Orbital Forcing Timescales and Cyclostratigraphy
Geological Society Special Publications Series Editor
A. J. FLEET
GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 85
Orbital Forcing Timescales and Cyclostratigraphy EDITED BY
M. R. HOUSE Department of Geology The University, Southampton, UK
A. S. GALE Department of Palaeontology Natural History Museum, London, UK
1995 Published by The Geological Society London
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Contents Preface HOUSE, M. R. Orbital forcing timescales: an introduction
vii 1
KELLY, S. B. & SADLER,S. P. Equilibrium and response to climatic and tectonic forcing: a study of alluvial sequences in the Devonian Munster Basin, Ireland
19
HOUSE, M. R. Devonian precessional and other signatures for establishing a Givetian timescale
37
WEEDON, G. P. & READ,W. A. Orbital-climatic forcing of Namurian cyclic sedimentation from spectral analysis of the Limestone Coal Formation, Central Scotland
51
MOSES, G. P. G. Calibration, analysis and interpretation of depositional cycles in the Early Toarcian of Yorkshire, UK
67
WATERHOUSE,H. K. High-resolution palynofacies investigation of Kimmeridgian sedimentary cycles
75
VALDES, P. J., SELLWOOD,B. W. & PRICE, G. D. Modelling Late Jurassic Milankovitch climate variations
115
COTILLON,P. Constraints for using high-frequency sedimentary cycles in cyclostratigraphy
133
GIRAUD, F., BEAUFORT,L. & COTILLON,P. Periodicities of carbonate cycles in the Valanginian of the Vocontian Trough: a strong obliquity control
143
QUESNE, D. & FERRY, S. Detailed relationships between platform and pelagic carbonates (Barremian, SE France)
165
GALE, A. S. Cyclostratigraphy and correlation of the Cenomanian Stage in Western Europe
177
FISCHER, A. G. Cyclostratigraphy, Quo Vadis?
199
Index
205
Preface The discovery in the 1970s that Pleistocene climates and especially ice-age development were controlled by identifiable orbital parameters in the Milankovitch Band, confirming the views of Milutin Milankovitch in the 1920s, is probably the greatest single advance in palaeoclimatology this century. Changes in temperature recorded by isotopes in calcite of deep-sea foraminifer tests provided a detailed record of ice-sheet advance and decay moderated by precession, obliquity and eccentricity cycles (19-400 ka). As a consequence of this work, the reality of orbital forcing of climate was established as a fact. An added bonus of this discovery is that the identification of individual Milankovitch frequencies allows construction of an orbital forcing timescale graduated by these frequencies. Earlier palaeontological research which gave evidence of the number of days in the year, and days in the lunar month in the past, contributed to calculations on how the precession and obliquity cycles may have differed in geological time. The possibility that Milankovitch cyclicity could form the basis for orbital timescales throughout the geological column was soon recognized, although G. K. Gilbert had seen the same possibilities in the Cretaceous of the Western Interior Basin of North America a century ago. The geological record contains abundant bedding cycles, the duration of many of which fall within the Milankovitch Band. The question as to which of these diverse bedding features record climatic cycles, the processes by which they were generated, and the identification of individual cycle frequencies is a lively and exciting area of research and discussion. The holy grail of this work is the construction of a pre-Pleistocene orbital timescale graduated in Milankovitch units. Bundling of precession cycles within short eccentricity cycles is a readily identifiable feature which is particularly valuable in the development of timescales. This volume is a product of a meeting on orbital timescales and cyclostratigraphy that we organized at the Geological Society of London apartments on 25 and 26 March 1993, and its diverse papers reflect the wide range of sedimentological, palaeontological, geochemical and stratigraphical research presented at that meeting and of the discussions generated also during a field excursion to the Dorset coast which followed that London meeting. M. R. House and A. S. Gale
Orbital forcing timescales: an introduction MICHAEL
R. H O U S E
Department of Geology, The University, Southampton SO17 IBJ, UK Abstract: A brief review is given of orbital patterns affecting the Earth which may be of use in
establishing, for long or short periods, orbital forcing timescales (OFT). The metronomic variations of the Earth-Moon system and of the Earth-Sun orbital patterns produce gravitational and temperature effects which alter the physical environment on the Earth's surface. These give an interpenetrating effect of forcing cycles ranging from twice daily tides, day-night alternations, various tidal patterns and the annual solar pattern. All of these have been used palaeontologically to give precision to short-term age determination in the past. It is cycles of the Milankovitch band which are showing promise of enabling new practical timescales to be established for parts of geological time. These depend on changes in the Earth-Sun distance (perihelion and precession cycles of 19 and 23 ka at the present time), changes in the tilt of the Earth's axis with respect to the Earth's orbit round the Sun (the obliquity cycles of 41 and 54 ka), and changes in the geometry of the Earth's orbit around the Sun (eccentricity cycles of 106 and 414 ka). Since the number of days in the year have changed through time; so have the periods of the perihelion and precession cycles. There is increasing evidence that small-scale sedimentary rhythmic couplets, often grouped into bundles, may represent the effect of some of these; often the precessional couplets are grouped into bundles of five or so within the lower eccentricity period. The disentangling of the interpenetrating cycles to produce an OFT is an exciting problem and challenge for palaeobiology and sedimentology. These should enable numerical dates to be given to biostratigraphic and chronostratigraphic timescales and eventually enable many earth processes to be analysed in real time. 26 Ma oscillations related to the Cosmic Year (c. 260 Ma) have been invoked to explain periodic mass extinctions in the fossil record. But evidence is presented to suggest such extinctions are not, in fact, periodic.
The purpose of this contribution is to provide an introductory review of those orbital patterns which have such an effect on the environment of the Earth's surface that they give potential for the establishment of orbital forcing timescales (OFT) for parts, perhaps eventually much, of Earth history. That the establishment of time in geology, for record of its past events and in the establishment of rates for processes is of major importance is self evident. At present we rely on biostratigraphic scales for the Phanerozoic, but these are relative, not absolute, scales. For the late Mesozoic and Tertiary, in suitable circumstances, the radiometric scales are extremely important, if incomplete; for pre-Cretaceous rocks, however, their increasing sparseness and unreliablity make them of limited practical use. The exciting possibility is that new timescales can be constructed using microrhythmic sequences which may show the effects of particularly precession, obliquity and eccentricity orbital patterns, over frequencies usually referred to as of the Milankovitch band, may provide timescales of considerable refinement. Such cycles affect the solar energy reaching the outer atmosphere because the Earth-Sun distance is changed during them, or seasonal distribution of insolation. The way in which outer atmosphere changes are reflected in local changes on the Earth's surface is undoubtably
complex, and in many ways poorly understood. However, it is thought that resultant sedimentary microrhythms result from changes of sea level, and changes in the pattern of vegetation and erosion on adjacent land areas which are mainly driven by climate. Figure 1 & Table 1 give the range of orbitally forcing frequencies which may contribute to the development of timescales. The recognition of the potential of orbitallyforced microrhythms for the construction of timescales was first most clearly stated by G.K. Gilbert (1895, 1900a, b) and developed further by Barrell (1917). Such ideas followed naturally from the laws of planetary motion established in 1609 and 1618 by Johannes Kepler, and the later recognition, by Newton, of the role of gravitational attraction between planetary bodies. Adhrmar (1842) and Croll (1875) gave an early summary of such views and Charles Lyell considered them in detail in the later editions of his Principles of Geology. However, it was the calculations of Milutin Milankovitch (1920, 1941), using climatic effects of orbital patterns to explain the ice ages, which was the major turning point; but such ideas were not well received at the time. There followed a long period when microrhythmic sequences formed the basis for mathematical studies of series analysis, but with little attempt to invoke the real
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescales and Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 1-18.
