by Melfi & Carvalho 1 983; Banerji 1986). Perhaps the most detailed documentation of its occurrences along the west coast, however, is given by Fox (1923, 1 936). Within western India laterites may be recognized in a series of different geological-geomorphological settings but, at the broader scale, the most wide spread occurrences are found in two distinct geomorphological zones: 1 capping the elevated basalt mesas of the Western Ghats (16°-18°N), where they are developed upon the Deccan lavas (Sahasrabudhe & Deshmukh 1981; Widdowson & Cox 1 996; Widdowson 1997b ); 2 an extensive, semicontinuous belt lying to the west of the main Western Ghats escarpment be tween Bombay and Mangalore (c. 12°-18°N) cap ping the coastal plateaux of the outer Konkan plain (Widdowson 1990; Gunnell 1996a; Fig. 2). Within the region of the Deccan lavas, those lat erites occurring in the Western Ghats are restricted entirely to the flat mesa tops of the highest peaks (900-1500m) along the Ghats ridge (Fig. 3), form ing indurated cappings. North of Bhor (73°52'E, 1 8°09'N), however, there is little evidence of fur ther mesa-cappings along the Ghats crest zone and Kalsubai peak (73°43'E, 19°35'N), the highest point of the Deccan (1650m), is composed entirely of eroded basalt masses. By contrast, isolated laterite capped basalt mesas are found along the Ghats ridge as far south as the southern limit of the lavas near
Belgaum (for detailed description of high-level lat erite occurrences see Widdowson (1997b )). Beyond Belgaum the laterites pass on to the Archaean and Proterozoic rocks that form the southern continua tion of the Western Ghats chain, although they occur in the Karnataka uplands mainly as buried rafts of indurated laterite in palaeosol profiles (Gunnell 1996b) rather than as the well-defined mesa-cappings of the Deccan. The aim of the current work, however, is to focus upon the lateritic materials of the coastal lowlands, explore their evolution, and investigate in detail those materials developed upon the Deccan lavas and pre-Deccan basement within this single geomor phological zone. In this zone the most northerly lat erite occurrences are found in the coastal regions of Kutch and Kathiawar (Valeton 1983), and they are developed extensively along the coastal lands of Maharashtra, Goa, Karnataka, and as far as south central Kerala, beyond which occurrences become progressively less common towards Cape Comorin. These coastal laterites are developed in the north upon the Deccan lavas and continue extensively southwards upon Archaean and Proterozoic litholo gies as well as upon rare outcrops of Mio-Pliocene sediments. The coastal lowlands of western India lying between the Arabian Sea and the foot of the Western Ghats escarpment are known as the Konkan in Maharashtra, and Kanara in Karnataka. The coastal laterites of those regions occur as clusters of plateaux over large tracts of these lowlands (Fig. 2). Plateau elevations show a remarkable uniformity in height from Maharashtra to Kerala, and this con sistency is reinforced by a systematic pattern of varia tions in elevation from west to east, because the laterite plateaux as a whole slope seaward away from the Western Ghats escarpment toward the coastline (Fig. 3). The origin of this coastward slope is an important aspect of interest and will be discussed in a later section. One of the most remarkable aspects of Konkan lat erite is its continuity and ubiquitous development on a wide range of lithologies. On initial examination there appears little difference in the general form, thickness and geomorphological setting of laterite developed upon Archaean and Proterozoic gneisses or metasediments and the chemically uniform Deccan basalts to the north. This belies a more complex relationship between different laterite types, lithology and geomorphological setting, which is duly explored in this paper.
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Fig. 2. Distribution and typology of indurated laterite outcrops along the Arabian Sea margin. Note the much more extensive development of autochthonous laterite on the northern Deccan basalt lithology, and the inner piedmont corridor devoid of indurated laterites in both the Deccan trap region (l6°-l8°N) and the Mangalore hinterland (c. 12°20'-13°10'N). Inset shows location of Fig. l l .
Fringe of allochthonous, lower level laterite plateaux along estuaries
Data acquisition
The work presented is a synopsis of data collected during a series of extensive field studies conducted between 1986 and 1994. General mapping of the lat-
erite distribution has been achieved using a combina tion of field observation and satellite image pro cessing techniques. Accuracy of the satellite-derived information was verified by detailed 'ground truth' survey in key localities.
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Fig. 3. Cross-section of the Deccan coastal margin (Konkan) and inland Ghats escarpment at 18°N showing main topographical divisions and associated coastward-dipping anticlinal-monoclinal geological structure (Mitchell & Widdowson 1991). Westerly dip of the lava sequences across the outer Konkan is c. 0.5°-0.7°. Note the presence of: 1, high level laterites on the highest mesas of the Ghats; 2, the low-lying inner Konkan corridor devoid of laterites; and 3, the more elevated outer Konkan plain characterized by an extensive development of a coastward-dipping low-level laterite ramp. West of the Ghats escarpment erosion has removed c. 1-1.5 km ( e) of basalt, exposing progressively older formations towards the anticlinal-monoclinal core, and consequently the low-level laterite-capped laterite profiles lie with angular unconformity upon exposed basalts of the Poladpur and Ambenali Formations, which form the outer Konkan plain. (Note: vertical exaggeration is c. 20 x.)
Sampling for chemical analyses of lateritic materi als was given a high priority in the Deccan basalt region as a technique for demonstrating that most of these laterites were formed by the in situ break down of basalt lava, and as a method of evaluat ing long-term, macroscale landscape development. The resulting weathering profiles are capped by a highly indurated laterite, which forms the uppermost levels to the residuum (Fig. 4). Preservation and exposure of entire profiles are not common, but can occur in precipitous cliffs around the margins of the high-level mesas of the Western Ghats and along the tops of the incised Konkan river valleys. The valley sections, however, are usually degraded by mass movement and dense vegetation, which means that details of the weathering profile are often partially masked.
Further south the great variety of parent litholo gies comprising the basement rocks (Fig. 1) militates against engaging in intricate geochemical characteri zation of the laterites of the southern Konkan and Kanara. Examination of the available literature indi cates that the major emphasis of previous geochemi cal investigation of these laterites has been largely towards viability assessment of potential iron ore and bauxite deposits. Geochemical investigations in the states of Goa and Karnataka, which together extract some 65 % of Indian iron ore exports, are cases in point. For the purposes of this study, priority was given to the macromorphological and physical aspects of lateritic weathering profiles, with the aim of ascertaining whether all laterite profiles were developed directly from the varied lithologies of the pre-Deccan rocks, or whether their protolith might in
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251
e
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Fig. 4. Generalized section through the lateritized weathering profile based upon study of numerous basalt weathering profiles. A variety of different lateritic textures may be identified in these profiles but all boundaries within this 'laterite stratigraphy' are essentially gradational, with one texture blending into the next. A broad zonal division can be made as follows: (a) hard, unaltered basalt displaying no visible signs of weathering (limited alteration of pyroxene and plagioclase may be apparent in thin-section); (b) basalt showing clear evidence of alteration, typically lighter in colour and often soft (incipient saprolite); (c) saprolite -some basalt textures are retained but become less readily identified as Fe-mottling increases up-profile; (d) laterite -often unindurated, but can show evidence of incipient induration in upper levels, and typically comprising pisolitic and vermiform textures; (e) indurated laterite -typically vermiform textures and commonly exposed upper levels of laterite.
fact be detrital Tertiary sediment lying uncon formably on this ancient basement. Given the loca tion of the laterites of the southern region in a high-energy environment adjacent to the foot of a high-relief escarpment, this latter assumption seemed reasonable and a sedimentological and geo morphological emphasis was thus warranted. Prob lems arose because many of the duricrusts are too thick (c. 20m), or sections too shallow, to reveal the lower portions of the profiles. Investigation, however, was greatly aided by access to roadside quarries and water wells. Fresh trenches along the new Konkan Railway line also offered satisfactory exposure conditions. Laterite types
Terminology The term 'laterite' has often been used very loosely since it was first suggested by Buchanan (1807), and
this subsequently has led to considerable confusion. Many researchers have since debated the problem with particular reference to the Indian examples (e.g. Oilier & Rajaguru 1989; and references therein), and more generally within the wider field of laterite research (Aleva 1994 ). Problems of co-ordinating lat erite description stem not only from it being the subj ect of investigation by a variety of scientific disci plines (e.g. geomorphologists, geologists and pedolo gists), but also from the difficulty in reconciling extensive anglophone and francophone terminolo gies. As the present work incorporates the region described by Buchanan it is not possible to avoid this contentious issue entirely. We therefore believe it important to outline briefly the inherent difficulties in terminology, and to clarify the usage of the term 'laterite' in the following sections of this paper. Fox (1936) correctly noted that Buchanan's definition strictly described an iron-rich material that indurated on exposure to air in the region of Angadipuram in central Kerala. Fox subsequently
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M. Widdowson and Y. Gunnell
proposed to requalify the laterites of the proto type region as a 'lateritic lithomarge'. This lateritic lithomarge, extensively observed in this region of India, corresponds to a horizon on granitic-gneissic rock that is harder and richer in iron than the lithomargic clays of the mottled zone, but softer and poorer in haematite than the cuirasse, which repre sents the indurated end member of lateritization. Tardy ( 1994) defines such a relatively soft accumu lation of kaolinite, goethite, and haematite, often bearing large proportions of quartz, as a carapace. This formal distinction between 'carapace' and 'cuirasse' is clearly validated in the field by the fact that laterite brick quarries are exclusively cut into this particular type of carapace material, whereas the harder cuirasse horizon is normally avoided. Contrast between this carapace-type laterite (sensu Buchanan) and the indurated cuirasse-type laterite mesas is also made apparent through differences in vegetation cover; carapace laterite is typically soft and moist enough to support fairly dense, although secondary, evergreen forest and jungle, but the edaphic aridity of the cuirasse-type laterites capping the mesas restricts vegetative growth to an open for mation of grasses and sparse shrubs. This contrast translates as distinctive false colour signatures on satellite images, which proved useful in the mapping process. From the observations made during the current study, laterite (sensu Buchanan) thus appears to correspond to carapace (sensu Tardy). As even this simple comparison may ultimately prove con tentious, however, for the sake of simplicity we shall use the term laterite in a generic fashion throughout the remaining text thereby obviating the neces sity to qualify it repeatedly as either cuirasse- or carapace-type. Lateritic profile development From the field studies it is evident that the majority of laterites developed in both the Konkan and Kanara regions can be described adequately in terms of two genetically distinct types. The first of these are essen tially a residuum, and have evolved by processes involving the in situ weathering and breakdown of the local lithology; the second group of laterites differ in that they clearly incorporate physically trans ported materials (i.e. clasts derived from adjacent lithologies or pre-existing laterites that originate beyond the immediate vicinity of laterite develop ment). These two types have been termed elsewhere in the literature as 'primary' and 'secondary' laterites.
To avoid any terminological conflict regarding the usage of primary and secondary originating from the type of historical precedent outlined above (see McFarlane 1976; pp. 95-96), however, we prefer here to use the more neutral terms of autochthonous and allochthonous to describe the nature of Indian coastal laterite occurrences. This broad, twofold dis tinction of laterite types is made on the following basis: 1 Autochthonous laterites are those that have formed directly from the breakdown of materials in their immediate vicinity. In other words, they do not contain any identifiable materials that have been transported physically from the wider environment and incorporated into the evolving laterite profile. These types of laterite are typically manifest as the uppermost layers of in situ weathering profiles. Where these profiles are fully exposed, taking for instance those examples developed on Deccan basalt, they consist of the following textural progression. Topmost is a highly indurated vermiform laterite, which then passes downward through softer layers in which the vermiform tubes become progressively less well developed, and then further downward into horizons in which individual pisolith structures may be recognized. Below the vermiform and pisolitic horizons, unindurated lithomarge layers occur, which eventually pass down into saprolitic and less weathered horizons in which structure and crystal pseudomorphs of the parent basalt may be recog nized (Fig. 4). 2 Allochthonous laterites are those laterites that incorporate materials not considered indigenous to the immediate locality in which the laterite is, or has formed. In most cases the materials transported can be identified readily as pebbles or clasts derived from adjacent lithological terranes or from the indurated layers of early generations of laterite. In principle, this grouping also should be extended to those lat eritic materials with constituents that have been aug mented substantially by the precipitation or capture of elements and compounds from solutions and fluids derived from the breakdown and mobilization of materials existing beyond the immediate locality of laterite formation. In practice, the degree of influence and the allochthonous nature of such fluids is often difficult to establish. Nevertheless, within the Konkan and Kanara, allocththonous laterites often can be dis tinguished by the fact that they typically display an obvious discordance with the lithologies upon which they are developed (i.e. there is no progressive
Passive continental margins weathering profile and the laterite effectively sits upon relatively unaltered bedrock). Moreover, they do not usually display the well-developed horizontal stratification of the autochthonous profiles, and typi cally occur either as lateritized accumulations within topographical depressions (e.g. as observed in the Konkan region), or as more extensive Iateritized fanconglomerates or alluvial cones that lie uncon formably upon the pre-Deccan lithologies of the Kanara. In effect, it is the allochthony of the con stituent materials of the laterite that justify the ap pellation because, in these examples, the actual lateritization of these materials may be either syn- or post-depositional. Finally, it is important to note that the above broad distinction was developed as a simple scheme to help distinguish the genetic types of laterite observed in the field. In detail, these groupings are not always mutually exclusive because a range of intermediate types do occur. For example, allochthonous laterites, once formed, can continue to evolve in response to prevailing climatic and groundwater conditions; such evolved allochthonous laterites may, over time, begin to exhibit some of the structural and textural features considered endemic to autochthonous weathering profiles. In future work such examples may warrant description of a third category of laterites, which are characterized by post-depositional alteration and lat eritic overprint of essentially allochthonous materi als. Conversely, the role of allochthonous ground waters together with lateral or downslope transport of materials cannot always be excluded in the development of in situ autochthonous laterite profiles. Such refinements, however, are beyond the scope of the current work.
THE C OA S TA L LAT E RI T E B E LT
Ubiquity of the lateritization process
Laterite is developed along the entire west coast irrespective of lithologies, although the indurated capping becomes scarce north of Bombay and south of Trivandrum. Generally, outcrops exhibit very similar materials, which typically are high amounts of Fe, and often AI, and low amounts of mobile elements (e.g. Na, Rb, Mg, Ca, Sr, etc.), but in detail they do in fact differ both chemically and mineralogically. For instance, owing to the nature of the protolith min eralogy there is no free quartz in those developed in the Deccan trap region. This general similarity of
253
weathering products indicates that, in the past, a fundamental geomorphological determinant has controlled laterite development irrespective of lithology. From north-central Konkan (c. l 9°N) to south central Malabar (c. l l 0N), the Western Ghats escarpment is preceded by a ramp of dissected lat erite-capped mesas that slope gently from altitudes of c. l80-200 m in the east of the coastal plain, to c. 80-lOO m adjacent to the coast. A second generation of laterite mesas is particularly prominent in estuar ine regions and occupies a lower elevation of 60-70 m. The rivers that cross the coastal lowland and which dissect this laterite belt, originate near or at the Western Ghats face, in both the Deccan trap and pre-Deccan basement areas, as well as from the more distant interior Karnataka plateau between 13°30' and l5°45'N. This intense dissection confers a good deal of relief to the landscape (i.e. up to 180 m) in the form of steep-sided meandering valleys. A remark able feature of the lower 20-40 km stretches of the rivers and estuaries that cut through the laterite tableland is that their bed gradients are impercep tible almost up to the knickpoint at the foot of the Ghats escarpment (Dikshit 1976; Bruckner 1989). Consequently, irrespective of whether they occur on 65 Ma basalt or on the Archaean-Proterozoic litholo gies, the tide during the dry season reaches far inland (several dozen kilometres) despite the dominant microtidal regime. So what are the controls that ensure this wide spread development of indurated laterite on the one hand, and the important differences observed between trap and pre-Deccan basement exposures on the other? First we shall examine the likely con trols on its ubiquity. From the wealth of previous studies on lateritization (see McFarlane (1976) and Thomas (1994) for detailed summaries) the key con trols are known to be as follows. 1 A favourable geomorphological environment, i.e. low runoff, lack of aggressive erosion, and the fact that palaeosurfaces are often closely associ ated with subsequent residual deposits. In the present case we are dealing with a low-lying, low-relief pediment formed by the recession of the Ghats escarpment. 2 Favourable climatic conditions: current annual rainfall is very high, more so in Kanara than in the Konkan or Malabar (Pascal 1982), and can reach averages of 5000mm at the foot of the Ghats. Yearly averages never drop below 2000 mm. There is reason to suggest that a monsoon climate probably offers the
254
M. Widdowson and Y Gunnell
ideal conditions for ongoing lateritization, as may also be inferred from the work of Maignien (1958) in Guinea and of Bourgeon (1989) in southern India. Kaolinite-forming weathering conditions are guaran teed by high humidity, and seem to have been so for the better part of the pre-monsoon times of the Ter tiary and Cretaceous (see discussion). 3 Relative tectonic stability: it is likely that uplift along the margin is an iterative phenomenon gener ated initially in the early Tertiary as a response to post-rift magmatic and thermal factors, and latterly in response to denudational unloading (Widdowson 1997b; Widdowson & Mitchell, in press). The com bination and temporal overlap of these processes has the potential to provide periods of apparent or dynamic stasis marking a transitional period between the conflicting effects of post-rift thermal subsidence and later, erosionally induced isostatic rebound along the continental margin during which lateritization of the coastal 'pediplain' could proceed (Widdowson & Cox 1996). Climate does not, however, explain the seaward tilt of the laterites, nor does it account for the intense dissection and water-table and base-level lowering, which are necessary to lateritic profile deepening and surface induration. Clearly, the consistent altitudes and the consistent seaward slope along the Konkan and Kanara are not simply fortuitous, and these points will be discussed in depth. Most probably the laterite was far more laterally extensive prior to its incision and erosion by westward-flowing rivers, although it may never have developed as a continu ous cover over the entire coastal plain. Gaps in the laterite distribution pattern, other than those caused by subsequent erosion, are prominent, always low lying (i.e. less than the 50-70 m level), and often blanketed by later autochthonous laterite profiles. Clusters of indurated laterite plateaux give way to softer laterite (sensu Buchanan) in two major loca tions: first, several stretches along the coastal lowland and most notably in South Kanara and, second, a remarkable low-lying corridor that separates the Ghats escarpment from the laterites of the outer coastal plain (Figs 2 & 3). This corridor feature par ticularly occurs in regions where the Western Ghats run at some distance from the coastline, for instance in the south-western Deccan basalt province (inner Konkan) and the deep embayment of the Mangalore hinterland (inner Kanara, see Fig. 2). Such significant interruptions in the coastal belt need to be addressed in greater detail.
LAT E R I T E S
OF THE D ECCAN
VO LCANIC P R OV I N C E AND P R E - D ECCAN BA S E M E N T: A C O MPARAT I V E S T UDY
Clearly, convincing arguments can be put forward regarding the ubiquity of lateritization along the Indian margin and hence the relative importance of the geomorphological and associated hydrological setting, lithology, tectonic stability and climate dur ing their formation. Detailed observations, however, do reveal important differences. Perhaps the most obvious are the differences between those materials developed on the basalt and those developed upon the pre-Deccan lithologies. As the two geological ter ranes lend themselves to different methodological approaches (i.e. geomorphological and geochemical data for northern profiles, and geomorphological and/or sedimentological data for southern profiles), the presentation will conveniently be divided into two parts. D istribution and characteristics of the Deccan coastal laterites
Much of the outer Konkan of the south-west Deccan from Srivardhan (73°01'E, 18°03'N) to Devgarh (73°23'E, 16°22'N) comprises a series of semicontinu ous laterite-capped plateaux. These descend to the coast, where they often form a characteristic capping to the basalt cliffs (e.g. at Ratnagiri). The Konkan laterites tend to form fiat, poorly vegetated expanses, but the topography of the surface of these plateaux often has a gently undulating character (Fig. 5). These plateaux are particularly common immedi ately inland from the coast, where they form a con tinuous, although dissected, laterite belt typically of 20-30 km width, but narrowing to about 1 5 km width in the north (Figs 2 & 3). South of the Shashtri river (c. 17°15'N) laterite can be easily discerned on satellite imagery (Widdowson 1990) as forming huge tracts of semicontinuous outcrop (c. 2000 km2) broken only by the deep meandering gorges of the main westward-flowing rivers (Fig. 6); such deep entrenchment of the Konkan drainage indicates a relative base-level fall subsequent to the lateritiza tion of the outer Konkan. Laterite development upon the basalts is most extensive between l 6°15'N and l7°l5'N, but the laterite-capped plateaux generally extend no further inland than 73°40'E, and have a quite marked easterly boundary with the inner
Passive continental margins
Fig. 5. Undulose surface of lateritized outer Konkan 10 km inland of Ratnagiri. Elevation is c. 200 m. Note the gently undulating form of this lateritized plateau. The patchy, semicontinuous exposure of the Fe-rich indurated laterite and vegetation die-back during the dry season are the main factors that permit accurate mapping from satellite images.
Konkan. The inner Konkan comprises a tract, 2030 km wide, which lies at a lower elevation between coastal laterite plateaux and Ghats foothills (Fig. 3). Indurated laterite plateaux are documented as far north as the Bombay area, for example, at Matheran and Tungar Hills (73°16'E, l8°59'N; 72°50'E, l 9°24'N; Wilkins et a!. 1994), and southwards of l 6°30'N they are widely developed on the pre-Deccan lithologies of Goa and North Kanara (see later sections). The coastal belt of laterite-capped plateaux lie at concordant elevations over large tracts of the Konkan, indicating that originally they comprised a lateritized palaeosurface of regional extent prior to river incision (Widdowson 1997b ). However, inspec tion of heights throughout the area studied as a whole reveals systematic east-west and north-south varia tions in elevation (Fig. 6, and inset). A gentle west wards slope can be recognized across many of the larger Konkan plateaux, as their elevation 15-25 km inland generally lie some 90-120 m higher than those adjacent to the coast. The height data also suggest a southerly decline in elevation, because the elevations of the northern plateaux lie c. 230 m near Srivardhan, at lOO m inland of Ratnagiri, and descend to 70 m in the southern Konkan, inland of Devgarh. Inspection of the elevation data (Fig. 6) also reveals embayments in the contours, especially along the courses of major rivers (e.g. Shashtri, Kajvi, Muchkundi). The origin of these is uncertain but they may simply represent an original lateritized topography comprising inter fluves and shallow valleys associated with the me andering rivers that drained the pediment prior to its
255
incision or, alternatively, may represent a series of closely spaced, subparallel lateritized benches or ter races that developed in response to an iterative base level fall (see later discussion). Although low-level laterite mesas tend to form the more elevated regions of the outer Konkan plain, they do not always form the highest topography. A number of unaltered basalt spurs and ridges may be recognized, sometimes extending from the Ghats escarpment, and rising above the general level of the Konkan laterite plateaux. Field investigation reveals no evidence of eroded laterite blocks or other lat eritic detritus upon the flanks of these ridges, which might otherwise indicate that they had originally been lateritized or were the remnants of a more ele vated surface. Rather, it seems that the laterite devel oped around the bases of the ridges, leaving them as unaltered basaltic 'islands' upon the Konkan. Inspec tion of satellite imagery and topographic maps confirms that although such east-west trending basalt ridges are widely spaced they occasionally extend to the coastline. These basalt outliers, together with the undulatory surface of the Konkan laterite profiles are characteristic of a pedimented region, and are entirely the result of easterly recession of the Ghats escarpment. It is worthy of note that early authors (e.g. Fermor 1909) speculated that a large proportion of the Indian low-level (i.e. coastal) laterites were detrital (allochthonous) in origin, and cited the high-level indurated material capping the summits of the Western Ghats crest as a possible source (Fig. 3). This belief may have originated from preferential investi gation and interpretation of examples of lateritized allochthonous accumulations found at some locali ties along the foot of the Ghats escarpment. It is clear from the current work, however, that such accumula tions are more common in those regions south of the Deccan outcrop (see later sections). For the bulk of the Konkan laterite, which is extensively developed upon the Deccan basalts, this allochthonous interpre tation is clearly incorrect. Autochthonous laterites For the most part, the majority of laterites which comprise the coastal belt of the Konkan are autochthonous. This fact can be readily demonstrated by two independent lines of evidence: first, the obser vation that where laterite profiles are exposed it can be readily seen that the laterite forms the uppermost part of an in situ weathering profile displaying a
73° 20'
73°30'
1 7° 1 0'
1 7 °00'
1 6°50'
1 JC 30'
1 6°40'
1 JCOO'
1 6°30'
73° 1 5'
73°45'
Fig. 6. Distribution of indurated laterite (heavy stipple) in the Deccan Konkan coastal plain (contour data in metres). Outcrop pattern has been determined using a combination of suitably processed LANDSAT images and extensive field survey. Major rivers: (a) Savitri; (b) Vashishti; (c) Shashtri; (d) Kajvi; (e) Muchkundi; (f) Vaghotan. Inset: detail of distribution of the laterite-capped plateaux inland of Ratnagiri (light stipple regions). This is one of the most extensive regions of indurated laterite, forming a semicontinuous blanket across the gently undulose topography. Deeply entrenched (80-120 m), westward-flowing river meanders (e.g. Shastri, Kajvi and Muchkundi river systems) have now dissected this laterite belt. Contours display the general westward slope of the lateritized surface. Note also inland deflection of the contours along the courses of these major rivers.
Passive continental margins gradual progression from unaltered basalt through to an indurated residuum (Fig. 4). In such examples there is no evidence of detrital stone lines or similar sedimentological evidence between the saprolite and the overlying laterite (e.g. Oilier and Galloway 1 990) that might suggest them to be anything other than in situ alteration profiles. Second, investigation of the geochemistry demonstrates the material sampled from these weathering profiles is entirely consistent with derivation from a basaltic precursor rather than from other possible sources represented by the adjacent pre-Deccan lithologies (e.g. Archaean Peninsular Gneisses, Dharwar supracrustal se quences, and later Archaean-Proterozoic granitoids) available in neighbouring terranes to the south (Fig. 1 ) . Moreover, Widdowson & Cox (1996) demonstrate that these weathering profiles, together with their associated laterite cappings, retain trace element geo chemical fingerprints characteristic of the basalt for mations upon which they lie and that currently form the Konkan coastal plain (i.e. Poladpur and Ambenali Formations). Clearly, if they were simply recemented detrital accumulations, originally derived from erosion of the high-level profiles (Fig. 3), then a geo chemical fingerprint typical of the more elevated Ghats formations (e.g. Panhala Formation) would be expected. Allochthonous laterites Allochthonous laterites may be found at some locali ties, but field investigation indicates them to be a relatively minor occurrence compared with the widespread autochthonous laterite described above. These allochthonous forms may be divided broadly into two types: first, there are those that are obviously formed by mechanical accumulation of up-slope debris; this type is virtually absent in the Deccan coastal region, and the only positively identified example occurs in the region of Phonda Ghat adja cent to Ghats escarpment. At this particular locality the unconformity between the overlying basalts and pre-Deccan basement has been exhumed and so this particular example of allochthonous laterite has the appearance of a polymictic conglomerate containing both altered quartz-rich (granitic and/or gneissic?) and basaltic pebbles bonded together with a lateritic cement. Geomorphologically, it is similar to the fanglomerate or glacis-type accumulations that are described in greater detail in later sections dealing with laterite types typical of those developed upon the pre-Deccan basement terrain. Second, there are
257
restricted occurrences of lateritic materials that appear to have formed largely by processes involving downslope precipitation and capture of materials from ground waters that have passed through the more elevated autochthonous laterites. These tend to be restricted spatially because their existence is confined to those depressions, or slopes bordering the estuaries, that lie below the general level of the exten sive, autochthonous laterite plateaux that comprise the coastward-sloping ramp. Examination of exposed profiles of these particular allochthonous forms indi cates that, through a process of in situ maturation, they can achieve a superficial resemblance to the autochthonous types because they can display both incipient pisolitic and vermiform textures. Neverthe less, even where such maturation has occurred, they can still be distinguished by an absence of a well defined weathering progression, and comparison of mineralogical and chemical characteristics reveal fundamental differences between these two geneti cally distinct laterite types. For instance, mineralogi cal analysis of these allochthonous laterites reveals a predominantly goethitic rather than a haematitic matrix; a generally more ferric-rich (and hence lower AI) composition together with higher concentrations of the more mobile elements (e.g. Ba, Mg, Ca, N a, K, Rb and Sr) , and an elevated Mn content (average 1 .02 %) compared with the in situ alteration profiles which are Mn depleted (average 0.17%) with respect to underlying basaltic composition (Table 1 ) . A s the autochthonous versus allochthonous argu ment is crucial to understanding the geomorpho logical evolution of coastal peninsular India, the geochemical characteristics of the Deccan Konkan laterites are worthy of further discussion. Geochemical characteristics The autochthonous Deccan laterites provide an excellent opportunity to study the patterns of rela tive element depletion and enrichment during basalt weathering and lateritization. The huge lateral extent and great thickness of Deccan basalt lavas ensures that the vast majority of weathering products, includ ing any elements and materials mobilized and trans ported within the groundwater fluids, are derived ultimately from the breakdown of tholeiitic basaltic precursors, which themselves display a very limited compositional range (Table 1). In examples of later ite profiles documented elsewhere in the world, the importance of allochthonous materials and solute laden ground waters derived from sources at higher
258
M. Widdowson and Y Gunnell
Table 1. Average chemical composition of the Deccan basalts and indurated examples of autochthonous and allochthonous laterites of the northern Konkan
Autochthonous laterites (N = 21 )
Basalt (N > 400) Element (wt%)*
Allochthonous laterites (N = 15)
Average
Maximum
Minimum
Average
Maximum
Minimum
Average
Maximum
Minimum
Si02 Ti02 Al203 Fe203 MnO MgO CaO Na20 K,O
48.83 2.50 13.72 14.79 0.21 6.21 10.62 2.35 0.30 0.23
52.45 4.04 18.44 18.3 0.67 1 1.49 14.16 2.95 0.87 0.4
45.64 1 .27 1 1.78 1 1.12 0.12 4.22 9.08 1.68 0.01 0.13
14.35 3.13 33.56 48.03 0.17 0.12 0.06 0.09 0.25 0.25
23.80 5.45 43.95 57.58 1.94 0.29 0.10 0.42 0.46 0.53
10.12 1 .72 19.07 37.80 0.00 0.02 0.03 0.00 0.02 0.09
18.10 2.30 23.53 53.65 1 .02 0.26 0.10 0.12 0.47 0.46
22.80 3.24 33.03 59.10 2.81 0.45 0.16 0.16 0.98 0.87
10.84 1 .47 18.40 45.39 0.08 0.14 0.05 0.06 0.29 0.21
Ba Co Cr Cu Nb Ni Pb Rb Sr
106 51 109 217 11 85 3 10 227 358 36 105 150
321 70 443 425 31 308 7 32 442 477 95 157 273
32 39 31 76 2 41 0 0 106 251 22 66 66
123 31 813 117 18 77 27 23 18 1 101 8 41 218
312 171 2421 538 26 159 62 34 38 2452 20 81 322
56 11 393 39 13 31 9 7 6 658 2 21 174
520 129 988 209 15 160 42 44 22 919 15 96 174
1571 226 1956 448 18 242 73 74 59 1 1 22 26 190 201
119 47 521 62 12 99 17 32 11 632 8 54 142
r;o5
v y
Zn Zr
* Rare earth elements are given in p.p.m.
elevations is considered a key issue because they are thought to influence significantly the chemical composition and mineralogical evolution of low level or downslope laterites (e.g. Bowden 1997). Such allochthonous influences can make determination of protolith composition particularly difficult, especially if a range of parent lithologies are involved. The pattern of chemical change characteristic of Deccan basalt alteration is consistent with obser vations regarding in situ lateritization of mafic pro toliths found elsewhere in the world. Briefly, the data demonstrate rapid loss of the more mobile elements (e.g. Ca, Na, Mg, K, Sr, etc.) in the earliest stages of the advance of the weathering front, followed by a decrease in silica content facilitated initially by the sequential breakdown of the autochthonous rock forming (i.e. basaltic) minerals, and subsequently by the breakdown of neoformed clay minerals (i.e. kaolinite) during the latter stages of alteration (Fig. 7, and inset). These losses result in a concomitant rela tive increase in the concentration of the less mobile elements within the developing laterite profile, these
being chiefly Fe, A!, and Ti, which typically are consid ered as being residual. Throughout the Deccan Konkan, as in other typical lateritic weathering systems, the chemical evolution of the weathering profile then becomes characterized by a loss of silica and a concomitant increase in Fe and AI, which are amongst the least mobile of the major elements. Importantly, in the present study, Fig. 7 demonstrates that silica loss observed in the middle and upper levels of the laterite profile results in a near-parallel increase in both Fe and AI during the early and middle stages of alteration. In effect the Al/Fe ratio remains near unity irrespective of degree of altera tion. Scatter beyond these roughly equal proportions of Fe and AI, which may be considered typical of the initial basaltic composition, occurs only in relatively few laterite samples, where more extreme silica depletion has taken place (i.e. Si02 < 20 % ) . In those cases where relative Fe enrichment becomes domin ant, this divergence seems to be related to processes that begin to operate only during advanced stages of induration or, in the case of relative AI enrichment,
259
Passive continental margins
Projection of average basalt onto Si02-Fe203-AI203 plane
q
� �I �
'
50%
<> <> <> ��<> <> <> <><><>0� 50%
Bauxites
�ml Fe203
Laterites
Fig. 7. Ternary diagram showing range of composition of all altered materials (grey diamonds) collected during sampling of the Deccan Konkan autochthonous laterite profiles. Black spot shows composition of average Deccan basalt (Table 1). Note that the apices Si02, Fe203 and Al203 are equivalent to those of the base of the tetrahedron (inset). Brackets represent approximate range of compositions found in a typical alteration for: I, weathered basalt and unindurated saprolite; II, unindurated/semi-indurated pisolitic and vermiform laterite; Ill, indurated and highly indurated vermiform laterite. Inset: tetrahedron designed to show the average composition of Deccan basalt and weathering products in terms of four components - (i) Si02, (ii) Fez03, (iii) AlzO, and (iv) other major elements; i.e . .E(Ti02 + MnO + MgO, etc.) -and to illustrate their relative changes during development of the lateritic profiles. Unaltered Deccan basalt is composed mostly of iron, aluminia and silica, but prior to alteration 20-25% of the rock comprises titanium, manganese and the alkalis + alkali-earth elements (Table 1). Bases are rapidly leached from the system, and weathering products lie upon the shaded plane. Arrow shows main enrichment trends during alteration.
are the result of bauxitization processes, which were observed to occur in some localities. Figure 8 demonstrates the fact that within the Deccan Konkan, weathered materials are derived only from a basaltic precursor, and that any influence from other non-Deccan lithologies effectively can be discounted. Here, the composition of unaltered basalt and the sampled alteration products are pre sented, together with those compositions typical of lithologies exposed in the coastal plain south of the Deccan basalt outcrop (i.e. Peninsular Gneiss, Dharwar supracrustals, and Archaean-Proterozoic granitic bodies). The Al/Fe ratio of the analysed al teration products clearly indicates a protolith compo-
sition that presented initially roughly equal propor tions of both Fe and AI. It is evident that the Deccan basalt composition represents the only suitable can didate (Table 1) because the range of alteration products define a trend that confirms the Deccan basalt composition as an end member. Moreover, any saprolitic or weakly lateritized material derived from non-basaltic materials would be characterized ini tially by both higher silica values and Al/Fe ratios. Therefore, if these Konkan weathering products had been derived from, or influenced by, other protolith lithologies, then they would define weathering trends with end members represented either by more acidic lithologies, or else corresponding to compositions
M.
