Palaeoweathering, Palaeosurfaces and Related Continental Deposits (Special Publication 27 of the IAS)

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 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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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 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:

-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 � ·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:
c.* - s (z)

(8 s - 1 .00448 - 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 - ) = 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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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

5R 412 silty

5R 412 silty

claystone

claystone

0

Bloomsburg

(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

pedogenic slickensides

reduction mottles Fe glaebules/

•

concretions

I A. 9

ee

�

-

'7

Sil.)

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:::: .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 !"--
... )-'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

----

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---------

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

DO

N4

oo

5YR3/4

o

DO

5R3/4

DO

10R3/4

o

5Y4/1 10R2/2 5GY5/1

3

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2

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7 8 9

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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'- '... '!I PL I O-PLEISTOCENE (>:-:-:-:· :) R K BED OC

i_ R BQ i �; D CLAY SEDIMENTS

/ "

'

CLAS TIC SEDIMENTS . · . . . .'

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

Mg - SMEC Tl TES

--f- SA LI NE OPEN L A K E ­ (ephemeral or per en niol)

SUBSIDENCE/ SEDIMENTA­ TION RATES

HIGH HIGH

( stttvttnsiftt, saponiftt "- . .j_ 0_ig!!.tt!_ �I'!!!�ondifions)

KEROL/TE / SMECTITE MIMD- L AYERS _

ZEOL I TES

_

_

MAINLY DETRITAL C L AYS

AUTHIGENIC IL L I TE

8'!.�11 77'7'/

li2l.

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

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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.

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SuRDAM, R.C. & SHEPPARD, R.A. (1978) Zeolites in saline, alkaline lake deposits. In: Natural Zeolites: Occurrence, Properties, Use (Eds Sand, L.B. & Mumpton, F.A.), pp. 145-174. Pergamon, New York. TALBOT, M.R., HoLM, K. & WILLIAMS, M.A.J. (1 994) Sedi­ mentation in low-gradient desert margin systems: a com­ parison of the Late Triassic of north-west Somerset (England) and the late Quaternary of east-central Australia. In: Paleoclimate and Basin Evolution of Playa Systems (Ed. Rosen, M.R.). Geol. Soc. Am. Spec Pap. , 289, 97-117. TARDY, Y., CHEVERRY, C. & F RITZ , B. (1974) Neoformation d'une argile magnesienne dans les depressions inter­ dunaires du Lac Tchad. Application aux domaines de stabilite des phyllosilicates alumineux, magnesiens et fer­ riferes. C. R. A cad. Sci. Paris, 278, 1 999-2002. TETTENHORST, R. & MOORE, G.E. (1978) Stevensite oolites from the Green River Formation of Central Utah. J Sedi­ ment. Petrol. , 48, 587-594. TIERCELIN, J.J. (1991 ) Natural resources in the lacustrine facies of the Cenozoic rift basins of East Africa. In: Lacus­ trine Facies Analysis (Eds Anad6n, P., Cabrera, Ll. & Kelts, K.), Spec. Pubis int.Ass. Sediment., No. 13, pp. 3-37. B lackwell Scientific Publications, Oxford. TRAUTH, N. (1977) Argiles evaporitiques dans Ia sedimenta­ tion carbonatee continentale et epicontinentale tertiaire. Bassins de Paris, de Mormoiron et de Salinelles (France), Jbel Ghassoul (Maroc). Sci Geol. Mem., 49, 195 pp. TRIAT, J.M. & TRAUTH, N. (1972) Evolution des mineraux argileux dans les sediments paleogenes du bassin de Mormoiron. Bull. Soc. Fr. Mineral. Cristallogr. , 95, 482494. TRuc, G. (1978) Lacustrine sedimentation in an evaporitic environment: the Ludian (Paleogene) of the Mormoiron basin, southeastern France. In: Modern and Ancient Lake Sediments (Eds Matter, A. & Tucker, M. E.), Spec. Pubis int. Ass. Sediment., No. 2, pp. 189-203. Blackwell Scientific Publications, Oxford. VALLERON, M.M., DULAU, N., POURZAHED, P. & SAUGRIN, T. (1983) Calcitizations et opalitisations dans !'Eocene du Sud-Est de Ia France. Comparaison avec des facies ana­ logues d'Alsace et de Touraine. Bull. Soc. Geol. France, 25, 1 1-18. WATTS, N.L. (1980) Quaternary pedogenic calcretes from the Kalahari (South Africa): mineralogy, genesis and dia­ genesis. Sedimentology, 27, 661-686. WEBSTER, D.M. & JoNES, B.F. (1994) Paleoenvironmental implications of lacustrine clay minerals from the Double Lakes Formation, southern Great Plains, Texas. In: Sedi­ mentology and Geochemistry of Modern and Ancient Saline Lakes (Eds RENAUT, R. & LAST, W.), Spec. Pub!. Soc. Econ. Paleont. Miner., Tulsa, 50, 159-172. WRIGHT, V.P. & SANDLER, A. (1994) A hydrogeological model for the early diagenesis of Late Triassic alluvial sediments. J geol. Soc. London, 151, 897-900. YAALON, D.H. & WIEDER, M. (1976) Pedogenic palygorskite in some arid brown (Calciorthid) soils of Israel. Clay Mineral. , 11, 73-80. YuRETICH, R.F. (1 979) Modern sediments and sedimentary processes in lake Rudolf (lake Turkana), eastern rift valley, Kenya. Sedimentology, 26, 313-331.