2
M.R. HOUSE FREQUENCY
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Fig. 1. Logarithmic table of orbital periods which exert gravitational effects on Earth, or which which exert orbitally forced changes in the temporal energy distribution reaching the outer atmosphere of the Earth from the Sun. time dimension, with the exception of some elegant discussions on long records, such as in the late Trias of the Newark Basin (van Houten 1964; Olsen 1984; Anderson et al., 1984) dealing mainly with sub-Milankovitch band effects. The modern phase was undoubtably reached with the calculations of possible orbital forcing to produce documented evidence in ocean cores of temperature changes that really established such theories indisputably (Hays et al. 1976; Imbrie & Imbrie 1979; Covey 1984). The urgency to use such tools to improve the timescale was pressed by House (1985a, 1986a, b) and is part of the theme of this symposium. There have been many symposia and reviews in past years, both of biological rhythmicity (Rosenberg & Runcorn 1975), sedimentological and astronomical aspects (Merriam 1964; Einsele & Seilacher 1982; Berger et al., 1984; Fischer & Bottjer 1991; De Boer & Smith 1994; Smith 1990a, b), and methods of mathematical analysis of microrhythmic sequences (Weedon 1993; Schwarzacher 1964, 1975, 1987). The term cyclostratigraphy has been coined for sedimentary phemomena, but there has been little concentration on techniques to improve the geological timescales. On the scale of daily, monthly and annual effects, the causation of tides essentially followed the recognition of the laws governing planetary motion. A turning point was the classic paper by
Wells (1962), who recognized daily and annual banding in Devonian rugose corals and was able to estimate the number of days in the Devonian year at rather over 400. The recognition of lunar effects followed shortly after (Scrutton 1964), which enabled the periods of the Earth-Moon orbits to be estimated for the Devonian. Since such motion controls the perihelion and precession cycles, it has subsequently been shown by Berger et al. (1989a, b) how these cycles have changed through geological time. The recognition of daily, lunar and annual effects in the shells of bivalves (House & Farrow 1968) was followed by many studies (Scrutton 1978). The annual changes in tree rings have long been known and dendrochronology is now a discipline in its own right extending back over several thousand years. Frequencies of orbital forcing cycles have been divided into the calendar band, solar band, Milankovitch band and galactic band (Fig. 1). Imbrie (1985) used another system, which may be modified here by the inclusion of the highest frequencies as follows: daily band (0-25 h), monthly band ( 25 h-0.5 a), annual band (0.52.5 a), interannual band, (2.5-10 a), decadal band (10-400 a), millenial band (400-10 000 a), Milankovitch band (10 000-400 000 a) and tectonic band 400 000+ a).
Annual and lesser orbital cycles (< 1.0 a) These frequencies have been named the calendar band (Fischer & Bottjer 1991). The principal lowerorder cycles may be separated into the tidal, whose effects result primarily from gravitational changes in the Earth-Moon system and the solar, which result from changes in the energy received from the Sun resultant upon daily to annual changes. The second are well known and merit little attention here, although it should be pointed out that for organisms, as for sedimentation, the interpenetration of these effects can be complex. Emphasis here will be on such factors as contribute to OFT criteria. Cycles at frequencies of < 1 a are recognizable in both the sedimentary record, where they are embraced in the term rhythmites, and in the fossil record, where they show as growth banding where the accretion of tissues reflects environmental rhythms: under suitable circumstances these may be preserved in both plant and animal tissues. It is unlikely that evidence from this source will ever be integrated into a continuous timescale for the past. Nevertheless, for short periods, the documentation of tidal, daily, monthly, equinoxial and annual cycles have already contributed much to short-term environmental analysis, quite apart from the contri-
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bution to factual knowledge on the Earth-Moon orbital parameters in the past.
Tidal effects The orbit of the Moon around the Earth (Fig. 2) attracts a tidal wave around the Earth with a period of approximately once a day, the actual period (24 h 54 min) being termed lunar day, since it is not quite of the same period as the day-night cycle, or solar day. It is fortuitous that, at the present time, the period of the Earth's rotation around it's spin axis is only slightly faster than the orbit of the Moon around the Earth giving such similarity between the solar day and the lunar day: this was not so in the past. However, as Newton demonstrated, because of the centrifugal force resultant upon the rotation of the Earth, not only is the strongest tide developed around the Earth where the plane of the Moon's orbit crosses the Earth (giving the equilibrium tide of Fig. 1), but tides are developed at opposite sides of the Earth in the form of water heaping (tidal nodes) at points closest to and furthest away from the Moon. Thus, given the rotation of the Earth, the typical tidal cycle is formed which gives about two tides a day, a pattern which it is convenient to call semidiurnal (half-day, or half-solar day) but which in reality is half of the lunar day (Fig. 3A). Usually, the two semi-diurnal tides are not equal, that is the alternation of levels reached by successive tides is not the same. Hence, there can be a lower high tide (low high water) followed by a higher high tide (high high water) (Fig. 3A). This diurnal inequality of the tide results from the fact that the orbital plane of the Moon around the Earth
makes a low (and changing) angle with the axis of the Earth's rotation (Fig. 2). This angle, or declination, is not constant, and obviously the inequality is greatest when the declination is greatest. At the time when a tidal node is below the Earth's Equator on one side of the Earth it will be above the Equator on the other side. At a place in mid-latitudes (illustrated in Fig. 2), therefore, the higher high water will be when a point is closest to the Moon. Local geographical effects of coastal shape and seafloor morphology, and of storms, often far away, will modify these simple cycles, often considerably. There is a distinct solar tide which has an interval of 12 h. Although the gravitational attraction due to the Sun is 177 times stronger than that of the Moon, the solar tides are smaller in effect than lunar tides because there is a fundamentally different relationship between the two; the Earth is in orbit around the Sun and hence a state of nonweightlessness occurs because the radial component of force due to the Earth's orbital momentum is almost cancelled out by the gravitational pull of the Sun. It is easy to consider tidal effects solely in terms of the sea, and the effects on the seashore, and limited to the changing distance between high and low tide marks. This would be a mistake. Undoubtedly of greatest importance for tectonics is the continual stresses on the deformable solid earth by tides ranging from the semidiurnal tides to the tidal node cycles of c. 20 a. Rocks may give a less obvious response, but the continual stresses produced in this way will be an important factor in their ultimate relief by fracture, jointing, faulting and folding, and perhaps at much larger scales too. Similarly, the effects caused by cyclical movements
4
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of the atmospheric gases will modify wind movements and pressure systems. But these we have no means, at present, of documenting.
Semidiurnal (0.00135 a) The typical twice-daily rise and fall of sea level can follow several patterns from one in which there is an extreme of semidiurnal change (twice daily) to one which is essentially daily, or circadian, but in reality reflects the period of one lunar day. Both cases have been recorded from sedimentary sequences and fossil shells from the past (see below). This is a means of obtaining precise
information on local tidal regimes and on the timing of astronomical controls. The founder of petrography, H.C. Sorby, over a hundred years ago, appears to have been the first to recognize petrographical differences between the ebb and flood tidal characteristics in the Jurassic Forest Marble of England. He was able to distinguish different directions of depositing flow between the ebb and the flow sedimentary regimes. Examples of such rhythmites from the German coast were reported by Reinecke & Wtinderlich (1967) (Fig. 4B), who found that fine sand was deposited at times of tidal current flow and clays were deposited from suspension during times of
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i~'ig. 4. Semidiurnal signatures in sediments.(A) Semidiurnal inequality and the progressive time progression between high-water slacks and low-water slacks (darker bands): an example from the Upper Carboniferous of Francis Creek, Illinois, USA. Photograph kindly provided by E Broadhurst. (B) A similar example, but from Present day North Sea sediments (modified from Reinecke & Wtinderlich 1967).
minimal current flow at either high or low slack tides. At that locality the ebb tide was much stronger than the flood tide and hence more sand was deposited on the ebb. Note that such records only occur when the environment allows continuous deposition. Modern examples have been discussed by Allen (1981, 1982). An excellent example (Fig. 4A) from the Carboniferous of Illinois has been illustrated by Broadhurst (1988), in which a typical semidiurnal tide pattern is seen where progressive changes in the timing of ebb and flood results in a systematic changes during the neap-spring-neap cycle (Fig. 4A). For fossils, elegant contributions on fine-scale banding in bivalve shells was given by Evans (1972) and Pannella & MacClintock (1968). Evans demonstrated for the Recent cockle, Clinocardium nuttalli from the coast of Oregon, that there was a fine-scale cyclicity in the growth of the ostracum which matched the semidiurnal cycle and which
shows the effects of changing diurnal inequality in the tides (Fig. 5). A rather more systematic study of such changes has recently been published by Ohno (1989). An example in Fig. 6 illustrates an acetate peel made from a polished radial section of a specimen of Cerastoderma edule from mid-tide level of the Bury Inlet, South Wales. This shows a dominant diurnal pattern with only slight evidence of a semidiurnal effect.