260
Widdowson and Y. Gunnell Post-Deccan alteration products
1 00
Pri mary saprol ites, laterites, ferricretes
Pre-Deccan Basement lithologies
90
Sedimentary
Gran ite/g neiss
80
OJ
Dharwar g reywackes
[I)
70
�
1 gN (f)
6 •
Deccan basalt
60
·�
�"'
OJ
50
0>
40
c ·u; "' � (.)
b
c
30
Archaen Banded I ron Formation (BIF) Archaen ( D h a rwar) g reywackes a n d phyll ites
Igneous/metamorphic
<) v 0 0
20
Archaen Fe-Mn formation ( a renites and a rg i l l ites)
Granito-gneiss clasts from Phonda conglomerate D h arwar acidic volcanic rocks (average) Peninsular g neiss Late Archaean-Early Proterozoic g ra n ites (Goa)
10 0 1
0.1 Min
I
10
100
Max Al20iFe203 ratio
Fig. 8. Discrimination diagram showing the composition of Deccan basaltic material and that of adjacent basement lithologies compared with the weathering products sampled from the Konkan autochthonous laterite profiles. General compositional fields for Deccan basalt, Dharwar greywackes and granite-gneiss compositions are shown by oval shaded regions. Vertical dotted lines are Al20iFe203 maxima and minima as determined from basalt protolith composition. The current data (grey diamonds) show a concomitant increase in AI and Fe such that the Al/Fe ratio typically approaches unity. Importantly, none of the sampled alteration products contain silica content greater than that of the basalt proto lith, and the Al203 and Fe203 data reveal a weathering trend entirely consistent with a basaltic composition. Arrows a and b are projected alteration vectors for basement greywacke and granite-gneiss compositions, assuming a pattern of Al203 and Fe203 behaviour similar to that observed in the basalt weathering profiles. Compositional data: Archaean Fe-Mn formation lithologies after Manikyamba & Naqvi (1995); greywackes and phyllites (North Dharwar craton) after Naqvi et al. (1988); Dharwar acidic volcanic rock (average) and Peninsular gneiss ( Chitradurga) after Taylor et a/. (1984 ); Archaean-Proterozoic granites (Goa) after Dhoundial et a/. (1987) and Widdowson & Fernandez (unpublished data).
lying on the mixing lines between Deccan basalt and the available Archaean-Proterozoic lithologies. Clearly non-basaltic influences have taken no part in the development of the autochthonous Konkan laterites. Laterites in the Archaean-Proterozoic (pre-Deccan basement) terrain
Both autochthonous and allochthonous laterite types occur in this region, but the situation is more complex than the configuration outlined for the Deccan occur rences. Other factors, including lithological variation and a more varied geomorphological environment
combine to produce a series of laterite plateaux and benches. Autochthonous laterite profiles In the region of Kanara and northern Kerala characterized by a variety of predominantly ancient crystalline rocks, the distribution pattern of autoch thonous laterite reveals an important degree of lithological control. Generally, the in situ deve lopment of weathering profiles was verified in the field by the observation of undisturbed quartz veins running through the underlying lithomarge, mottled zone and/or laterite carapace.
Passive continental margins The Archaean-Proterozoic lithologies may be divided broadly between iron-rich rocks, which are more conducive to laterite development (i.e. mafic types such as meta-greywackes and meta-argillites in North Kanara and charnockite in Kerala), and iron poor types, such as granite-gneiss in South Kanara. It is important to note, however, that even these iron rich lithologies do not develop to the same extent as on the laterally continuous and apparently homoge neous laterites of the northern basalt terrane, and the overall result is that laterite plateaux relief is less widespread than in the basalt province (Fig. 2). This broad distinction, however, should not overshadow a much more intricate picture of laterite variation within the southern terrane. Beginning in the southernmost reaches of the current study area, the upper level laterite plateaux of the hinterland around Mangalore (74°50' 12°53') are developed on gneiss. They are detached entirely from the Ghats escarpment by a broad depression blanketed by ferrallitic soils and non-indurated lat erite (sensu Buchanan), in a geomorphological posi tion that bears much similarity with the situation found in the basalt province (Fig. 2). A little further north, still in South Kanara, these upper level laterite mesas give way to a low relief plateau of non indurated laterite, which has developed on granite gneiss terrane (c. 13°20'-15°20'). This laterite is virtu ally devoid of any indurated capping and the area constitutes the most notable gap in the coastal belt of indurated laterite within the region studied. Never theless, in common with all other landforms of the coastal region, this plateau slopes coastward from c. 100 to 130 m at the foot of the Ghats escarpment, to between 25 and 35 m at its coastal edge, and is occa sionally dominated by residual bedrock bornhardts and domes. These are geomorphological equivalents of the basalt spurs, ridges and hills described for the northern trap region. There is, however, one notable exception to the absence of indurated laterite in this region, because at Manipal (Fig. 2), 10 km in the hin terland of Udipi (74°45' 1 3°20'), an isolated laterite capped mesa culminates at 100 m. TI1is indurated laterite is the topmost section of a weathering profile that has developed upon a large isolated mafic intru sive and, moreover, the resistant capping has pro tected this lithologically controlled landform from the effects of vertical erosion suffered by its immedi ate surroundings. It therefore can be interpreted as a marker horizon and indicates that since lateritiza tion the general topographical lowering process has proceeded by c. 70-80 m in this region of the Western
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Ghats piedmont. Further north, similar litholo gical contrasts involving amphibolite bodies in the region of Gokarn (74°19' 14°33') substantiate this observation. In the region of North Kanara, other lithologically controlled laterite types occur. South of Honavar (74°26' 12°23'; Fig. 2) an archipelago of laterite capped mesas of uniform appearance occur where a belt of Archaean ferruginous phyllite reaches the coast. Owing to its iron-rich composition, rocks such as these inevitably have been affected by a form of duricrusting, because they represent compositions approaching those of the lateritized alteration prod ucts described elsewhere (see Archaean Fe-Mn-rich lithologies, Fig. 8). Many of these mesas are currently exploited as open-cast iron mines. TI1e range of lithological controls on the nature of the laterite outcrops is probably best summarized by the case of Goa province, which displays the richest variety of rock outcrops of the west coast. It is imme diately apparent that none of the felsic rocks, which predominate notably in the form of various granites in the southern two-thirds of the territory, are capped by laterite. By contrast, in the northern third of the state, the predominant outcrop of meta-greywacke and meta-argillite seems to explain the pervasive presence of laterite, which even drapes fairly steep slopes between plateau benches. In addition to this broad division of lithologies and associated later ite development, interior Goa is ribbed with low NNW-SSE striking (i.e. Dharwar trend), Mn-rich Archaean ferruginous schist and quartzite ridges broadly similar to lithologies associated with Banded Iron Formations (BIF) and which are detailed else where in the Dharwar craton (Mankiyamba & Naqvi 1995). Where these ridges have not been exploited by huge open-cast mining craters, duricrusting fea tures again prevail. Importantly, it appears that this Archaean ironstone has acted as an upstream supply source for the allochthonous laterite, found near the coast, via the main rivers, which follow the NNW-SSE strike of the hinterland before bending across to the coast as a succession of closely spaced estuaries. The pattern of autochthonous laterite in this region thus reveals itself as a mosaic of primary occurrences that are genetically distinguishable according to parent bedrock type. None of these subtle differences can be borne out from the cursory examination of satellite pictures and topographical maps, and the available geological maps are too inaccurate to adequately guide investigations of this
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detail. Difficulties with existing geological map data are compounded by the fact that 'lateritic material' (itself an example of the existing confusion between autochthonous and allochthonous types) is treated and mapped as a single geological formation despite obvious inherent differences in the relative ages and evolution of the coastal laterites. Clearly, although laterite may be regarded legitimately as a type of rock, this type of oversimplified mapping treatment does not help in understanding the different genetic relationships and lithological controls that are instru mental in its formation. Allochthonous laterite profiles Unequivocal occurrences of allochthonous laterite on the pre-Deccan lithologies of the southern terrane are to be found as low-lying mesas fringing estuaries. These latter always appear 'countersunk' within the autochthonous laterite level, although rivers are still entrenched even into these lower levels, by about 60 m. As in the basalt province, these allochthonous lat erites constitute a second, younger generation of indurated laterite developed at a lower topographical level. Moreover, not only is there an elevation differ ence between these upper level and lower level laterites, there is also a clear sedimentary character because, whether in Mangalore city itself or at Panaji (73°50' 15°30') in Goa, this lower level laterite is developed on Tertiary alluvial material. Texeira ( 1965) was the first to report that the basal layers of the 60-70 m high ironstone caps of Goa were com posed of quartz pebble beds; our experience can confirm the findings of Texeira, as well as those of Fox (1936), who made explicit mention of pebble beds in Kanara. This fact seems to have been overlooked by the numerous subsequent studies of laterite in Goa (Sriram & Prasad 1979, 1980). As a result, much inter pretation of the laterites of western India has missed the fundamental point that the so-called Warkalli beds of terrigenous Mio-Pliocene sediment, well documented in Kerala, continue all the way up the coast into Goa at least. Our observations of sand and pebble beds in the Mangalore area as well as north ern Kerala, confirmed by the descriptions of Manga lore beds provided by Subrahmanya et al. (1991) and Subrahmanya & Rao (1991), verifies the distinctions we wish to underline. The original sedimentary nature of the material becomes apparent only through detailed scrutiny. The fact that this duri crusted alluvium occurs as palaeodeltaic 'pockets' at
the locus of present-day estuaries suggests that there may exist a direct relationship between the prove nance of this alluvial material and the late stages of scarp retreat in the Western Ghats. Considering the above, we suggest that a younger generation of allochthonous laterite formed in the coastal lowland region of western India. The observa tion is further substantiated by the presence of this laterite-capped alluvium in a region where the more elevated autochthonous laterite is absent from the landscape. This region of North Kanara, between 14°N and 15°N is exceptional in that the Ghats come very close to the coastline, thereby constricting the coastal lowland (Fig. 2). The region between Baindur (74°37' 13°52') and Kumta (74°25' 14°26') exhibits sloping plateaux, with a genetic relationship with the Ghats escarpment that is not obscured by the usual intervening autochthonous laterite mesas and the aforementioned corridor, which elsewhere (e.g. the Deccan Konkan) separates them from the escarp ment face (Fig. 3). Roadside sections, as well as the distal cliff face 1 km away near Baindur, reveal that the material comprises sand and rounded quartz pebble beds (Fig. 9). The fact that the glacis-type relief rests against the escarpment suggests a fan glomerate or alluvial fan (see Figs 1 1 & 12, p. 264), which became lateritized and indurated at a later stage. Quartz pebbles in these allochthonous later itized accumulations are up to 15 em in diameter, implying relatively high-energy transport capacities. A similar type of lateritized allochthonous accu mulation may be observed in a section through weakly cemented sandstones just north of Mangalore
Fig. 9. Detail of deeply weathered quartz pebbles in sandy detrital formation ( Baindur fan) .
Passive continental margins
Fig. lO. Exposed section of Mangalore beds, Gurpur river estuary, Mangalore, resting unconformably on coastal basement. An iron accumulation (i.e. hardpan) occurs at the unconformity between the permeable overlying sandstone and impermeable underlying lithomargic clays, which are developed below on in situ basement rock. The unconformity is interpreted in detail as partly erosional, and the deposit is similar to the Baindur detrital material of Baindur (Fig. 1 1 ) and elsewhere. This generation of indurated allochthonous laterite, which forms benches close to the estuaries, is countersunk below the higher autochthonous laterite tablelands situated further inland.
harbour (Fig. 10). This also reveals cross-bedding fea tures indicating turbulent flow during its deposition. The pebbles within these sandstones, as in similar examples elsewhere in Mangalore or Goa, are most often in a highly advanced state of decay and can be instantly reduced to powder by rough handling. Clearly, they could not have originally been trans ported in such a state of decomposition and must therefore reflect the intensity of the lateritization that has occurred since their deposition. In an earlier study of the Baindur laterite, Khanadali & Devaraju (1987) and Devaraju & Khanadali (1993) demonstrated the presence of moderate quantities of gibbsite in the shallow quar ries at the surface, thereby clearly establishing that the degree of weathering was indeed advanced and thus supporting the observations made here. It should be noted, however, that their interpretation of the Baindur laterite differs from the fanglomeratic model forwarded here. Instead, these authors favour an essentially autochthonous origin for the laterite, and cite observed mineralogical and chemical grada tions within the profile that are considered to be more consistent with an in situ weathering of the
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underlying Peninsular (tonalite) Gneiss. Careful re examination of the distal cliff described above, reaffirms that much of this laterite is compatible with derivation from transported material. Nevertheless, further detailed examination of the Baindur laterite mesa is clearly warranted because the feature may represent a more complex geomorphological situa tion than has been hitherto appreciated. One al ternative is that it could comprise a lateritized palaeosurface that combines elements from both models; for example, a pediment cut directly into the gneissose terrain by scarp recession, and then later partially covered by a glacis-type fanglomerate deposit. Such a composite structure would explain the occurrence of both autochthonous and allochtho nous laterite types across the Baindur mesa. Around Bhatkal (74°33' 13°59'), Honavar (74°27' 14°18') and Kumta, similar alluvial or fanglo meratic materials are found in quarry sections and include graded bedding sequences 5-1 0 m thick displaying coarser quartz pebble beds overlain by upward-fining sand and silt material. This material frequently is a pinkish colour, suggesting a pro portion of bauxite content, and bauxite was indeed extracted up to quite recently from open-cast quarries. These mixtures of sand, clay, and pebble material which are common at the base of the allochthonous laterites strongly suggest a phase of erosion in the Western Ghats and adjacent coastal plain, during which existing weathering profiles of autochthonous laterite were apparently scoured down, at least locally, to the weathered rock, and quartz veins or stone lines were washed away. In addition to the fan glomerate illustrated in Figs 1 1 & 12, all other alluvial occurrences are located invariably at the outlet of major drainage breaches in the Ghats, between 14°N and 15°N. The connection between river piracy of the eastward-flowing system east of the Ghats caused by such breaches, and the sudden influx of sediment deposited at base level is one possible explanation of their development. However, tectonic, eustatic or cli matic causes also need to be granted close scrutiny (see discussion) . Finally, northern Goa displays at least two tiers of lateritized pebble terrace levels at c. 40 m elevation and lower along the Terekhol and Chapora rivers, and which are sunk below the elevated (60-70 m) autochthonous laterite level. As development of these terraces appears unassociated with breaches affecting the Ghats escarpment their origins are more
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M. Widdowson and Y. Gunnell Glacis l a ndform (fang lomeratic) Sloping, lateritized mesa \
:::::
Metabasalt
& banded
ironstone (easterly d i p )
Escarpment face
+665 Spot height ( m )
Arabian Sea
Fig. ll. Map of the Baindur laterite mesa. The fanglomerate is clearly the result of a more recent (late Tertiary Quaternary?) erosional discharge resulting from the Western Ghats scarp recession. Induration, river incision and mesa development occurred subsequently. Note the photographs of Figs 8 & 1 2 are taken from a roadcut along the coastal highway through the Baindur mesa, and from the spot height at 60 m on the mesa surface, respectively.
Fig. 12. The Baindur mesa and Western Ghats. Looking due west from the sea towards the escarpment (visible in the background), the indurated glacis fan slopes away from the Ghats. Note detail of dark blocky lateritic surface in foreground.
probably related to river incision (Bruckner 1989), to the point that downcutting, terrace deposition and their subsequent lateritization seem never to have ceased since the upper coastal surface of autochtho nous laterite was first established. To summarize the diverse findings relevant to the crystalline coastal subregion: 1 autochthonous laterite concerns the whole of the upper glacis level (200-SO m) in northern Kerala, the Mangalore hinterland, and interior Goa; in terms of geomorphological setting and genesis, it correlates well with the upper level Konkan laterites of the basalt province. 2 allocththonous laterite is found largely seaward of, and at lower levels than, the elevated autochthonous level. The 'fanglomerate' form is represented best where the coastal region narrows and the Ghats escarpment comes close to the coast.
Passive continental margins DISCUS SIO N Laterite evolution in the Indian coastal lowlands
Having detailed both the distribution and typology of laterites along the coastal lowlands comprising the Konkan and Kanara, it is necessary to discuss their origins and evolution and, importantly, develop a model that accounts adequately for the observa tions outlined regarding the various autochthonous and allochthonous forms. This can be achieved best in terms of a morphotectonic model that combines the fundamental controls of tectonics of the Indian margin, its post-rift geomorphological evolution and the changing climatic regime.
Evolution ofthe Deccan Konkan lowlands Considering first the northern coastal zone compris ing the Deccan Konkan, it is clear that the predomin ant form of laterite development is autochthonous, in situ weathering profiles capped by an indurated layer. These are derived from a basaltic protolith (i.e. Poladpur and Ambenali Formations), which forms the underlying lithology of the coastal plain (Fig. 3). From established stratigraphical relationships, evolu tion of the coastal plain here is demonstrably the result of the removal of between 1 and 1.5 km thick nesses of basalts, which originally lay west of the current Ghats escarpment: the coastal strip therefore constitutes a pedimented tract or pediment palaeo surface. The mechanism for this type of erosion is by the process of scarp retreat, and although such con vincing stratigraphical arguments cannot be applied so readily to the pre-Deccan basement terrain lying to the south owing to its characteristically complex lithological relationships, the fact that the Western Ghats continue uninterrupted as a geomorphological feature makes it entirely reasonable to assume that the southern extension of the coastal plain (i.e. south ern Konkan and Kanara) is similarly the product of scarp retreat. This view is consistent with the models of King (1953) who observed that this type of parallel scarp retreat was largely unaffected by variations in rock type or structure; this therefore would account for the preservation of the Ghats escarpment as a linear, coast-parallel feature developed in both the Deccan basalts and varied Archaean-Proterozoic lithologies. In general, average elevations across pediments are controlled by the regional base level and their surfaces typically comprise low-amplitude topogra-
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phy characterized by sheet-flow, sluggish drainage, very low rates of erosion and, as a consequence, are often associated with deep weathering effects and widespread lateritization (Mabbutt 1966). All these features accord with those observed for the Konkan. For instance, the undulating nature of the lateritized surface probably represents the original pediment topography, whereas the unaltered basalt hills and ridges that rise above the level of the Konkan may represent outlier hills and ridges remaining behind as a consequence of scarp retreat and largely escaping the lateritization affecting the surrounding pediment (Fig. 13a,b). What is also apparent in the northern Konkan, however, is the fact that such a widespread development of the indurated capping laterite could have resulted only from pro longed alteration of the exposed pediment. This itself implies a protracted period in the geomorphological evolution of the western Indian margin characterized by an unprecedented degree of tectonic and climatic stability. The causes of this stability remain the source of debate, but Widdowson & Cox (1996) suggest a period of 'dynamic stasis' transitional between post-rift thermal decay and its associated subsidence, and later isostatic rebound resulting from onshore erosion. It is also evident that, at a later stage, stability of the Indian margin was interrupted by a relative base-level fall or regional (epeirogenic?) uplift leading to incision of the westward-flowing Konkan drainage (e.g. Vashishti, Shashtri, Kajvi, Muchkundi and Vaghotan) into this coast-fringing laterite apron. The courses of these rivers, which derive their headwaters along the foot of the current Ghats escarpment, have been preserved as a series of deeply incised meanders, suggesting that to some extent they retained their courses (i.e. those devel oped originally across the low-relief pediment) and, moreover, suggests swift entrenchment following a phase of rapid base-level change. It is also likely that the uplift that caused this entrenchment subsequently disturbed the established hydrology and morphodynamics of the coastal plain and, as a result, also curtailed the further maturation and development of the Konkan autochthonous laterite profiles. The exact timing and duration of the uplift is difficult to ascertain but palaeomagnetic data (Schmidt et al. 1983) record a middle-late Tertiary age (Miocene?) for the Konkan laterites. These data demonstrate that further development and resetting of magnetically susceptible oxyhydroxides within the Konkan profiles essentially halted after this date, and it is logical to assume that this was a
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M. Widdowson and Y Gunnell U n l ateritized residual Ghats scarp outliers Low-lying ped iment formed by eastward recession of G h ats scarp l i n e and cut to preva i l i n g sea level
V a l l e y c u t b y d r a i n a g e breach i n G h ats escarpment
A l l uvial fan glacis developing at mouth of drainage breach
(a) Mid-upper Tertiary
Widespread autochthonous laterite forming over much of the Konkan coastal p l a i n
Extensive tracts of autochthonous latarite preserved on u p l ifted rem n a nts of dissected ped i p l a i n
Meandering, westward flowing rivers
Low-lying i n n er Konkan corridor
Western G hats escarpment
Localized al lochthonous laterite developing i n topographic lows
Valley cut by drainage breach i n Ghats escarpment
U n i nterrupted ped i ment to G h ats ( i n n e r Konkan corridor absent)
(b) Tertiary to Quaternary
Lateritized river terraces cut below pediment level
Deeply entrenched river meanders M u ltip le generation a l luvial fan glacis prograding from pre-established drainage breach
Fig. 13. (a) and (b) Block diagrams depicting the evolution of the main geomorphological elements comprising the Konkan-Kanara coastal lowlands during the late Tertiary-Quaternary. Distance d is degree of scarp recession occurring since uplift of lateritized pediplain. Key: autochthonous laterite- light grey; allochthonous laterite -medium grey; alluvial fans/fanglomerate -dark grey.
consequence of the base-level change described above. If these relationships are correct, then the indurated laterite-capped plateaux of the outer Konkan comprise the remnants of a middle-late Ter tiary pedimented palaeosurface. Moreover, the time elapsed since uplift and fossilization of these later ites would have provided ample opportunity for the
development of younger generations of autochtho nous and allochthonous laterites at lower topogra phical levels. The scattered occurrences of laterite within localized 'lows', along valley floors, and fring ing estuary mouths therefore may represent later generations of autochthonous laterite upon topo graphical benches such as river terraces or, al ternatively, allochthonous laterites derived from
Passive continental margins downslope precipitation and mechanical accumula tion from the pre-existing Konkan laterites capping the surrounding interfluves. Indeed, multiple genera tions of laterite terraces along the entrenched river valleys could provide a satisfactory alternative expla nation for the contour embayments coincident with these river courses (Fig. 6). During the course of the current study it proved beyond the resolution of the satellite imagery used in the laterite mapping of this region to distinguish between individual narrow strips or laterite terraces developed at discrete eleva tions along these river valleys. In the field, the natural vegetation and land-slippage along the valley sides largely obscures evidence that might otherwise point to multiple generations of lateritized river terraces. If future investigation should confirm the presence of such terraces, their relative position and age will prove crucial to determining the chronology and style of uplift across the Konkan in the period follow ing dissection of the Konkan pediplain and its as sociated laterites (i.e. late Tertiary morphological evolution). Significantly, the case for later genera tions of laterite has been demonstrated already in the pre-Deccan basement terrain to the south, where their existence and stratigraphical relationship with localized occurrences of Tertiary sediments may be determined more readily. Evolution of the southern Kon.kan and Kanara lowlands Lateritization presumably was continuous through out the entire Cenozoic in a tropical climate char acterized by seasonal contrasts, as suggested by palaeobotanical (Meher-Homji 1989), palaeo weathering (Bruckner & Bruhn 1 992) and numerical modelling evidence (Fawcett et al. 1994). Rainfall and seasonality contrasts probably intensified when the present-day monsoon circulation, first recorded by palaeoenvironmental data at c. 8 Ma in South Asia (Prell & Niitsuma 1 989) , set in and thus sus tained conditions appropriate for ongoing profile deepening and laterite generation. The main point remains that at least one important post rifting phase of uplift needs to be held responsible for the two tiers of laterite landforms observed in the coastal landscape; the hot and humid climate being a necessary backdrop to account for the pervasive kaolinitization seen throughout the region. The major phase of uplift affecting the coastal plain seems to have occurred in the late Tertiary, as confirmed by the massive influx of prograding ter-
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rigenous sediments on the continental shelf from middle- to late Miocene (Whiting et al. 1994; Gunnell 1997). It is then likely that the detrital material, now capped by the lower level allochthonous laterites (Fig. 10), represents the emerged thin end of the off shore wedge of late Tertiary clastic sediments. If such is the case, the relative chronology of the landscape evolution may be refined further. The allochthonous laterites just mentioned are likely to be equivalent to the Ratnagiri, Warkalli, and Mangalore Beds described in the literature (Saxena et al. 1992; Bruckner 1 989). These incorporate a series of plant-rich clays and sands that have been found to be of Late Miocene age. This logically implies that the lateritization of these formations occurred later (i.e. in the Pliocene). This finding then suggests that the dissection of this 'allochthonous laterite' 60 m level may have occurred during the Quaternary, and therefore possibly could be related to glacio-eustatic and/or tectono-isostatic base-level changes. Such an apparently recent and rapid formation of laterite in this situation may be explained by the fact that the unconsolidated parent material already consisted of abundant primary ingredients of laterite, kaolinite, iron and quartz clasts. If this interpretation is correct, then the more elevated laterite mesa level, lying between the elevations of 100-200 m, is probably Miocene (i.e. middle-late Tertiary) in age, as is independently confirmed by the palaeomagnetic ages obtained by Schmidt et al. (1983) from southern Konkan exposures. This level, before being dis sected in the late Tertiary by river entrenchment, therefore reflects an earlier stage of post-rift pedimentation. Evolutionary models
Scarp recession of the Western Ghats The origin of the Western Ghats escarpment is difficult to determine, but given the evidence of east ward recession derived from drainage patterns (Wid dowson 1997b ), it follows that over geological time its position has changed with respect to the rifted edge of the continental margin. The Western Ghats scarp may have had its origin 65 Ma at the nascent Indian continental margin, possibly as an emergent seaward facing fault-scarp generated as a consequence of the rifting process. If correct, this structurally defined proto-escarpment initially would have been adjacent to, or near, the locus of rifting, but once exogenic processes became dominant its subsequent geomor phological evolution resulted in an easterly recession
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of the scarpline across the continental margin. However, owing to the varied character of the lithologies encountered by the recessing scarpline, together with changing climatic factors, it should not be assumed that rate of recession has always remained constant over time or indeed along its length. Nevertheless, an estimation can be made given the current relative positions of the south-west Deccan Ghats scarpline and the offshore edge of the continental margin (c. 200 m isobath; Naini & Talwani 1 983), which suggests that c. 120-180 km of scarp recession has occurred since the Seychelles rifting event c. 65 Ma. Simple calculation gives a recession rate of 1 .8-2.8 km Myr-l , which is remarkably similar to estimates of recession determined for the Aus tralian Great Escarpment (Pain 1985; Siedl et al. 1996), and the Drakensburg of South Africa (King 1967). It should be noted, however, that the offshore shelf edge is somewhat closer (i.e. between 70 and 100 km) to the southern Konkan and Kanara coastline (1 3°-16°N), which could be taken to indicate that recession in the pre-Deccan basement terrain has here proceeded at a slower rate. Whatever the case, if the model is correct then both the Konkan and Kanara coastal plains must be notably diachronous across their width, because their most juvenile reaches should lie adjacent to the present Ghats escarpment. The latter point is of some importance because of the current debate regarding the origin and evolution of great escarpments, and their effect upon megascale drainage at passive continental margins (e.g. Bishop & Goldrick, in press). A range of evolutionary variants are certainly possible, including those advocating an uplift-controlled asym metrical downwearing, which do not necessarily
require significant scarp recession; these latter have been cited as a plausible explanation for the macro geomorphology exhibited at other passive margin examples (Summerfield 1991). Clearly, a demonstra tion of the diachronous nature of passive margin coastal plains would confirm scarp recession as the fundamental macromorphological control. Such proof, however, awaits the application of techniques such as surface-exposure dating (e.g. by cosmogenic isotope techniques) and detailed exploration of the exhumation record across these coastal plains (e.g. by fission-track methods). Nevertheless, the scarp-recession model certainly accords well with the geomorphological and geologi cal observations of western India made here and else where (e.g. Widdowson & Cox 1996; Oilier & Pain 1997; Widdowson 1997b). Moreover, such a model also provides a plausible explanation for the origin of the low-lying, unlateritized inner Konkan: using the rates of recession outlined earlier, estimates indi cate that a further 25-30 km of retreat would have occurred during the late Tertiary. This value agrees well with the current width of the inner Konkan, which lies between the inner edge of the coastal lat erite apron and foot of the Western Ghats escarp ment. We therefore suggest that the low-lying, inner Konkan and Mangalore embayment evolved dur ing late Tertiary times and were overdeepened in response to the lowering of base level during uplift. In this region a protective laterite capping did not develop, possibly as a result of more intense stream dissection, or else changes in climatic or tectonic sta bility which may have worked to further enhance the erosion and overdeepening process. If this interpreta tion is correct, then the well-defined easterly margin
Fig. 14. (Opposite.) (a, b & c) The geomorphological effects of scarp recession and coastward cambering of the Deccan coastal lowlands. Figures show evolution of the Konkan-Kanara laterite belt and development of the low-lying inner Konkan-Kanara corridor. (a) Post-rift subsidence and concomitant onshore uplift produces initial cambering of the margin and basalt stratigraphy. Scarp recession produces a low-lying pediment cut to the base level existing during mid-Tertiary times. Evolution of this pediment, and its subsequent widespread lateritization, would have required a period of protracted stability along the continental margin. Given the prevailing morphotectonic environment, this stability probably could have been achieved only during a phase of 'dynamic stasis', when post-rift thermal subsidence was balanced by denudationally driven isostatic compensation resulting from ongoing scarp recession (Widdowson & Cox 1 996). (b) Widespread lateritization of the exposed coastward-dipping basalt stratigraphy is terminated by uplift in mid-upper Tertiary (Miocene?) times. Meandering coastward drainage begins to etch into the lateritized pediment (see also Fig. 13b ), and ongoing scarp recession continues by cutting to a deeper level in response to the relative base-level fall; thus the inner Konkan-Kanara corridor begins to form. The easternmost edge of the autochthonous laterite belt marks the position of the mid-upper Tertiary scarp position while its westernmost reaches become increasingly cambered toward the coast. (c) Coastward cambering of the lateritized lowlands continues as the axis of uplift migrates inland in response to scarp recession. The autochthonous laterite belt now exists as a series of extensively dissected laterite-capped mesas displaying a general coastward dip (see also Fig. 6). The low-lying inner Konkan-Kanara corridor continues to widen in response to the relative lowering of the base level. Conditions in the corridor region remain unsuitable for the development of indurated autochthonous laterite profiles.