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

; .......

� ; �
--' 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

�

10

1 090

11

1 089

12 saprolite 1 3

B,

I ;

·H

1 087 - 1 085

An

I

� � ,

Bax

(

\

l : · � Ill Ij ' ·� .. I � : ·� � \\ -' .lo. ... I ;.

� I ''

. JJ.

�.ttl··. I

I

I

� )

1 00 0 20 40 60 0 50 o joint filling Si02 •AI203 •Fe203 M i neral Composition (wt. %) •

2 0 • Ti02

4

(parent rock: amphibole­ biotite-gneisses) Otz An

(Pit) 7-3-P3

..

I� . I t t ...

·� ·"·

.·(..

;,

·�r\ .( . ·:• . .·. .� �� � • I#

�: .'

� ).

- 1 080 - [ 1 079] - 1 078 - [ 1 077] - 1 076

,<J

!--
f

"'

cl=P D

I '"'

d.

DD

� D

I

D

I

cid + -a

I

' )(

D

)(

)(

0

D

)(

"' )( �

)(

)( )(

)(

)(

)(

0

}

I' (x, n=1 2) - basic rocks II' ( x , n= 1 2) intermediate Ill' (x, n=13) to IV' ( x, n=19) acid rocks

AMPHIBOLITE (POINT 5+6) @ Parent rock "' Bauxite (·parent rock 'group I') 1' Saprolite x Joint filling .. Soil AMPH.-BT. - GNEISSES ( POINT 7+8)

0

I )(

)(

) 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

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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

215

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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-

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

0

of the Masssif Pliocene

Miocene

u

Central

I

fluviatile deposits

-

subsidence

0 N >0 a: z 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



Shallow kaolinitic saprolitcs below Cambrian covers

..

Other remnants of kaolinitic saprolites

"..

Kaolinization interpreted as hydrothermal

•

Clayey/sandy saprolites

o

Weathering breccia with limonite Gravelly weathering

'� � = I .._ Oo _ = = .,� 2 01lkm

go

I

� Heavy glacial erosion c:::::J Tertiary relief rs:::::sJ �

C2J [=:J

Sub-Cretaceous etchsurface

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

Lidmar-Bergstrom et a!. m

60 WNW 50

40

- -

Upper Cretaceous limestone

- -

....

boulder

weathering front

--j x

0

50

_ _ _ _ _ _ _

X

20

1 x

I I

core stones _ _ _ _ _ _ _ _ _ _ _ _ _ _ _ _ _ _ _ _ _ _

X

X X

30

X

X

- x I I X I

X

10

I_ � _ _ _x _ _ - �

L�k�l�d��f�c� _ _ _ _ _ _ _ _

_ _ _ _ _ _ _ _

100

200 m

0

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

200

300

�

100 0

B

,.._, - C..., ...

I

SOUTH SMALANO

l 1 200 m

PENEPLAIN (SSP)

N SOUTH St.4A.LAND PENEPLAIN

300

200

I

SSP

I

200 m

L

I

200 m L

A

1!

E

I

150

100

B

LEVEL

200

s

200 m LEVEL

LAKE VATTEFIN

100

•

300 200

,\------+ 100

100

Kaohnitic saprolite remnant

c

500 �

300 200 100

SUB - CAMBRIAN PENEPl,t,IN

E

-400 300 200

+-----------�----------���--��� -==�==�==�: 100

=�

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

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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.