Solar day (0.0028 a) Because the lunar day so closely corresponds, at the present time, to the solar day, it is understandable that it is often difficult to disentangle the effects of day-night rhythmicity from daily tide rhythmicity, in the past especially when modulation by a semidiurnal tide effect is not apparent. Thus, the estimation of days in the lunar month, or Earth year, are dependent on whether tidal or insolation effects are dominant. The authors of many such estimates
6
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Fig. 5. (A) Progressive semidiurnal tidal signatures for Empire, Oregon, USA, for a period in June and July 1970. (B) Correlated growth pattern record of Clinocardium nuttalli from the Oregon coast for the same period (modified from Evans 1972).
either do not mention this difficulty or assume that the present day situation of near-identity held in the past. Scrutton (1978) drew attention to this problem. The importance of a circadian biorhythmicity to living things is well known and well documented (Harker 1964; Neville 1967; Clarke II 1974). It operates physiologically, behaviourly and, when approriate, in the deposition of organic tissues. Geological information is only obtainable when fossils record circadian growth increments in preservable hard parts.
Spring tides (0.038 a) Spring tides occur approximately every 14 days when there is an approximate alignment of the Sun, Moon and Earth (Fig. 2). This occurs either at conjunction or opposition of the Moon and Earth with respect to the Sun. Neap tides occur between these when the Sun and Moon are in quadrature with the Earth. Tidal records show this well (Fig. 8A). Sedimentological evidence of this effect in ancient rocks has been recognized and discussed especially by Williams (1989a, b). In the pattern of growth increments in modern
shells such patterns have been well documented (Figs 5 & 6), but there are not many good fossil examples partly because many shells, being made of aragonite, recrystallize after fossilization. In the case of littoral shells, position relative to mean sealevel is important since this controls modulation of shell growth and low tidal forms show the phenomenon least (Farrow 1971, 1972).
Lunar month (0.081 a) At the full Moon alignment of the Earth, Moon and Sun is most perfect and this gives a dominant tidal effect leading to enhanced spring tides. The period of the lunar month at present is 29.53 days, and there are 12.37 lunar months in the year. Tidal gauges show the greater amplitude of spring tides at this time (Fig. 8A). Evidence for the lunar month in the past has been claimed from the Precambrian to the Recent and attempts have been made to show how the number of lunar months per year has increased in the past, reaching in excess of 13 in the Precambrian (Williams 1989a, b). Data has been drawn from corals, bivalves, stromatolites and other groups. The first such study was by Scrutton (1964),
Fig. 6. Semidiurnal, diurnal, lunar and annual (winter) signatures in the growth of the cockle Cerastoaerrnaeaute (Linnaeus) from the mid-shore level in the Bury Inlet, South Wales, UK, based on an acetate peel of a polished radical section, • 80 [by Farrow (1972, pl. 8B)].
ORBITAL FORCING TIMESCALES who claimed lunar month periodicity in Devonian using corals collected by the writer from the Bell Shale of Michigan (Scrutton 1964, Fig. 7). The calculation was based on fine-scale increments thought to be daily. But, as Scrutton (1978), pointed out, whether this referred to the lunar day or solar day is not clear. Thus, the calculation of the lunar month could be in error. Figure 7 shows a well-developed banding of this type. Pompea & Kahn (1979) made similar observations in nautiloids. The cause of the lunar banding is not clear; is this caused by tidal interference at extreme spring tides or is it a lunar-related or biological rhythm connected with spawning? Pannela et al. (1968) produced a graph of days per month showing a rise from the present period of 29.53 to a plateau just below 30 days during the Mesozoic and rising to c. 31.5 in the Cambrian. The more recent review of Williams (1989b) shows greater error bars on the data, but a similar trend. Calculations of the period of the Milankovitch band precession and obliquity in the past depend on accurate knowledge of the period of the lunar month in the past.
7
Equinoxes (0.5 a) These are the two times in the year when the Sun is exactly above the Equator and day and night are of equal duration. The vernal (spring) equinox for the northern hemisphere is about March 21 and the autumnal equinox is about September 23: for the Southern hemisphere it is the reverse. Associated with this are the equinoxial tides which form their annual maxima at this time and storms are often associated with climatic changes.
Annual cycle (1.0 a) This is the frequency which has received most study. There are 365 solar days in the Earth year. Environmentally it represents the complete cycle of the dominant extremes of solar radiation reaching the outer atmosphere of the Earth. There are, however, lag effects in how this operates to affect climate at the surface of the Earth where highest or lowest temperatures are delayed. For sediments the classical work of de Geer (1928), on annual varving was on sediments in
Fig. 7. Diumal rhythms and lunar monthly bands shown on the epitheca of a rugose coral from the Bell Shale, Michigan, USA, collected by the author. (Photograph kindly supplied by Dr. C. T. Scrutton and figured Scrutton 1964).
8
M.R. HOUSE
lakes near glacial regimes where spring melts and summer organic debris classically comprise a couplet or two of laminae resulting in a single varve. Varves have led to the establishment of a post-Glacial chronology. A more general term for laminated sediments is laminites since their annual period is not often demonstrable. Lacustrine varves through time have been reviewed by Anderson & Dean (1988). There are many examples of ancient varved sediments. For the Tertiary, the Eocene Green River Formation Colorado (Bradley 1929, 1931; Fischer & Bottjer 1991). The Jurassic Todilto Formation (Anderson & Kirkland 1960) has long sequences of varves. Hallam (1960) has claimed occurences in black shales in the Jurassic. For the Trias long sequences have been established in the Lockatong Formation of the Newark Basin (van Houten 1962). For the Carboniferous Kvale et al. (1989) have recognized annual and higher-frequency cycles. Devonian annual varves have been claimed in the Ireton Shale of western Canada (Anderson 1961) and in the Achanarras Limestone of Scotland (Rayner 1963; Trewin 1986). Precambrian varving has been claimed for the Mid-Proterozoic (c. 1.75 Ga) by Jackson (1985). The Elatina laminites (Williams 1989a, b) suggest that there were then c. 400 solar days to the year. Organisms commonly reflect the annual cycle in their deposited tissues. Annual growth rings in plants have established dendrochronology as a discipline (Creber 1977; Fritts 1976a, b) giving a chronology going back over five millenia [Suess (1970), following fundamental contributions of A.E. Douglas in his monumental Climatic Cycles and Tree Growth (1919-1936)]. Studies have been attempted on Mesozoic material (Creber & Chaloner 1985). Similarly for animals, annual bands have been used for dating fish otoliths and bivalves, and such groups, and in addition brachiopods, stromatoporoids and stromatolites, but the time spans involved a very small. An early well-documented study demonstrating annual banding in the cockle Cerastoderma edule was by Orton (1926), in which growth is very restricted during winter-forming winter bands (Fig. 6).
Orbital cycles between annual and the Milankovitch band (1.0 a-10.0 ka) Frequencies within this span have been refered to as the solar frequency band because solar phenomena and atmospheric and magnetospheric reactions to them dominate (Fischer & Bottjer 1991, p. 1065). But several gravitational elements continue to be important. For convenience these are again treated in order of increasing wavelength.
Chandler Wobble (1.17 a) In 1891 S.C. Chandler noted that latitude variations contained two components with periods of 428 days (14 months) and 1 a. Known as the Chandler Wobble, it is considered to be a free wobble, the period being extended from an expected 304 days (on a rigid-body theory) by the response of the deformable Earth to rotational forces (Munk & MacDonald 1960; Smylie & Mansinha 1971; Chinnery 1971). This seems to have no OFT relevance.
El Ni~o or ENSO (1.0-9.9 a) Off the coast of Peru the cold Humboldt Current from the south usually gives way around Christmas (hence E1 Nifio) to tropical and warmer waters from the north. Changes in rainfall, sedimentation and ecology result. Stronger changes occur about every four years and there may be other peak frequencies up to 9.9 a (Quinn & Neal 1987). More extreme E1 Nifio effects than over the previous century occured in 1982 and 1984. It is now recognized that there are global effects climatic of an E1 Nifio type. Whilst these effects may be initiated by orbital factors their operation is complex Sedimentary cycles at about this period are increasingly being interpreted as ENSO effects. Cycles at 4.8-5.6 in the Eocene Green River Formation (Ripepe et al. 1991) may be of this type, confirming their recognition by Bradley (1929), and there may be a weaker periodicity at 33 a.
Lunar perigee (8.85 a) The distance between the Earth and Moon varies and it is closest, at perigee, every 8.85 a. It has been calculated for the Precambrian Elatina Formation and Reynella Siltstone (65 Ma) at 9.7 + 0.1 a (Williams 1998a, b).