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of the upper laterite pediment marks the position of the Ghats escarpment during mid- to late Tertiary times (Fig. 14a & b). With respect to the low-lying, fanglomeratic lat erites of North Kanara, there seems little doubt that they belong to the same, Plio-Quaternary generation as the other allochthonous laterites identified else where in the region. The question arises, however, as to the disappearance in this area of the inner Konkan depression and its equivalent, the Mangalore embay ment. An explanation may be given as follows: first, it may be that the more variegated lithologies of North Kanara have obscured the neater pattern of more elevated occurrences of autochthonous laterite, which is continuous both on the basalt and on the South Kanara gneiss. Second, if it is accepted that allochthonous laterites reflect a burst of renewed erosion associated with a late Tertiary scarp retreat phase, it must follow that sediment supply occurred following two different modes: in the more readily weatherable rock, such as basalt and gneiss, the Western Ghats retreated physically as described pre viously. When, however, the immediate hinterland was barred by the major Ghats-parallel greenstone belts, scarp recession was seriously impeded or even halted altogether. The response to uplift thus mani fested itself by major streams cutting back into the Deccan plateau and, eventually, as clearly evident from topographical maps, capturing entire portions of the upland drainage network. The well-exposed fanglomerates are associated intimately with major drainage breaches across the continental divide of the Western Ghats and reflect this particular form of sediment discharge (Fig. l3b ). They thus appear to be geologically more recent, post-retreat features of the Western Ghats, perhaps related to Pliocene(?) river piracy and breaching of the escarpment. Importantly, some of them lodge against the current Ghats escarp ment and have hardly detached from the scarpline. Excepting the Phonda conglomerate, the existence of Iateritized fanglomerate cones in northern Kanara and southern Konkan are seemingly absent (or unidentifiable) in the volcanic region. This absence of younger lateritized clastic sediments on the Deccan Konkan is entirely consistent with the lack of any major drainage breaches across the Ghats divide in this region. Seaward tilt ofWestern Ghats lateritic foreland Another observation that requires addressing is that of the coastward dip of the laterite belt. This char-
acteristic is ubiquitous in both the Deccan and pre-Deccan basement terrains, but the preceding chronology leaves unanswered the origin of this westerly slope. It is evident that the Konkan and Kanara autochthonous laterites along the continen tal margin of western India have undergone uplift, but the timing and underlying causes of this phe nomenon remain a source of debate. Dynamic and thermal effects of hot-spot plume activity are known to be the cause of initial surface uplift along the flanks of newly rifted continental margins, but it is evident from both geological and geomorphological observa tion that uplift effects in western India have con tinued long after such plume effects should have decayed. With this in mind, two viable models are suggested. 1 The slope is a simply geomorphological feature (i.e. a pediment); its coastward dip is a natural geo morphological consequence consistent with palaeo surface evolution resulting from recession of the Ghats escarpment, and later uplift has since caused the established drainage to etch itself into this pediment. 2 The slope itself is the result of uplift, the axis of which is roughly coincident with the Western Ghats. Such coast-parallel uplift is consistent with numerical modelling of denudational isostatic adjustment along rifted passive margins, which indicate a warping of the margin in response to the removal of thicknesses of 1-2 km of material by scarp recession. The first possibility is based purely on the common knowledge that pediments and associated alluvial fans at the foot of mountain ranges normally slope away from the scarp face and that, according to the King model of scarp retreat, pediments widen concomitantly with free face recession. The seaward sloping Ghats pedimented palaeosurface could thus be interpreted as an ancient geomorphological feature, preserved in its primitive form, and the entrenched river valleys have had time to grade their beds to the present-day base level. This, together with post-Pleistocene sea level rises (Kale & Rajaguru 1987) affecting the west coast, could explain why the tide comes so far inland through the laterite apron toward the escarpment. The second possibility, favoured by Widdowson & Cox (1996) and Widdowson (1997b ), emphasizes the alternative view that the slope is a consequence of flexural (monoclinal) arching of the margin, and follows the general model originally promulgated for south-west Africa by Summerfield (1988, 1991), Gilchrist & Summerfield (1990) and Gilchrist et al.
Passive continental margins (1994) . Although the first hypothesis suggested that the lateritic cappings were subsequent to pediment and slope achievement, this second scenario postu lates that the slope is acquired and increases over time, including the period after achievement of the laterite capping. Its evolution is governed instead by a regional tectonic response rather than by wholly geomorphological processes. The natural con sequence of this denudationally driven model is a net uplift across the inner Konkan and adjacent escarp ment, matched by relative subsidence of the offshore continental margin and adjacent coastal regions in response to offshore sedimentary loading. In effect, the outer margin is progressively drowned concomi tantly as its inland reaches are extended inland by scarp recession, thus providing an alternative explan ation for the drowned estuaries and associated tidal influences described above. This latter model, which predicts coastward cambering of the Deccan Konkan lowlands, is supported by the fact that the underly ing basalt stratigraphy also exhibits a westerly dip, which is steeper than that observed across the laterite plateau (note the angular disconformity between the basalt stratigraphy and overlying weathering profiles shown in Fig. 3). Clearly, in order for a disconformity of this nature to have evolved, coastward cambering of the basalts must have begun prior to the develop ment of the pediplain and associated autochthonous laterites, and as such points to the existence of a long term uplift mechanism operating along the entire western margin of peninsular India. This being the case, it follows that the laterite-capped detrital sedi ments along the coast in the Archaean-Proterozoic terrane, most of which are clearly tilted towards the sea, should be considered as tilted river terraces rather than dissected alluvial fans. The lateritized alluvial fans lodged against the base of the escarp ment away from the coastal reaches, however, must testify to later stages of sediment lateritization, which affected alluvial cones fed originally by westward flowing drainage reversals and associated breaches of the Western Ghats divide (Fig. 13b).
C O NC L USIO N S
This work describes the distribution and typology of lateritic materials developed upon the low-lying coastal plain of western India, and places their devel opment within the context of modern ideas regarding the long-term morphotectonic evolution of rifted passive continental margins.
271
Laterite and laterite development is ubiquitous along the Konkan and Kanara lowlands irrespective of obvious lithological controls and, at the broadest scale, clearly points to an underlying morphotectonic control. During the evolution of the continental margin different phases of lateritization can be rec ognized. Each phase corresponds to a period during which the topographical, tectonic and climatic condi tions were conducive to the development of wide spread laterite. The most widespread phase appears to have been a consequence of the development of a coastal pediment cut by scarp recession simultane ously into both the Deccan basalt sequences and Archaean-Proterozoic basement lithologies alike, and which subsequently provided a topographical site suitable for lateritization to proceed during middle-late Tertiary times. The product of this phase are the extensive autochthonous laterites of the Konkan and Kanara. In the pre-Deccan terrain of the southern Konkan and Kanara the array of laterite types and their dis tribution appears complex initially. A series of autochthonous laterite-capped mesas is evident (e.g. inland of Mangalore) and these may be considered as equivalent to the extensive autochthonous laterites of the coastal Deccan. The elevated position and geo morphological environment in which these laterite capped plateaux are currently located leaves little doubt that a major and permanent uplift affected the western Indian margin during the late Tertiary (Figs 13a & b & 14a & b). Uplift resulted in entrenchment of drainage across the coastal plain and, together with scarp-recession effects, initiated erosion of the autochthonous laterites and adj acent lithological ter ranes. This ultimately provided the material for the later generations of allochthonous laterite, which are now manifest as downslope accumulations and river terraces that developed at elevations below the autochthonous laterite mesas, and as lateritized fan glomerates lodged adjacent to the escarpment in those locations where coastal drainage systems suc cessfully breached the Ghats divide. The uplift mechanism is interpreted here to be the consequence of the isostatic response to onshore denudational unloading operating through eastward retreat of the Western Ghats escarpment, and a concomitant offshore sedimentary loading, which has produced a lithospheric flexuring of the entire margin. Importantly, because denudational unload ing is independent of plume dynamics and post-rift thermal effects, it provides a long-term mechanism that permits the generation of permanent and contin-
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uing uplift over geological time. Therefore, this mor photectonic regime provides a suitable framework within which the development of the autochthonous and allochthonous laterites of the coastal lowlands of western India may be adequately explained.
AC K N O W L E D G E M E N T S
The work presented was achieved through grants GT4/85/GS/83 and GS9/963 given by the NERC (MW), and CNRS-URA 1562 (YG). Geochemical analyses were performed at Department of Earth Sciences, Oxford University, and this XRF work owes much to Keith Parish, Steven Wyatt and Roy Goodwin for their invaluable assistance. Drafting facilities were made available by the Department of Geological Sciences, University of Durham. We are grateful for the logistical help and advice offered in India by K.V. Subbarao, F.B. Antao, O.A. Fernan dez, and Gerard Bourgeon from the French Institute in Pondicherry. We wish to thank Michael Thomas, Keith Cox, T.C. Devaraju, and the Editor, Medard Thiry, all of whom made useful suggestions toward improvement of this manuscript.
R E F E R E NC E S
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Spec. Pubis int. Ass. Sediment. (1999) 27, 275-301
Relief features and palaeoweathering remnants in formerly glaciated Scandinavian basement areas
K . L I D M A R - B E R G ST R O M * , S . O L S S O N t and E . R O A L D S E T :j: *Department ofPhysical Geography, Stockholm University, S-1 0691 Stockholm, Sweden; t Department of Quaternary Geology, Tornaviigen 13, S-22363 Lund, Sweden; and +Department of Geology and Mineral Resources Engineering, Norwegian Institute ofScience and Technology, N-7034 Trondheim, Norway
A B S T RACT
The relationship between saprolite remnants and palaeosurfaces in Scandinavia is examined. Thin kaolinitic saprolites are associated with the sub-Cambrian peneplain and its contact with the Cambrian cover, whereas thick kaolinitic saprolites are associated with Mesozoic cover sediments and exhumed sub Mesozoic (sub-Jurassic and sub-Cretaceous) undulating hilly relief. Characteristics of clayey-sandy and gravelly saprolites from several sites are described. Furthermore, mineralogical and chemical analyses of Weichselian deposits show that pre-weathered constituents may make up a substantial fraction of these deposits. Altogether this shows that deep weathering has been an important process in the shaping of relief even within formerly glaciated basement areas.
IN T R O D UCTI O N
almost entirely of Precambrian and Caledonian base ment rocks (Fig. 1). The Precambrian basement is sur rounded by Vendian to Lower Palaeozoic rocks. Palaeozoic strata also top the Proterozoic sedimen tary sequence in the Gulf of Bothnia and rest directly on the Precambrian basement surface within small areas in southern Sweden and Norway and along the overthrusted nappes of the Caledonides. The Precambrian basement was rifted in the Late Palaeo zoic, forming the Permo-Carboniferous Oslo Rift (Ramberg & Spjeldmes 1 978; Ziegler 1 990; Neuman et a/. 1992). In the west and south Mesozoic sedimentary strata rest directly on Caledonian and Precambrian base ment. Jurassic and Cretaceous outliers are encoun tered within the basement in the vicinity of Mesozoic sedimentary basins (e.g. Brekke & Riis 1987; Wikman & Bergstrom 1987; B¢e & Bjerkli 1989; Ziegler 1990; Dore 1991; Riis & Fjeldskar 1992; Lidmar-Bergstrom 1 996). Rifting outside western and northern Norway occurred in the Mesozoic and Cenozoic, mainly along prominent fault systems that date back to Permian time (e.g. Ziegler 1988, 1990; Gabrielsen et al. 1990). The major fault zones have been reactivated several
Fennoscandia i s characterized b y glacially polished bedrock surfaces that were produced during the Pleistocene Epoch. In seeming contradiction, the gross forms constituting the landscapes generally are pre-glacial in origin. It has long been thought that pre-glacial saprolites were totally removed by the Quaternary glaciations. However, an increasing number of pre-glacial saprolite remnants have been encountered in situ and redeposited in Quaternary strata. The in situ saprolites occur on palaeosurfaces of various ages and extensions. These surfaces are still insufficiently mapped and studied. High elevations in the west exhibit uplifted remnants of palaeo plains. Widely distributed exhumed palaeosurfaces are found in south and central Sweden and southern Norway. Our contribution is a compilation of previ ous and ongoing studies on the palaeorelief and the associated saprolite remnants.
T E C T O NIC S E T TIN G
The Scandinavian peninsula of Sweden and Norway is the western part of Fennoscandia, which consists
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
275
K.
276
Upper Vendian Palaeozoic
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Area with Lower Palaeozoic remnants
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Lidmar-Bergstrdm et al.
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Permian
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times throughout the Mesozoic and Cenozoic. In the North Sea and the Norwegian Sea the build-up of thick coarse clastic sediment sequences throughout the Triassic and Jurassic indicates significant weath ering and erosion on the Norwegian mainland (e.g. Gabrielsen et at. 1990; Steel 1993). During the Creta ceous the offshore basins subsided and the sediment influx decreased, indicating more quiet conditions on the mainland (Rundberg & Smalley 1989). The edges of the offshore basins substantially overstepped the present-day coastline of Norway (e.g. Sigmond et at. 1 984; Ziegler 1990; Dare 1991; Lidmar-Bergstrom 1 995; Riis 1996) to an extent that is not yet clear. The Scandinavian Mountains were created largely by uplift and warping of an old landsurface in the late Tertiary (Reusch 1901a; De Geer 1 910; Ahlmann
Fig. I. Tectonic setting of Scandinavia, compiled from different sources. SSD = South Swedish Dome.
1 919; Rohrman et at. 1995) and by its subsequent dis section by large valley systems (Ahlmann 1919; Peulvast 1985a). From late Oligocene and onwards the sedimentary supply to the offshore basins increased in response to erosion caused by the uplift (e.g. Stuevold et at. 1992). The base of the Neogene wedge is Rb-Sr dated to 33 Ma (Rundberg & Smalley 1989). The Vattern graben, a late Proterozoic feature (Kumpulainen & Nystuen 1 985), may have been reactivated in connection with uplift of central south Sweden during the Tertiary (Lidmar Bergstrom 1996). The almost complete absence of Mesozoic and Cenozoic sedimentary rocks onshore together with the existence of large volumes of Mesozoic and Cenozoic sediments along most of the Norwegian
Retieffeatures and pataeoweathering Margin are considered to be the result of several phases of erosion caused by regional uplift of Scandi navia and parts of the continental shelf (0. Holtedahl 1953; Rokoengen & R0nningsland 1983; Peulvast 1985a; Rundberg 1990; Green 1991; Dore 1992; Riis & Fjeldskar 1992; Jensen & Schmidt 1993; Stuevold & Eldholm 1996). The correlation between on- and off shore geology has the potential to improve under standing of the major denudation surfaces occurring on land in Scandinavia.
PALAE O C LIMATE AND C O N DITI O N S F O R D E EP W EATH E RI N G
Palaeogeographical reconstructions indicate that Scandinavia was at equatorial latitudes in early Cambrian times, south of the equator in middle Ordovician times, close to the equator from Silurian to Carboniferous times, before drifting northwards in Permian times, and reaching between 20°N and 35°N in early Triassic times (Toivakka 1995; based on Ziegler 1 978, 1988; Smith et at. 1981). In northwest Europe the Triassic climate was warm and dry, and from latest Rhaethian age through the rest of the Mesozoic Era it was humid tropical to subtropical (Brinkman 1969; Hallam 1 975; Frakes 1979; Lidmar Bergstrom 1982). This was a period for formation of deep kaolinitic saprolites (e.g. Stbrr 1975; Pulvertaft 1979; Lidmar-Bergstrom 1982;Thiry & Jacquin 1993). During the Tertiary Period Scandinavia experienced a variety of climatic conditions, including warm temperate, subtropical, tropical and cool conditions, as well as humid and arid periods (Spjeldmes 1975; Daley 1972; Buchardt 1978; Lidmar-Bergstrom 1982). The arid periods witnessed at least partial stripping of the old weathering mantles, which is obvious from the sedimentary record (e.g. Tank 1963; Spjeldmes 1975; Storr et at. 1977; Thiede et at. 1980). Thus for a long time Scandinavia was subject to weathering in tropical to subtropical and warm temperate climates, with alternating humid and arid periods (see Fig. 11). Cooling started in the Late Miocene (e.g. Hall 1985; Rundberg 1990). An age of 2.57 Ma is inferred for the onset of the major ice sheet expansion and small-scale glaciations may have occurred from 5.45 Ma (Jansen & Sj0holm 1991 ) . PALAE O R E L I E F
The palaeosurfaces of Scandinavia (Fig. 2) belong to two different categories, exhumed surfaces and sur-
277
faces in stepped sequences developed across uplifted basement areas. The exhumed surfaces have been identified by remnants of cover rocks and character istic saprolites. They are of two main types and ages, namely an extremely flat sub-Cambrian peneplain and hilly sub-Mesozoic etch-surfaces. The surfaces in Sweden have been mapped in comparatively great detail. The sub-Cambrian peneplain was recognized early by Hogbom (1910) and the first map of it was produced by Rudberg (1954) . The low South Swedish Dome has turned out to be an important key area for understanding relief evolution on a shield. It has Cambrian cover rocks directly on basement in the north and east and Mesozoic cover rocks directly on basement in the south and west (Lidmar-Bergstrom 1988b ). The mapping of its surfaces included three steps. First the surfaces with their characteristic relief and characteristic saprolites were identified in con nection with the cover rocks. Second the aerial dis tribution of different relief types and remnants of saprolites and cover rocks were analysed. Third the mapping was performed on contour maps by follow ing the characteristic relief in computer-drawn profiles out from the cover rocks. The exhumed surfaces are inclined and they end abruptly where they are cut by younger more horizontal surfaces. These latter surfaces were mapped by identifying the steps between them in the profiles and by following the steps in the contour maps. The mapping was extended to all of Sweden (Lidmar-Bergstrom 1994, 1996) . Surfaces in stepped sequences, with low inclination and without relation to cover rocks have since been identified in northern Sweden (e.g. Wrak 1908; Rudberg 1954), whereas the central parts are occupied by an undulating hilly relief (Rudberg 1960). The Norwegian parts have never been mapped in detail and only a rough outline is presented. The sub-Cambrian peneplain
The exhumed sub-Cambrian peneplain makes up the relief in large parts of south-eastern Scandinavia, along the Gulf of Bothnia, and in two areas along the Caledonian front in Norway (Fig. 2) (Rudberg 1954; 0. Holtedahl 1960a,b; Schipull 1974; Peulvast 1978; Sigmond et at. 1984; Lidmar-Bergstrom 1994, 1 996) . The sub-Cambrian peneplain rises to about 1300 m in south Norway. In its well preserved parts it is extremely flat with a relative relief of less than 20 m. Shallow kaolinitic saprolites are found associated with Cambrian covers in central Sweden (references in Elvhage & Lidmar-Bergstrom 1987; Lidmar Bergstrom et at. 1997) and are also reported from
1 2°
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..
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o
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Sub-Jurassic etchsurface Sub-Cambrian peneplain in summits Sub-Cambrian peneplain
Fig. 2. Generalized map of palaeosurfaces within Scandinavia (from Peulvast 1978; Rudberg 1984; Lidmar-Bergstri:im 1994) with weathering residues (compiled from different sources). SSP = South Smaland Peneplain. TZ = Tornquist Tectonic Zone.
Relieffeatures and palaeoweathering Finnmark, north Norway (A. Bj¢rlykke, personal communication). Sub-Mesozoic etch surfaces
Exhumed sub-Mesozoic etch surfaces with a relative relief up to 200 m are encountered in southernmost Sweden (Lidmar-Bergstrom 1982, 1988b, 1989, 1994, 1996). Most of them are sub-Cretaceous, but some are sub-Jurassic. They are associated with many rem nants of cover rocks and clayey saprolites up to 60 m thick (Fig. 2). The Mesozoic weathering front is exposed in an abandoned quarry on the Ivo island, where the Quaternary cover was first removed, then the Cretaceous limestone used, and thereafter the underlying kaolin exploited (Fig. 3). In this example it is well demonstrated how deep weathering and strip ping of the saprolite have shaped the steep slopes of the hills in the area. Within the exhumed sub-Meso zoic surfaces the present relief is virtually identical to the stripped weathering front. The deep weath ering and the subsequent stripping of saprolites have created landscapes with joint aligned valleys or undu lating hilly relief (Fig. 4), depending on the length of time the areas were exposed to the Mesozoic warm and wet climates (Lidmar-Bergstrom 1995). This relief type (Fig. 5) can be followed north wards along the west coast of Sweden into southern Norway and into central and north-eastern Sweden. It is here interpreted tentatively as belonging to the exhumed sub-Cretaceous etch surface, based on the relief and evidence of kaolinization (Lidmar Bergstrom 1995). The undulating hilly relief in the Trondheim area also is interpreted tentatively as a sub-Mesozoic landscape. This view is supported by the occurrence of a downfaulted Jurassic outlier (B¢e & Bjerkli 1989). Tertiary plains with residual hills
The exhumed old surfaces have become slightly tilted after formation and are truncated by subhorizontal plains. Such plains extend over large areas in north ernmost Sweden, where they are characterized by an abundance of residual hills (Rudberg 1988). They constitute two general levels about 300-400 m and 400-550 m a.s.l. In the literature they are known as the Muddus Plains (Wrak 1908; Lidmar-Bergstrom 1994, 1996). In south Sweden similar plains have developed after uplift and warping of the sub-Cambrian pene plain, which resulted in the South Swedish Dome
279
(SSD; Figs 2 & 5a & b; Lidmar-Bergstrom 1988b, 1993, 1996). The main surface, the South Smaland Peneplain (SSP), lies between 125 and 175 m a.s.l. and carries only few residual hills. The Muddus Plains and the SSP are considered to have developed during the Tertiary, as they truncate the exhumed sub-Mesozoic inclined hilly surfaces (Fig. 5; Lidmar-Bergstrom 1982, 1996). This observa tion suggests that Mesozoic surfaces have not been preserved intact unless they had a protective cover of sedimentary rocks until late in the Tertiary. Late Mesozoic(?) to Tertiary high plains of the Palaeic surface
Large parts of the Scandinavian Highlands are char acterized by fjeld plains, slightly undulating surfaces, with shallow valleys. They are the Palaeic surface of Reusch (1901a). A new interest in the Palaeic surface has arosen from all the new information about the offshore geology and a desire to correlate the Palaeic surface with the offshore sediments. As the notion 'Palaeic surface' is used differently by different authors it is somtimes difficult to compare the inter pretations. The word was first used by Reusch (1901a) to distinguish the high plains in southern Norway from the deeply incised valleys. As the interpretation of the true nature of the Palaeic surface is not yet conclusive, we present some of the different views. The Palaeic surface with a relief amplitude of more than 1500 m is regarded as one surface, which developed during the Cretaceous and early Tertiary by deep weathering and subse quent stripping (Gjessing 1967). Dore (1992) relates the Palaeic surface with the base Tertiary surface, whereas Stuevold & Eldholm (1996) consider it equi valent to the late Oligocene unconformity. The Palaeic surface was interpreted to be com posed of two different levels by Reusch (1901a), Str¢m (1948), and Peulvast (1985a). Str¢m (1948) suggested Miocene and Pliocene ages. Peulvast (1985a) regarded the summit surface with some high residuals (e.g. Jotunheimen) as the result of Late Palaeozoic-Mesozoic peneplanation and the lower generation to be the result of erosion initiated after a Palaeocene uplift, with possible development into the Neogene (Peulvast 1985a). From a generalized contour map of summit relief by Nesje (in Riis & Fjeldskar 1992) four main levels can be identified. A fission-track study suggests that the highest summits of the Palaeic surface might date back as far as the Cretaceous (Rohrman et al. 1995) and Riis (1996)
K.
280
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60 WNW 50
40
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- -
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Fig. 3. Profile in the quarry at lvd, south Sweden. The Quaternary cover, the Upper Cretaceous rocks and some residual kaolin have been removed. Sub-Cretaceous bedrock surfaces and the Mesozoic weathering front are exposed.
suggested even a Jurassic age for the summits, based on a correlation with the offshore sediments. Since uplift, the Palaeic surface has experienced continued denudation (etching, stripping and inci sion) during about 3 0 myr before the onset of glacia tions. It is commonly thought that the main forms are of early Tertiary age, although features in the surface forms and associated pre-glacial weathering residues could be much younger, particularly as the climate turned cooler and wetter in the Pliocene, which pro moted stripping and incision. Mountain glacia tions have destroyed the Palaeic surface in certain settings (Kleman & Stroeven 1997; Rudberg 1984). The Palaeic surface is separated from the lower sur faces in the east by a zone of incised valleys. The strandflat
A marked feature of the west and north-west coast of Norway is the low, uneven and partly submerged rock platform extending seawards up to 60 km from the coastal mountains. Where it is typically developed it consists of fiat forelands at the foot of steep mountain and hill sides, with a belt of thousands of low islands and skerries on the seaward side. It was termed the 'strandfiat' by Reusch (1894). He and others consid ered it to be a pre-glacial feature formed after uplift of the Palaeic surface. Ahlmann (191 9) regarded the strandfiat as a peripheral base-levelled plain, the last stage in the Davisian fluvial cycle, whereas Biidel ( 1977) argued that tropical -subtropical deep weath ering was an important process in its formation.
Nansen (1904, 1 922) regarded the strandfiat as a Qua ternary feature, the moulding of which was strongly influenced by frost processes. He stressed the impor tance of a glacially dissected coast as a prerequisite for strandfiat formation and was of the opinion that specially favourable conditions for its planation by shore erosion had been prevalent during the cold periods preceding glaciations. 0. Holtedahl (1929) and H. Holtedahl (1958, 1960) favoured a glacial origin. Larsen & Holtedahl (1985) consider the strandfiat to have formed during the last 2.5 Ma, and the main process for its formation to be frost-shatter ing in combination with sea-ice transportation and planation during glacial stages. Peulvast (1985b) noted the importance of deep weathering for the for mation of the strandfiat in the Lofoten-Vesteralen area, but in contrast to Biidel (1977), who argued for tropical conditions, Peulvast concludes that the com position of the saprolites does not imply warmer con ditions than now. The saprolites are interpreted to be pre-Weichselian and maybe pre-glacial in origin (Peulvast 1985b ).
PALA E O W E AT H E RI N G
Palaeoweathering within formerly glaciated b asement areas
Nathorst (1879) discussed the possibility that the abundant lake basins in southern Sweden originated from glacial stripping of old weathering mantles. The
Relieffeatures and palaeoweathering
281
(b)
(a)
(c)
Fig. 4. Cretaceous palaeosurfaces and saprolites. (a) Exhumed sub-Cretaceous hilly relief, south-east Sweden, with Ivii island in the background. The quarry is marked. (b) Ivii quarry. The steep lower part of the rock is stripped from its Mesozoic saprolite in parts of the quarry. Before they were quarried the Cretaceous strata rested directly on fresh basement on the top of the rock. Location see Fig. 3. (c) Dalhejaberg, south-east Sweden, an exhumed sub Cretaceous hill with the characteristically steep slopes. (d) Dalhejaberg. Note the weathered fractures, which contain remnants of a kaolinitic saprolite.
idea originated from Pumpelly ( 1879), who had noted the existence of weathering residues on the Laurent ian shield. In 1898 Chalmers wrote: 'That they (the sedentary beds the saprolites) form a very im portant member of the superficial deposits in the glaciated areas of Eastern Canada at least, and one from which the bowlder-clay and all the other overly=
(d)
ing stratified deposits have been mainly derived, no geologist will now attempt to deny.' The erosive effect of glaciers was later stressed by other researchers, e.g. Shepard ( 1937), and the interest in the dynamics of glaciers during the 1950s and onwards contributed to the ignorance of saprolite remnants in formerly glaciated areas for a long time. Mattsson (1962),
282 A
K. Lidmar-Bergstrom et al. w
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300
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Fig. 5 . Profiles across south Sweden showing the relationship between the sub-Cambrian peneplain, Mesozoic etch surfaces, and subhorizontal Tertiary plains. (A) East-west profile across the South Swedish Dome (SSD). (B) North-south profile across the SSD. (C) East-west profile from the sub-Cambrian peneplain on to the undulating hilly relief, interpreted tentatively as a Mesozoic etch surface. For location see Fig. 2.
however, gave many examples on deep weathering in Sweden and pointed out that weathering residues are more abundant than commonly thought. During the 1970s and 1980s several pre-Wisconsinan saprolites were discovered also in eastern Canada (see refer ences in Godard 1989). Later, field meetings in Finland (Fogelberg 1985) and Sweden (Lidmar Bergstrom 1988a) put the saprolite remnants of Fennoscandia again into focus. Works in northern Europe and Canada by a French research group (Godard 1989) and in Scotland by Hall (e.g. 1986) have discussed the weathered mantles over basement rocks of high latitudes. Extension of weathering remnants in Fennoscandia
Compilations of known weathering residues have been presented previously by Lundqvist (1985) and Elvhage & Lidmar-Bergstrom (1987) made a litera ture inventory of sites with saprolite remnants in Sweden. Figure 2 shows the distribution of known saprolite remnants on the different palaeosurfaces. Shallow kaolinitic saprolites occur below Cambrian covers (Fig. 1 ) . Remnants of thick kaolinitic sapro lites occur in connection with Cretaceous and Juras-
sic cover rocks in south Sweden and in combination with the undulating hilly relief extending out from these covers (Fig. 1) (Lidmar-Bergstrom 1989). Some sites with kaolinitic saprolites in Norway, e.g. Tangen, south-east Norway (Follestad 1974) and Sunnm0re, western Norway (Dahl 1954) are probably also part of this Mesozoic weathering and some of the sapro lites labelled 'clayey/sandy saprolites in general' in south Sweden might belong to the root zone of the Mesozoic saprolites. A few kaolinitic saprolites are encountered within the Palaeic surface (e.g. Gjems 1963). In addition, kaolinitic and smectitic saprolites are reported from Norway by Reusch (1901b, 1903), Rosenqvist (1952) and Bergseth et al. (1980). Kaolin ized basement is also encountered below Jurassic sedimentary strata offshore Norway. Kaolinization along fracture zones is described from some sites in Sweden (Ljunggren 1955; Frietsch 1960). This type of kaolinite occurrence has, without supporting evidence, been described in a routine fashion as hydrothermally formed, but formation simply by weathering processes caused by groundwater circula tion cannot be excluded. A special group of weathering residues are the so-called soft ores (Geijer & Magnusson 1926, 1944;
Relieffeatures and palaeoweathering Lundqvist 1985) or earthy ores (Vivallo & Broman 1993). The iron ore bodies or zones of weakness close to them are weathered to a depth of at least 200 m. The ores consist of complex masses of mainly iron oxides/hydroxides and carbonates, with kaolinite usually present and sometimes also smectite. Also the surrounding rocks are weathered to kaolin rich masses. Concerning the soft or earthy ore at Garpenberg, Vivallo & Broman ( 1993) conclude: 'It was formed long after the primary ore deposition and clearly postdate the Svecokarelian orogeny (2000-1750 Ma). The alteration of the sulphide ore occurred as a nearsurface process at low temperature and pressure. It was caused by ground water circula tion through the orebodies, channelled by faults and fractures.The water reacted with the iron sulphides in the ores and produced an initially acid solution which altered the host-rock silicates into clay minerals.' The ore district of the Svekocarelian province extends eastwards to Harcas, where the bedrock is cut by the sub-Cambrian peneplain. The ore bodies are not weathered to soft ores. Therefore it is suggested that the weathering did not occur until the former cover of Lower Palaeozoic rocks was eroded away, which might have occurred in the Late Mesozoic (Lidmar Bergstrom 1995). A weathering breccia with limonite within the Kristineberg ore district is described by Grip (1944) and weathering in copper ores in Finn mark, north Norway is described by Gjelsvik (1956). Kaolinitic weathering residues are found in frac tures in the Stockholm area. Here j oint-aligned valleys are etched out from the sub-Cambrian pene plain, which still is recognizable in the summits. It has been suggested that this relief was formed in the Late Cretaceous and caused by uplift, erosion of the Palaeozoic cover, and subsequent deep weath ering (Lidmar-Bergstrom 1995). Some tills in this area contain high amounts of kaolin (see below). Saprolites, consisting mainly of gravel-sized frag ments and with a very low clay content, are encoun tered in many places in Scandinavia. In south east Sweden they are comparatively well mapped (Lidmar-Bergstrom et a/. 1 997). The situation in other parts of Scandinavia is much different. Information on such weathering residues is found in many of the geological map descriptions and other papers (e.g. G. Lundqvist 195 1 ; Mattsson 1962; J. Lundqvist 1969, 1987; Hillefors 1985). Internordic projects performed by the Geological Surveys of Finland, Sweden and Norway have revealed many sites with gravelly saprolites (Hirvas et al. 1988). The sites in northern central Sweden are always covered by glacial
283
deposits and the depth of weathering amounts to at least 5-6 m (M. Sund, Geological Survey of Sweden, personal communication). The gravelly saprolites are widespread but in five areas they seem to be more abundant, namely south east Sweden, central Sweden, and two areas in north Sweden (areas around Lycksele and areas in the far north). Some types of gravelly saprolites in the Gate borg area on the Swedish west coast, and on the east coast of central Sweden are ascribed to particular characteristics of the bedrock (Samuelsson 1973). The Scandinavian block fields with depths up to slightly over 1 m are produced mainly by frost shat tering and frost heaving (Stromquist 1973; Malm strom & Palmer 1984). In recent works they are considered to pre-date at least the Late Weichselian (Nesje et al. 1988; Kleman 1994). It also has been suggested that block fields with tor-like forms and subrounded boulders originate from pre-glacial saprolites (Reusch 1 878; Hoverman 1949; E. Dahl 1961; Roaldset 1978; Malmstrom & Palmer 1 984; Nesje et al. 1988). Laboratory analyses of interstitial fines from boulderfields in north Norway were per formed by Rea et al. (1996). They found very high amounts of clay and silt, 30-70% , and classified the samples as clayey gruss. Clay minerals such as vermi culite, chlorite, kaolinite and gibbsite were identified. The fines were regarded as frost-sorted residues of an original weathering profile and an origin in Tertiary climates for the chemical weathering was proposed. Characteristics of saprolites
It is difficult to summarize the characteristics of saprolites from the literature because the methods used are different and the reports are of greatly dif ferent age. An integrated study of saprolites from the different surfaces on the flanks of the South Swedish Dome therefore has been performed. Samples from the following types of saprolites were collected and analysed: 1 clayey-sandy saprolites from below Cambrian and Mesozoic covers; 2 clayey-sandy saprolites from exhumed Mesozoic etch surfaces; 3 gravelly saprolites. Laboratory analyses included grain-size analysis according to standard methods, SEM studies of the surface texture of quartz grains, and X-ray diffraction analysis of the clay (< 2 Jlm) and, in some cases, silt fraction (60-2 Jlm). Semiquantitative estimates of clay minerals are based on peak areas. Detailed
284
K. Lidmar-Bergstrom et a!.
Table 1. (a) Mineralogical composition of the unweathered Vanga granite, Ivo (from Kornfalt & B ergstrom, 1990) and (b & c) chemical composition of the weathered granite, Ivo (C. Bristow, personal communication)
(b) Chemical composition of samples < 5 Jlm (a) Minerals (% vol.) Quartz Plagioclase, including sericite Potassium feldspar Biotite Muscovite Chlorite Epidote Phrenite Pumpellyite Titanite Zirkon Apatite Fluorite Topase Calcite Opaques
25 17 51 5 +
+
Si02 Alz03 Fe203 Ti02 CaO MgO KzO Na20 LOI
Grey
Red
48 37 1.1 0.03 0.22 0.08 0.79 O.Q7 13.1
46.3 35.0 2.73 0.2 0.16 0.49 0.88 0.12 14.1
(c) Modal composition
Kaolinite Mica Quartz Montmorillonite
Grey
Red
88 11
88 10
1
2
+ + 1 + +
descriptions of the methods are given in Lidmar Bergstrom et at. (1997). In the following account of Scandinavian saprolites those from south Sweden are in separate sections. A sub-Cambrian saprolite from south Sweden Despite an almost complete argillization of the feldspars, the deep-weathered gneiss at Lugnas (Fig. 2) generally is coherent owing to cementation by silica, calcite and iron oxides, which has prevented disintegration of the rock. The fine fractions of the altered rock are dominated by kaolinite of a fairly well-ordered type (Fig. 6:1) but also include minor amounts of illite and smectite-dominated mixed layered minerals. Cemented lumps of kaolinitic material are incorporated in the overlying Cambrian arkose, together with pebbles of unaltered gneiss (Hadding 1929), which suggest that a phase of sil crete/ferricrete formation preceded the Cambrian transgression. Sub-Mesozoic sapralites ofsouth Sweden Samples were collected of saprolites in granitoid rocks from below Jurassic and Upper Cretaceous cover rocks in Scania, south Sweden (Fig. 2: Ivo).