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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,

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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

�

�

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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M. J Kraus and A. Asian

304

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

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: "' "'

"0 ;:,

E

"0 c: "' "'

"0 ;:,

E

u

A

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A

Bw

Bt

c

I

-1 -lu -1

utl -r-tl -1

Ab Ab/Btj

Bt2

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Cb Cb/C

c

Composite

Compound

(a) Non-steady

"0 ;:,

"0 ;:,

"0 ;:,

E

E

E

I

"0 c: "' "'

"0 ;:,

E

Ag

Ag

Bg

Bg/Bkg

Cg

Cg Completely mottled

Completely mottled

(b) Steady

Key Stratification

( tJ

"\

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

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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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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.

315

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

316

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.

M. J. Kraus and A. Aslan

318

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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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

323

324

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-

PEDIMENT

FLUVIAL ENVIRONMENT floodplain

� low water table

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LAKE marginal

channel belt

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Fig. I. Distribution and main characteristics of the carbonate-rich palaeosols.

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� -1

champh

_ p�lu�ri�e

calcic soils

� main

well drained

poorly drained except levee deposits

colour banding

colour mottling

light colour mottling

rootlet traces CaC03 glaebules, nodules, rhizoliths

rootlet traces

rootlet traces nodulization of lacustrine muds

weathering of unstable minerals

Fe, Mn nodules

desiccallon/dissolution features

vadose cementation

no drainage

phreatic cementation

f e a t u r e s

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

THANET11EN

Marseille

� Palaeogene

_.1..-

Thrust

-- Fault Measured section

0

- - ·? - -

DANO..MONTlEN

IOkm

.___.__,

- - -? -

� floodplain deposits

lacustrine/palustrine � t::I:::::I: limestooes

®

-

Fig. 2. (a) Map of the study area showing of the Upper Cretaceous-lower Palaeogene formations (CA, Castellas; GR, Griffon; OLI, Oliviere; RH, Roques Hautes; RI, Ribas; VA, Valabre; VT, Vitrolles). (b) Hypothesized stratigraphical pattern of the Maastrichtian-Palaeocene deposits.

Carbonate-rich palaeosols

Roques Hautes

327

Ribas

Calcretes Lac./palustrine limestones

Castelias

� lo .I

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

A 8

A

�

?

ss

III I I , , II I


gastropods

(

burrows

whitish greyish pink yellow/green orange red greenish

Pedogenic features 0 ..

I

mottling carbonate nodules rhizomorphs

v pseudo karst ss

slickensides

t rootlets

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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(1994) Lacustrine facies model for non

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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palaeoclimatic interpretation. Evidence of palaeo­ surfaces and weathering from neighbouring areas (e.g. Israel, Jordan, Syria, Iraq and Saudi Arabia) is included in the discussion. In his comprehensive overview, King (1962) has mentioned relics of planation surfaces in northern Sudan, without, however, attributing them to parti­ cular weathering or erosion processes. Since then a wealth of data concerning palaeoweathering have been collected, which now allows a more detailed picture to be drawn of palaeoweathering surfaces in north-east Africa. Indications of old land surfaces in north-east Africa now can be related to Early Palaeozoic (Germann et al. 1993), Carboniferous (El Sharkawi et al. 1990a), Cretaceous (Bowitz 1988; Fischer 1989; Germann et al. 1990) and Tertiary (El Aref 1993; Schwarz 1994) weathering periods. Asso­ ciated sediments in many cases yield material derived from eroded weathering surfaces.

PALAEOSURFACES AND ASSOCIATED SED IMENTS

The development of palaeosurfaces and subsequent destabilization, stripping and reworking of cratonic

regolith, resulting in the accumulation of weathering products in continental and marine sedimentary basins was first described from the lower Tertiary 'Siderolitique' of Switzerland and France (Fleury 1909; Kulbicki 1956). In West Africa, a comparable association of rocks, known as ' Continental Terminal' (Kilian 193 1 ; Faure 1966; Lang et al. 1986, 1990), was formed by the alternation of deep weathering and reworking of weathering crusts during the Neogene Epoch. Here, lateritization processes on the contin­ ent led to the formation of ferruginous weathering crusts (ferricretes), whereas subsequent erosion and reworking led to associated continental sediments rich in ferruginous components. Millot (1970) was one of the first authors to also systematically corre­ late marine deposits with weathering processes on the continent by applying the classic model of 'biorhexistasy' of (Erhart 1955). In north-east Africa a variety of palaeosurfaces and associated sediments occur (Germann et al. 1994). Residual deposits have been formed by weath­ ering in places where intensive chemical in situ alteration of magmatic rocks, gneisses and siliciclastic sediments led to the formation of deep weathering crusts. It is this friable material that provided the basis for the development of vast planation surfaces