Solar year (c. 11.0 a) Sunspots are localized vortices on the surface of the Sun thought to be produced by magnetic activity: spots often appear in pairs with opposite magnetic polarity. The spots reduce luminosity of the Sun and hence the solar energy received by the outer atmosphere of the Earth (Dickey 1979). In 1843 S.H. Schwale discovered that there is a cyclicity when spots reach maximum of every c. 11 a, and this is termed the solar year (a term sometimes used, misleadingly, also for the earth year); spots move latitudinally during the cycle. Higher periods
9
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Fig. 8. (Top) Showing variation in the daily tidal height from January 1 to April 10 1966 for Townsville, Queensland, Australia, showing fortnightly tides of differing amplitude and the corresponding lunar month period (data from Flinders Institute for Atmospheric and Marine Sciences; after Williams 1989b). (Below) Annual mean tidal range for Boston, Massachusetts, USA, showing the 18.6 a lunar nodal cycle (modified from Kaye & Stucldey 1973; Williams 1989b).
up to 60-120 a have been thought to be due to similar activity but the 22 year cycle (Hale cycle, see below) is the best attested. These are not, of course, controlled by orbital factors but they have some time relevance. Anderson (1965) reviewed Recent, Pleistocene and pre-Pleistocene records of this cycle. It has been, in particular, long recognized in the Eocene Green River Formation, where Bradley (1929) recognized an 11 year cycle: Crowley et al. (1986) calculated a 10.4-year cycle and Ripepe et aI. (1991) a 10.4--14.7-year cycle. In the Jurassic Todilto Formation, Anderson & Kirkland (1960) recognized this cycle, and also higher wavelength cycles. The cycle has been recognized in the Precambrian (Williams 1981). For the midProterozoic Jackson (1985) considered a 7.0-11.0year cycle to be recognizable.
Lunar nodal (18.61 a) The lunar orbital plane relative to the ecliptic plane (Fig. 2) is not constant but precesses with a present day period of 18.61 a. This is the period of the lunar nodal cycle. There is a wobble on the plane of the Moon's orbit around the Earth such that the direction of tilt changes through 360 ~ in longitude on the celestial sphere over 18.61 a. Although the orientation of the tilt changes, the angle between the plane of the Moon's orbit and the ecliptic is constant at e. 5 ~ (Kaye & Stuckley 1973). The effect of this is to give very enhanced tides at this frequency (Fig. 8B). Williams (1989a, b) has claimed to recognize this cycle in the Proterozoic Elatina Formation and Reynella Siltstone (650 Ga) when, he suggests, the period was 19.5 + 0.5 a.
10
M . R . HOUSE
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Hale cycle (c. 22.0 a) A sunspot cycle twice that of the solar year has been been recognized both in historic time and in geological time. It is named after G.E. Hale who, in 1908, recognized the magnetic character of sunspots (Mitchell et al. 1979). Early records at this frequency were reviewed by Anderson (1965). This is the dominant periodicity in the Miocene Sicilian anhydrites (Fischer 1986) and has a powerful signature in the Devonian Ireton Shale (Anderson 1961).
Cycles of the Milankovitch band (10 ka to 1.0 Ma)
Introduction The complex orbital patterns of the EarthMoon-Sun system presumably result from chance patterns on coalescence during their origin. These orbital effects operate through changing the seasonal distribution of insolation and the distance between the Earth and the Sun from time to time. Such changes alter the amount of solar energy, or insolation reaching the outer atmosphere of the Earth. They are appropriately named after Milutin Milankovitch who used such orbital changes to produce a coherent explanation of the ice ages
(Imbrie & Imbrie 1979; Imbrie 1985). The main cycles of precession, obliquity and eccentricity (Fig. 9), although almost sinusoidal, combine to give quite complex patterns (Fig.10), and changes of insolation flux of c. 5%, and of up to c. 100W m -2. The reflection of these outer atmosphere energy changes at sea level in climatic changes is undoubtedly very complex, not only affecting climate, but by thawing or freezing ice, in changing sea level, and in changing climatic regimes, altering erosional and weathering processes so that the nature and rate of sedimentation will change. These various frequencies vary in relation to the Equator. This geographical control is apparently the reason why at various times in the geological past, when obliquity or precession influences may dominate as the major control of orbitally forced microrhythms, the precessional effects are greater at low latitudes.
Precession cycle (19-23 ka) Precession is usually used for the combination of the precession of the equinoxes and the movement of the perihelion. These cycles (or pseudocycles) refer to the movement of the axial projection of the axis of rotation of the Earth with regard to the stars. It was first recorded in 129 BC by Hipparchus. The
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projected axis today lies close to the Pole Star, Polaris, but through time it precesses through a celestial ellipse with a period of c. 26 ka. This wobble of the Earth's axis is caused by the pull of the Moon and Sun on the Earth's equatorial bulge. Other planets introduce a slight additional precession. The fundamental periodicity varies relative to the ellipse of the orbit, i.e., to the equinoxes or perihelion, with two peaks at the present time of 1900 and 23 000 a. The former, the time off
perihelion, is indicated in Fig. 10, and is an important high-frequency modulator.
Obliquity or tilt (41 ka) The projection on to the sky of the Earth's Equator is the celestial equator. This plane lies at 23.5 ~ from the vertical to the ecliptic, the plane of the Earth's orbit. This angle varies by c. 3.5 ~ fluctuating from 21.5 ~ to 24.4 ~ and back with a period of about
12
M.R. HOUSE
Fig. 11. Photograph of a Jurassic sequence in the Atlas Mountains north of Rich showing the small scale couplets groups as bundles of about five couplets and representing the interpenetration of orbital forcing due to the eccentricity and precession cycles (photograph M.R. House).
41 000 a. This changes insolation energy received by the outer atmosphere in two ways; it changes the intensity of the seasonal cycle and it alters the poleto-equator insolation gradient on which climatic and ocean circulation depend. For a given latitude and season, typical departures from present day values are of the order of • (Imbrie 1985).
Eccentricity (54, 106 and 410 ka) There are several wobbles on the orbit of the Earth-Moon system around the Sun (Fig. 9). Some are true eccentricity factors, but also the orbit itself changes from near-circular to an ellipse. The effect of the latter is considerable, and variations in insolation, which are minimal when the orbit is circular, can reach 30% of the total flux at extremes of the ellipse. The most powerful of these are the 106 and 410 ka cycles. Figure 12, based on the analysis of a 2.5 Ma record of oxygen isotope records from ocean cores, shows the signature (labelled 100 ka).
Interpretation Sedimentary rhythmicity is common in the stratigraphical record (Fig. 11) and is much debated. The only striking suggestion that such patterns are orbitally controlled is when couplets of rhythms are combined into groups of about five couplets as bundles (Fig. 11), since this
suggests the operation of an approximate 5 : 1 ratio in agreement with the approximate ratio between the 23 ka precessional and 106ka eccentricity cycles. Although orbital forcing was invoked as a causation of some sedimentary rhythmicity earlier, there is no doubt that it was only the analysis of the oxygen isotope records in ocean cores, now going back over 2.5 Ma (Fig. 12), which has convinced most sceptics of its importance. Precession and obliquity are dependent upon the character of the Earth-Moon system. Since, as demonstrated earlier, there is good evidence that the lunar day and lunar month have changed with time, then so will both the precessional and obliquity periods. A calculation to correct these periods through geological time has been provided by Berger et al. (1989a, b) (Fig. 13). The eccentricity cycles are not changed by modifications of the Earth-Moon system. Many examples of the operation of Milankovitch frequencies below 106 ka have been documented (see especially Berger et al. 1984; Fischer & Bottjer 1991, De Boer & Smith 1994). Particularly valuable have been the correlations between rhythmic sequences and oxygen isotope measurements (Ditchfield & Marshall 1989) confirming temperature dependence and correlations with carbonate percentages (Weedon & Jenkyns 1990), which might be expected to have a temperaturecontrol effect. The frequency of metre-scale rhythms in the stratigraphic record, in many facies,
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was emphasized by House (1986a, b), who argued that a common interpretation was required. Anderson & Goodwin (1990) considered such an allocycle a fundamental stratigraphic unit, building on their earlier work on Devonian rhythms (Anderson et al. 1984; Goodwin & Anderson 1985). Similar cycles have been recognized in terrestrial sequences (Astin 1990), and in bituminous and evaporitic sequences (BougersmaSanders 1971). For operation of longer-term cycles a fascinating study lies ahead. For example, Heckel (1986) recognized a long series of Pennsylvanian sedimentary cycles which he estimated might be 0.5 Ma each, and related to transgressive pulses due to ice-melting. If these are the 410 ka cycles then the framework for a precise timescale is, established. It is unlikely that a continuous timescale at the Milankovitch band level is possible unless long
term-cycles of the sort recognized by Heckel (1986) can form the basis for a framework, and we are clearly not at that stage yet. Meanwhile, the best approach may be to concentrate on stage duration, and the duration of zones within stages, as has been commenced for the Givetian (House 1992 and this volume) and Cenomanian (Gale 1989 and this volume). For this purpose sequences from basinal or pelagic regions with little asymmetricity in rhythms and with good biostratigraphic control are required, i.e. sequences of the sort illustrated in Fig. 11. Sections such as those of the British Lias or Kimmeridge Clay suffer from too many hiatuses and are from too shallow a facies to be very helpful.