These saprolites typically consist of massive, soft bodies with approximately equal proportions of sand (2-0.06 mm) and silt (0.06-0.002 mm) and 20-30% clay. Macroscopic grains of any remnant minerals other than quartz are sparse. Several features in the shape and surface texture of the quartz grains (subrounded grains, dulled surfaces from dissolution/reprecipitation of silica, orientated V-shaped etch pits, etc.) indicate intense chemical etching (Lidmar-Bergstrom et at. 1997). Chemical data on the weathered granites (Table 1) display a considerable cation depletion (Ca, K and Na), and XRD analysis confirms that both the plagioclase and the K-feldspar of the parent rock (originally c. 70%) have been replaced more or less completely by kaolinite. The modal composition of the kaolinized rock (material < 0.005 mm) corresponds to a kaoli nite content between 80 and 90% . In addition, the fine fraction generally contains 2 : 1 phyllosilicates (mica/illite and/or expansible, irregularly interstrati fied clay minerals) as minor constituents (Fig. 6:2). The very high ratio (2 : 1) of kaolinite to phyllosili cates shows that the sub-Mesozoic and older sapro lite remnants have reached an advanced stage in the evolution towards a single-phase assemblage of kaolinite, which generally develops during the 'final' stage of weathering, virtually irrespective of
Relieffeatures and palaeoweathering
285
1 AD I G
4
2 D
I
3.5
I
4.26
A
Fig. 6. Characteristic X-ray diffractograms of the < 2 �m fraction from saprolite samples of different origin. For location of sites see Fig. 2. (1) (Lugnas), kaolinite-rich saprolite from a sub-Cambrian site; (2) (Ivo), kaolinite-rich saprolite with some mica from a sub-Cretaceous site; (3) (Snallerod), smectite-kaolinite-rich saprolite close to the weathering front from an exhumed sub-Jurassic site; ( 4) (Hunnebostrand), kaolinite-smectite-rich sample with some haematite (peak at 2.69 A) from a fracture zone within the exhumed sub-Cretaceous relief; (5) (Knasekarret), an immature gravelly saprolite with a mixture of clay minerals (kaolin minerals, vermiculite, illite) and also quartz and feldspars. AD = air-dried; EG = ethylene glycol solvated; K = potassium saturated; HCl = digested in hot 2M HCI. CuK"-radiation. Samples 4 and 5 scanned on a goniometer with an automatic slit.
286
K. Lidmar-Bergstrom et a!.
parent-rock composition (review in Hall et al. 1989; Weaver 1989; StOrr 1993). However, a certain com positional variability, which probably reflects a verti cal zoning in weathering regime (Velde & Meunier 1987; StOrr 1993), can be found among stratigra phically and geographically connected sites. This is exemplified by the sample from Snallerod (Fig. 2), taken at a position closer to the weathering front. Although friable and disintegrated, the bedrock here has the primary structures of the parent rocks preserved. The decomposed rock material is coarse grained with fairly solid pebble-sized fragments in a gravelly-sandy matrix. Remnant feldspars and quartz are common, as they are also in the finer fractions. The clay fraction of such samples is often smectite-dominated, has kaolinite as the second most abundant mineral and illite as a minor component (Fig. 6:3) . The smectite is of a type that has a high chemical stability, as is suggested by its high resis tance in hot 2 M HCL Impeded drainage conditions in the relatively rigid structure of the relict rock could explain that positions close to the weathering front represent environments where smectite is one of the stable products of weathering. Kaolinized fracture zones and pockets with clay weathering within the sub-Cretaceous relief of south Sweden Argillized rock can be found in depressions and fracture zones in the bedrock along the western (Halland, Bohuslan) and south-eastern (Blekinge) coasts of Sweden (Fig. 2), even in positions below the Weichselian highest shore-line. In granitic rocks these fracture fillings consist of friable, white-mottled gravel- and pebble-sized rock fragments in a sandy matrix, with a clay content ranging from very low up to 8-10 % . Remnant quartz and feldspars are common in the coarser (> 0.06mm) fractions. The SEM data available for one of the sites show that quartz grains from the matrix are subrounded, with a high frequency of surface textures indicative of intense chemical etching (Lidmar-Bergstrom et al. 1 997). Fine fractions generally are dominated by kaolinite (Fig. 6:4) but clay mineral assemblages can be somewhat variable, with smectite sometimes being more abundant than kaolinite. The variability may simply reflect different weathering regimes at the weathering front of formerly more extensive saprolite bodies, and the available data give little support for distinguishing these weathering residues from the previous group.
Sub-Mesozoic saprolites in Norway and offshore Unconformably beneath Mid-Jurassic and probably also older sediments at And¢ya (Fig. 2) a 32 m-deep fossil tropical weathering profile is preserved in a downfaulted position (Dalland 1974, 1 975; Sturt et al. 1 979) . The weathered zone consists mainly of kaolin ite and quartz. A high feldspar content at the base disappears upwards. The clay fraction contains up to 90 % kaolinite, with a small amount of illite. The clay mineral assemblage thus resembles that of the Swedish sub-Mesozoic saprolites. The weathering profile is unconformably overlain by a sandy lime stone of unknown age, with sandstones and shales of Mid-Jurassic age above. The information from the offshore exploration in the northern North Sea indirectly relates the depo sition of the shelf sediments to the contempora neous weathering and erosion processes of western Norway. Several exploration wells drilled in the northern North Sea and Norwegian Sea have pene trated Jurassic sediments resting on weathered crys talline basement. The Jurassic Froan basin (Fig. 2) is a downfaulted graben structure off the present coast in the M¢re-Tr¢ndelag Fault zone. In this basin clays with extremely high contents of kaolinite, commonly as aggregates (up to 80% in the bulk sample), strongly indicate deposition close to a deeply weathered land surface. Microtextural studies suggest that the kao linite has been transported as sand-sized aggregates. The high kaolinite content may suggest short trans port and deposition in shallow marine conditions or a marsh-coastal plain environment. A weathering profile located in quadrant 35, south west ofVags¢y (Fig. 2), not too far off the Stad penin sula (see below), appears to be overlain by Lower Jurassic sediments (Dunlin Group) (Riis 1 993, 1996). Samples of weathered crystalline basement rocks and overlying sediments in well 35/9-1 have been studied (Riis 1 993, 1996; Roaldset et a/. 1993). The weathering profile is only cored in the upper parts, as the sonic and resistivity logs indicate the weathered zone to be about 12-1 5 m thick. The altered basement rock, which originally was an amphibolite, or pos sibly greenschist, contained considerable amounts of kaolinite, smectite, quartz and albite, however, no gibbsite could be observed. The feldspars and amphi boles have weathered to kaolinite and smectite. After deposition of marine sediments above, calcite pre cipitated in fissures and fractures of the weathering profile.
Relieffeatures and palaeoweathering Clayey and sandy saprolites in Norway The weathering remnants may consists of gravelly sand and whitish or yellowish to red clayey material, rich in quartz, smectite and hydromicas, but also kaolinite, aluminium- and ferric oxides and/or hydroxides and siderite (e.g. Goldschmidt 1928; Barth 1939; Isachsen & Rosenqvist 1 949; Lag 1963; Englund & J¢rgensen 1975; Roaldset et al. 1982). At Kvitebekk (White Creek), Seljord, south-east Norway (Fig. 2) a zone with a whitish clay formed from Precambrian metasediments has been known for a long time. The weathered material consists almost exclusively of kaolinite and quartz with traces of haematite and sericitic illite. The fraction < 2 f.J.m consists exclusively of kaolinite and illite with a chemical composition close to kaolinite (Si02 = 46.3 % , Al2 03 = 35.4% , K20 = 0.52% ) . I n the north-western part o f south Norway, weath ered crystalline basement rocks are preserved below the late Quaternary till (Longva & Larsen 1979). At Stad (Fig. 2) a weathering profile 2-3 m thick is over lain by basal tills and solifluction deposits (Roaldset et al. 1982). The profile has developed on a palaeo surface, which today lies 400-450 m a.s.l. The upper layers contain vemiculite and illitic minerals, gibbsite, goethite, K-feldspar, plagioclase, quartz and amphi bole, indicating that the deposit is a mixture of glacially abraded-till material mixed with weathered material by solifluction. The saprolite is characterized by high amounts of gibbsite, some quartz, minor amounts of goethite, smectite and illite. The silt frac tion of the weathered granitic gneiss contains up to 45 % gibbsite and the clay fraction up to 90% . The weathering remnants at Stad have the characteristics of tropical palaeosols and bauxite. At Vags¢y just south of Stad, weathered gabbro with corestones is exposed below till. The weathering is more than 5 m deep and exhibits a gradual transi tion into unweathered rock. The section is located about 430 m a.s.l. (Roaldset et al. 1982). The clay frac tion consists mainly of vermiculite and illitic miner als, smectite and trace amounts of plagioclase, amphibole and quartz. The silt fraction contains the same minerals except for higher contents of plagio clase, quartz and amphibole. At Tingvoll, north-east of Stad and Vags¢y, rem nants of Precambrian gneiss weathered to soft clay occur below till. The clay is almost monomineralic, consisting of an aluminium rich smectite of the montmorillinite-beidellite type, with some K and only traces of Na and Fe.
287
The age of these saprolites is not clear. Favourable conditions for lateritic/bauxitic weathering have not been available since Miocene times (Roaldset et al. 1982), which indicates that the Stad profile is at least as old as that. If the Vags¢y profile corresponds in age with the Stad profile, it may represent the deeper part of a weathering profile, from which the possibly more kaolinitic upper parts have been removed. The Tingvoll profile has not yet reached the kaolinite stage. There are two possibilites for the age of these saprolites on the west coast of Norway. They can belong to either an exhumed sub-Mesozoic surface or the profiles were formed after exhumation of the Mesozoic cover rocks, following uplift in the late Oligocene Epoch, but before the onset of the cold climates. Gravelly saprolites on the South Swedish Dome, south-eastern Sweden In some inland areas on the South Swedish Dome in south-eastern Sweden (Smaland) glacial erosion was limited during the Weichselian glaciation because glaciers were mainly cold-based (Lagerlund 1987). Saprolite remnants are comparatively widespread within this region, as demonstrated by a recent mapping, which has documented more than 35 sites. Nine of the sites have been analysed in detail (Lidmar-Bergstrom et al. 1997). The saprolites have thicknesses varying from < 0.5 m to > 10 m and are developed in plutonic and volcanic parent rocks, ranging in composition from basic to acid. Although most of the weathered rocks are highly friable and more or less completely disintegrated, macrostruc tures, such as banding and dykes of mafic rocks, are preserved in bedrock exposures, confirming the in situ position of the weathered material. When devel oped in granitoid rocks, the saprolite material con sists of rounded core stones with concentric surface layers and gravel-size angular rock fragments in a reddish brown, sandy-silty matrix. The matrix derives its colour from grain coatings of secondary iron oxides and/or oxyhydroxides. Between 60 and 85% of the material is > 1 mm, and the clay content seldom exceeds 5 % . Profile studies in one of the thicker (> l O rn) saprolites show that neither the grain-size distribution nor the mineralogy varies much with depth. At all the sites the cover units consist of thin Weichselian glacigenic deposits only, with thick nesses seldom exceeding 2 m. The gravelly saprolites studied are developed in granites and granodiorites. They form a rather
288
K. Lidmar-Bergstrom et al.
heterogeneous group with respect to their clay min eralogy but all are multiphase associations (Fig. 6:5). Remnant quartz and feldspars are ubiquitous in size fractions > 2 !liD. Quartz grains are angular subangular and the surface has few chemically pro duced features but a high frequency of conchoidal fractures, arcuate steps and breakage blocks (Lidmar-Bergstrom et a/. 1997), i.e. features that indi cate mainly mechanical breakage of the rock. Optical studies of macroscopic biotite flakes show a decrease in refractive index with grain-size and a colour change from greenish brown to pale brownish yellow. Silt-sized biotite flakes are strongly bleached and have a brassy submetallic lustre. The XRD analyses of fine fractions show that one or other of the follow ing phases is predominant among the phyllosilicates: vermiculite, low-charge vermiculite, interstratified vermiculite-smectite or smectite. The ease with which these secondary phases dissolve in acids sug gests that they are all Fe-rich (trioctahedral) minerals of low chemical stability. They most likely formed as interim products in the continuous weathering of Fe-mica to smectite, which starts with the release of K and decrease in layer charge by oxidation and loss of structural Fe. The formation of grain coatings of Fe-oxides, such as haematite, lepi docrocite and goethite, probably was associated with and is a result of the vermiculitization of the Fe bearing micas. Hydrological factors may explain that at one third of the sites investigated, predominantly those situated on hill crests, the association of secondary minerals also includes kaolin minerals, which may contribute as much as 40% of the clay mineral content in the < 2 !lm fraction. Diffractogram 5 (Fig. 6) is an example of the mineralogy of deeply weath ered rock in such a position (Fig. 2: Knasekarret). Kaolin minerals occur also in the coarser fractions, as a 'powder' along the contact surfaces in multiphase aggregates, but kaolin is still a minor constituent (1-5 % ) of the bulk samples. The expansion behav iour of the kaolin minerals on formamide treatment (Churchman et al. 1984) suggests that both 7 A hal loysite and kaolinite may occur. Gravelly sapralites in Norway Several generations of palaeosols within Quaternary strata have been found in Finnmark, northern Norway (Fig. 2; Hirvas eta/. 1988; Olsen 1995; Olsen et al. 1996). The lowermost palaeosol is developed directly on the weathered Precambrian bedrock and
is kaolinitic. Olsen (1995) suggests the weathering in the bedrock to be of Tertiary age. Deeply weathered rocks are widespread also in the Lofoten-VestenUen area (Vogt 1912; Rekstad 1 915). The relationship between morphology and saprolites has been studied by Peulvast (1978, 1985b, 1 989) . The main morphological features are peaks between 600 and 1200 m, plateaus between 300 and 400 m, valleys and wide basins, and the strandflat. Saprolites 0.2-0.8 m thick are encountered on the plateaus. In lower areas saprolites are between 4 and 6 m thick and saprolites over 6 m thick occur on the strandflat. They are situated in areas where glacial erosion has been weak. The saprolites are gravelly or sandy with a silt content of less than 5 % and only traces of clay material (0-1.9%). The samples contain hydro biotite, illite, vemiculite, some smectite and in one sample traces of kaolinite. The clay minerals were believed to have been produced mainly by the alter ation of biotite. The saprolites are clearly pre-Weich selian and maybe Pliocene in age (Peulvast 1 985b, 1989). S0rensen (1988) reported sites close to the coast in Vestfold, south-east Norway, with intensive disinte gration of the rock. The weathering was assumed to be related principally to the microtexture and miner alogy of the Permian magmatic rock. The saprolites contain less than 4% silt and clay, 30% gravel, and the remainder sand. Vermiculite, illite, chlorite and smec tite were identified by X-ray diffraction. S0rensen (1988) did not suggest any particular age for these saprolites, but acknowledged their similarity with the saprolites described from Lofoten-Vesteril.len. Landforms and gravelly saprolites
In positions where glacial erosion has been limited, tors are left after stripping of gravelly saprolites, particularly in sheeted and fractured granites. This phenomenon is not very well documented but is described from northern Sweden (Fig. 2: Lycksele; Ivarsson & Zale 1 989) and also occurs in south-east Sweden (Fig. 7a & b). Dahl (1966) regarded the tor like forms in the Narvik mountains to be entirely post-glacial in origin, whereas Kleman & Stroeven (1997) conclude that such features were preserved below ice sheets frozen to the ground and thus are pre-glacial. The weathering into gravel and the subse quent stripping may, however, have been a continous process during the ice-free periods of the Pliocene and Pleistocene (Peulvast 1985a; Lidmar-Bergstrom et a/. 1 997) and these forms may, in their details there-
Relieffeatures and palaeoweathering
289
(a )
Fig. 7. Granitic tor with residual gravelly weathering along fractures and sheet planes, near Malexander south-east Sweden. (a) Although overriden by ice the shape of the hill is governed by the fractures. (b) Weathering along sheet planes and incipient core stones.
fore, not totally pre-date the Quaternary but are cer tainly non-glacial features that pre-date overriding by the last ice sheet. The Revsund granitic area, central Sweden (Fig. 2), is in a border region between the undulating hilly relief and the Muddus plains. Landforms controlled by granular weathering are preserved. Gravelly saprolites with core stones are of common occur rence (Lundqvist 1988). The saprolites commonly are overlain by Weichselian strata and are at least of Eemian age but probably much older. Rock surfaces with glacial striae and overlain by till are still fresh and thus little weathering has occurred during the Holocene. Lundqvist ( 1988) states that it is improb-
(b)
able that the saprolites were produced during the rel atively short Eemian interglacial and concludes that they must be older.
C O N T RI B UTI O N O F P R E - G LACIAL W EATHE RIN G C O MP O N E N T S T O G LACIAL D EP O SITS
The majority of the loam and clay sediments within the crystalline bedrock regions of Scandinavia was deposited during the last 12 000-10 000yr. The mineralogy of these deposits is typically dominated by rock-forming minerals, such as illite, chlorite,
290
K. Lidmar-Bergstrom et al.
feldspars and quartz (e.g. Snail et al. 1979; Brusewitz 1 982; Olsson 1991). It was therefore generally con sidered that the material for these sediments was derived from the granitic-gneissic bedrock, mainly by glacial comminution during the last glaciation followed only by subsequent erosion, sorting and redeposition. Chemical and mineralogical alterations were regarded as unimportant. This was the pre vailing view for a long time, until detailed modern studies, combining mineralogical and chemical analy ses, showed that the pre-weathered constituents may make up a substantial fraction of the otherwise unweathered Weichselian deposits. Examples from Sweden
Reports on pre-weathered constituents in the unweathered Weichselian deposits in Sweden are not infrequent (Collini 1 956; Rosenqvist 1975a,b; Snail et al. 1979; Brusewitz 1982; Olsson 1 991; Stevens & Bayard 1 994). Of special interest is the admixture of kaolin minerals in these sediments, because condi tions for kaolin formation have been uncommon at high latitudes during the Pleistocene and Holocene Epochs (Wilson et al. 1984; Weaver 1 989) . It is there fore normally assumed that kaolin minerals in Pleistocene sediments are inherited from an ancient regolith, and this must certainly be suspected in regions where deep-weathered bedrock is common today. Situations where the mineralogical evidence is in accord with these assumptions are easy to find within the area of the sub-Cambrian peneplain in south ern central Sweden, where kaolinite is a ubiquitous, although generally minor, component of the Weich selian, glacigenic sediments (e.g. Olsson 1991; Stevens & Bayard 1 994). Diffractogram 1 (Fig. 8) shows the kaolinite-rich < 2 f..Lm fraction of a clayey silt, which was deposited during the retreat of the Weichselian ice in the near-shore environment of a glacial basin, situated 1 km south of the sub-Cambrian saprolite at Lugnas (Fig. 2). Chlorite and clay mica tend to increase in abundance with increasing distance from the saprolite, whereas kaolinite decreases (Fig. 8:2) . The compositional trend might be attributable to size-sorting, but also to the diminished influence of minerogenic matter of local provenance in favour of material dominated by suspended matter, the prov enance of which is mixed. Kaolinite is ubiquitous in the glacial sediments of southern Sweden, where kaolinitic saprolites as well as kaolinitic Palaeozoic and Mesozoic cover rocks
are quite common. Kaolinite, however, also can be quite a significant component of Weichselian sedi ments in regions where no source rocks for kaolinite are known to exist today and in situations where kaolin formation through Holocene pedogenesis can be excluded. Snail et al. ( 1979) investigated subsoil samples of tills within a 600-km2 area in eastern central Sweden, with a predominantly granitic-gneis sic bedrock composition (Fig. 2: Katrineholm). The kaolinite content of the < 2 f..Lm fraction varied from > 50-0 % . Although variable bedrock composition in the direction of the last ice movement seemed to exert strong influence on the type and abundance of other phyllosilicates, it did not explain the variable kaolinite content. This was therefore believed to be determined by inheritance from a former, kaolinitic regolith (see above). The distribution of kaolin minerals across the till-saprolite boundary was examined at three of the sites with gravelly saprolites in south-eastern Sweden. Gross mineralogical trends are discontinu ous over the saprolite-till boundaries and the strati graphical distribution of kaolin and Fe-chlorite can be described as a reversed 'evolutionary trend' (Fig. 8:3). Fe-chlorite, which is highly susceptible to weath ering, is one of the major phyllosilicates in the till (and common also in the fresh bedrock), but has not been detected in the gravelly saprolite. The kaolin minerals have a contrasting distribution. These trends show that saprolite formation was a process distinct from Holocene soil formation. Examples from Numedalen, south Norway
The Numedalen project was launched with the spe cific aim to explain the origin and formation of the unconsolidated sediments in Norway (Rosen qvist, 1 975a,b ). In addition to references given below, results are published in Roaldset (1975, 1 979, 1980), Korb01 & J0rgensen (1973), Rueslatten (1976), R0n ningsland (1976), Ormaasen (1977) and Rueslatten & J0rgensen (1977). The drainage basin of the River Numedalslagen (Fig. 2) represents a typical Norwegian valley system and the direction of the main valley corresponds to the direction of the last regional ice movement (Vorren 1977; J0rgensen et al. 1977). Within the drainage basin various types of glacial, fluvial and marine sediments were deposited during and after the Late Weichselian glaciatjon. The post-glacial sedi ments are the result mainly of fluvial reworking of till, which is the main sediment in the drainage area.
291
Relieffeatures and palaeoweathering
1
3.5
4.26 5
7
10
14
A
3
3.5
4.26 5
7
10
14
A
Fig. 8. X-ray diffractograms of the < 2 f.tm fraction of some Weichselian sediments: (1) illite-kaolinite-rich glacial clayey silt 1 km south of Lugnas; (2) illite-chlorite-rich glacial clay 8 km south of Lugnas; (3), till, overlying gravelly saprolite at Knaseklirret (see Fig. 6: 5). GLY = glycerol solvated after Mg-saturation; other abbreviations as in Fig. 6.
Grain size, mineralogy and major elemental composition The sand and silt fractions of the tills of the Numedal valley are considerably richer in quartz than the bedrock below (Korb¢1 1972; Dekko 1 973; Lien 1973; Rosenqvist 1975a,b ) . Dekko (1973) found that sand in the various coarse deposits corresponds to a mixture of approximately half unweathered bedrock
and half pure quartz. It was considered that if quartz is residual from earlier weathering, then a minimal amount of bedrock equal to or greater than the volumes of the present unconsolidated deposits has been fully weathered and the residual material mixed with fresh rock during the Quaternary glaciations. The relatively low content of clay minerals in the tills indicates a considerable separation of minerals during the formation of the till. The content of
K.
292
Lidmar-Bergstrom et a!.
expanding minerals is low or absent in most clays (as also is the case for all other Pleistocene-Holocene clays in Norway). The mineralogical and chemical composition of the < 2 ).lm fraction differs from the crystalline bedrock, supporting the idea that the < 2 ).lm fraction cannot be derived from the bedrock by mechanical grinding processes alone (Roaldset 1 972, 1974; Rosenqvist 1975a). The clay minerals of the Quaternary deposits of the Numedal Valley have higher values of Al, Fe, Mg, K and loss-on-ignition, and lower values of Si and Na than the underlying Precambrian rocks (Table 2). They are considered to represent degraded, primary phyllosilicates of the ancient metamorphic rocks (Roaldset 1972, 1973a, 1978). Rare earth elements The distribution pattern of the rare earth elements (REE = yttrium and lanthanoides (Ln)) was also investigated within the Quaternary deposits of the Numedal Valley (Roaldset 1 970, 1 978; Roaldset & Rosenqvist 1971 a,b, 1973b; Rosenqvist 1975a). The REE content of tills is strongly enriched in the finest fraction and impoverished in the silt and sand fractions (Roaldset 1978). The REE content in the clay fraction of tills is four to five times higher than in the crystalline rocks. It has been calculated that the time needed to release such amount of ions from rock forming minerals would be at least 0.11 Myr. Because REE are specifically and strongly adsorbed by clay minerals, the REE content in the clays seems to be an indicator of the minimum amount of weathering. Seventeen clay till samples in the Numedal area gave an average of 527 p.p.m. total REE (:ELn=
437 p.p.m.). Up to 80-90% of total REE was in the exchange position (Table 3). The glacial and post glacial marine clays in the lower part of the valley had an average of 335 p.p.m. total REE (56 samples). After removal of adsorbed REE the average of clay tills and marine clays decreased to 168 p.p.m. (:ELn= 140 p.p.m.) (Table 3; Fig. 9). The Precambrian gneisses and granites in upper Numedalen have undoubtedly been exposed to glacial erosion, but some slightly weathered gneisses seem to represent deeper parts of a pre-glacial weathering profile. Where the rock shows signs of weathering, the content of REE is much higher in the mica fraction than in the rock itself. The extreme case represents a slightly weathered mylonite gneiss with an overall content of 525 p.p.m. total REE, whereas the light micas (degraded muscovite) contain 3755 p.p.m. and the dark micas (chlorite and vermiculized biotite) contain 1584 p.p.m. In another slightly weathered granitic rock the chlorite had 1 808 p.p.m. total REE (:ELn= 1254 p.p.m.), whereas the rock itself had 363 p.p.m. (:ELn = 336 p.p.m.). In these cases most of the REE in the micas was extractable. These rocks were assumed to represent deeper parts of the pre-glacial weathering profile (Roaldset & Rosenqvist 1971a,b ). Ln adsorption on clays is a highly pH-sensitive process and the adsorption is almost complete from neutral solutions, whereas the adsorbed Ln ions can be desorbed/extracted by coming into contact with more acid water (Brown et al. 1955; Amphlett 1958; Aagaard 1 973). Apparently when clays with high con centrations of REE are transported into marine water an ion exchange process and an adjustment towards the normal pattern of marine sediments takes place. Some distribution curves for the lan thanoide elements are shown in Fig. 9. The observed
Table 2. Average chemical composition of Numedal rocks and of fractions < 500 Jlm and <2 Jlm of Numedal tills (wt % ) (Data from Roaldset, 1972, and Rosenqvist, 1975a)
Si02 Ti02 Alz03 Fez03 (total) MgO CaO Na20 KzO LOI
Numedal rocks (149 samples)
Numedal tills (< 500)lm fraction)
Numedal tills (< 2 )lm fraction)
70.42 0.61 12.98 4.57 1.43 2.25 2.88 3.36 1.01
77.42 0.64 10.14 3.18 0.93 1.74 1.86 2.96 1.18
50.90 0.86 17.39 10.58 3.99 2.97 2.11 4.61 5.92
293
Relieffeatures and palaeoweathering N U M EDAL CLAYS
N U M E DAL ROCKS I M I N ERALS
Moraine clay
5
2
I.Ln
=
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20
437ppm
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EDTA treated clays
Lln
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Lln
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------------�--�--L a Ce Pr N d Pm Sm E u Gd T b Dy He Er T m Y b Lu 57 58 59
60
61
62 63 64
65 66
67 68 69 70
Rock-forming phyllosilicates I.Ln
Marine mud (E83-E85) Lln
0
71
Lanthanide
=
283ppm
5
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Precambrian basement
2
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=
336ppm
La Ce Pr Nd Pm Sm Eu Gd Tb Dy He Er Tm Yb Lu 57
atomic
58 59
60
61
62 63 64 65 66
67 68
69
70
71
number
Fig. 9. Lanthanide comparison diagram relative to the composite of North American Shales (Haskin et a/. 1966) for the clay fraction of tills, river water, marine mud, basement rocks and rock-forming phyllosilicates from the Numedal area, Norway (Roaldset 1978).