Weathering surfaces, north-east Africa on the Gondwana continent (Millot 1983). Depend­ ing on the parent rocks and as result of vertical and lateral differentiation, lateritization processes could produce deposits of, for example, in situ kaolin (saprolite), bauxite or ferricrete (Fig. 1 ) . On carbon­ ate rocks subaerial exposure has led to karstification. Although in the in situ deposits residual enrich­ ment of stable elements and minerals is the dominant process, the elements released into solution or miner­ als set free by mechanical erosion were transported into the continental depositional realm. Secondary accumulation could be accomplished by mechanical processes, leading to alluvial placers, sedimentary kaolinites or oolitic ironstones, or by chemical pro­ cesses, which are responsible for the formation of, for example, alunite, secondary silica minerals and hydromagnesite.

369

1 the mainly marine Palaeozoic cycle (Cambrian to Early Carboniferous); 2 the purely continental Karoo cycle (Late Carbonif­ erous to Early Jurassic); 3 the marginal marine to continental Nubian cycle, initiated at the beginning of the disintegration of Pangea and extending from Late Jurassic to latest Cretaceous times (Fig. 2). During all three cycles subaerial exposure and periods of non-deposition led to the development of palaeosurfaces on both older sediments and on base­ ment rocks. Although Palaeozoic surfaces are pre­ served in rare cases only, weathering surfaces of Mesozoic and Tertiary age are ubiquitous in north­ east Africa.

EARLY PALAEO ZOIC LATERITIC WEATHERI N G GEOLOGY

The study area of north-east Africa comprises Egypt, northern Sudan and parts of Ethiopia. In contrast to western Africa, this area in the past experienced only little more than very broad geological survey work. Only quite recently could it be demonstrated that in north-east Africa, in addition to the Cenozoic Era, nearly the whole Phanerozoic Eon is documented in both marine and continental sediments (Klitzsch 1989, 1990). During the geological evolution of north-east African intra- and epicontinental basins the struc­ tural, palaeogeographical and palaeoclimatic condi­ tions for the formation and destruction of weathering surfaces repeatedly changed. Thus this area provides a wide variety of examples of different types of sur­ faces and their associated sediments. The Precambrian basement, best accessible in the central and eastern parts of the study area, displays a variety of different parent rocks, such as old gneissic terranes of pre-Pan-African age, Pan-African volcano-sedimentary assoCiations with granitoids and metasedimentary belts with ophi­ olite complexes (see, e.g. Schandelmeier et al. 1988, 1990). The stratigraphical subdivision of the overlying sedimentary strata, investigated since the mid-1970s, is now known in much detail (Klitzsch & Squyres 1990). Stemming from the reconstruction of tectonics and palaeogeography of north-east Africa, these former 'Nubian' strata now can be subdivided into three cycles:

Following the Pan-African orogeny, an erosion surface developed during the Cambrian Period that is characterized by a low relief with elevation differences rarely exceeding 1 0 m in Jordan (Wolfart 1981). Block faulting towards the end of the Cambrian led to the development of a NNW­ trending relief. Following this structural pattern, Palaeozoic transgressions penetrated the continent from the north-west (Klitzsch 1987). In north-west Sudan, south of the Jebel Rahib in the Jebel Tawiga-Jebel Tageru area an extensive weathering surface with the sole complete lateritic weathering profile known at present from north-east Africa was preserved underneath these transgressive Palaeozoic sediments on top of Precambrian basement rocks (Fig. 3). Over an area of about 1000 km2 a weathering crust, consisting of both kaolinitic saprolite and overlying bauxitic laterite, is developed (Fig. 4). It rests on strongly deformed metabasalts and metapelites, which belong to the southern extension of a late Proterozoic ophiolite complex in the Jebel Rahib fold-and-thrust belt (for details see Germann et al. 1993). The weathering profile with its maximum thick­ ness of about 25 m is overlain by shallow marine Skolithos-bearing sandstones of late Ordovician to early Silurian age, according to trace fossil assem­ blages with Cruziana acacensis and Cruziana cf. ancora (Seilacher 1991), for example. The unusual good preservation of this Palaeozoic weathering crust can be related to the sheltering effect of the marine sandstone blanket.

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