Long period orbital cycles (>1.0 Ma) P e r i o d i c extinctions
There has been much interest in recent years in the possibility that orbital events of very long frequencies may have periodically affected the environment of the Earth. The reappearance of Halley's Comet every 76 a appeared to be a trigger for such thoughts, and especially the return in 1985 and 1986. It was the publication of the claim by Raup & Sepkoski (1982, 1984) that extinction of fossil groups occurred at regular intervals during the post-Palaeozoic that led to the proposal of many hypotheses to explain this. The period suggested by Raup & Sepkoski was c. 26 Ma, but 29 and 31 Ma periods were also proposed. Fischer & Arthur (1977) had made similar suggestions earlier but had
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not assigned a period. Associated in timing with these ideas was the claim by Alvarez e t al. (1980) that the Iridium Clay associated with the Cretaceous/Tertiary (K/T) boundary was caused by cometary dust clouds. On the presumption that Raup & Sepkoski's evidence was correct, many theorists suggested that the Earth periodically passed through regions of cosmic dust clouds. The simplest models invoked the cosmic year, and suggested that the solar system oscillated above and below the galactic plane at the required frequency (Fig. 14), taking it in and out of vulnerable space (Rampino & Stothers 1984). Others invoked a periodically returning Companion Star (Whitmire & Jackson 1984), named the Nemesis Star (Davis e t al. 1984), or the periodic effects of a Planet X (Whitmire & Matese 1985); these speculations have continued (Crawford 1985; Clube & Napier 1982). But from the beginning, there were criticisms of the basic conclusions of Raup & Sepkoski. Some claimed falsity in their statistical tests (Noma & Glass 1987; Stigler & Wagner 1987), or bias in the selection of data (Patterson & Smith 1987), or the timescale definitions (Hoffmann 1985)9 That the K/T boundary was caused by cosmic events was also questioned (Clemens e t al. 1981; Hallam 1987). The crudity of Raup & Sepkoski's data, with only 55 data points used to assess extinctions, and the assumption that the stages they used were of equal duration, raised problems. Accordingly the writer assembled data, at 2 Ma time intervals, for
a period of 320 Ma, with 162 data points, for the history of the marine Ammonoidea families from their origin in the Devonian to their extinction at the KIT boundary (House 1989, 1993). This may well represent the most detailed and closely documented such analysis ever undertaken for the fossil record. A Fourier transform of the data (House 1993) gave no evidence whatever of a 26 Ma, or other, long frequency extinction pattern (Fig. 15). Hence the Raup & Sepkoski proposals must be rejected. Were a regular orbital effect proven then it would be of enormous advantage in calibrating those time scales using the Milankovitch band but regrettably there appears to be no such modulator. Cosmic
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The period for the solar system to move around the Milky Way galaxy (Fig.15) has been variously estimated, but is thought to be c. 220-250 Ma. The suggestion has been made that major glaciation periods in the past might be related to this cycle. The Pleistocene (1.6 Ma), Permian (c. 250Ma), Ordovician (c. 440 Ma) and Vendian (600 Ma) do not appear to be separated by equal amounts on current radiometric data, so this suggestion is highly speculative. This contribution, given here as an introduction to this volume on O r b i t a l F o r c i n g T i m e s c a l e s a n d Cyclostratigraphy, is based essentially on the Special Invitation Lecture given by the author to the Geologists' Association in March 1990.
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Fig.
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Equilibrium and response to climatic and tectonic forcing: A study of alluvial sequences in the Devonian Munster Basin, Ireland S E A N B. K E L L Y 1 & S H A U N
E SADLER 2
Department of Geology, University College Cork, Ireland 1present address: Geochem Group, Chester Street, Sattney, Chester, Cheshire CH4 8RD, UK 2present address: Simon Petroleum Technology, Exploration Services, Llanududno, Gwynedd LL30 1SA, UK Abstract: The Upper Palaeozoic Munster Basin is characterised by thick (> 6 km), Late
Devonian alluvial successions (the Old Red Sandstone) that are well exposed throughout southwest Ireland. These represent the deposits of large-scale, complex terminal fan systems which operated under a semi-arid climate. A coarse biostratigraphic subdivision of these sequences has been combined with established radiometric evidence for Devonian stage durations in order to estimate a sediment accumulation rate for this part of the Munster Basin (0.38-0.46 mm a-l). This rate has then been used to estimate the duration of cyclic variations in a range of sedimentary parameters that have been identified throughout both proximal and distal basin-fill sequences. 'Spot'-data derived from variables, such as coset thickness, maximum bed thickness and sandstone content, were used to generate time series which have been subjected to routine Fourier analysis. The application of the estimated sediment accumulation rate to the three most prominent cycle thicknesses (31-175 m) and the assessment of relative frequencies has been used to establish the approximate cycle periodicities, all of which appear to lie within the Milankovitch Band. It is suggested that these features represent the stratigraphic expression of climatic perturbations related to orbital forcing (particularly eccentricity variation). It is further suggested that the recognition of these cycles throughout the basin-fill may, in future, form the basis of high resolution relative timescale and allow the correlation of proximal and distal sections through the basin-fill at scales comparable to the high frequency cyclicities (c. 40 m). The stratigraphic expression of these cycles and their interpretation as evidence of climatic rather than tectonic variation are considered to be a function of their relationship to the equilibrium time constant ( Teq ) of the Munster Basin 9The establishment of Teq requires an assessment of diffusivity which is largely controlled by the rate of water supply to the depositional systems. The latter is assessed by palaeohydrological reconstructions of the scale and discharge characteristics of the major fluvial distributary systems within the Munster Basin. Two methods have been applied. Method I uses a series of established empirical relationships to estimate palaeohydrological variables and Method II uses the thickness of sets and cosets of cross-strata within alluvial sandbodies to assess discharge and channel depth. The resulting estimates of water flux allow the assessment of diffusivity which suggests a Te_ for the Munster Basin of the order of 2 x 106 a. The Milankovitch Band (105 a) periodicities ~73 of the cycles observed in the basinfill are therefore considered to represent 'rapid' variations ( T < < Teq) in water/sediment flux associated with relatively high frequency climatic perturbations. However, these cycles are superimposed on kilometre-scale, sand-prone 'wedges' that dominate the basin-fill architecture. These are considered to record the basinward progradation and subsequent retreat of the terminal fan systems. The 1.5-2 km thicknesses of these units represent time intervals greater than Teq for the Munster Basin. These features are therefore considered to record 'slow' variations (T > Teq) in base level related to changes in tectonic subsidence rate and/or sediment flux.
This paper presents data f r o m the Late Devonian Old R e d Sandstone (ORS) o f southwest Ireland. The ORS has been interpreted as the deposits o f terminal fan complexes which operated under a semi-arid climate and which occupied the western side o f an intracratonic half-graben (the Munster Basin; see Fig. 1). C o m p r e h e n s i v e sedimentological evaluations o f these sequences are available
in previous accounts (e.g. G r a h a m 1983; Williams et al. 1989; M a c C a r t h y 1990; Sadler 1992; G r a h a m et al. 1992; Sadler & Kelly 1993; Kelly & Olsen 1993) and only after a brief s u m m a r y o f the proposed depositional systems is provided here. However, recent analysis o f extensive log sections through this thick (> 6 km), well-exposed alluvial basin-fill succession indicates the presence o f a
From HOUSE,M. R. & GALE, A. S. (eds), 1995, OrbitalForcing Timescalesand Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 19-36.