Ln variations within the Numedal area are con sistent with accumulation of the Y and Ln elements in hydrolysate sediments during subaerial weathering, and with subsequent leaching of surface-adsorbed ion by humic acids or acid river water. Mass balance calculations show that the minimum amount of adsorbed REE in the Pleistocene and Holocene sediments are at least 6 x 105 t (Roaldset & Rosenqvist 1 971a ) . More detailed sampling in 1971-1973 resulted in a calculation nearer to 2 x 106 tons, corresponding to an entire REE content of 1 x
1010 tons of average bedrock. This is the minimum amount of rock that has to be weathered in order to supply the REEs adsorbed on sediments inside the present shore line (Rosenqvist 1975a) . The present weathering rate for the Numedal watershed is 6 x 104 t of average rock per year. By taking into con sideration composition of unweathered rocks and the annual weathering rate for the Numedal area (Jensen 1972) , it was concluded that more than 100 000years would be needed to provide the adsorbed REE in these sediments. Based on additional mapping and
294
K. Lidmar-Bergstrom et al.
geophysical investigations of sediment thicknesses (e.g. Gr¢nlie & J ¢rgensen 1 974) the time-span for the wearthering was considered to be much longer (Rosenqvist 1 975a). Heavy minerals The comparison of the heavy mineral content of tills and crystalline bedrock can throw light on the inter mixing of pre-glacial deeply weathered material into the glacial sediments, as the heavy minerals are dif ferent with respect to weathering resistance. Among the least resistant minerals are amphiboles, garnets and apatite, and among the most resistant are
Table 3. Average contents (p. p.m.) of yttrium and lanthanides in untreated and EDTA-treated Numedal tills (fraction < 2 !-t.m) (Roaldset, 1973b, 1978) together with data for the crustal average (Rosier & Lange, 1965)
Numedal tills fraction < 2 J.Lm (average of 17 tills) y
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
87.5 77.6 202.3 19.8 57.7 19.4 4.8 19.3 3.0 23.7 3.9 12.9 * 16.0 *
Total
527
EDTA treated (average of 12 tills) 28.4 22.2 56.8 6.8 22.8 6.6 1.5 5.6 0.9 7.0 1.0 3.3 5.5 168
Crustal average 31 29 58 8.1 30 7.0 1 .25 7.8 1 .2 4.2 1 .4 3.4 0.52 3.0 0.65 186
* The element is present in small, but observable amounts.
haematite, magnetite, ilmenite, cromite, rutile and zircon (e.g. Valeton 1972; Grim 1973). Stiberg (1983) hypothesized that if the products of pre-glacial weathering are intermixed with glacial deposits, one could expect the content of heavy minerals to be reduced, the least stable often more than the others. The ferrimagnetic oxides, especially, were investi gated. Maghemite y-Fe203 may develop by alteration of titanomagnetite. A minimum age of 106 years was estimated for the complete oxidation of titanomag netite in sufficient amounts to account for the maghemite in the tills. In the areas with Precambrian bedrock the amount of zircon shows an increase in the silt fraction of tills relative to the bedrock. A minimum of 46% of all heavy minerals was estimated to have disappeared relative to the bedrock (Table 4). The heavy minerals and the content of Zr, Cr and V are closely related to the composition of underlying bedrock. In the upper part of the Numedal Valley, amphibole has been reduced the most, then epidote and iron minerals. Some of the resistant minerals, such as zircon have been enriched in the till relative to the underlying bedrock. The reduction in total heavy mineral content, the changes in heavy mineral assemblage relative to the bedrock, the increase in Zr and Cr, the absence of apatite and presence of maghemite, strongly in dicate that the glacial sediments include material formed under warm, maybe even tropical, climatic conditions. Conclusions ofthe Numedal study One of the main results was a verification of the assumption that the glacial deposits were a mixture of material produced by mechanical grinding of the ice sheet and pre-glacially weathered material (Fig. 10). For the finer silt and clay fractions, large differ ences were found relative to the underlying rocks. The observed mineralogical and chemical fractiona-
Table 4. Heavy minerals in Numedal bedrocks, and in the silt fraction of Numedal tills (wt%) (Data from Stiberg, 1983)
Bedrock Till (C horizon) Change Relative change ( % )*
Amphibole
Epidote
Ertz
Other
Heavy mineral content (total)
9.5 3.1 -6.4 -64
4.4 2.3 -2.1 -48
4.6 3.5 -0.5 -12.5
1.7 1.7 0 0
19.6 10.6 -9.0 -46
* Relative change adjusted for rock volume.
295
Relieffeatures and palaeoweathering
0 >
�
�
E -� :::J t .� � 0
Q)
N -
0 "0
"'Q) "-
::::!' ·E
Saprolite
Weathering i n warm and h u m i d cli mate
\ Tropical sub-tropical ' to temperate ' ' cli mate
Fluvial and eolian erosion
U n a ltered crysta l l i n e rock
3.2m .y.B.P.
I
Q) c
Repeated glaciali nterglacial periods
"0Q)
1i)
·;;;
I
Glacial erosion and transport
I
t
t
0:::
' 75000.y.B.P.
'
,
Fluvial erosion redeposition soil formation .......
_
_
_
I
_
_
......
/ .....
G l acial erosion ( m ech anical)
/
/
Weichsel glaciation 1 0 000-12 000.y.B.P. Fluvial erosion and redeposition soil formation Q) c Q)
"0
0 I
Post-glacial
Glaciofluvial and fluvial sand and G ravel deposits
l
Lacustrine silts and clays
Marine silts and clays
Fig. lO. Stages in the formation of the Scandinavian Quaternary deposits (modified from Roaldset 1978) illustrating the mixed origin of the Quaternary sediments from glacially abraded crystalline rocks mixed with pre-glacial weathering products.
tions can be explained only by the existence of deeply weathered material. It is difficult to quantify the amount of pre-glacial saprolitic material mixed into the Pleistocene tills. In the coarser fraction the content of weathered material is low, whereas in the silt and sand fractions it is up to 25 % , possibly even higher. In particular the presence of maghemite in the coarse fractions indicates that the glacial sedi ments in central Norway have incorporated a consid erable amount of pre-glacial saprolites. It was also concluded that the tills in the Numedal area were of local origin. The coarsest fractions of the tills are transported less than 5-10 km (Dekko 1973; Stiberg
1 983), whereas the finest clay fraction can be trans ported over long distances.
C O N C L USI O N S
Clayey, kaolinitic saprolites are associated with exhumed denudation surfaces; shallow saprolites with fiat sub-Cambrian surfaces and deep saprolites with hilly sub-Jurassic and sub-Cretaceous surfaces. Other saprolites are difficult to date. Clayey-sandy saprolites could date back to the early and middle Tertiary. Gravelly saprolites and associated tor forms
296
K. Lidmar-Bergstrom et al.
Relief, saprolites, and correlative sediments in southern Fennoscandia, Permian - Pleistocene Age 1 .7 5 24
37
55
66
Time
Pleistocene Pliocene Miocene Oligocene Eocene Paleocene
Cli mate humid arid
cold cool
humid
warm
arid
cool
humid
very warm
arid drier?
cool
Correlative sediments � chlorite � u illite o -o
-� � � lij
E "'
Type of saprolite
gravelly
weathering
smectitic illitic and
- - -? - - - - - - ? - - -
t
I
covered
)
cool
humid
141
t
all
warm Cretaceous
1
kaolinitic clays
limestone
kaolinitic
?
plains with residual hills
?
and chalk
tors
t
kaolinite illite
smectitic clays
Relief in basement
)
deep
etch
and
clayey
sur-
illitic
kaolinitic
faces
clays
tropical Jurassic
to
quartz
sand
sub
210 Triassic
250 Permian
291
) j
arid
tropical
smectitic
clays
r
1
shallow
arkoses
are ascribed mainly to a late Tertiary to Pleistocene age (Fig. 1 1 ) . Where a relationship to datable cover rocks does not exist it is only the characteristics of the saprolites that give some hint of the age. The block fields, which bear witness to severe periglacial processes but also to no glacial erosion, have in certain locations been interpreted as col lapsed saprolites. Their origin is still not fully under stood.
pediplains
j j
Fig. ll. Relief, sa pro lites and correlative sediments in southern Fennoscandia, Permian Pleistocene ( modified from Lidmar-Bergstrom et a/. 1996).
Within the Palaeic surface of Norway, remnants of more advanced weathering, gravelly weathering, and block fields occur, which a dresses the question of the relationship between denudation surfaces and sapro lites. On the exhumed surfaces it is easy to ascertain the type of saprolite associated with the creation of the surface. Surfaces that have been exposed for long times can to a large extent retain their original geom etry, but their associated saprolites might have gone
Relieffeatures and palaeoweathering (Lidmar-Bergstrom 1 982). Thus saprolites of widely different ages can occur on old surfaces that have been exposed for a long time. The study of the Numedal Quaternary deposits revealed that pre-glacially weathered material has contributed considerably to the Quaternary deposits, which thus strengthens the idea of a saprolitic mantle over the fresh bedrock before the glaciations. Relief features and saprolite remnants in Scandi navia bear witness to deep weathering as a funda mental process in the shaping of relief also within the formerly glaciated basement areas. Much of the relief is of etch-surface character. It is therefore necessary to separate the effect of deep weathering and subse quent stripping from glacial erosion of the fresh base ment when evaluating the long-term denudational processes.
AC K N O W L E D G E M E N T S
This study was supported b y grants from the Swedish Natural Science Research Council. The drawings were made by Lazlo Madarasz and Karin Weilow, Stockholm, and Ann Iren Johansen, Trond heim. We also thank the reviewers Alain Godard and Dale Leckie for many valuable comments on the first draft of the paper and Medard Thiry for good advice concerning the arrangement of the paper.
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Spec. Pubis int. Ass. Sediment. (1999) 27, 303-321
Palaeosol sequences in floodplain environments: a hierarchical approach
M . J. K R A U S * and A . AS L A N t *Department of Geological Sciences, University of Colorado, Boulder, C O 80309-0399 USA; and t Bureau of Economic Geology University Station, Box X, University of Texas at Austin, Austin, Tx 787 13, USA
ABSTRACT
Floodplain soils and palaeosols are considered at four spatial and temporal scales. The processes and factors that influence floodplain systems vary depending on the scales considered. Short-lived, local processes (e.g. a local influx of coarse sediment related to channel crevassing) give rise to small-scale spatial variability, whereas longer-lived autogenic and allogenic processes are responsible for intermedi ate- and large-scale spatial variability. Various floodplain soils and palaeosols illustrate the different scales and show that recognizing and analysing these different scales are important for evaluating how land scapes evolved over time and for assessing the relative significance of the various autogenic and allogenic controls on landscape evolution in alluvial basins. Spatial changes in palaeosol properties are commonly studied at the channel/floodplain scale (e.g. catenas and pedofacies that extend hundreds to thousands of metres). At this mesoscale, autogenic processes (e.g. lateral channel migration, crevassing and overbank flooding) that operate over timespans of 1-102yr influenced soil formation by controlling both patterns and rates of short-term sediment accu mulation and soil hydrology. Embedded within mesoscale changes are microscale changes in soil morphol ogy (tens to hundreds of metres in lateral extent), which formed in response to geological processes that operate over days to months. For example, a flood can locally erode and deposit sediment, producing subtle grain-size and topographical irregularities on the floodplain that influence pedogenesis by creating slightly different drainage conditions. Macroscale changes involve stratigraphical thicknesses of a few tens of metres and lateral changes over kilometres to several tens of kilometres. Such changes can represent a combination of autogenic and allo genic processes, including avulsion, tectonism and climatically controlled floodplain incision and aggrada tion. These processes probably operated over intervals of 10L104yr. Megascale palaeosol variability, which involves hundreds of metres of alluvial deposits and extends over an entire basin, is generally con trolled by global or regional climate change, sea-level fluctuations and regional tectonics, processes that influence palaeosol development over 10L 107 yr.
INTRODUCTION
and climatic changes (e.g. Fastovsky & McSweeney 1987), estimate accumulation rates in alluvial basins (e.g. Retallack 1983; Kraus & Bown 1993a) and deci pher patterns of plant and animal evolution (e.g. Retallack 1983, 1985; Bown & Beard 1990). Soils are an integral part of the landscape, and geomorphologists have turned increasingly to soil landscape studies to interpret landscape evolution (e.g. McFadden & Knuepfer 1990; Gerrard 1993). Following their work, some floodplain palaeosol studies have used palaeosol-landscape relationships to improve interpretations of ancient landscapes and
An important advance in fluvial sedimentology during the past 15yr has been studying floodplain palaeosols to provide a more complete and correct understanding of the depositional history of allu vial deposits. Analyses of floodplain palaeosols are improving our understanding of the three dimensional geometry of alluvial successions (allu vial architecture) and the autogenic and allogenic factors that control the alluvial architecture of partic ular alluvial deposits (e.g. Retallack 1986; Kraus 1987; Platt & Keller 1992). Studies of floodplain palaeosols are also helping to interpret past climates
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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palaeoenvironmental conditions (e.g. Bown & Kraus 1987; Besly & Fielding 19 89; Platt & Keller 1 99 2).The value of palaeosol-landscape studies is that they produce a clearer, more complete picture of the envi ronmental conditions and processes operating across ancient floodplain surfaces. Additionally, palaeosol landscape analysis can provide insight as to how ancient landscapes evolved over time and what processes controlled that evolution. This is espe cially true of thick alluvial successions because they contain vertically stacked palaeosols that record a series of landsurfaces. Because the landscape is a nested hierarchy of landform systems and subsytems (e.g. Haigh 1987), palaeosol-landscape associations can be studied at different spatial and temporal scales. For example, spatial changes in palaeosol properties can be exam ined at the scale of a channel and its associated flood plain (e.g. Bown & Kraus 19 87; Kraus & Asian 199 3; Kraus 1996). Alternatively, changes among groups of palaeosols can be examined at the scale of a sedi mentary basin (e.g. Atkinson 19 86; Alonso Zarza et al. 199 2). In this paper, we discuss: 1 different temporal and spatial scales of alluvial palaeopedogenesis; 2 how these different scales of palaeosol develop ment relate to physical processes and landscape evolution; 3 how analysis of floodplain palaeosols contributes to a more complete understanding of the deposi tional history of alluvial rocks and the autogenic and allogenic factors that controlled the alluvial stratigraphy. We emphasize aggradational alluvial systems in which sequences of vertically stacked palaeosols formed. In addition to examples provided by the growing literature on floodplain palaeosols, we have drawn a number of examples from our work (and that of colleagues) on Palaeogene palaeosols in the Willwood Formation of Wyoming, USA.
FLOODPLAIN ARCHITECTURAL PROCESSES
Floodplains are related genetically to the channels that construct them, and, because alluvial channels are highly variable, so too are the associated flood plains. The most complete classification is that of Nanson & Croke (1992), which is based primarily on stream power and the cohesive or non-cohesive nature of the alluvium and secondarily on fluvial
processes (e.g. lateral accretion and vertical accre tion). From a stratigraphical perspective, the sub division of floodplains into those formed primarily by lateral versus vertical accretion is most useful. Lateral accretion involves the deposition of coarser grained sediment (gravels and sands) as bar deposits during episodes of channel migration or shifting. In contrast, vertical accretion of finer grained floodplain sediment (fine-grained sandstones, siltstones and claystones) occurs during overbank flooding of the trunk channel (e.g. Allen 1965). Although floodplain palaeosols are described from ancient braided river deposits (e.g. Turner 1993) and, in some cases, make up a significant volume of those deposits (e.g. Bentham et al. 199 3), they are recognized more commonly in the fine-grained component of strati graphical successions attributed to meandering or anastomosed rivers (e.g. Retallack 1986; Bown & Kraus 19 87; Smith 1990; Nadon 199 4). Consequently, palaeosols are generally associated with floodplains in which overbank deposition was important. Recent studies of the Saskatchewan River indicate that, for some rivers, fine-grained floodplain alluvium is deposited by a combination of channel avulsion and overbank flooding (Smith et al. 1989; Smith & Perez-Arlucea 199 4) (Fig. 1). In this model, avulsion begins with crevassing of the trunk channel and continues as splay systems expand into low-lying floodplain areas (Smith et al. 19 89). With continued development, older splay systems are abandoned and the flow is gradually concentrated in fewer but larger channels in the avulsion belt. Eventually, a new meander belt develops, which occupies only a portion of the old avulsion belt. Only after avulsion, and once the new trunk channel is established, does true over bank deposition take place on levees and in flood basins. The avulsion deposits are dominated by silty clays and silts that encase sandy splay-channel and thin sheet deposits (Fig. 1). The avulsion belt is an additional and important floodplain landform, and palaeosols developed on avulsion deposits have been described from both meandering systems (Kraus & Asian 1993; Kraus 1996) and braided systems (Bentham et al. 1993).
ALLUVIAL PALAEOSOL ANALYSIS: A HIERARCHICAL APPROACH
Floodplain soils and palaeosols can be considered at a variety of spatial and temporal scales. As noted by Haigh (1987) and DeBoer (199 2), geomorphological
Palaeosol sequences in floodplains
305
Pre-avulsion surface
Fig.l. Hypothetical cross-section and plan-view diagram of mature avulsion-belt deposits based on the Saskatchewan River example of Smith et al. ( 1989 ) . The avulsion deposits consist of fine-grained sediment with ribbon sands deposited by crevasse channels. Pre- and post-avulsion deposits are overbank deposits from the trunk channel. The old trunk channel is abandoned and the new trunk channel locally truncates both avulsion and pre-avulsion deposits. (Modified from Smith et al. 1989. )
'-.
_;
�
\. ,.--.--,
.......
---Old Trunk Channel
New trunk channel &
Overbank deposits
associated sand body
't��:jt:;p�/ Overbank deposits
Table 1. Hierarchy of spatial and temporal scales in alluvial systems ( After Summerfield, 1991 )
Spatial dimensions Linear (m )
Areal ( km2)
Temporal range (yrs )
Micro: topographic irregularities
10L102
10-L1
Days to months
Meso: catena, pedofacies
10L103
10-103
10L104
10L105
Spatial scale
Macro: partial alluvial successions
( 1-10km)
Mega: formations, basin fill
(>10 km)
Stratigraphical thickness ( m)
Autogenic and allogenic processes/factors
<1
Flood Minor floodplain topography
1-102
>1
Lateral accretion Vertical accretion Crevassing
10L103
10L104
>10
Avulsion Local-regional climatic change Neotectonics
10L105
105-107
>100
Global climate change Regional subsidence/tectonics
systems incorporate both time and space, and the processes and factors that influence floodplain systems will vary depending on the scales considered (Table 1). The following sections discuss different scales of alluvial soil formation and show that inter pretations of alluvial palaeosols and soil-forming processes depend on the spatial and temporal scale of study. For example, short-lived, local processes (e.g. a local influx of coarse sediment related to channel crevassing) give rise to small-scale spatial variability,
whereas longer lived autogenic and allogenic processes are responsible for intermediate- and large-scale spatial variability. An understanding of large-scale variability in alluvial palaeosols will prob ably first require a thorough knowledge of the factors that control the smaller scales of variability. Analysis of alluvial palaeosols from the Palaeo gene Willwood Formation demonstrates the impor tance of studying palaeosols at a variety of scales. Lateral changes in palaeosol morphology that occur
M. J Kraus and A. Aslan
306
over distances of tens to hundreds of metres (microscale, Table 1) are embedded within mesoscale changes. The mesoscale variability in palaeosols is, in turn, nested within larger scales of variability, up to and including the entire alluvial basin. Upward changes between individual palaeosols may reflect autogenic factors; upward changes in assemblages of multistory palaeosols generally are the result of allo genic factors (e.g. Kraus 1987). Thus recognizing and analysing different spatial and temporal scales of floodplain palaeosol variability are important for evaluating how landscapes evolved over time and for assessing the relative significance of autogenic and allogenic controls on landscape evolution. Floodplain landforms such as channel bars, natural levees and floodbasins are characteristic of many modern and ancient river systems and soils devel oped on these landforms provide a unifying theme and an appropriate starting point for discussing allu vial pedogenesis. The floodplain landforms have been described extensively, and we refer the reader to those sources (e.g. Allen 1965; Lewin 1978) rather than providing a synopsis here. The literature on floodplain soils is also considerable (see Gerrard 1987, 1992, and references therein), and the following discussion summarizes the variability of alluvial soils and palaeosols observed at the scale of a channel and its adjacent floodplain.
MESOSCALE ALLUVIAL PEDOGENESIS
In aggradational settings, floodplains generally extend hundreds to thousands of metres on either side of the channel and soils vary systematically in grain size and drainage with distance from the trunk channel (Fig. 2). Floodplain soil variability at this scale is related closely to autogenic processes, such as lateral channel migration, crevassing and overbank flooding, which operate over time scales of 1-100yr (Table 1). These processes, along with water-table fluctuations, influence floodplain soil formation by controlling both patterns and rates of short-term sediment accumulation and soil hydrology. Floodplain sedimentation
Floodplain sedimentation is sporadic, with relatively long periods of inactivity between episodes of depo sition or erosion. The developmental history of any soil or palaeosol generally reflects the balance between sediment accumulation rate and the rate of pedogenesis (Fig. 3). Depending on short-term sedi mentation and erosion, a variety of floodplain soils can form (e.g. Morrison 1978; Kraus & Bown 1986; Marriott & Wright 1993; Wright & Marriott 1996). If erosion is insignificant and sedimentation is rapid
Channel
Floodplain Distal--
--
Proximal
Pedofacies Relations ------ Sediment Thickness Decreases �------ Accumulation Rate DEcreases ------ Paleosol Maturity Increases
Paleocatena �------ Grain Size Decreases �---- Elevation Decreases
------- Soil Drainage Decreases
Fig. 2. Schematic diagram showing changes in various floodplain properties with distance from the active channel. Palaeocatenas arise because of changes in grain size and topography. Pedofacies are characterized by increasingly mature palaeosols with increasing distance from the active channel, and they form because short-term accumulation rates decrease away from the channel. See text for more details. (Modified from Bown & Kraus 1987.)
Palaeosol sequences in floodplains
3 07 Pedogenesis
Sedimentation
"0 c: "' "'
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u
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c
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"0 ;:,
"0 ;:,
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E
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E
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Ag
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Bg/Bkg
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(b) Steady
Key Stratification
( tJ
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u
Bioturbation Features
-1
Peds with Clay films
•
Nodules
�
Slickensides
Fig. 3. Vertical profiles of alluvial sediments and soils (palaeosols) reflecting varying rates of pedogenesis and sedimentation for (a) non-steady and (b) steady depositional conditions. Compound palaeosols are likely on natural levee and crevasse splay deposits. Weakly developed cumulative profiles may form in interchannel areas of avulsion belts; better developed cumulative profiles form on overbank deposits in floodbasins. Ab, buried A horizon;Ag, gleyed A horizon; Bg, gleyed B horizon; Bw, B horizon showing colour or structure development but little if any ill uvial accumulation; Bt, B horizon showing accumulation of clays;Btj, incipient development of a Bt horizon; Cg, gleyed C horizon; Cb, buried C horizon. (After Morrison 1978 and Bown & Kraus 1981.)
and unsteady, compound palaeosols can form (Morri son 1978) (Fig. 3). These represent weakly developed, vertically stacked profiles that are separated by mini mally weathered alluvium. Morrison also described composite palaeosols, in which vertically successive profiles partly overlap. These form when the rate of pedogenesis exceeds floodplain accretion. In con trast, if erosion is insignificant and sedimentation is steady, thick cumulative soils can form (Fig. 3). These profiles reflect the deposition of successive, thin increments of floodplain sediment accompanied by
bioturbation and mottling. At the opposite end of the spectrum from the alluvial palaeosols described above are truncated profiles, which can develop if erosion removes the upper part of a developing profile. At the scale of the floodplain, the processes of floodplain construction (lateral accretion, overbank flooding and crevassing that may or may not be related to channel avulsion) strongly influence the distribution, morphology and composition of flood plain soils. Lateral accretion deposits tens of centime-
308
M. J Kraus and A. Asian
tres up to a few metres of sands and coarse-grained floodplain sediments per year on bars adjacent to the active channel (e.g. Fisk 1 944; Lewin 1978; Lapointe & Carson 1986). Bar deposits will thus either show little evidence of pedogenesis or contain compound soils (Fig. 3). As the channel continues to migrate over time, composite or well-expressed soils with Bt horizons may form. In contrast to lateral accretion, overbank deposi tion is generally slow but steady, commonly on the order of 1-10 mm yr-1 depending on proximity to the channel (e.g. Kesel et al. 197 4; Walling et al. 1992; Nicholas & Walling 1 995) . Aggradation is more rapid on natural levees and decreases towards floodbasins, which leads to the formation of an alluvial ridge (e.g. Bridge & Leeder 1979; Pizzuto 1987). Overbank deposits also systematically thin and decrease in grain size away from the channel that sourced the overbank flow (Guccione 1 993; Weerts & Bierkens 1993). Thus, compound soils with weakly expressed profiles form on natural levees, whereas cumulative profiles develop in floodbasin areas distal to the active channel (Fig. 3). Avulsion deposits, which are deposited by crevass ing, accumulate quickly because the transfer of flow from the old to the new channel appears to be completed instantaneously in a geological sense (e.g. Tornqvist 1994 ). Smith et al. (1989) found 3 m of sedi ment deposited in only 100 yr. Compound soils are thus likely to occur in crevasse splay sands and silts, whereas cumulative profiles that are very weakly developed may form in muddy interchannel areas of the avulsion belt (Fig. 3) . Depositional processes also control broad compo sitional differences among floodplain sediments (Schumacher et al. 1988). Quartz, feldspar and lithic fragments characterize the sands and coarse silts that accumulate on channel bars and natural levees of the alluvial ridge. In contrast, clay and fine silt, which typically accumulate in distal floodbasins, consist primarily of clay minerals such as smectite, illite, kaolinite and chlorite (e.g. Asian 1994). Studies of young alluvial soils have suggested that soil chem istry is also controlled by parent material grain sizes (e.g. Sidhu et al. 1977; Hayward 1985). For example, in soils along the Mississippi River, Aslan (1994) found that down-profile variations in Fe203 and Al203 weight percentages resemble the down-profile changes in clay content and he attributed the chemi cal changes to differences in soil parent materials rather than to weathering. Grain size differences across the floodplain can also influence rates of
mineral weathering. Cronan (1985), for example, found that the mineral weathering rate of soils is inversely proportional to mean grain size. Floodplain hydrology
Precipitation patterns and river stage influence water table levels and fluctuations, and affect soil moisture in floodplain settings. Alluvial soils that are saturated for several months of the year can undergo gleying, in which iron and manganese are reduced and mobil ized (e.g. Duchaufour 1982; Bridges 1973; Vepraskas 1994). As water-table levels fall and the soil dries, the iron and manganese may be leached from the soil or concentrated in more oxidized areas, either within peds or along ped faces and soil channels as mottles and/or nodules (e.g. Duchaufour 1982; Fanning & Fanning 1989). These processes produce redoximor phic features that can be observed in the field or in thin-section (Vepraskas 1994). Redox depletions, caused by iron removal, include grey soil matrix and grey root mottles (Fig. 4). Redox concentrations, formed by the re-precipitation of iron in the better oxidized areas, include various iron oxide nodules and mottles (Figs 4 & 5). Another pedogenic feature related to fluctuations of the water table are slicken sides, which are formed by the shrinking and swelling of clays (Fig. 6). In floodplain settings, soil saturation and gleying may involve surface and/or ground waters. Where clay is abundant, seasonal rains and flooding pro duce perched water tables. Subsequent surface water
Fig. 4. Photomicrograph of branching, grey root mottles or redox depletions (G). These are rimmed by intensely red stained matrix, which is a redox concentration feature (r). B horizon of a moderately well-drained palaeosol from the Willwood Formation. Frame length is 2.5 mm; plane polarized light.
Palaeosol sequences in floodplains
Fig. 5. Photomicrograph of an iron oxide nodule (redox concentration feature) with a grey halo (redox depletion feature). Lower B horizon of a poorly drained palaeosol from the Willwood Formation. Frame length is 2.5 mm; plane-polarized light.
Fig. 6. Photomicrograph of slickensided mudstone with conjugate sets of orientated clay (arrows), orientated approximately perpendicular to one another. Frame length is 4mm. Cross-polarized light.
gleying produces grey horizons and mottles as a result of poor drainge and these horizons overlie better drained brown horizons. In contrast, ground water gleying, caused by seasonal or periodic satura tion of soil materials by ground waters, is expressed by a downward increase in grey soil colours, reflect ing proximity to the ground-water table. In many instances, floodplains undergo a combination of surface and groundwater gleying (e.g. Duchaufour 1 982; Fanning & Fanning 1989). PiPujol & Buurman (1994) found that the effects of groundwater and surface-water gleying can be distinguished in the palaeosol record on the basis of
309
micromorphological features. They noted, how ever, that palaeosol studies generally fail to distin guish between the two kinds of gleying, despite the fact that making this distinction is important for palaeoenvironmental interpretations. A complicat ing issue for reconstructing the hydrological regime of floodplain palaeosols is what Retallack (1991) termed 'burial gleization'. As floodplains aggrade and soils are buried, the soils are submerged beneath the low seasonal water table. In the presence of sufficient organic matter and reducing conditions, ground waters can produce gley features in the buried soils. Similar to sedimentation, hydrological effects on floodplain pedogenesis are variable, as evidenced by lateral changes in the quantity and distribution of soil organic matter, matrix and mottle colours and soluble soil constituents (e.g. carbonate, gypsum) across alluvial floodplains (Bridges 1973; Duchau four 1 982; Vepraskas 1994) (Fig. 7). In general, hydro logical influences on pedogenesis (soil hydromorphy) correlate with soil texture and floodplain topography. Hydromorphy is greatest in clayey, poorly drained floodplain depressions (e.g. floodbasins) and least in sandy, moderately to well-drained alluvial ridge soils. For example, sandy and silty soils formed on natural levees and channel bars commonly have dark brown to brown A and Bw horizons and low organic matter contents, which reflect soil oxidation and moderate drainage. Grey soil colours are more abundant in Bg and Cg horizons and reflect reduced conditions and greater proximity to the ground water. In contrast to the natural levee soils, poorly drained flood basin soils are grey, have clayey textures, higher organic matter contents and contain abundant mottles, iron nodules and slickensides (Fig. 7). The presence of a black, organic-rich and mottled Ag horizon and a thick, grey Bg horizon with many mottles and nodules indi cates relatively prolonged saturation and poor drainage throughout the profile. The low topographi cal position and clayey texture favour poor soil drain age, anaerobic conditions and the accumulation and preservation of organic matter. Water-table fluctuations cause intersecting slickensides in the clayey sediments. Alluvial palaeosol-landscape relationships
At the scale of the channel and its associated floodplain, two important palaeosol-landscape rela tionships are topographically controlled, catenary relationships and pedofacies, which are controlled by
M. J. Kraus and A. Asian
310 River channel
Narural levee
Flood basin
Seasonal high water table Zone of water
table fluctuation
l
Seasonal low water table Increasing waterlogging and clay content
within soil profiles
[]] Brown silt with grey mottles
��tll' b��: �&�:�
Q Grey silty sand • Black orgru)ic-rich
silty day with grey mottles
bJ Grey silty clay
• Fe-Mn nodule
. . 5 Slickens1de
lateral vanat1ons in sediment accumulation rate. The two associations are not mutually exclusive and floodplain palaeosols can show a combination of the two. Palaeo catenas
The catenary relationships (or toposequences or hydrosequences) observed between contemporary alluvial ridge and floodbasin soils also have been described in alluvial palaeosols (e.g. Fastovsky & McSweeney 1 987; Platt & Keller 1 992). Fastovsky & McSweeney (1 987) recognized a catena in which palaeosols that formed in higher topographical positions show an oxidized zone. At lower landscape positions, the palaeosols show increased gleying and an 0 horizon characterizes the most poorly drained and topographically lowest part of the flood plain. In a second example, Arndorff (1 993) found that Jurassic palaeosols developed on natural levees and crevasse splays were more leached than palaeosols developed on backswamp deposits because they were sandier and better drained. The backswamp palaeosols, which formed in depressions, were dark-coloured silty claystones interpreted as ancient examples of gleyed alluvial soils (gleysols). In contrast, the levee palaeosols formed in sand and sandy silts and were light brown in colour with a yel lowish to orange subsurface horizon, reflecting better drainage.
Fig. 7. Schematic cross-section of a modern river floodplain showing the effects of topography, ground water table fluctuations, and texture on soil profiles from a levee and the floodbasin. A moderately or well-drained soil generally forms on the alluvial ridge, although subsurface horizons can be gleyed because of proximity to the groundwater table. Poorly to very poorly drained soils are more typical of the flood basin.
Pedofacies
Bown & Kraus (1 987) introduced the concept of pedofacies. They observed lateral changes in palaeosol type, as defined by the stage of maturity, relative to a coeval channel sandstone body and attributed these changes, in large part, to decreasing accumulation rates away from a channel (Fig. 2). Ped ofacies relationships have been recognized in other ancient alluvial sequences (e.g. Wright & Robinson 1 988; Smith 1 990; Alonso Zarza et al. 1 992; Platt & Keller 1 992); however, the pedofacies model, as it is currently understood, does not explain satisfactorily the lateral relationships in all floodplain palaeosol successions (e.g. Wright 1 992). In Neogene deposits of Pakistan, for example, Behrensmeyer et al. (1 995) found no systematic changes in palaeosol maturity relative to a channel sandstone, nor did they describe any catenary relationships. Our studies in the Willwood Formation suggest several limitations to the pedofacies model. For instance, variable sediment accumulation rates appear to limit the recognition of pedofacies. Study of the Willwood Formation in different parts of the Bighorn Basin, Wyoming shows that pedofacies rela tionships are readily observable in stratigraphical intervals with relatively rapid sediment accumulation rates (between 0.6 and 0.7 mm yr-1 ). In a stratigraphi cal interval with sediment accumulation rates of only 0. 3-0.4 mm yr-1 (and more mature cumulative
Palaeosol sequences in floodplains
palaeosols) , pedofacies changes are obscure. Rela tively slow sediment accumulation rates and the attainment of steady-state conditions may be respon sible for the absence of pedofacies. As Yaalon (1971) and Birkeland (19 84) have pointed out, many soil properties attain a steady-state condition. Because of the unconsolidated parent material and warm climate, the Willwood soils were probably able to reach steady-state relatively quickly. Consequently, with sufficient time, profiles at different locations in the ancient alluvial landscape may have reached steady-state conditions and thus erased pedofacies variations. Second, some processes of floodplain construction may not lead to pedofacies relationships. In the Will wood Formation, only overbank deposits from a trunk channel show the systematic decrease in sedi ment accumulation rate that promotes the devel opment of pedofacies relationships. We have not observed any systematic changes in palaeosols that developed on fine-grained sediment deposited on an avulsion belt. In the Willwood Formation, only about half of the fine-grained deposits are overbank deposits, which is probably similar to many other alluvial sequences and the pedofacies model is thus limited to only part of the fine-grained deposits. The studies of Willis & Behrensmeyer (1994) and Behrensmeyer et al. (199 5) show that not all fine grained floodplain sediment is deposited by either of the processes identified in the Willwood Formation. Although they recognized a vertical alternation of
�Paleosol Fig. 8. Schematic cross-section through Miocene fluvial deposits in Pakistan. The floodplain consists, in large part, of crevasse-splay deposits, which filled in low areas of the floodplain and on which palaeosols developed. Mudstones are shown in white and sandstone bodies are stippled. This diagram shows that some floodplain construction may be the result of crevasse-splay deposition that was not associated with avulsion of a trunk channel. (Modified from Behrensmeyer et al. 1995.)
311
weakly developed and more intensively developed palaeosols, similar to that in the Willwood Formation, they concluded that slow aggradation of an alluvial ridge followed by episodic and rapid avulsion was not responsible. Rather, both studies suggested that floodplain construction was mainly the result of the deposition of laterally extensive crevasse-splay lobes, which filled in low areas of the floodplain (Fig. 8) . Crevasse-splay deposition was not necessarily associ ated with avulsion of a trunk channel. These studies, like that of Smith et al. (19 89), indicate that a mech anism other than overbank deposition can lead to floodplain aggradation, including the deposition of significant quantities of fine-grained sediment.
MICROSCALE ALLUVIAL PALAEOSOL VARIABILIT Y
Microscale pedogenic variability is embedded within mesoscale variations and involves changes in soil morphology that occur over distances of tens to hun dreds of metres and involve geological processes that operate over time intervals of days to months. A flood event that lasts for a period of days or weeks will produce subtle topographical irregularities on a floodplain surface by locally eroding and depositing sediment. These topographical as well as probable grain-size differences, in turn, influence floodplain pedogenesis by creating slightly different drainage conditions (e.g. Fanning et al. 1973; Sobecki & Wilding 1983; Knuteson et al. 19 89). For instance, a soil developed adjacent to a floodbasin distributary or tributary stream may show evidence of better drained conditions than a soil formed 100m from the same channel, owing to minor flooding, overbank deposition and the construction of levees. In the Willwood Formation, individual palaeosols can show changes in their degree of development or hydromorphy over distances of tens to hundreds of metres. Asian (1990) described a change from a mod erately well drained to a poorly drained palaeosol over a lateral distance of lOOm. Textural similarities between the two profiles indicate that the drainage differences were not related to significant differ ences in the parent material grain sizes (Fig. 9). Field relationships show that the more poorly drained palaeosol had a restricted, elliptical distribution and was surrounded by the better drained palaeosol. The more poorly drained palaeosol is associated with the deposits of a small floodbasin channel, probably a crevasse-splay channel, that was abandoned and
M. J Kraus and A. Asian
312 Ag Bgl
Bg2
IICg
200m
--rT1andstone I �dtstone mudstone 0 Calcite Nodule e Iron-oxide nodule > Slickensides
filled. These relationships indicate that the palaeosol formed in a small floodplain depression, which favoured reducing conditions, whereas the better drained palaeosol developed on slightly elevated margins of the depression. Recognizing that palaeosols can change, in some cases significantly, over short distances as a result of local controls is important when evaluating up section changes through vertical sequences of palaeosols. Changes between vertically successive palaeosols commonly are attributed to an autogenic mechanism, such as channel avulsion (e.g. Kraus 19 87). Small-scale changes in the Willwood Formation suggest that differences in palaeosol char acteristics may, in some cases, reflect a combination of local topographical and hydrological controls on palaeosol development rather than larger scale auto genic or allogenic factors. Especially where expo sures are laterally restricted or of poor quality, caution should be used in interpreting up-section changes.
MACROSCALE ALLUVIAL PALAEOSOL VARIABILITY
Vertical changes in alluvial palaeosols involving stratigraphical thicknesses of a few tens of metres and lateral changes observed over distances of kilo-
Fig. 9. Mapped distributions of two stratigraphically equivalent, hydromorphic cumulative palaeosols and representative profiles through each palaeosol. The central area is a local floodplain depression and has a more poorly drained palaeosol than the surrounding area, which is slightly elevated. Ag, gleyed A horizon; Bg, gleyed B horizon; Bkg, carbonate-enriched and gleyed B horizon; IICg, gleyed C horizon developed in a different parent material (sandy) compared with overlying horizons. (Modified from Asian 1990.)
metres to perhaps a few tens of kilometres, can rep resent a combination of autogenic and allogenic processes (Table 1). Possible processes involved with this scale of alluvial palaeosol variability include avulsion, local tectonism and climatically controlled floodplain incision and aggradation. Based on Qua ternary studies, palaeosol variability related to these processes probably occurs over time intervals of 1QL1Q4yr. Avulsion
Avulsion influences floodplain pedogenesis by re moving areas of the floodplain (e.g. alluvial ridges) from the locus of deposition for periods of time that are probably of the order of 103yr, which is the peri odicity of avulsion (e.g. Bridge & Leeder 1979). This autogenic process produces a well-developed soil profile with a well-expressed Bt horizon and suban gular blocky structure on the abandoned alluvial ridge during the period of little or no sediment influx (Schumacher et al. 19 88; Ferring 1992; Asian 199 4) (Fig. 3). Over time, old alluvial ridges may be buried by floodbasin muds and this can lead to a buried allu vial ridge soil overlain by a surface soil developed in the floodbasin muds or, if floodbasin deposition occurs soon after alluvial ridge abandonment, a cumulative profile may form. This type of soil would be characterized by two parent materials:
Palaeosol sequences in floodplains 1 silts and sands in the lower half of the profile repre senting the alluvial ridge parent materials; 2 muds in the upper half, which would reflect renewed floodbasin sedimentation (Asian 199 4). In the Palaeogene Willwood Formation, episodic avulsions have produced up-section variability at two thickness scales. First, stratigraphical intervals that are metres thick consist of two parts: 1 avulsion-belt deposits that are characterized by very weakly developed compound or cumulative palaeosols 2 overbank deposits that are characterized by more strongly developed cumulative palaeosols (Kraus & Aslan 199 3) (Fig. 10). Vertical sequences consist of repetitions of these two types of floodplain deposits and their associated palaeosols, and no allogenic controls need be invoked
313
to generate such an alternation. A similar bipartite subdivision of floodplain palaeosols has been described from Miocene rocks by Willis & Behrens meyer (1994). They concluded that the weakly devel oped palaeosols had formed on overbank deposits that filled local depressions on the floodplain. The process of filling could have been the result of short lived avulsion or of continual crevassing. These alternations of immature and more mature palaeosols can be nested within what Kraus (1987) termed 'pedofacies sequences' in the Willwood Formation (Fig. 11). These larger scale sequences are tens of metres thick and are bounded below and above by major channel sandstones. A sequence shows upward changes in the maturity of the cumu lative palaeosols caused by episodic avulsion. The maturity of a particular palaeosol in the vertical sequence indicates the relative distance between the palaeosol column and the stratigraphically equivalent channel sandstone body. Thus, the upward changes in palaeosol maturity should reflect the pattern or style of avulsion, for example, whether the channel moved in fairly regular, step-wise jumps (as shown in Fig. 11) or in a series of random steps (see Mackey & Bridge (199 5) for a discussion as to how different avulsion styles can develop). Kraus & Bown (1993b) suggested that the upward changes in matur ity are potentially useful in petroleum exploration because they can be used to predict the direction the channel and its resulting sandstone body in which have moved through time. Local tectonics
Fig.lO. Field view of the Willwood Formation showing alternations between brightly coloured, cumulative palaeosols that developed on overbank deposits (C) and avulsion-belt deposits on which drab, weakly developed compound palaeosols formed (A). Vertical sequence is about 15 m thick.
Non-alluvial, or allogenic, processes can also influence depositional patterns and the pedogenic history of floodplain soils. For example, floodplain tilting caused by local fault activity can produce floodplain lows with soils that show the effects of rapid sediment accumulation and poor drainage (Alexander & Leeder 1987). In contrast, raised areas will have soils that reflect slower aggradation and better drainage. Local tectonic activity can also influence soil development through its control of channel diversion. For example, the Gandak River in India has shifted eastward three times over the past 5000yr as a result of episodic tectonism (Mohindra et al. 1992). This eastward shifting is recorded as a westward increase in the pedogenic development of the floodplain soils. Other examples of channel migration in response to active tectonics are pro vided by Schumm (1986).
M. J. Kraus and A. As lan
314 \ \
\
\
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e �
8
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I
I
---
Extent of avulsion deposits
8 Fig.ll. Schematic diagram showing upward changes in palaeosol maturity caused by channel avulsion. The stratigraphical column shows a stacked sequence of floodplain palaeosols with major channel sandstones at the top and bottom of the column. The Arabic numerals indicate palaeosol maturity (1 is immature and 3 is mature) for cumulative palaeosols, which alternate with immature palaeosols that formed on avulsion-belt deposits. Roman numerals indicate the successively younger channel sandstones associated with each cumulative palaeosol. The lateral extent of the avulsion-belt deposit underlying each channel sandstone is indicated by the heavy black line. Horizontal scale of the diagram is of the order of 5-lO km. See Fig. 3 for symbols. (From Kraus & Bown 1 993b.)
Climatically controlled floodplain incision and aggradation
Floodplain incision and aggradation caused by climate fluctuations and changes in the discharge and type of sediment transported by rivers significantly influence alluvial pedogenesis. For instance, impor tant consequences of floodplain incision on alluvial pedogenesis are the changes in soil hydrology and geochemical conditions that accompany incision. Floodplain soils formed initially in poorly drained floodplain settings are oxidized and leached follow ing floodplain incision and lowering of water tables (Bettis 1992). Climatically controlled increases in peak flood discharge also can cause floodplain aggra dation and produce cumulative soils on floodplains (Schumm & Brackenridge 1987; Brakenridge 19 88). Climate changes and episodes of floodplain aban donment and aggradation also influence alluvial palaeosol stratigraphy. Numerous studies of late Quaternary alluvial deposits in the USA show that soil and palaeosol-bounded allostratigraphical units represent climatically controlled episodes of flood plain abandonment and/or aggradation that have occurred over time intervals of 103 yr (e.g. Knox 198 3; Schumm & Brackenridge 1987; Chatters & Hoover
19 88; Hall 199 0; Autin 1992; Ferring 1992; Blum & Valastro 199 4) . In other examples, down-valley changes in alluvial palaeosol characteristics and stratigraphy reflect the interplay between climatic and base-level (sea-level) influences on alluvial pedo genesis. For instance, late Quaternary alluvial soils and palaeosols located in the bedrock-confined valley of the Colorado River in south Texas differ significantly from those located on the alluvial plain near the coast. Within the bedrock-confined valley, alluvial soils and palaeosols with leached E and/or Bt horizons are present beneath inset alluvial terraces, and the soils and palaeosols bound allostratigraphical units that formed in response to changes in Colorado River flood hydrology (Blum & Valastro 199 4). In contrast, stratigraphically equivalent and morpho logically similar palaeosols that are located less than 100km down-valley are buried by 10-15m of pedo genically modified Holocene muds. The Holocene muds are cumulative soils that formed in response to Holocene sea-level rise. Evidence of floodplain incision in pre-Quaternary deposits probably is recognized most easily over outcrop distances of tens to hundreds of metres, where scour surfaces and overlying channel fills can be observed in the field. Floodplain incision involving
Palaeosol sequences in floodplains
large floodplain areas, comparable to those observed in Quaternary fluvial systems, is probably recorded in the older floodplain deposits, but recognition of these types of events is difficult because the margins of incised floodplains and topographical relief may not be apparent or well exposed. Even in floodplain deposits where detailed biostratigraphical data indi cate the presence of a significant unconformity, Kraus & Bown (1993a) found scant evidence of macroscale floodplain incision in the Willwood Formation. Simi larly, Wright (1992) concluded that although flood plain incision and terracing is common in Quaternary fluvial systems and probably was common in the ancient record, pre-Quaternary terraces generally are inferred rather than observed directly. Despite these problems, a few examples of flood plain incision and palaeosol development, resulting from both autogenic and allogenic controls, are described in the literature (e.g. Retallack 1986; Kraus & Middleton 1987a) (Fig. 12). One of the best exam ples is that of Marriott & Wright (1993), who found significant differences in upper Silurian to lower Devonian floodplain palaeosols, which they sug gested related to the stability of the fluvial systems. Cumulative palaeosols developed where sedimen tation was slow but relatively continuous. Complex truncated palaeosols were attributed to periods of erosion, possibly triggered by changes in climatic conditions or vegetation cover, which resulted in incision and local truncation of palaeosols. This example is particularly important because it shows how careful examination of the palaeosol record can
Fig.12. Floodplain incision in the Triassic Chinle Formation, Arizona. Scour surface (arrows) truncates deeply coloured palaeosols and is filled by sediment on which light-coloured, immature palaeososols formed. Incised hill in left background is about 1 1 m high.
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lead to a richer, more thorough understanding of the complex processes responsible for a particular allu vial succession.
MEGASCALE ALLUVIAL PALAEOSOL VARIABILITY
Palaeosol variability involving hundreds of metres of alluvial deposits and extending over an entire alluvial basin generally is controlled by factors such as global or regional climate change, sea-level fluctuations and regional tectonics, processes that influence alluvial systems over time intervals of 10L107yr (Table 1). The relative importance of these factors is controlled, in part, by location of the fluvial system (Shanley & McCabe 1994). Eustasy is a significant control on coastal plain rivers (e.g. Shanley & McCabe 1993, 1994); however, its effects decrease away from the sea. In alluvial systems distant from the sea or in closed basins, base-level changes are locally con trolled, and climate or regional tectonic activity are the major allogenic controls on rivers (e.g. Blum & Valastro 1994;Blum 1994). Because smaller-scale changes in floodplain palaeosols are nested within the basinal-scale vari ability, larger packages of floodplain palaeosols must be analysed to assess the relationship between large scale geomorphological systems and palaeosols. Furthermore, because the evolution of large-scale geomorphological systems is controlled by allogenic processes that operate at long time-scales, thicker sequences of palaeosols need to be examined to determine whether changes in those processes have influenced the alluvial basin. Not surprisingly, exam ining palaeosol variability at this scale can be difficult because it depends on widespread exposures, and, if those exposures are not relatively continuous, a reli able method of establishing the time-equivalence of the palaeosols. For example, the Eocene Capella For mation, studied by Atkinson (1986) at a basinal scale, changes from 300m to c. 140m over a distance of 28 km. Approximate time equivalence of the formation across this distance was established from a marine intercalation. Kraus (1992) examined differences in palaeosols in the lower 150m of the Willwood Formation across a study area of 2000km2. Because exposures are not continuous across this area, bio stratigraphical data were used to establish strati graphical equivalence. In those examples where large-scale variability in floodplain palaeosols has been studied, it has been
M. J. Kraus and A. Asian
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attributed generally to tectonically controlled varia tions in relief and differences in subsidence rates, which control accumulation rates. Several studies have observed pronounced palaeosol variations in a down palaeoslope direction. In his study of the Capella Formation, Atkinson (1986) described more mature and better-drained palaeosols proximal to the source area. Progressively less mature and more poorly drained palaeosols are found with increasing distance from the source. Atkinson attributed these changes to a decline in topographical relief towards the sea, in which the fluvial system drained, and to increased accumulation rates away from the source. Miocene floodplain palaeosols studied by Platt & Keller (1992) also show changes in maturity and hydromorphy down palaeoslope. In this example, both palaeosol maturity and hydromorphy increased downslope over a 7 5-km distance. Associated with those changes, the stratigraphical interval thins downslope, indicating slower sediment accumulation. Similarly, in part of the Gangetic Plain, soil maturity increases away from the Himalayas over a distance of 160 km (Srivastava et al. 1994). The control is differ ential subsidence, which has led to decreasing rates of sediment accumulation away from the mountain front. A different, although still tectonically controlled, situation was described by Kraus (1992) in the Willwood Formation. She used remote sensing data to map the distribution of four lithofacies that cover areas ranging from 150 to nearly 500 km2. Different types of palaeosols in the facies reflect variable drainage conditions across the study area. The geographical distribution of the different types of palaeosols suggests that east-west faults, which extend into the alluvial basin from moun tains flanking its east margin, were active when the Willwood Formation was deposited. Kraus suggested that movement along these basement controlled faults generated topographical gradi ents that helped produce variable drainage conditions, which, in turn, affected Eocene soil development: Also important is the rate of basin subsidence, which influences the relative importance of channel migration and overbank deposition. For example, models of alluvial stratigraphy (e.g. Allen 1978, 1979; Bridge & Leeder 1979) suggest that, when basin sub sidence is relatively rapid and causes rapid sediment accumulation, overbank deposits have a high preser vation potential and the floodplain is dominated by ·
fine-grained alluvium. In contrast, when basin sub sidence is slow, channels have the opportunity to rework older floodplain deposits and to produce a floodplain dominated by channel deposits. An inter esting study by Mack & James (1993) showed that basin symmetry also can influence the preservation of fine-grained deposits and floodplain palaeosol development. In a study of Plio-Pleistocene fluvial deposits filling asymmetrical and symmetrical basins in the Rio Grande rift, they found that symmetrical basins contain more numerous and more mature palaeosols than do asymmetrical basins. The preser vation potential of floodplain deposits was low in the asymmetrical basins because they had narrow allu vial plains, which led to the reworking of floodplain deposits. In contrast, the wider floodplains of the symmetrical basins favoured preservation of the floodplain deposits on which the soils developed and also allowed for periods of inactivity during which mature soils formed. Although palaeosols from floodplain settings were not the main focus of Alonso Zarza et al. (1992), this study is worth mentioning because it is truly basin-wide in scale. The authors described Miocene palaeosols from alluvial fan and lake margin settings as well as floodplains. Stratigraphical intervals c. 100 m thick were examined in two contrasting areas of the Madrid Basin. This study emphasizes the impor tance of palaeosols for better interpreting ancient landscapes because it clearly links different kinds of palaeosols to their positions in the ancient landscape. The authors also show that palaeosols from the two parts of the basin show differences that reflect not only different rates of sediment accumulation but also different climatic conditions. In vertical sections, changes in palaeosol drainage or maturity have been used to interpret changes in regional palaeolandscapes over time. These changes have been attributed to allogenic mechanisms including climatic change, tectonic activity and eusta tic change. One of the best examples of climatic control is provided by the distinct change in flood plain palaeosols across the Cretaceous-Tertiary boundary described in different parts of Montana (e.g. Fastovsky & McSweeney 1987; Retallack et al. 1987). The palaeosol record indicates that the earliest Palaeocene floodplains were more poorly drained than those of latest Cretaceous age. Although several allogenic mechanisms (e.g. sea-level rise) could have caused this regional-scale palaeoenvironmental change, palaeobotanical evidence suggests that a
Palaeosol sequences in floodplains
change to a more humid climate was responsible (Fastovsky & McSweeney 1987). Changes in basin subsidence have been invoked to explain a major change in palaeosol maturity in a thick stratigraphical sequence of palaeosols (e.g. Kraus 1987). Kraus found that cumulative palaeosols of latest Palaeocene age in the northern Bighorn Basin are less mature than overlying cumulative palaeosols of earliest Eocene age. This change corre sponded to a change in sandstone body architecture, and both changes suggested that sediment accumula tion rates declined from latest Palaeocene to earliest Eocene time in the northern part of the basin in res ponse to slowed subsidence of the basin.
FUTURE DIRECTIONS: A L LUVIAL ARCHITECTURE STUDIES
The development of quantitative models of alluvial architecture (e.g. Allen 1978, 1979; Bridge & Leeder 197 9; Mackey & Bridge 1995) has had a major impact on fluvial sedimentology over the past 20yr. These models examine the geometries and gross arrange ment of channel sandstones and non-channel deposits within thick alluvial packages as well as the autogenic and allogenic controls that produce particular arrangements. The quantitative models have been applied to ancient alluvial successions to evaluate the autogenic and allogenic processes that were important in a particular alluvial basin (e.g. Blakey & Gubitosa 1984; Kraus & Middleton 1987b; Shanley & McCabe 1 993). Alluvial architec ture studies, either computer models or field studies, are important because they provide a temporally and spatially broad perspective on alluvial deposits and the factors that controlled the development of those deposits. Although the focus of alluvial architecture models is the gross arrangement of the channel sandstones and surrounding fine-grained deposits, detailed study of palaeosols developed on the fine-grained deposits can provide a much more detailed picture of the allu vial stratigraphy of a particular stratigraphical unit (e.g. Alonso Zarza et al. 1992; Kraus & Asian 1993). In one of the earliest studies of alluvial architecture, Allen (1974) used floodplain palaeosols to help develop models of alluvial architecture that he then tested against actual field examples. Like Allen, we believe that the analysis of floodplain palaeosols has much to offer alluvial architecture studies, both
317
field studies and computer modelling. Floodplain palaeosols are sensitive indicators of change in the fluvial system, and changes in the size or other char acteristics of the major sandstone bodies in an allu vial deposit should be associated with a change in the palaeosol record. Alluvial architecture models do not yet incorporate information from floodplain palaeosols, and an important area of future research is to better integrate palaeosol-landscape analysis with both quantitative and field studies of alluvial architecture. Here we use the elegant model of Mackey & Bridge (1995) as an example of the potential value of integrating palaeosol analysis with the analysis of channel sandstone bodies when studying alluvial architecture. Their model is three-dimensional and thus a significant improvement over earlier models in that the positions of avulsed channels can be pre dicted more realistically; however, it has yet to be field tested. This model predicts 'avulsion sequences' in which (i) the point at which avulsion occurs moves progressively upstream and (ii) the time between avulsions progressively decreases. One avulsion sequence ends and the next begins when the avulsion point reverts to a downstream position again. An avulsion sequence develops because, upflow of an avulsion point, the alluvial ridge continues to aggrade and increases the probability of avulsion there. Each alluvial package produced by an avulsion sequence should have a thicker sandstone at the bottom with progressively thinner sandstones at stratigraphically higher positions (Fig. 13). If avulsion sequences are present we predict that the maturity of vertically suc cessive palaeosols should decrease up-section as the avulsion frequency increases and as sandbody thickness decreases. An unusually mature palaeosol should develop in association with the thick sand stone at the base of a new avulsion sequence. This mature palaeosol should overlie the least mature palaeosol in the sequence.
ACKNOWLEDGE MENTS
Research contributing to this paper was supported by National Science Foundation Grant EAR-9303959 to MJK. P.D. Gingerich provided invaluable logistical support for field work in Powell and field assistance was provided by Brian Gwinn. Constructive reviews were provided by Drs Isabelle Cojan, Medard Thiry and Cesar Viseras.
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ALLUVIAL SUCCESSION
AVULSION SEQUENCE
Schematic diagram showing avulsion sequences in which a thicker sandstone forms at the bottom and progressively thinner sandstones form at stratigraphically higher positions. An avulswn sequence forms because (1 ) the pomt at wh1ch avulsion occurs moves progressively upstream and (2) the time between avulswns progressively decreases. One avulswn sequence ends and the next begins when the avulsion point reverts to a downstream positiOn agam.Tius d1agram has a . horizontal scale ranging between 101 and 103 km and a vertical scale ranging between 101 and 103 m. (From Mackey & Bndge Fig. 13.
1995.)
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J.E. (1987) Sediment diffusion during overbank flows. Sedimentology,34, 301-317. PLATT, N.H. & KELLER, B. (1992). Distal alluvial deposits in a foreland basin setting-the Lower Freshwater Molasse (Lower Miocene), Switzerland: sedimentology, architec ture and palaeosols. Sedimentology, 39,545-565. RETALLACK, G.J. (1983) A paleopedological approach to the interpretation of terrestrial sedimentary rocks: the mid Tertiary fossil soils of Badlands National Park, South Dakota. Geol. Soc. Am. Bull. , 94, 823-840. PIZZUTO,
(1985) Fossil soils as grounds for interpret ing the advent of large plants and animals on land. Philos. Trans. R. Soc. London, B309, 105-142. RETALLACK, G.J. (1986) Fossil soils as grounds for interpret ing long-term controls on ancient rivers. f. sediment. RETALLACK, G.J.
Petrol., 56, 1-18. G.J. (1991) Untangling the effects of burial alteration and ancient soil formation. Ann. Rev. Earth Planet. Sci. , 19, 183-206. RETALLACK, G.J., LEAHY, G.D. & SrooN, M.D. (1 987) Evi RETALLACK,
dence from paleosols for ecosystem changes across the Cretaceous/Tertiary boundary in eastern Montana. Geology, 15, 1090-1093.
B.A., DAY, W.J., AMACHER, M.C. & MILLER, B.J. (1988) Soils of the Mississippi River alluvial plain in Louisiana. Louisiana Agric. Exper. Sta. Bull. ,
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S.A. (1986) Alluvial river response to active tec tonics. In: Active Teqtonics, pp. 80-94. National Academy Press, Washington, D.C. SCHUMM, S.A. & B RACKENRIDGE, G.R. (1987) River responses. In: North American and Adjacent Oceans During the Last Deglaciation (Eds Ruddiman, W.R. & Wright, H.E.), pp. 221-240. Geology of North America, K-3, Geological Society of America, Boulder, CO. SHANLEY, K.W. & McCABE; P.J. (1993) Alluvial architecture in a sequence stratigraphic framework: a case history from the Upper Cretaceous of southern Utah, USA. In: ScHUMM,
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(Eds Flint, S. & Bryant, I.D.), Spec. Pubis int. Ass. Sedi ment., no. 15, pp. 21-56. Blackwell Scientific Publications, Oxford. SHANLEY, K.W. & McCABE, P.J. (1994) Perspectives on the sequence stratigraphy of continental strata. Bull. Am. Assoc. Petrol. Geol. , 78, 544-568.
P.S., SEHGAL, J.L. & RANDHAWA, N.S. (1977) Elemen tal distribution and associations in some alluvium-derived soils of the Indo-Gangetic Plain of Punjab (India).
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M.A. (1991) Global Geomorphology: an Introduction to the Study of Landforms. Wiley, New
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(1994) Middle and late Holocene avulsion history of the River Rhine (Rhine-Me use delta), Nether lands. Geology, 22, 7 1 1-714. TURNER, B.R. (1993) Paleosols in Permo-Triassic continen tal sediments from Prydz Bay, East Antarctica. I. Sedi ment. Petrol. , 63, 694--706. VEPRASKAS, M.J. (1994) Redoximorphic features for identi fying aquic conditions. North Carolina Agric. Res. Serv. Tech. Bull. , 301,33 pp. WALLING, D.E., QUINE, T.A. & HE, 0. (1992) Investigating contemporary rates of floodplain sedimentation. In: Lowland Floodplain Rivers: Geomorphological Perspec tives (Eds Carling, P.A. & Petts, G.E.), pp. 165-184. Wiley,
Chichester. H.J.T. & BIERKENS, M.F.P. (1993) Geostatistical analysis of overbank deposits of anastomosing and mean dering fluvial systems: Rhine-Meuse delta, TI1e Nether lands. Sediment. Geol. , 85,221-232.
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B.J. & BEHRENSMEYER, A.K. (1994) Architecture of Miocene overbank deposits in northern Pakistan. f. Sedi ment. Res. , 64B, 60-67. WRIGHT, V.P. (1992) Paleopedology: stratigraphic relation ships and empirical models. In: Weathering, Soils and Paleosols (Eds Martini, I.P. & Chesworth, W.), pp. 475-499. Elsevier, Amsterdam. WRIGHT, V.P. & MARRIOTT, S.B. (1996) A quantitative approach to soil occurrence in alluvial deposits and its application to the Old Red Sandstone of Britain. I. geol. Soc. London, 153, 907-913. WRIGHT, V.P. & ROBINSON, D. (1988) Early Carboniferous floodplain deposits from South Wales: a case study of the controls on palaeosol development. I. Geol. Soc. London, 145, 847-857. YAALON, D.H. (1971) Soil-forming processes in time and space. In: Paleopedology (Ed. Yaalon, D.H.), pp. 29-39. Israel University Press, Jerusalem. WILLIS,
Spec. Pubis int. Ass. Sediment. (1999) 27, 323-335
Carbonate-rich palaeosols in the Late Cretaceous-Early Palaeogene series of the Provence Basin (France)
I . C OJAN Ecole Nationale superieure des Mines de Paris, CGES-Sedimentologie, 35 rue S1-Honore, 77300 Fontainebleau, France.
ABSTRACT
Carbonate-rich palaeosols are numerous in the continental formations of the Provence Basin. In the allu vium floodplain deposits, they developed in reddish silty mudstone and are characterized by an oblitera tion of the primary sedimentary structures, the presence of root moulds, an extensive colour banding and the existence of carbonate nodules or rhizoliths of varied sizes and density. In the lake marginal environment, this palaeosol occurs in the palustrine facies. The soils developed in the lacustrine carbonate mud and common features are root traces, a faint but distinct mottling, an in situ brecciation and numerous recrystallizations related to periods of high water table. The analysis of the bulk rock composition and the clay mineral assemblages showed a replacement of the detrital clay assemblages by authigenic ones in both facies. Despite the different host sediments and facies associations, many similarities may exist on the macro scopic scale between the very mature floodplain calcretes (coalescent nodules) and the nearly pure car bonate palustrine facies, or between the palustrine facies with a high terrigenous input and the red aluvium with scattered carbonate nodules. The Provence Basin series offer a well preserved sequence of vertically aggrading carbonate-rich palaeosols characterized by alternating periods of development of floodplain calcretes and palustrine facies.