19
20
S.B. KELLY • S. P. SADLER
Fig. 1. Location and general geological setting of study sections within the northwestern part of the Munster Basin, southwest Ireland.
hierarchy of apparently cyclic fluctuations in a range of sedimentary parameters (Kelly 1992, 1993; Sadler & Kelly 1993). Two of the most important factors influencing aluvial sedimentation over geological timescales are tectonism and climate. Although regular 'cycles', 'megacycles' or 'megarhythms' have been identified in other Devonian alluvial sequences (e.g. Steel 1976), climate is often discounted as a possible driving mechanism. For example, in an assessment of the cyclic sediments of the Hornelen Basin, Bryhni (1978, p. 298) stated that, 'The recurrence of similar megarhythms throughout the entire stratigraphic succession would require a regularity of climate which is rather unlikely'. However, recent work suggests that the influence of regular cyclic fluctuations in climate, which may be related to orbital forcing, is recorded in a number of Devonian continental basins (Hamilton & Trewin 1988; Olsen 1990; Astin 1990; Kelly 1992). Differentiation between the influence of tectonism, climate and other factors, both allocyclic (e.g. eustasy) and autocyclic (e.g. channel switching) in alluvial sequences remains generally difficult. However, the isolation of these separate effects is facilitated if a workable timescale can be applied using, for example, magnetic polarity reversals or the dating of interbedded tuff horizons. An accurate timescale allows the assessment of various rates of processes (e.g. sediment accumulation) and their influence on alluvial stratigraphy or architecture. In this way, Johnson et al. (1986, p. 74), studying fluvial sediments of the Andean foreland basin and the Siwaliks molasse of Pakistan, were able to conclude that, '...the deciding factors in determining exact stratigraphic form are climate and hydrology and not tectonism'. However, in the absence of high resolution
magnetic polarity or radiometric time frameworks for the Munster Basin, previous workers have assessed sediment accumulation rates by reference to a coarse biostratigraphy and accepted Devonian stage lengths (Graham 1983, p. 481). The results of this study indicate that the recognition of Milankovitch Band cyclicity within the Munster Basin allows the development of a high resolution relative timescale. Estimates of sediment accumulation rate allow the application of diffusivity models for alluvial basin-fills (e.g. Paola et al. 1992). These assess the implications of major variables (subsidence rate, sediment flux, gravel fraction, water supply) for the grain-size distribution and architecture of basin-fill successions. One of the main difficulties with diffusivity models is constraining water supply. This paper attempts to resolve the problem by the palaeohydraulic estimation of flow rates in the ancient river systems. Numerous quantitative (empirical) relationships which describe a range of geomorphological and hydrological variables observed within modern fluvial systems are used to evaluate palaeohydrological parameters for the ORS river systems of the Munster Basin. The approach is not a new one and was first applied in detail to the Devonian Wood Bay Group of Spitsbergen by Friend & Moody-Stuart (1972). The objective is to provide estimates of sediment and water flux from the drainage basin (Ad) into the depositional basin (Bb). The basin equilibrium time or 'equilibrium time constant' (Teq) is then calculated using the model developed'by Paola et al. (1992). Teq is the timescale over which a basin naturally responds to cyclic variations in factors such as subsidence and water/sediment flux. In summary, this paper has five main objectives: (1) to provide an estimation of sediment accumulation rates within the Munster Basin; (2) to demonstrate the presence of a hierarchy of cyclicities within the Munster Basin fill and their interpretation as the expression of orbitally-forced climatic variations; (3) to investigate the palaeohydrology of alluvial systems within the Munster Basin; (4) to use palaeohydraulic reconstructions to determine an equilibrium time constant (Te.q) for the Munster Basin; (5) to use T~, to evaluate ttie relative rates of climatically- and "~ tectonically-driven alluvial processes and their expression within the basin-fill.
Location and geological setting The data presented in this paper were recorded from locations in west Co. Cork and south Co. Kerry, Ireland (Fig. 1). Widespread glaciation of this relatively remote upland region has resulted in extensive coastal and inland exposures. These provide thick and locally continuous sections
SEQUENCES IN THE DEVONIAN MUNSTER BASIN through the basin-fill succession which have been logged on a 1 : 5 0 scale. Specific study locations are concentrated in the northeast of the Beara and Iveragh Peninsulas (Fig. 1). Southwest Ireland represents a westerly extension of the Variscan fold belt in Europe. The area is characterized by thick, conformable Late Devonian and Early Carboniferous sedimentary sequences which have been subject to intense deformation and shortening. The structural framework is now dominated by a series of northerly verging, regional anticlines and synclines (Cooper et al. 1986; Price & Todd 1988; Meere 1992). However, shortening has also been accommodated by the development of parasitic, subregional and mesoscale asymmetric fold couplets, a pervasive pressure solution cleavage, thrust propagation on a variety of scales and widespread strike-slip and oblique-slip faulting (Sanderson 1984; Sadler 1992). In addition, subsequent relaxation of deformational stresses and the effects of Jurassic and post-Mesozoic erosional unloading (Keeley et al. 1993) have resulted in the pervasive development of joint sets (Meere 1992). Recent studies have evaluated the distribution and orientation of major compressional faults with respect to regional stratigraphic patterns. Systematic relationships indicate that the current structural configuration of this region is a product of the compressional inversion of a pre-existing
,
..
ll~ Mlddle-Ul~er Devonian
~
21
extensional tectonic framework (Price & Todd 1988; Williams et al. 1989; MacCarthy 1990). The thicknesses of the conformable sedimentary sequences throughout southwest Ireland (locally > 6 km) testify to active sediment accumulation on a subsiding crustal block. The extensional framework is therefore considered to have been related to the development of an intracratonic half-graben the Munster Basin. Subsidence patterns appear to have been largely dictated by extensional offset on a major, southerly dipping, listric fault system that is now traceable in a zone running from Dingle Bay in the west, across to the Galtee Mountains in the east (Williams et al. 1989). However, regional stratigraphical and sedimentological analyses suggest that the extensional framework of the Munster Basin was more complex, with the development of a prominent antithetic structure in the west (Williams et al. 1989) and localized crustal buoyancy associated with proximity to a major granitic pluton (the Leinster Granite) resulting in a complex eastern closure (Price & Todd 1988). Subsidence within the Munster Basin is thought to have been initiated in the late Middle Devonian (Clayton & Graham 1974, 1988) and to have persisted throughout the Late Devonian and into the Early Carboniferous. However, an apparent southerly translation of the locus of crustal extension in the Late Devonian resulted in the superposition of a genetically related sub-basin which
~,'.;,;,71
t.,"-~~
-~-','I
DINGLE
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.:--
.~ ~.~.~.~.~".:-: : :.:.: :iiiiii,
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.
:................',iiiiiiiii!ii!iiiiii':ililiiiiiiil, .
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.:.:::i~;.~.:i:ii!ili iiii::iiii~iiiiiiiiii~ii~ ~ ~ , " , ~ i :~.ii!iiii:iiiiiiiiiiiiii!i!iiiiiiiii!i~ ~i.,iii~i~----,'~,.~'x\,~,~'~"~',~xx~:
:IWS;:;~ ~i:ii~!iiii~i::i~iiii:ii~i!i:?:iiiii:iiiiiiiii~i!i ~
~ , - ~ ' , , \ x , , \ . ' b , , . ~ , ~ ,
r:~,~-.\x,x\\'x\\~ ~ , ~ . ' , % % ~ , 10W
8W
~:,I 6W
Fig. 2. Schematic paleogeography illustrating the main depositional systems identified within the Munster Basin (after Graham 1983; Williams et al. 1989). Note the presence of several large distributory systems which are either transverse (from the north) or axial (from the east and west). This study concentrates on the 'Northern System' which deposited the GPE Also indicated are major tectonic features which are thought to have influenced both ORS and subsequent sedimentation.
22
s.B. KELLY t~ S. E SADLER
dominates the south of the region (the South Munster Basin) (Price & Todd 1988; Sadler 1992). Initially, basin subsidence resulted in the deposition of a thick alluvial succession (Graham 1983), with subsequent marine incursion during the Late Devonian restricted to southeastern areas (MacCarthy 1987). However, marine influence persisted and spread throughout the Late Devonian, resulting in the inundation of the whole of the South Munster Basin by Early Carboniferous times. Subsequently, the continuation of this transgressive regime resulted in the drowning of the northern reaches of the Munster Basin and their transformation into a Carboniferous limestone province (Williams et al. 1989). This contrasted with the marine clastic sedimentation characteristic of the South Munster Basin. Palaeontological data constrain the onset of ORS sedimentation in the west of the Munster Basin to the late Middle-early Late Devonian (see below). Conformable relationships with the overlying, fossiliferous marine sequences suggest marine inundation of this area in latest Devonian ('Strunian') times. Lithostratigraphic subdivision of these thick alluvial sequences is commonly based on the regular presence or absence of medium- and coarse-grained sandstone bodies within a pervasive background sequence characterized by the decimetre-scale intercalation of purple very fine/fine grained sandstones and siltstone. This principle has allowed the establishment of a coarse lithostratigraphy which defines the basin-fill architecture. The vertical distribution of mediumand coarse-grained sandstone bodies defines
A
two, kilometre-scale clastic wedges [Chloritic Sandstone/Slehany Formation (CSF) and the Purple Sandstone/Gun Point Formation (GPF)] (Fig. 3) These thicken to the north and appear to amalgamate adjacent to the prominent basin margin fault. A similar coarse-grained alluvial "tongue' has been identified in the south of the region [Sherkin Sandstone Formation (SSF)]. All three project into the finer-grained background sequence (Valentia Slate/Caha Mountain/Castlehaven Formations) (Fig. 3).