INTRODUCTION
1 987; Lehman, 1 98 9; Smith, 1 990; Wright & Alonso Zarza, 1 990; Alonso-Zarza et al., 1 992; Rodas et al., 1 994;Wright & Platt, 1 995). Over the last 1 5 years there has been an increasing interest in floodplain deposits and associated palaeosols (e.g. Bown & Kraus, 1 987; Retallack et al., 1 987; Sigleo & Reinhardt, 1 98 8). According to the climatic conditions, different types of palaeosols can be present on the stratigraphical column of a basin; on the other hand, within a period of climatic stabil ity, the succession of palaeosols and their maturity stages is a powerful tool for studying channel migra tions (Bown & Kraus, 1 981; Kraus & Bown, 1 988; Retallack, 1 990; Platt & Keller, 1 992). In contrast to the channel facies, the original stratification in the floodplain is most often obliterated by bioturbation and pedogenic horizons. Channel facies, however, are of limited extent, whereas palaeosols can be
Amongst the various palaeoweathering surfaces that are preserved, the carbonate-rich ones are certainly those for which much debate has arisen (e.g. Gile et al., 1 966;Esteban & Klappa, 1 983;Goudie, 1 97 3). This type of carbonate-rich horizon is discussed based on the Upper Cretaceous-Lower Palaeocene continen tal series of the Provence Basin which display numer ous horizons of fossil soils that developed in both alluvium and palustrine environments. The present study describes these palaeosols, interprets the processes responsible for their formation and shows the correlation potential that can be developed from these horizons in continental formations where the frequent paucity of fossil remains prevents establish ment of a solid biostratigraphical framework. Fossil soils can be used as marker horizons to reconstruct palaeosurfaces, palaeoenvironments or stratigraphi cally equivalent palaeoclimatic units (Bown & Kraus,
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
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I. Cojan
traced over large distances (Kraus & Bown, 1988). The main characteristics for identifying a palaeosol are the presence of a soil horizon, root traces and soil structures. Quite often, lacustrine environments also are associated with fluvial systems. The marginal lake environments are regularly exposed to subaerial conditions during low lake-level stages. During these periods, soils develop on the marginal lacus trine facies and frequently can be correlated with the soils that developed on the floodplain alluvium (Cabrera et al., 1985; Sanz et al., 1995). From the experience gained from study of the Provence carbonate-rich palaeosols, this paper attempts to present criteria, based on both macro scopic description and mineralogical analysis, to distinguish the palaeosols developed on floodplain alluvium from those on the marginal lacustrine mudstones. Carbonate-rich soil profiles
Amongst the carbonate-rich soil profiles, 'calcimor phic soils' in the USA soil classification, two major groups can be identified: 1 soils developed on floodplain alluvium that display a carbonate-rich accumulation horizon; 2 calcic soils that developed on a parent material that was already very rich in carbonate, such as palustrine facies (lacustrine and/or ephemeral pond mudstones), marine marginal deposits, or carbonate basement rock. Nodular or micritic alteration of subaerially exposed marine limestones, dolostones or other carbonate basement rock will not be discussed as these do not occur in the Provence series. Despite the differences among the various environments where the two major groups of soils can be observed, a palustrine facies may resemble, on the macroscopic scale, the mature calcisoils from the floodplain. Calcic soils in the alluvial environment
All sediments exposed permanently or periodically to subaerial conditions are affected to varying degrees by pedological processes. The substrate is altered by physico-chemico-biological processes, the intensity of which depends greatly on the tempera ture but also on the rainfall pattern, the drainage area, the chemical instability of the parent rocks, the granulometric characteristics of the sediments and the density of the organisms living or colonizing the substrate. Duration of the palaeopedological
activity is also important in the degree of maturation (Duchaufour, 1982). A soil is characterized by a profile that records the mineralogical transformations of the substrate during pedogenesis. The downward succession com prises, below the humic horizon: (i) the A horizon dominated by leaching processes, followed by (ii) the accumulation horizon (B, including the carbonate enriched layer Be) and then (iii) the unmodified substrate. Profiles develop in the soil as a progressive downward migration of the leached elements and their precipitation in the accumulation horizon. The calcium carbonate is able to move downwards in the profile in climatic zones where an alternation of dry and humid conditions dominates (Duchaufour, 1982). In this context, the soils or palaeosols display ing a calcic (petrocalcic) horizon are often called cal cretes. The term calcrete has been the subject of much debate because it has been broadly used to describe facies resulting from different types of processes. In the scope of this volume, we shall use it to describe soil profiles with a carbonate-rich horizon, the most widespread use of this term. Excellent reviews of the different meanings of this term have been published by Esteban & Klappa (1983), Wright & Tucker (1991) and Milnes (1992). Pedogenic calcretes cannot be considered as a soil type but are part of a soil profile. In ancient series, the upper part of a palaeosol (A horizon, plus upper part of B horizon) most often has been truncated by erosion. In the horizon preserved, the B-Bc bounda ries are usually gradational. In the profile, the calcic horizon constitutes a threshold to the erosion and represents the part of the soil profile that has the greatest chance of being preserved. In ancient sequences it is therefore difficult to appreciate the original depth of the accumulation horizon in the soil because: 1 the upper part of the soil profile has not been preserved; 2 most of the soils are cumulative or compound, as defined by Kraus & Bown (1988). Calcretes constitute widespread pedogenic horizons, both in ancient sequences (e.g. Freytet, 1973; Buurman, 1980; Lang et al., 1990; Djurdjevic-Colson, 1996) and in modern ones (Watts, 1980; Arakel, 1986). Several classifications exist for the description of calcretes. These are based mainly on the macroscopic features of the calcic horizon development profile, and the morphology of the carbonate cement in this accumulation horizon ( Gile et al., 1966; Goudie, 1973;
Carbonate-rich palaeosols
Netterberg, 1 980). Some others take into account the role of the parent-material characteristics (grain size and mineralogy) that influence the rate of profile development (Machette, 1 98 5; Wright & Tucker, 1 991). Early stages correspond to small carbonate coatings, discrete soft to very hard con cretions of carbonate that will pass through time to honeycomb and hardpan facies. The predominant macrofeatures in calcretes are: colour banding, nodules of varied size (from a few millimetres to several centimetres) and rhizoliths (filling of dead roots) or rhizomorphs (cementation around roots by vertically stacked nodules).
bonates, largely as a result of the high rate of carbon ate production in the littoral zone (Freytet & Plaziat, 1 982; Freytet, 1 984; Platt & Wright, 1 991; Platt, 1 992). Inorganic precipitation is favoured by warm temper atures and bio-induced precipitation is governed by algae (charophytes), rooted plant activity and encrusting of carbonate encrustation on the vegetation (Klappa, 1 980; Wright & Robinson, 1 988). Another source of biogenic carbonate is the remains of calcareous organisms, the most common are molluscs, ostracods and, among the plants, the charophytes. The importance of charophytes as major contributors of calcium carbonate in the littoral envi ronment has been long recognized (Freytet & Plaziat, 1 982). The marginal lake area also represents the area where most of the detrital sediment is trapped by vegetation and deposited where the stream energy and sedimentary gradient diminish on reaching lake base level. The detrital sediment may dilute the car bonate content of the littoral carbonate if it consists mostly of terrigenous grains, or increase the carbon ate content if it is rich in carbonate debris. Palustrine facies belong to areas where the lacustrine carbonates are subjected to soil processes during subaerial exposure, but are still saturated by water some of the time. Most common pro cesses are colonization by land plants (reflected by rootlets, mottling), dessiccation during dry periods (resulting in the formation of glaebules, nodules, circum-granular cracking), dissolution and carbon ate precipitation during high water-table stages (Fig. 1). In this environment, a 'continuous spectrum' of facies from lacustrine carbonates to palustrine
Palustrine carbonates
In the continental realm, precipitation of calcium carbonate is restricted mainly to environments with carbonate-rich substrate rocks and marginal areas of lakes or ephemeral ponds. In the littoral parts of lakes, evidence of pedogenic modification within the carbonate-rich muds is common. These facies were named palustrine by Freytet & Plaziat (1 982), a term that is now in common use (Cabrera et a!., 1 985; Platt, 1 98 9; Platt & Wright, 1 992; Valero Garces et al., 1 994). Soils formed on carbonate substrates involve the same processes as for the development of a soil on an alluvial sediment, except for the fact that the calcium carbonate content is higher and the granulometry and texture of the rock are different. This paper will focus on the pedogenic processes taking place on carbonates of marginal environments of shallow, unstratified lakes (Fig. 1). The littoral lake sediments usually contain a high percentage of car-
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Fig. I. Distribution and main characteristics of the carbonate-rich palaeosols.
325
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I. Cojan
326
carbonates and pedogenic calcretes can be observed (Esteban & Klappa, 1983).
DEPOSITIONAL SETTING O F THE AIX-EN -PROVENCE BASIN
Regional setting
The Provence Basin lies in the southern part of France, some 20km north-northeast of Marseille (Fig. 2). The Upper Cretaceous-lower Palaeogene continental formations are well exposed throughout the present, E-W orientated Aix-en-Provence syn cline which is limited on its southern and northern borders by thrusted massifs. Throughout Upper Cretaceous and lower Palaeo gene times, the palaeogeography corresponded to a dominantly braided fluvial system, which drained a low-relief floodplain and then flowed into a perman ent lake (Fig. 3) (Cojan, 1993). The shorelines of the lake migrated over large surfaces in this low-relief landscape, in response to lake-level fluctuations. Cli matic reconstructions from the pollen record show that the climate was tropical to subtropical over that period of time (Medus et al., 1992). The Maastrichtian-Palaeocene strata, averaging 400 m in thickness (Fig. 2), largely consist of interbedded mudstones and siltstones, with some lenticular sandstones. These fluvial deposits are interfingered with some lacustrine, palustrine and pond carbonates.
The fluvial facies
The floodplain deposits, consisting mainly of reddish mudstones, represent in volume the largest amount of sediment cropping out in the region. Owing to bioturbation and the development of abundant palaeosol horizons (pedoturbation), no original stratification structures are preserved in this clayey-silty material. The palaeosols occur mainly on the overbank deposits and were identified on the basis of root traces, mottling, colour banding, presence of carbonate accumulations and, in the most clay-rich facies, slickenside features. The channel belt area is characterized along the stratigraphical column studied by a higher, but still relatively low, sand/shale ratio of around 40% . The sandstone bodies correspond to sandbars, plus some filling of braided stream channels. The mud stone deposits resulted from periodic overbank floodings but no primary structures are preserved. The lacustrine facies
The lacustrine carbonates are composed of nearly pure micritic limestone containing gastropods, ostracods and charophytes. Dolomitic facies, devoid of fossils are attributed to playa environments. The terrigenous input was very low, as quartz and clay contents do not exceed a few per cent. These sub aquaeous facies do not show any feature in relation to water stratification, but quite often display desicca tion features and root traces, indicating that these
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Carbonate-rich palaeosols
Roques Hautes
327
Ribas
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Castelias
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limestones Fluvial network Alluvial fans
Fig. 3. Schematic block diagram illustrating the palaeogeography of the Provence Basin. Extension of the main lake corresponds to that of a low lake level (from Colson et al., 1998).
lakes were of variable shallow depth and that the associated deposits were periodically emergent, probably in relation to seasonal water fluctuations. In 196 4, Pierre Freytet proposed the term 'palustrine carbonate' to distinguish these facies from per manent swamp deposits. Intercalation of palus trine facies within the lacustrine deposits is quite common (approximate lacustrine/palustrine ratio of 2 : 3), demonstrating the shallow nature of these environments.
PALAEOSOL PRO FILES: DESCRIPTION
The floodplain carbonate-rich palaeosols
The palaeosols that display carbonate accumulation horizons are the most frequent and widely distrib uted in the Provence formations. Although these soils represent different stages of maturity, they all underwent similar pedogenic processes. They can be described from the distribution of the soil horizons, the root traces and their structure (Figs 4a, 5 & 6). Soil horizons
The thickness of the fossilized horizons can vary from 0. 5m to 3 m according to the rate of alluvial sedimen tation (Figs 4a & 5). Most of these palaeosols are cumulative soils, which implies a displacement of the accumulation horizon that is migrating upward in the profile through time. In the palaeosols of the
Provence Basin, only B and Be horizons are preserved. In the field, the macroscopic identification of the Be from the B horizon is based on the occurrence of the carbonate glaebules. A typical feature of these palaeosols is a colour banding as a result of iron hydroxide accumulation in the B horizon (Fig. 5). Despite the fact that the top part of these ancient soils has never been preserved in these series, the fol lowing succession of colours is commonly observed: a downward increase in redness (with gradation from yellow-orange to dark-red) (see Fig. 8). Below the darker horizon a reverse sequence is observed down to the grey-greenish parent material that has not been modified. In alluvium subjected to incipient pedogenesis, the palaeosol exhibits only a light colour mottling of yellowish-orange colours with diffuse boundaries. Root traces
They represent one of the most common features in the Provence palaeosols. They correspond to the best preserved part of the plants in this basin, where no macroplant debris has ever been found. Root traces are generally vertical, irregular in width and show numerous rootlets (Fig. 6a,b ). Their size does not exceed a few millimetres in diameter and can reach up to 1m in length. Rootlets are most often horizon tal. In early stages of pedogenesis the root traces are surrounded by light coloured halos (whitish), which represent the chemical transformations in a microenvironment associated with living roots
I. Cojan
328 A) alluvial floodplain facies
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Fig. 4. Simplified horizonations for carbonate-rich palaeosols that developed in the Provence Basin. (a) Floodplain alluvium. Facies sequences 1-3 correspond to stages of increasing maturity. (b) Palustrine facies. Facies sequences 1 and 3 illustrate palustrine facies with low terrigenous input, in contrast to sequence 2.
(fungi, microorganisms, soil water, etc.) (Buurman, 1980; Retallack, 198 5; Pipujol & Buurman, 1993). In these soils, diffuse calcium carbonate concentration is the main feature observed. Soil structures
In outcrop, the palaeosol always gives a rugged appearance. In the clay-rich B horizons, a common type of cutan is clay skin (Fig. 5d), formed when the clay was washed down in the cracks within the soil (illuviation argillans; Brewer, 1976). Slickensides are associated with these clay cutans as a result of the shrinking and swelling of the clays during alternating wet and dry periods. Another typical feature is the local concentration of calcium carbonate. Combined with the colour banding, the carbonate-enriched layer (Be) generally is developed within the more reddish horizon (see Fig. 8). It is characterized by diffuse carbonate accu mulations or distinct carbonate glacbules that can coalesce to form massive horizons (Fig. 4a). When pedogenesis is incipient, carbonate accumulation is diffuse. It is restricted to halos around the root traces (Fig. 6a,b). In the clay-rich facies the ped shapes are quite platy with subangular boundaries (Fig. 6d).
The carbonate peds are found in more mature profiles, where colour banding is marked. Their shapes range from irregular rounded to columnar, prismatic and blocky polyhedral. Irregular shapes are very common in slightly calcic horizons where the nodules are scattered, their size being less than 2 em in diameter. Well-developed nodules tend to be coalescent (Fig. 6c) (K horizon of Gile et al., 1966). Vertical fabrics are quite common, the stacked calcium carbonate nodules form columnar structures around the root traces and are typical of highly rooted palaeosols (Klappa, 1 980) (Fig. 6c). Prismatic structures have been rarely observed in these series, and where present in the lower part of the Be horizon, they are associated with blocky polyhedral carbonate-rich nodules (Fig. Sf). In the Maastrichtian deposits, numerous dinosaur egg clutches have been found in these alluvial deposits. Most of them are preserved in the accumulation horizon (Be) (Figs 4a & 6). Determination of the relationship between the eggshell and the carbonate nodules shows that the egg was laid before the carbonate nodule development. Successive layers with egg clutches within the Be horizon of the same palaeosol suggest periods of low sediment input between successive overbank alluvium deposits.
Carbonate-rich palaeosols
329
influenced by adjacent mineralized water bodies such as the main lake, a playa pond or infiltrated lake brines. The typical features of the palus trine facies affected by pedogenesis can be grouped in the same categories as those defined for the carbonate-rich palaeosols in the floodplain alluvium. Soil horizons
Fig. 5. Superposition of two palaeosols with well developed petrocalcic horizons (RH section). Identification of the B horizon in the lower palaeosol is based on the reddish colour of the sediment and the presence of slickensides.
On the basis of these macroscopic features, the cal cretes observed in the Provence Basin mainly corre spond to stages 2 and 3 in Machette's classification (1985). In the floodplain deposits, the calcretes can be identified easily on the basis of their macroscopic fea tures, but as will be shown later, some of the palus trine facies do look like these. Palustrine facies
The calcium carbonate palustrine facies are the most abundant in the Provence Basin. Other types of marginal lake facies correspond to evaporitic facies containing moulds of gypsum crystals or to phreatic calcretes or dolocretes (Colson & Cojan, 1996). Mineral distribution in these latter facies is
The palustrine facies also display horizonation. The original parent material (C horizon), a calcium carbonate-rich mud, can be distinguished from the modified sediment by its whitish colour. It generally grades up into the pedogenic horizon, which is char acterized by a faint but common mottling, defined mostly by pink, purple and yellow (Fig. 7d). With a substantial amount of terrigenous input, the lacus trine carbonate mud is mixed in its upper part with a sediment containing a higher amount of clay and sand grains. In this instance, the B horizon does resemble a floodplain calcrete, with carbonate nodules that are scattered in a reddish material (Fig. 7c). A clear lower boundary of the palustrine facies may constitute a good criteria for distinguishing these calcic soils from those in the floodplain alluvium. The gradual transition from the lacustrine facies to the palustrine facies indicates a relative low ering of the lake level. The numerous sequences of this type that are vertically stacked, however, suggest an aggradation with a permanent rise of the water table that was able to catch up with the sedimentation rate. This pattern seems to be exactly contrary of that described by Sanz et al. (1995) in the Madrid Basin, where short-lived lakes were flooded by alluvial deposits. Root traces
Root traces generally are vertical. Incipient pedogen esis is marked by vertical fissures, probably root moulds that give a prismatic fabric to the palustrine bed (Figs 4b & 7a). Horizontal cracks, 'the sheet cracks' (Freytet & Plaziat, 19 82) filled with sparry calcite, are often associated with the roots. In some facies, some large vertical structures, 2-20 em in diam eter and up to 70 em in length are present (Fig. 7b). These are interpreted as pedogenic stacked carbon ate nodules that precipitated around the roots. Similar structures are common in the Cameros Basin and were interpreted as roots of larger plants (Platt, 1989). This type of facies can easily look like a mature
330
I. Cojan
(a)
( c)
(d)
(f)
Fig. 6.
Carbonate-rich palaeosols developed in floodplain deposits. Location of sections are shown on Fig. 2. (a) Incipient pedogenesis in a silty sediment (CA). Main features are downward root mottles. The greyish pigmentation of the root mottle contrasts with the orange colour of the siltstone parent material. Frequent horizontal cracks, often filled with sparry calcium carbonate surrounding the vertical root mottles, are interpreted as probable rootlet moulds (black arrow). Hammer is 40cm in length. (b) More pronounced pedogenesis (CA). The downward fossil root traces show laterally branching rootlets. The vertical tubular structures are filled with marly sediment and the surrounding horizontal fractures are surrounded by a whitish halo, corresponding to a diffuse carbonate accumulation. Hammer is 40 em in length. (c) Calcrete with colour banding in a shaly silty material (RH). The palaeosol shows a well-developed petrocalcic horizon. The stacked carbonate nodules had coalesced vertically to build irregular tubules (rhizomorphs) around the greyish root traces. The filling of some of the root tubules has been reworked by burrowing organisms. Hammer is 40 em in length. (d) Pedogenesis in a clay-rich material (OLI). Subangular peds are outlined by darker surfaces enriched in clay (black arrow). Small darkened clasts are present throughout the sediment. Scale bar is l O cm in length. (e) Detail of Fig. 6(b): the scattered carbonate nodules weathered out in relief in the muddy sediment. Presence of the carbonate peds inside the dinosaur eggshell (black arrow) provides evidence that the egg was laid in the floodplain sediment before pedogenesis. Scale bar is 20cm in length. (f) Crudely polygonal peds at the base of a well developed petrocalcic horizon with many coalesced nodules (OLI). Note the presence of cracks filled with calcite. Scale bar is 5 em in length.
Carbonate-rich palaeosols
3 31
Fig. 7. Palustrine facies. Location of sections is shown on Fig. 2. (a) Palustrine limestone showing vertical prismatic structures (CA). The vertical cavities filled with greyish material are interpreted as probable root traces. Hammer is 40cm in length. (b) Rhizomorphs that formed from vertically coalescent stacked carbonate nodules (CA). Host palustrine sediments that contain more clays, weather out more rapidly (lower relief in the picture). Hammer is 40cm in length. (c) Outermost part of a shallow lake system, palustrine facies (CA). Note the sharp boundary between the palustrine deposits and the reddish silty sediments (black arrow). The original lacustrine mudstone has been modified extensively by pedogenesis and exhibits numerous nodules (1-5 em). TI1e upper part of each calcium carbonate palustrine bed is progressively enriched by terrigenous supply in fine-grained siliciclastic material. It can be identified by its higher clay content and its more pronounced reddish colour. Scale bar is l O cm in length. (d) Incipient pedogenesis in a palustrine facies with a low terrigenous content (CA). Mottling is subtle, displaying yellowish and pinkish colours. Root traces are filled with darker sediment and are outlined by whitish halos as well as horizontal cracks (black arrow). Scale bar is 5 em in length. (e) Polished slab through a palustrine mudstone moderately transformed by pedogenesis (VA). The pink colour mottling surrounds the vertical traces of the roots and contrasts with the whitish colour of the parent sediment and of the intraclasts. Bleached haloes developed around the vertical cracks, which are filled with calcite. Scale bar is 2.5 em in length. (f) Polished slab through a palustrine mudstone highly affected by pedogenesis (CA). A large pocket is filled with intraclasts formed by desiccation ofthe lacustrine mud and pedogenic nodules, which often contain darkened intraclasts (black arrow). Scale bar is 2.5 em in length.
calcrete Be horizon. Presence of roots within the sediment helps the mechanical dislocation of the sediment fabric during the drying-wetting cycles. Dead roots create vertical conducts that favour water circulation and dessication processes.
Soil structures
A distinctive feature of the palustrine facies is cer tainly the pseudokarst structures (Freytet & Plaziat, 1 982) (Figs 4b & 7e,f). These are typical of a sub-
I. Cojan
332
aerially exposed environment with low terrigenous input. They correspond to different stages of the dess iccation of a carbonate mud affected by pedogenesis. Early stages are characterized by root colonization and cracks (vertical and horizontal) that facilitated an in situ brecciation of the lacustrine mud (Fig. 7e). Later, deep dessiccations cracks, favoured by dead plant conduits, are filled with intraclasts that fell from the wall of the fissures (Fig. 7f). They correspond to a more pronounced pedogenesis under climatic condi tions that favoured long periods of prevailing low lake levels. Clast sizes range from 0. 5 mm to a few centimetres. Many clasts contain darker intraclasts (Fig. 7f), for which several origins have been pro posed and are still the subject of debate (see review in Platt, 1 992). The cavities are filled only partially with the intraclasts. The open voids are filled with muddy sediment or sparry calcite, which was deposited during further rise of lake level. Most often the carbonate accumulation is diffuse (Fig. 7d). When carbonate nodules developed, the palustrine facies can be identified on the basis of the matrix around the nodules, which is composed of a grey/ greenish marly mudstone or a pinkish carbonate-rich mudstone. When the terrigenous supply is high enough, the carbonate nodules are scattered in a reddish clayey-silty sediment difficult to distinguish from the floodplain alluvium. As for the floodplain facies, several stages of coalescence of nodules can be observed and these palustrine facies are then difficult to tell apart from the floodplain calcretes.
PALAEOSOL MINERALOGY
In an attempt to assess the amount of mineralogical transformation in the soil profiles, a mineralogical study based on their bulk rock composition and clay mineralogy has been carried out. Channel-belt facies underwent slight pedogenic modification so that they have been considered as a good proxy for the unweathered substrate. Time-equivalent, well developed palaeosols on the floodplain or palustrine facies are compared with these (Fig. 8). Method
Bulk rock compositiOn and clay mineralogy were determined by X-ray diffraction analysis (XRD), using a Philips PW diffractometer with Cu Ka radia tion. Four sets of diffraction patterns were used: air-
dried, glycolated, hydrazin saturated and one heated at 500°C for 3 h. The estimation of the clay mineral content in the < 21-1.m fraction was determined by comparison of the main peak surfaces of the glyco lated XRD diagram. Estimated error for both bulk rock and clay mineralogies is = 1 0% . The parent material: the floodplain alluvium deposits
The mineralogical composition of the floodplain allu vium deposits is fairly stable. Samples from the silty material of the channel-belt area, where the pedo genic imprint is faint, show an average content of 2 5% in calcium carbonate, around 1 5% in quartz and up to 60% in clay. Detrital grains are predominantly quartz ( 25%) and calcite silt ( 50%) with minor amounts of K feldspars and rock fragments (igneous rocks, schists, carbonate rock fragments). The clay assemblages are dominated by the presence of smectite (50%) and illite (30% ), with minor amounts of kaolinite ( 5%), chlorite ( 5%) and mixed-layer illite-smectite (1 0% ). Interpretation
The relatively low content of quartz is interpreted as characteristic of depositional dynamics dominated by suspension, sediment being brought into the flood plain during overbank flows. The abundance of detri tal carbonate constitutes a specific aspect of these series. Some carbonate clasts eroded from the surrounding Mezosoic massifs are found in the silty alluvium deposited against them. Within the sand fraction, the major part of carbonate is detrital. Within the Palaeocene, in situ or reworked microco dium prisms can be found. In the finer grained facies it is not possible to differentiate detrital matrix from later cement. The floodplain carbonate-rich palaeosols
In the calcrete profiles showing carbonate nodules that are not coalescing, both bulk rock and clay mineral compositions have been modified previously (Fig. 7b). The bulk rock is composed mainly of calcium carbonate (around 80% ), quartz ( 5%) and clay minerals (1 0-1 5% ). The clay assemblage is characterized by a significant increase of mixed-layer illite-smectite ( 25%) to the detriment of the illite content.
Carbonate-rich palaeosols CHANNEL BELT FACIES (RI)
colour lithology
samp.
bulk rock
333
FLOODPLAIN FACIES (RH)
clay minerals
composition
colour lithology samp. bulk rock composition (m)
PALUSTRINE FACIES (RH)
day minerals
to
colour lithology samp. bulk rock composition
--, 1 2 ·=-
clay minerals
j
10·� •·
..
\00
Sediment colour
Bulk rock
100
..
0
Clay minerals
composition
� brown
� red !Z2I orange � yellow c::J p;nk
E3 kaolinite ITllJ chlonte B elays E;:::J illite � quartz i D �i���a:c�r:e c=J calcite i c;:;) dolomite � �i��h���:· c=J smectite
I�
0
Pedogenic features
'
carbonate nodules m�ttling ss slickensides
1 00
..
0
..
100
Fossils
dinosaurs eggshells { burrows
..
100
Fig. 8.
Channel-belt facies and palaeosols developed in floodplain alluvium and palustrine facies. Comparative vertical facies succession and mineralogical distribution in each of the environments.
Interpretation
These results suggest that diffuse carbonate accumu lation is significant in the B horizon. The clay mineral assemblages also are modified. The detrital assem blage is developing into authigenic interstratified illite-smectite and smectite. In very mature profiles, smectite and palygorskite comprise the entire clay fraction (Colson et al. , 1998). The palustrine facies
In this study, we consider only the palustrine facies associated with large water bodies, because in pond or ephemeral lakes, the mineralogical association (bulk rock and clay minerals) reflect the chemistry of the pondwater more than the pronounced leaching favoured by the stability of the profile through time (Colson & Cojan, 1996).
Around the permanent lake area, palustrine facies are composed of 8 0% calcium carbonate and 20% clay. Detrital grains, such as quartz are rare. The clay assemblage in the < 2J..Lm fraction is dominated by smectite (50%), then mixed layers ( 2 5% equally dis tributed between illite-smectite and illite-chlorite) and illite (10% ) . Interpretation
In the palustrine facies, the original high carbonate content favoured the coalescence of the pedc,genic nodules over shorter periods of time than in the floodplain facies. The clay mineral assemblage, however, reflects a transformation that is similar to that observed in the calcretes with scattered nodules that developed on floodplain alluvium, and suggests that the duration of the pedogenesis was of compar able length in both environments.
I. Cojan
334 CONCLUSIONS
Carbonate-rich palaeosols developed in flood plain and palustrine environments are most often easy to distinguish on the base of macroscopic descriptions. Some stages of development, how ever, can be indistinguishable between the two environments: 1 well developed calcic horizons (Bc-K) in flood plain sediments do resemble palustrine facies that underwent pedogenesis under relatively short periods of desiccation; 2 palustrine facies with a high terrigenous input show scattered carbonate nodules in a reddish shaly sediment that looks like floodplain carbonate-rich palaeosols. A mineralogical study of the bulk rock composi tion and the clay assemblages of both facies shows similar processes in the B horizon: carbonate accu mulation and replacement of the detrital clay miner als by authigenic ones. In the Provence lacustrine and floodplain sediments, despite the development of multiple soil profiles, plus the erosion of topsoils, the system was dominantly vertically aggrading with time. The super posed sequences of carbonate-rich palaeosols can be considered as a good record to investigate the auto cyclicity of the fluvial-lacustrine system, the climatic changes through time and the tectonic evolution of the basin.
AKNOW LEDGEMENTS
N. Platt and J.P. Calvo are warmly thanked for their helpful and constructive comments during the review of the manuscript. I wish to thank also M. Thiry and J. Colson for our numerous discussions on these carbonate-rich palaeosols.
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spore assemblages of the Uppermost Cretaceous conti nental formations of south-eastern France and north eastern Spain. Cretaceous Res. , 13, 1 1 9-132. MILNES, A.R. (1992) Calcrete. In: Weathering, Soils and Paleosols (Eds Martini, J.P.& Chesworth, W.), Elsevier, Amsterdam, 309-348. NETTERBERG, F. (1980) Geology of Southern African cal cretes: 1 . Terminology, description, macrofeatures, and classification. Trans. geol. Soc. S. A/1:, 83, 255-283. PIPUJOL, M.D. & BuuRMAN, P. (1993) The distinction between ground-water gley and surface-water gley phenomena in Tertiary paleosols of the Ebro basin, NE Spain. Palaeogeog1: Palaeoclimatol. Palaeoecol. , 110, 103-1 13.
PLATT, N.H. (1989) Lacustrine carbonates and pedogenesis: sedimentology and origin of palustrine deposits from the Early Cretaceous Rupelo Formation, W Cameros Basin, N Spain. Sedimentology, 36, 665-684. PLATT, N.H. (1992) Fresh-water carbonates from the Lower Freshwater Molasse (Oligocene, western Switzerland): sedimentology and stable isotopes. Sediment. Geol. , 78, 81-99.
(1992) Distal alluvial deposits in a foreland basin setting-the Lower Freshwater Molasse (Lower Miocene), Switzerland: sedimentology, architec ture and palaeosols. Sedimentology, 39, 545-565. PLATT, N.H. & WRIGHT, V.P. (1991) Lacustrine carbonates: facies models, facies distributions and hydrocarbon aspects. In: Lacustrine facies Analysis (Eds Anad6n, P., Cabrera, Ll. & Kelts, K.), Spec. Pubis Int. Ass Sediment. , 13, 57-64. Blackwell Scientific Publications, Oxford. PLATT, N.H. & KELLER, B.
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Evidence from paleosols for ecosystem changes across the Cretaceousffertiary boundary in eastern Montana, Geology, 15, 1090-1093.
RODAS, M., LUQUE, F.J., MAS,R. & GARZON, M.G. ( 1994) Cal cretes, palycretes and silcretes in the Paleogene detrital sediments of the Duero and Tajo basins, Central Spain. Clay Miner. , 29, 273-285.
M.E., ALONSO ZARZA, A.M. & CALVO, J.P. (1995) Carbonate pond deposits related to semi-arid alluvial systems: examples from the Tertiary Madrid Basin, Spain,
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marine limestone within cyclothems in the Pennsylvanian (Upper Freeport Formation, Appalachian Basin) and its implications. In: Lacustrine Reservoir and Depositional Systems (Eds Lomando, A.J., Schreiber, B.C. & Harris, P.M.), pp 321-381. Core Workshop 19, Society of Eco nomic Paleontologists and Mineralogists, Tulsa. WATTS, N.L. (1980) Quaternary pedogenic calcretes from the Kalahari (Southern Africa): mineralogy, genesis and diagenesis. Sedimentology, 27, 661-686. WRIGHT, V.P. & ALONSO ZARZA, A.M. (1990) Pedostrati graphic models for alluvial fan deposits: a tool for inter preting ancient sequences. ]. geol. Soc. London, 147, 8-10. WRIGHT, V.P. & P LATT, N.H. (1995) Seasonal wetland carbon ate sequences and dynamic catenas: a re-appraisal of palustrine limestones. Sediment. Geol. , 99, 65-71. WRIGHT, V.P. & ROBINSON, D. (1988) Early Carboniferous floodplain deposits from South Wales: a case study of the controls on paleosol development. J geol. Soc. London, 145, 847-857.
& TuCKER, M.E. (1991) Calcretes: an introduc tion. In: Calcretes (Eds Wright, V.P. & Tucker, M.E.), Reprint Series int. Ass. Sediment., No. 2, pp. 1-22. Black well Scientific Publications, Oxford.
WRIGHT, V.P.
Spec. Pubis int. Ass. Sediment. (1999) 27, 337-366
Sedimentary infillings and development of major Tertiary palaeodrainage systems of south-central Australia
N. F. A L L E Y * , J. D. A. C L A R K E t , M. M AC P H A I L: !: a n d E . M . T RU S W E L L § *Primary Industries and Resources SA, GPO Box 1671, Adelaide, South Australia 5001, Australia; tCRC LEME, Australian Geological Survey Organization, PO Box 378, Canberra;ACT 2601, Australia; tResearch School ofPacific and Asian Studies, The Australian National University, Canberra, ACT 0200, Australia; and §Australian Geological Survey Organization, PO Box 378, Canberra, ACT 2601, Australia
ABSTRACT
Tertiary palaeochannels are widespread on the Australian continent. Their best preserved sedimentary infillings are found in the Eucla Basin and central Australian area. Palaeochannel development had its origins during earliest Cretaceous times in the south-western Eucla Basin and at least in possibly Late Cretaceous times in the central continent. Major phases of sedimentary infilling occurred in Palaeocene-Eocene, late Oligocene-Miocene and Pliocene-Pleistocene times. Marine influence extended several hundred kilometres up the Eucla palaeochannels during at least three major transgressions in the middle Eocene-late Eocene interval. Reduced marine influence occurred in some eastern Eucla channels during the Early Miocene Epoch. The sedimentary and geomorphological evidence indicates that no con nection existed between the Eucla and inland channels. Deep weathering was prevalent prior to deposition in the channels, and may be as old as early Meso zoic times. Later weathering was related to duricrust development. Ferricrete probably formed in early Mesozoic, late Oligocene-Middle Miocene and Late Miocene-Pleistocene times. Major phases of silicification occurred in late Eocene-Middle Miocene and Late Miocene-Pleistocene times, when significant groundwater silcrete formed. Temperate rainforest existed along the southern continental margin during earliest Palaeocene times. By the late Palaeocene to early Eocene interval, rainforest of megathermal aspect existed in central Aus tralia, indicating that conditions there were warmer than along the southern continental margin. In middle Eocene times, monsoonal-like conditions prevailed in central Australia and moister conditions in the south, where rainforest of meso- to megathermal aspect grew, here extending late into the Eocene Epoch. The ?late Oligocene-Miocene interval was a time of development of extensive shallow, alkaline lakes in parts of the palaeochannels and in two major depocentres in central Australia. Lakes in the inland area supported a diverse fauna, including crocodiles. Vegetation had changed to dry, open woodland through out the palaeochannel areas, with rainforest-like vegetation confined to wetter valley bottoms. By the Pliocene Epoch further drying had produced a chenopod shrub to open woodland environment, contain ing isolated pockets of forest in edaphically suitable sites.
INTRODUCTION
Tertiary strata occur throughout much of onshore and offshore Australia, occupying gently down warped basins and rifted troughs, infilling palaeo river channels, and occurring as widespread thin sheets in the interior (Fig. 1). Where exposed, sedi ments are often highly weathered, silicified and ferruginized.
The existence of extensive Tertiary palaeodrainage systems in Australia has been known for almost a century (Fig. 2). Their presence was first hinted at by Carnegie (1898), who suggested that should the rain fall be greater than present, the playas in southern Western Australia might form connected channels and flow to the Eucla Basin area. Not long after,
Palaeoweathering, Palaeosurfaces and Related Continental Deposits. Edited by Médard Thiry and Régine Simon-Coinçon © 1999 The International Association of Sedimentologists. ISBN: 978-0-632-05311-7
337
338
Fig. I.
N F. Alley et al.
Tertiary sedimentary basins in Australia referred to in the text.
Gibson (1909) concluded that the playas marked the courses of ancient rivers, although this was hotly dis puted and alternatives such as tectonic valleys or relict estuaries and drowned valleys were preferred. Now, however, the concept is well accepted and Cenozoic palaeochannels and their infillings have been recognized from many parts of the Australian continent (e.g. van de Graaff et al., 1977; Langford et a!., 1995). As a result of State and Federal Government mapping programmes (e.g. Pitt, 1980) and the search for Tertiary placer deposits, the extent, stratigraphy, sedimentology and geomorphological, palaeocli matic and weathering implications of the palaeo channels have become much better known. A consequence of the rifting of Australia from Antarctica was the initiation of marine transgressions
into the southern basins and the deposition of consid erable thicknesses of marine sediments. Thus, the Ter tiary succession here passes upwards and landwards from Early Tertiary temperate water limestones into marginal marine sediments and then palaeochannel infillings. Because the palaeodrainage system here was exorheic, sea-level changes had a significant influence on phases of channel infilling. Deposition of relatively thin, fluvial and lacustrine, carbonaceous and arenaceous sediments character ized the Palaeogene inland depocentres and related palaeochannels (Benbow et al., 1995a; Alley, 1998). Thin lacustrine argillaceous and carbonate mud stones were laid down during the Neogene Epoch. The landscape in which these continental sediments were deposited was generally subdued. Uplands were located in similar areas as now and, together with
Tertiary palaeodrainage systems
339
Fig. 2 . Distribution of palaeochannels and drainage divides in southern Australia. (Modified from Langford et al., 1995 and Alley & Lindsay, 1995a.)
the Great Dividing Range, were major sources of terrigenous sediment. Dating and correlating the non-marine sediments, particularly in the central continent, relies heavily on palynology. In this paper dating and correlation uses the palynological zones of Stover & Partridge (197 3, 1982) and Macphail et al. (1994). There is difficulty in dating the younger part of the Tertiary succession because of weathering, facies that do not preserve palynofloras well and the increasing regionalization of vegetation through the Tertiary Epoch, making correlation with the dated southern succession difficult. Dating Palaeogene sequences is also compli cated by the earlier first appearances of some species
in interior sequences than in the southern basins, causing some problems in correlation with dated . sequences (Alley & Benbow, 198 9; Alley & Beecroft, 1993; Macphail et al., 1994) The sedimentary infillings of the palaeochannels contain important evidence that bears on unravelling landsurface evolution, particularly the age of duri crusts, and the nature of palaeoclimate during the Tertiary Epoch. The aims of this paper are to: 1 discuss the distribution of the palaeochannels centred on the Eucla Basin (an exorheic system) and possibly an endorheic system in central Australia draining largely into the Lake Eyre Basin; .
340
N F Alley et al.
examine the sedimentology of the palaeochannels in three time slices (a) Palaeocene to earliest Oligocene (b) middle Tertiary (c) Pliocene-Quaternary; 3 elaborate on the age and development of duricrusts from the evidence found m the palaeochannels; 4 elucidate palaeoclimate and palaeogeography of the southern and central parts of the continent in the context of palaeochannel development.
2
EUCLA BASIN PALAEODRAINAGE
Tertiary sediments in the Eucla Basin occur in three broad settings: an offshore rift-margin area contain ing marginally marine terrigenous clastic deposits succeeded by mainly deep-water pelagic carbonate accumulations, a shallow-water platform on which neritic carbonate and inner platform non-marine to marine terrigenous sediments were deposited, and a vast region of palaeodrainage fringing the basin and preserving alluvial, lacustrine, evaporitic, aeolian, col luvial, marine and paralic sediments (Clarke, 1 993; Benbow et al., 1 995b). The palaeochannels around the platform margin, although partially obscured by a mantle of Quaternary sediments, are remarkably intact. The Eucla Basin palaeochannel infillings extend from the eastern basin (Gawler Craton area) to the western basin margin (Yilgarn Craton area), these areas having slightly differing stratigraphies (Figures 1-3). Three major phases of deposition are present in almost all channels, and probably are equiva lent to second-order cycles in the marine record. These phases are of Eocene (to possibly earliest Oligocene), Oligocene-Miocene and Pliocene Quaternary ages (Fig. 4). The tripartite stratigraphi cal division can be recognized even in minor tribu taries in elevated areas. Cretaceous strata are present locally in lower reaches of some channels of the Yilgarn Craton area. Middle Eocene to early Oligocene facies
The Eocene succession (Figures 3-5) overlies older rocks along an erosional disconformity and is present throughout the palaeodrainages. Although a basal Early Cretaceous infill is present in some palaeo channels in the western part of the basin (see below), the bulk of the channel sediments is Tertiary. Deposi-
tion of the oldest part of the Tertiary succession com menced at least in middle Eocene times, during the Wilson Bluff Transgression in the east (Alley & Beecroft, 1 993; Benbow et al., 1 995b) and Tortachilla Transgression in the west (Clarke, 1 994a), continuing through the Tuketja and possibly Aldinga Transgres sions. These eustatic events, along with climatic changes and tectonism, have produced a complex of marine to non-marine facies, differing slightly in the channels from east to west across the basin (Figs 3 & 4). Facies E-1 comprises non-marine fine to very coarse sand and gravel of the Pidinga and Werrilup Formations. Sand packages are commonly 20-30 m in thickness, but may be significantly thicker on central Eyre Peninsula owing to syndepositional subsidence. The sand typically fines upwards from basal cobble lags and may pass laterally and vertically into fine grained and lignitic facies with common fossil wood and leaves. Lignitic sediments (facies E-2) in the palaeo drainage systems reach 40-50m in thickness, and the lignites themselves can approach 20m in thickness (Elms et al., 1 982; Rankin & Flint, 1 991). The pres ence of rare to common dinoflagellates indi cates marine influence during deposition (Alley & Beecroft, 1 993). The lignitic facies commonly grade laterally into non-marine clastic sediments, but in the eastern basin grade occasionally laterally into marine carbonate and spicular-bearing carbonaceous clastic sediments; they invariably grade vertically into marine sediments. Palaeontological evidence indi cates that the lignitic facies were deposited during the Wilson Bluff and Tortachilla Transgressions in the eastern channels and mainly in the Tuketja Transgres sion in the west. The non-marine clastic sediments and lignites are interpreted as being deposited in an aggrading fluvial to estuarine plain as part of early transgressive systems tracts. Clastic sediments are the most common marine facies deposited during the Wilson Bluff and Tor tachilla Transgressions (facies E-3). Fossil content is generally low, but includes calcareous and sili ceous sponge spicules, forams, molluscs, bryozoans, dinoflagellates and fossil wood (Lowry, 1 970;Alley & Beecroft, 1 993; Clarke, 1 994a). Marine sediments in the Cowan palaeodrainage system pass vertically and laterally into shallow-marine limestone of the upper Norseman Formation, although similar facies are found sporadically in the Bremer Basin (Clarke et al., 1 996). In a few localities in the eastern Eucla Basin the marine clastic sediments may pass laterally and
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MESA
Schematic facies and stratigraphical relationships in palaeodrainage systems in the Eucla Basin.
vertically into thin Wilson Bluff Limestone inter bedded with carbonaceous clastic sediments (Alley & Beecroft, 1993; Benbow et al., 1 995b). Marine clastic facies dominate the sediments deposited during marine incursions into most palaeodrainage systems because of dilution of bio genic sediment by siliciclastic sediment trans ported down the palaeodrainage systems. Biogenic sediments normally accumulated only in those palaeodrainage systems lacking significant terrige nous influx, such as the Cowan palaeodrainage system in the western Eucla. Distinctive marine facies (facies E-4) comprise calcareous sandstone, coarse bryozoan grainstone, trough cross-bedded grainstone and fine bryozoan grainstone (Alley & Beecroft, 1993; Benbow et al., 1995b; Clarke et al., 1996). Fragments of bryozoans, coralline algae, molluscs, echinoids and forams are
the dominant constituent of the grains (Fig. 6). Car bonate and terrigenous muds are minor components, occurring as lenses with coarse and cross-bedded car bonates, in the deepest part of the palaeodrainage systems in the western Eucla, or in sheltered, near shore localities. The shallow-marine limestones of the Norseman Formation (Clarke et al., 1996) and Wilson Bluff Limestone (Benbow et al., 1995b) are interpreted to have been deposited in a tide-dominated embay ment with a depth range of 10-90m, as indicated by the abundance of shallow-water indicators, such as coralline algae and large benthic forams. In the eastern Eucla the limestone facies is intercalated with the lignitic facies of the Pidinga Formation, further indicating restricted, shallow water conditions. The shallow-marine spongolites (facies E- 5) of the Princess Royal Spongolite (western Eucla) and Bring
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Fine grainstone facies of Norseman Formation dominated by bryozoans and echinoids in the Cowan palaeochannel. Width 14mm.
Member of the Pidinga Formation (east) are domin ated by siliceous sponge spicules, varying in abun dance from 2 5 to 99%; occasional sponge body fossils are also present (Benbow, 199 3; Clarke, 1994a; Figs 7 & 8). An older spicular facies may be present in the eastern palaeochannels. In the western Eucla the spongolite is generally massive to finely laminated but cross-bedded, herring bone and unidirectional cross-laminated, and conglomeratic lenses are present locally, adjacent to bedrock highs. The pres ence of diatom faunas are indicative of mangrove and sea-grass environments (Thomas, personal com munication, 1989), which are present in the most upstream portions of the palaeodrainage systems in the western Eucla. The bioturbated spongolite facies has common Thallasanoides burrows and is present
344
N F Alley et al.
Fig. 7.
Princess Royal Spongolite from the Lefroy palaeochannel. Width 14 mm.
in the more seaward reaches of the palaeodrainage systems. A diverse fauna biofacies containing Mol lusca, together with forams, echinoids and bryozoans, is indicative of more open coastal environments. The spongolite facies always occurs in weakly tidal coastal embayments to estuarine settings in water depths typical of the photic zone, especially for diatom-containing spongolites. Palaeodrainage geometry suggests deposition at depths of 25-80m (Clarke, 1994a). Facies E-6 (Ooldea Sand) comprises aeolian quartz sand deposited during regression in latest Eocene to early Oligocene times (Benbow, 1990; Benbow et al., 1 995b ). The sand forms the large dunes of the Ooldea and Barton Ranges and spills over into some palaeochannels to interfinger with the carbona ceous and marine clastic sediments of the Pidinga Formation. Oligocene-Miocene? facies
Fig. 8. Cross-bedded, sponge-spicule bearing Bring Member of the Pidinga Formation. ( Photograph courtesy of M.C. Benbow. )
Overlying the Eocene sediments or bedrock along an erosional contact is an Oligocene-Miocene? succes sion comprising Garford and Narlaby Formations in the eastern Eucla, and Revenge and Gamma Island Formations and Cowan Dolomite in the west (Figs 3 & 4). These sediments comprise a number of facies ranging from coarse- and fine-grained clastic packages, dolostone, oolitic ironstone, carbonaceous clastic sediments, and marine clastic and carbonate accumulations. Together these are normally l O rn in thickness, except where there has been extensive channelling of the substrate, or in the Malbooma Kingoonya palaeochannels, where up to 8 0m may be present (Benbow et al., 1995b). Although the dating controls over the age of this part of the succession are not not well established, correlation of the dolostone facies with the Etadunna Formation of the Lake Eyre Basin (see below) and limited palynological evidence from the eastern palaeochannels, indicate an age range from latest Oligocene to earliest Pliocene (Harris, 197 3; Truswell & Harris, 1982; Benbow et al., 1995b;Alley, 1996). Facies 0-1 is coarse-grained sand and minor gravel. They are common along palaeodrainage margins and in tributary channels, where they rest on an eroded bedrock surface. These lithologies thin and fine towards the middle of the main channels. Within the larger palaeochannels, sand units are typically confined to small lenses. Most sand and gravel, espe cially where exposed along the margins of the lakes, are ferruginized. Marine glauconitic sand and silt are
Tertiary palaeo drainage systems
present in some channels incised into the Eocene suc cession. This unit probably passes laterally into shallow-marine Nullarbor Limestone. Silt and mud (facies 0-2) are absent from the valley sides but partly fill the main valleys and their tributaries. The sediments are massive to finely laminated and commonly contain concretions of haematite. In the Tarcoola area and Eyre Peninsula the mud may exceed 50m in thickness. Oolitic iron stone has been recognized in Lake Lefroy, there consisting of thin lenses of flattened microscopic goethitic ooids and peloids. The ooids have com monly nucleated about quartz grains or ferruginized plant fossils, and were then cemented by goethite and siderite. Minor lignite (facies 0-3) occurs in marginal areas of the western Eucla Basin (Frewster & Denman, 1 984) and thin carbonaceous lenses in Lake Cowan. In the Yaninee Palaeochannel near Port Kenny the facies is a carbonaceous grit and sand 1 0-20m thick containing abundant marine dinoflagellates indica tive of estuarine conditions (Alley, 1 996). Facies 3 may grade laterally and vertically into sand, grit, mud or dolostone facies. Dolostone facies (0-4) are extensive in the palaeochannels and are undoubtedly correlative with the Etadunna Formation of the palaeochannels of the Lake Eyre B asin (see below). In the eastern part of the Eucla B asin the dolostones are unnamed; in the west they are named Cowan Dolomite and Gamma Island Formation. A number of lithologies are present, but stratigraphical relationships are poorly known because of discontinuous exposure. Lithologies include massive microcrystalline dolo stone, pisolitic, stromatolitic, to oncolitic rudstones and floatstones, oolites and breccias. Locally, the dolostones pass laterally into dolocretes. The range of coarse- and fine-grained facies, dolo stones and ironstones are interpreted as fluvial and lacustrine depositional environments. The marine facies are regarded as indicators of shallow water environments where biogenic carbonate production was diluted by the influx of terrigenous sediments from the palaeochannels.
345
Polar Bar and Roysalt Formations in the west. Four facies are recognized, including evaporites, coarse and fine-grained sediments and poorly sorted pebbly muds. The basal part of the succession has been dated palynologically as earliest Pliocene in the Lefroy, Tay and Cowan palaeochannels (Bint, 1 981; Clarke, 1 993; C. B. Foster, personal communication, 1 994). The full age range of this part of the succession, however, is unknown. Facies P-1, the evaporites, vary in thickness up to at least 9m, but locally exceed 20m in areas of dune accumulation; in the eastern palaeochannels the fine grained facies are gypseous or capped by a gypsum crust. For the most part the facies is composed of flat bedded gypsum crystals, with an organic-rich, partly calcareous mud matrix (Fig. 9). The dune deposits are composed of cross-bedded and cemented, coarse
Pliocene-Quaternary facies
Sediments of this age are the thinnest found in the palaeodrainages and rarely reach l O rn in thickness, overlying the middle Tertiary units along an erosional surface (Fig. 4). They are found within the Ilkina and Munjena Formations in the eastern channels and
Fig. 9. Laminated and structureless muds and laminated gypsum (top) of the Ilkina Formation in the Narlaby palaeochannel, eastern Eucla Basin. (Photograph courtesy of M.C. Benbow. )
N F Alley et a!.
346
sand-sized gypsum crystals. Up to three dune build ing cycles have been observed in the Lefroy channel (Clarke, 1994b); polygenetic gypsiferous lunettes are widespread in the eastern palaeochannels. In the upper part of palaeochannels adjacent to the Mus grave Ranges non-marine limestone (P- 5; Mangatitja Limestone) may be equivalent to the evaporitic facies. Coarse-grained siliciclastic sand and gravel (facies P-2) occur in dune fields and alluvial channels. Fine-grained facies (P-3) consisting of silt and clay form ground-mantling deposits across broad areas of low relief. They are massive to poorly stratified and extensively bioturbated by soil organisms. Poorly sorted, structureless to crudely stratified pebbly mud, conglomeratic sand, clayey sand and breccia (facies P- 4) occurs as wedges and sheets sur rounding low hills. The clasts are all derived from the local bedrock. This facies is commonly mottled and sometimes partially cemented by dolocrete in the west; pedogenic silcrete occurs in the facies in east channels, and the presence of rounded silcrete clasts indicates reworking from older silcretes. Sediments deposited during this phase are indica tive of arid conditions. The evaporitic facies infer deposition in salt lakes and deflation of gypsum from these lakes to form dunes. Coarser clastic sediments are related to channel, alluvial fan and sand dune deposition, whereas finer sediments represent depo sition of wind-blown dust and overbank flooding on alluvial plains. The poorly sorted pebbly muds were deposited as colluvial aprons around hills and bedrock exposures.
CENTRA L AUSTRALIAN BASINS
Lake Eyre Basin palaeodrainage
Following deposition of the Winton Formation in the Cenomanian and up until late Palaeocene times, weathering and erosion appear to have prevailed for a large part of central Australia. During late Palaeocene times, tectonic subsidence in north eastern South Australia produced the large Lake Eyre Basin, in which episodic fluvial and lacustrine sedimentation has taken place up until the present (Krieg et al., 1990; Callen et al., 1995; Alley, 1998). Much of the basin surface is close to sea-level and the topography is very subdued. Outcrop of Cenozoic sediments is poor, capping Mesozoic strata in areas of uplifted and eroded domes, occurring as low cliffs
around salt lakes, or as discontinuous silicified palaeochannels that once flowed into the Lake Eyre Basin. The palaeochannels preserve alluvial, lacus trine, evaporitic, aeolian and colluvial sediments. There are three main phases of deposition in the Tertiary palaeochannels (Figure 1). The first phase was in early Tertiary (late Palaeocene to middle Eocene) times, when sandstone, carbonaceous clastic sediments and conglomerate were deposited. The second phase from (?)latest Oligocene to Miocene times, comprised clay, fine sand and carbonate, with lesser conglomerate. The third phase, ?Pliocene to Quaternary times, was characterized by deposition of sand and sandy clay. These phases of deposition were controlled largely by tectonism and palaeoclimatic changes. Late Palaeocene to middle Eocene facies
The lower Tertiary succession includes the Eyre For mation and its regional correlatives, overlying Mesozoic and Palaeozoic sediments and Precam brian bedrock along an erosional surface. Interplay of tectonic and climatic influences produced an array of alluvial and lacustrine facies in the Eyre Formation of the greater Lake Eyre Basin, but only alluvial facies were laid down in the palaeochannels. The dis cussion below deals largely with the palaeochannel fills, except where the greater basin fill sheds light on the channel facies and the palaeoenvironmental interpretations outlined in a later section. Palynological dating of the channel facies along the southern margin of the Lake Eyre Basin indicates that they are middle Eocene in age (Alley et al., 1996), although alluvial facies in deeper parts of the basin are as old as late Palaeocene in age (Wopfner et al., 1974; Krieg et al., 1991; Callen et al., 1 995; Alley, 1998). A distinctive basal pebble conglomerate set in a sand matrix (facies E-1) and measuring up to 2 0cm thick occurs throughout the basin (Figs 1 0 & 11). A diagnostic feature of the facies is the presence of highly polished, resistant lithologies, such as milky quartz, chert, jasper, agate, fossil wood and silcrete, with pebbles of basement rock in some localities. Facies E-2 is a cross-bedded quartz sandstone, con taining thin interbeds of silt and clay and occasional pebbles, that conformably overlies the basal gravel. In deeper parts of the basin the sandstone facies may exceed 35m in thickness, but in the palaeochannels is usually 5-1 0m thick. Although the sandstone in outcrop is usually weathered and often kaolinitic,
Tertiary palaeodrainage systems WEST
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Winton Fonnation: SlighHy fissile dark grey clay with fine siltstone/sandstone laminae
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_ _ _ _ _ _ _ _ _ _ _ _
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Channel fill of Eyre Formation overlying unweathered Winton Formation, Lake Eyre South area. ( Modified from Alley et al., 1996.)
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where fresh it may be carbonaceous and contain mummified leaves, twigs and wood. Facies E-2 may overlie, underlie or laterally interfinger with facies 3. The sandstone in Nelly Creek overlies facies E-3 (Fig. 10). In the Poole Creek palaeochannel the sandstone facies contains the remarkable 'silcrete floras' (see below). Intercalated, carbonaceous, medium- to fine grained sand, silt and clay (facies E-3) occur as lenses and broader, more elongate bodies in the palaeochannels and were deposited in backwater swamps and shallow lakes on the floodplains. Although considerably thicker in deeper parts of the Lake Eyre Basin, this facies is generally quite thin, but reaching 15m in thickness in the Nelly Creek channel. Large mummified logs, branches, twigs, leaves and fruits are common in this facies; in the lower reaches of Nelly Creek the facies contains a leaf flora of middle Eocene age (Christophel et al., 1992;Alley et al., 1996). This facies undoubtedly is the lateral equivalent of the sandstone with silcrete floras in lower Poole Creek palaeochannel. In the southern Callabonna Sub-basin, facies E-3 is a carbonaceous, pebbly medium sand forming the basal unit of the Eyre Formation in the palaeochan nels. In this area this facies is overlain by facies E-2, a moderately to poorly sorted medium sand with significant bodies of silt and clay (Callen, 1990; Curtis et al., 1990). Sedimentary evidence suggests that the Eyre For mation was deposited largely by braided streams, accompanying epeirogenic uplift of the Olary Block
_ _ _
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Fig. ll. Basal pebble layer ( c. l O cm thick ) of the Eyre Formation at its type section near lnnamincka, central Lake Eyre Basin. Disconformably overlying Cenomanian Winton Formation.
N F Alley et al.
348
and the Barrier Ranges together with subsidence in the Lake Eyre Basin (Krieg et al., 1 990) . Near the basin margin the channels are more discrete and exhibit meandering flow conditions. Some erosion of the Eyre Formation probably occurred during early Eocene times, followed by fluvio-lacustrine deposi tion in middle Eocene times (Wopfner et al., 1 97 4) . ? Latest Oligocene to Miocene facies
the Oligocene Epoch to the Early-Middle Miocene Subepochs (Stirton et al., 1 961; Woodburne et al., 1 98 5; Woodburne & Clemens, 1 986) . Palynofloras from the formations are sparse but suggest an age range from Miocene to ( ? ) Pliocene Epochs (Callen & Tedford, 1 976; Morgan, 1 977; McMinn, 1 981; Martin, 1 990; Callen et al., 1 995) . As correlative facies occur in the palaeochannels, the above age ranges are also applicable td the middle Tertiary channel facies. Although a variety of the middle Tertiary facies are well developed in deeper parts of the Lake Eyre Basin, they are only poorly represented in the palaeochannels, the best occurrences being along the southern margin of the Tirari and Callabonna Sub-basins. These facies lie on an eroded surface developed on Mesozoic and older strata and within channels eroded into Eyre Formation and the middle Tertiary Watchie Sandstone of the Billa Kalina Basin. Four main facies, together measuring up to 20m in thickness, are present in a few palaeochannels; the facies include sandstone, limestone/dolomite, clay stone and siltstone grading up to sandstone. Facies 0-1 is a sand that usually forms the base of the succession (Fig. 1 2) . This facies is silicified and contains silicified sandstone pebbles and boul-
( Figure 1 2)
The middle Tertiary dolomite and clastic sediments of the Etadunna and Namba Formations overlie an erosional surface developed on the Eyre Formation and Mesozoic sediments. The Etadunna and Namba Formations, as derived from evidence from deeper parts of the Lake Eyre Basin, may range from late Oligocene to Pliocene in age, and thus so might the Willalinchina Sandstone (see below) . An Rb-Sr date on diagenetic illite in the upper part of the Etadunna Formation approximates the Oligocene-Miocene boundary (Norrish & Pick ering, 1 98 3) , whereas a monospecific foraminiferal microfauna has been ascribed tentatively a late Oligocene age (Lindsay, 1 987) . Estimates of the age of fossil vertebrates from the formations range from
Mlddle Poole Creek Palaeochannel
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0
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100 90 80
70 60
common
00 �------�
70
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Fig. 12. Etadunna Formation filling
channel cut into Eyre Formation and Mesozoic sediments in the middle and upper reaches of Poole Creek palaeochannel, southern Lake Eyre Basin. (Modified from Krieg et al., 1991.)
Late Cainozoic (? Wipajiri equivalent)
Mesozoic
z 0
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Upper Poole Creek Palaeochannel
0-2
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0-1
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_ _ _ _ _
Lacustrine dolomite
___
Sand/sandstone _ _ _ _ _
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Plant fossils _ _ _ _ _ _ _ _ _ �
30
10 �------� SJ&-0927 MESA
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0-3 Clay and
60
20
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P-1 Gypsum_ _ _ _ _ _ _ _ _ �
.. . . � S1hc1fication - - - - - - - - - �
Precambrian
Carbonate. _ _ _ _ _ _ _ _ _ Bleached zone.. _ _ _ _ _ _ _
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Tertiary palaeodrainage systems
ders, some with leaf fossils; the coarser clasts are often present as a basal lag. The sandstone in Poole Creek palaeochannel is in part strongly cross-bedded, contains lenses of dolomite and is over-lain and interfingers with a more extensive dolomite. Plant fossils are common in the sandstone facies here. Laminated to cross-bedded channel sand (Willal inchina Sandstone) forming the Stuart Creek palaeochannel (Krieg et al., 1 991) may be a facies equivalent to facies 0-1 in Poole Creek, or be younger and equivalent in age to facies 4 (Callen et al., 1 986; Alley and Zang, in press). The sandstone facies in the Stuart Creek channel contains some of the most remarkable leaf fossils in the interior of the continent (Fig. 1 3). Limestone and dolomite (facies 0- 2), with siliceous nodules and gypsum casts are present in Poole Creek, including in the upper reaches of the
349
palaeochannel. The dolomite is undoubtedly cor relative with the extensive dolomite facies of the deeper parts of the Lake Eyre Basin and occupying a palaeochannel trending northwards from the north western edge of the Flinders Ranges (White, personal communication, 1 997). The dolomite is overlain by a silicified paly gorskite-rich clay (facies 0-3) containing dolomite nodules and sand lenses. This facies is widespread in the deeper basin and contains an extraordinary array of diverse fossil faunas that bear strongly on the middle Tertiary palaeoenvironmental interpreta tions (see below). Facies 0-3 is also present in the unnamed channel north-west of the Flinders Ranges. Facies 0-4 is a thin silicified siltstone grading up to sandstone with pebbles in some localities. This facies also contains plant fossils, but whether it is equivalent to facies 0-1 in the Stuart Creek channel is yet to be tested. Dolomite and clay may be present in the middle to upper reaches of some channels, indicating more extensive lacustrine conditions at times and that channel gradients were probably quite low. The Etadunna Formation was deposited after widespread doming of the Eyre Formation in deeper parts of the basin (Wopfner et al., 1 97 4). Similar structuring has been observed in some palaeochannels. White (per sonal communication, 1 997) notes that fiat-lying Etadunna Formation occupying the palaeochannel north-west of the Flinders Ranges has at least one channel side formed of tilted Eyre Formation. ?Pliocene to Quaternary facies
Leaf fossils, including aff. Brachychiton, in the silicified Willalinchina Sandstone, Stuart Creek palaeochannel.
Fig. B.
The final phase of Tertiary sedimentation in the Lake Eyre Basin resulted in the deposition of red-brown arenites and dark, fine-grained lacustrine sedi ments. These often are present in a series of broad, shallow meandering palaeochannels that drained from Queensland into the Lake Eyre area and which are entirely of latest Tertiary to Quaternary age (Callen et al., 1 995; Alley, 1 998). These facies are found in the Wipajiri and Tirari Formations. Facies P-1 is a fine-grained sand and silt over 1 m thick and found in palaeochannels cut into the Etadunna Formation in the Tirari Desert (Callen et al., 1 986). It contains coarser conglomeratic lenses in which vertebrate remains and Unionid shell impres sions are found. A clay in the facies at Lake Hydra contains well preserved palynofloras suggestive of a late Tertiary age (Callen et al., 1 995). The youngest
3 50
N F Alley et a!.
sandy gravel in the Poole Creek Palaeochannel may be equivalent to this facies (Fig. 12). Facies P-1 is characterized by shallow sand-filled channels, and is interpreted as a meandering fluvial system or a lacus trine delta-fan. Facies P- 2 is a widespread 4-5m thick red brown silt with sand lenses, usually cemented with massive gypcrete and extensively exposed in chan nels over the central Tirari Desert and northward into the Simpson Desert. Fossil vertebrates, including Diprotodon and the short-faced kangaroos, are found in the facies. The unit overlies possible facies P-1 at Lake Hydra and has been dated there as greater than 0. 40Ma by thermoluminescence methods (Callen & Nanson, 1992). The environment of deposition for facies P- 2 was probably one of intermittent stream flow under semi-arid conditions
Tertiary sedimentary basin Mesas ot dissected Tertiary sedimentary rock
___
_ _ _ __ _
...}J.'(.. l'-r--{-
Presentday drainage
'::::>--
___
SECTION A-A
Northern Territory palaeochannels
This area includes a series of elongate basins extend ing for several hundred kilometres to the north and south of Alice Springs (Figs 1 & 2). Some are struc tural features and others palaeochannels, many forming the upper reaches of those draining to the Lake Eyre Basin. In places the palaeochannel-fills reach over 200m in thickness, but outcrop is rare and most sediments are deeply weathered (Senior et al., 1994a,b ). Many of the valleys are partly flanked by mountain ranges. Most of the Tertiary succession overlies highly weathered Precambrian and Palaeo zoic bedrock. Correlation of units in this area is hampered by the degree of weathering and the resulting poor palyno logical control. In general, however, the lithostrati graphy of the channels and small basins is similar. This section focuses on the Hale, Ti-Tree and Waite Basins, where the stratigraphy and facies are better known (Fig. 14; Senior et al., 1994a,b). There are two main phases of deposition, the earlier from ?Palaeocene to middle Eocene (possibly late Eocene) and the latter in Miocene times, these being more or less equivalent to the early and middle Ter tiary phases in the Lake Eyre Basin.
D
HA.LE BASIN
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E-1 Kaolinitic clastics_ Ferricreta_
____
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Fig. 14. Tertiary basins and palaeochannel fills in the Alice
Springs area. Location of line of section shown on map. (Modified from Senior et al., 1994b.)
Early Tertiary facies
Together these facies are referred to as the Hale Formation and comprise largely quartz sandstone, siltstone and mudstone, grading to coarser clastic sed iments in the basin margins (Figs 3, 14 & 17). The basal unit is facies E-1 consisting of kaolinitic, argilla-
Fig. lS. Weathered Lower Cretaceous Bulldog Shale,
southern side of Stuart Creek palaeochannel.
Tertiary palaeodrainage systems
ceous, poorly sorted sandstone with intercalations of granule and pebble conglomerate (Senior et al., 1994a,b). It reaches up to 55m in thickness, overlies weathered and ferruginized bedrock and grades towards channel margins into conglomerate of facies E-2, and is overlain by middle and late Tertiary to Quaternary facies (Fig. 14). The upper sandstone facies, where separated from the lower sandstone by carbonaceous sediments (see below), is named the Tug Sandstone Member. In some areas the sandstone facies is overlain by, or contains lenses, about 4 m of carbonaceous shale and lignite (facies E-3; Ulgnamba Lignite Member of Stewart et al., 1980). In other localities facies E- 3 may include a lower mudstone and siltstone (Clar aville Mudstone Member of Senior et al., 1994a,b ). Mudstone (facies E- 4) in the Waite Basin may be up to 250 m thick, overlies weathered bedrock and is overlain by middle and late Tertiary to Quaternary facies. Although its age is only poorly known, the stratigraphical relationship of facies 4 to the overly ing middle Tertiary strata indicates a probable early Tertiary age. The facies may be equivalent to a basal mudstone facies underlying the Ambalindum Sand stone Member in the Ti-Tree Basin (Fig. 15; Senior 