Depositional m o d e l Recent sedimentological analysis of the ORS in southwest Ireland has established that the sandprone wedges which extend into the Munster Basin represent the deposits of large-scale terminal fans (Graham 1983; Williams et al. 1989; Graham et al. 1992; Sadler 1992; Sadler & Kelly 1993; Fig. 2). General models of terminal-fan sedimentation, with specific reference to Devonian examples, are provided by Friend (1978) and Kelly & Olsen (1993). The scale and grain-size distribution characteristic of depositional systems within the western part of the Munster Basin suggest that they comprised major distributary networks. These were marked by systematic downstream variations in channel size and sediment grade due to bifurcation, transmission losses and evaporation. Such trends are also characteristic of modern, semi-arid terminal distributary systems, such as those described by Parkash et al. (1983) and Abdullatif (1989). The main lithostratigraphic unit considered in
A~
2km
0
Fig. 3. Schematic north-south cross-section through the ORS in the northwest of the Munster Basin (predeformation), illustrating the main, sand-prone wedges that prograded from the north into finer grained, 'basinal' areas.
SEQUENCES IN THE DEVONIAN MUNSTER BASIN this paper is the GPE A detailed account of the sedimentology of this unit is provided by Sadler & K e l l y (1993). The proximal 'feeder' zone was characterized by large-scale, low sinuosity channels which transported sand and gravel through a low relief, muddy alluvial plain. Further downstream, widespread channel multifurcation in the 'distributary zone' initially resulted in the development of abundant, highly mobile, low sinuosity, sandy bedload streams. Downstream transmission losses due to infiltration and evaporation resulted in the replacement of variable but persistent flows by high energy, unconfined ephemeral floods. These events locally resulted in the scouring and episodic filling of transient, low sinuosity ephemeral channels in the distal reaches of the distributary zone. Eventually, flows reached a 'basinal zone' in which a persistently quiescent regime was episodically punctuated only by low magnitude, shallow sheetfloods. These mainly represented the residual expression of flood pulses originating in the distal distributary zone. T i m e - s e d i m e n t a c c u m u l a t i o n rates In order to assess the rates of various palaeohydrological processes, it is necessary to evaluate time in a stratigraphic context by reference to sediment accumulation rates. Measured or calculated sedimentation rates generally vary in inverse relation to the length of time over which the measurement is made (Sadler 1981). This is a function of the discontinuous nature of sedimentation within the alluvial environment, with accumulation typically punctuated by phases of erosionor non-deposition, and therefore rarely persisting for more than a few days or weeks at any one locality (Sadler & Strauss 1990). The significance of this inherent 'unsteadiness' will be a function of the relevant timescale. At timescales of > 105 a, alluvial units such as the GPF are likely to be stratigraphically 'complete' (Johnson et al. 1988). However, at c. 104 a level of time resolution, the GPF is likely to become 'incomplete', with sedimentary hiatuses introducing significant vertical fluctuations in sediment accumulation (Johnson et al. 1978). Miall (1978) showed that non-marine basins, in various tectonic settings, have sedimentation rates (averaged over time periods of the order of 106 a) of 0.03-1.5 mm a q, and a survey of recent literature suggests that sediment accumulation rates are most frequently in the range of 0.1-0.6 mm a -1 for foreland and successor basins (Miall 1978; Sadler 1981). It has long been recognized that many ORS basins experienced high sediment accumulation rates, primarily as a result of rapid subsidence. Periods of rapid subsidence typically appear to have been 5-10 Ma in duration and rarely persisted
23
for more than c. 20 Ma (Friend 1969), with sediment accumulation rates generally in the range of 0.2-0.6 mm a -1 (Friend 1969, p. 708). Sediment accumulation rates for the Munster Basin may be estimated using radiometric determinations of Devonian stage durations. Sparse fish debris and palynological data indicate that the oldest strata exposed in the Munster Basin are upper Givetian/lower Frasnian in age. Palynological dating of the overlying coastal/ shallow marine sequences indicates that nonmarine sedimentation was largely complete by the end of the Fammenian, although the Tournaisian marine transgression was strongly diachronous (Clayton & Higgs 1979; Sadler 1992). Devonian stages were not equal in length (Friend & House 1964). Although, many earlier workers simply assigned average durations of 7 Ma to each of the Devonian stages (Friend 1969), more recent estimates give durations of 4.6, 10.5 and 3.5 Ma for the Fammenian, Frasnian and Givetian stages, respectively (Harland et al. 1989). It may therefore be estimated that the 6 km alluvial succession observed in the depocentre of the Munster Basin accumulated over a period of 13-16 Ma. This yields a mean sedimentation rate (a) of 0.38-0.46 mm a-1, which is similar to the 0.4 mm a 1 estimated by Graham (1983). These estimates are comparable to sediment accumulation rates derived from other thick GivetianFammenian non-marine successions using a similar approach. For example, the c. 7 km cumulative sediment thickness in the East Greenland Basin indicates sediment accumulation rates of up to 0.5 mm a -1 [Friend 1978 (p. 534), 1981 (p. 153); Friend et al. 1983 (p. 41)]. Cyclicity Well-exposed sections through the basin-fill were routinely logged at a 1 : 5 0 scale. 'Spot'-data recorded include coset size, mean/maximum set size and bed thickness. In order to generate pseudotime series, frequency and percentage data (e.g. calcrete frequency, sandstone percentage) were grouped or averaged into 10m intervals. This allowed data to be analysed using a regular sample spacing. In general, the generation of time series for 'spot' data was based on the maximum value recorded within a 10 m interval. Before proceeding further, it is necessary to consider the significance of the sampling scheme in relation to the statistical characteristics of the palaeohydrological record (cf. Hirschboeck 1988). The preserved sedimentary record represented by the GPF suggests that the recurrence interval (r) of aggradational 'events' related to channel/flood activity and the magnitude-frequency distribution
24
S.B. KELLY & S. P. SADLER
of aggradational 'events' generally have a log normal character, which is typical of both sedimentologic and hydrologic time series (Wolman & Miller 1960) and the rapid response of sedimentary phenomena linked to sediment transport (e.g. bedform development) (Jackson 1975, p. 1528). The selection of a maximum value within a 10 m interval suppresses the filtering effect of stochastic sedimentary processes (Rachoki 1981; Howard 1982). The latter is thought to be minor, given that the sedimentary record is relatively complete (sensu Crowley 1984), considering the character of the depositional system and the timescales involved (see above). Assuming a sediment accumulation rate of approximately 0.4 mm a -1, the sample spacing is equivalent to a time period of c. 25 ka. This is greater than the timescale of any anticipated complex responses (cf. Howard 1982). Although the effective sampling is not uniformly spaced (the original data are discrete), the variation in spot data point spacing is minor in relation to the thickness of the stratigraphic intervals considered. By analogy with modern hydrologic records, we are effectively sampling the maximum recorded flood event in a period of c. 25 ka. The log normal distribution of the data requires the application of a logarithmic transform [x'n = log (Xn)] prior to any time series analysis (cf. Hinnov & Goldhammer 1991).
Individual 'time series' were subject to routine Fourier harmonic analysis and the results are illustrated in Fig. 6 as periodograms. The frequency range extends to 0.05 cycles m -~, which represents the Nyquist frequency for the 10m sample spacing. The majority of the plots illustrate simple analysis of complete data sets, allowing the detection of any regular, long-wave periodicities. Although the periodograms are often 'noisy' they clearly indicate the presence of several pronounced peaks or lines. Three dominant frequencies are definable in nearly all of the spectra from the GPF. The wavelength (m) and estimated duration (based on 'reasonable' values for the sediment accumulation rate; 0.38-0.46 mm a -~) of each of the frequencies are provided in Table 1. Inspection of the various logs (Fig. 4) also reveals the presence of the P1 and P3 cyclicities which correspond to intervals of c. 40 and 150m, respectively. Significantly, all cycle periodicities appear to lie within the Milankovitch Band. This assumes reasonable errors involved in the estimation of sediment accumulation rates using stage durations. However, this method of estimating sediment accumulation rates remains generally uncertain and, in the absence of an absolute timescale. The use of relative frequencies may yield information concerning the possible origins of cyclicity (Fischer
North
South
Derrincullig-Eskna brock
....
~
South Beara
:_----
I
i-,.
0
I
i Im
coset thickness (cm)
1000
1~0
1o0~
maximum sandstone bed thickness (cm)
Fig. 4. Correlation of main study sections within the GPF using variations in sedimentary parameters (all data have been smoothed using a spline function). Note the correlation of small-scale cycles (c. 40 m) and the presence of longer cycles (c. 150 m). Both of these periodicities are thought to reflect climatic changes related to Milankovitch Band orbital parameters, the 100 and 400 ka eccentricity components.
SEQUENCES IN THE DEVONIAN MUNSTER BASIN & Bottjer 1991). The ratio of the two most prominent cyclicities, P1 and P3, is c. 1 : 3.8, which is comparable to the 1:4.1 ratio of the 101 and 414 ka orbital eccentricity cycles (cf. Kelly 1992). If the P1 and P3 cyclicities do represent the expression of these eccentricity components, then the P2 periodicity may correspond to longer elements of the c. 100 ka eccentricity cycle, which are considered to have periodicities of 122-133 ka (Berger 1978; Hinnov & Goldhammer 1991). This conclusion is in accordance with the relative frequencies and apparent duration of the various periodicities based on estimated sediment accumulation rates (Table 1). Longer periodicities have also been detected, the most prominent of which is the c. 700 m P4 periodicity (see Fig. 5), (cf. Sadler & Kelly 1993). If the estimated sedimentation rates are of the correct magnitude then this periodicity may correspond to the 2100ka eccentricity component (Hinnov & Goldhammer 1991). Climatic perturbations, such as those predicted by the Milankovitch Theory, will influence both the sediment and water flux of a closed alluvial system through changes in precipitation and subsequent run-off. Recent studies have confirmed that cyclic changes in insolation are capable of inducing cyclic variations in precipitation (e.g. Rossignol-Strick 1985). The following summary of responses to climatic change is based on numerous studies of Quaternary fluvial systems (e.g. Williams 1970; Dorn et al. 1987; Bull 1988; Maizels 1990) and sedimentological observations of the cyclical deposits within the Munster Basin. Increased aridity is likely to result in a reduction in run-off from the drainage basin and a lowering of water tables within the depositional basin. This is associated with a decrease in streamflow, which may be supplemented by localized fluvial incision on fan surfaces. Conversely, an increase in run-off will promote degradation in the drainage basin and a proportional increase in sediment flux within the depositional basin. Increased run-off will also be reflected by increased discharge or streamflow and an associated increase in the frequency and magnitude of flooding.
Table 1. Cycle lengths (m) derived from power spectra (Fig. 6)
Cycle P1 P2 P3
Wavelength (m) 31-42 55-62 110-175
Estimated periodicity (ka) 78-104 138-155 275-438
The estimated duration of cycles uses a mean (compacted) sediment accumulation rate (a) of 0.4 nun a-1.
25
The rates and processes of calcrete formation are also influenced directly by climate through rainfall and subsequent changes in run-off, which influence the leaching regime, and indirectly by variation in the availability of airborne Ca 2§ (Machette 1985; McFadden 1988). Evidence from Quaternary calcretes suggests that climatic variations may produce systematic changes in profile maturity (Bryan & Albritton 1943). In semi-arid environments, calcrete formation is likely to be initiated when there is a sufficient decrease in effective precipitation (Bull 1991), possibly in association with an increased supply of airborne Ca 2+ (Machette 1985). The maturity of profiles will also increase when the sediment flux is reduced (Wright 1990), as anticipated during relatively arid periods. In spite of dating problems associated with the lack of detailed biostratigraphic and radiometric data within the Munster Basin, it is concluded that long-term sediment accumulation rates were probably between 0.3 and 0.5 mm a 1 and that some of the cyclicities recorded in these successions represent the sedimentary expression of the astronomical forcing of the Earth's climate. The recognition of this 'cyclo-stratigraphy' allows the correlation of the two main study sections within the GPF (Figs 4 & 5). When the Milankovitch Band cyclicity is combined with conventional lithostratigraphic criteria (based on detailed field mapping) the correlation of sediment packages at the same order as the P1 cycles (i.e.c. 40 m) is possible. This is remarkable considering the significant downslope distance between the two sections (30-40km following palinspastic restoration).
Palaeohydrology The interpretation of these cyclicities as the product of orbitally forced climatic changes rather than tectonic pulses may be clarified by estimating the rate at which the depositional basin (Munster Basin) is likely to have responded to imposed fluctuations in autocyclic variables. This involves an assessment of diffusivity (Paola et al. 1992) which is largely controlled by water supply. This parameter may be constrained by the palaeohydrological reconstruction of the Devonian fluvial systems. The estimation of palaeohydrological parameters from alluvial sandbodies is possible using a number of empirical relationships (e.g. Williams 1988), although many workers insist on the use of confidence intervals for statistical significance (Leeder 1973; Ethridge & Schumm 1978) and errors may be introduced by the improper algebraic manipulation of empirical equations (Williams &Troutman 1987; Williams 1988). In addition, such power functions do not
26
s.B. KELLY • S. E SADLER N
~
( ;lenlle.~kEsknabrock
~
South Reara
~
~
.....~z~ ~~~
34
~
S
",
~.~
/;ili
Itane % oleann
sandstone %
sandslone %
readily account for the large and generally unpredictable natural variability in relations between one site and another (Knighton 1975; Richards 1977; Rhodes 1977; Park 1977). Although it is generally accepted that the sedimentary basins of the ORS continent probably experienced warm to hot climates, this does not necessarily imply that discharge patterns in rivers which fed the depositional systems were strongly ephemeral or 'flashy'. A significantly different ('wetter') climate may have prevailed in the upland source areas where discharge is likely to have originated (Allen 1986). Consequently, data from drainage basins in a wide variety of climatic, topographic and tectonic settings have been used to develop empirical relationships from which the palaeohydrology of the Munster Basin may be estimated. Palinspastic facies maps provided by Williams et al. (1989) suggest that the terminal fan systems in the 'northern' part of the Munster Basin (CSF and GPF systems) had a maximum downstream extent of 90-110 km (Fig. 2). This assumes a N-S layer parallel shortening strain of c. 2 : 1 (e.g. O'Sullivan et al. 1986). The easterly and westerly limits of the fans are less well constrained but
Fig. 5. Correlation of basin-fill sequences as recorded in the two main study sections, illustrated by logged sandstone content (smoothed using a spline function). Note that the CSF variations in sandstone content are more 'rapid' fluctuations (P4 cycles) with a periodicity of c 700 m (in addition to the relatively short-term P1, P2 and P3 cycles illustrated in Fig. 4). These pulses are thought to reflect long-term climatic change due to cyclical variations in the eccentricity component with a periodicity of 2100 ka.
recent mapping suggests that the systems were c. 120 km wide. These dimensions are consistent with the data plotted by Heward (1978, p. 677, Fig. 3) which indicate that alluvial fans with areas ranging from 103 to 104 km 2 are characterized by fan lengths generally equal to, or frequently greater than, fan widths. The depositional area (Bb) covered by the large-scale 'northern' terminal fan system in the western Munster Basin was therefore c. 13 2 0 0 k m 2. In order to estimate the sediment/water flux of this system two approaches have been used. The first method uses quantitative relationships between factors such as drainage basin size, sediment yield and discharge observed in modern drainage basins. The second method uses direct estimates of channel depth and discharge based on sedimentary structures within alluvial sandbodies. The results of these palaeohydraulic reconstructions will be used to establish an equilibrium time constant for the Munster Basin. Method 1
Rates of denudation for particular present-day drainage basins have been calculated using the
SEQUENCES IN THE DEVONIANMUNSTER BASIN sediment transport rate in rivers or the rates of sediment accumulation in reservoirs (Langbein & Schumm 1958; Schumm 1963). As the Munster Basin was effectively closed, the volume of sediment within the individual depositional systems provides a direct measure of the amount of denudation that took place in the drainage basin (cf. Friend & Moody-Stuart 1972). The known area of the depositional system (Bb) and the estimated rate of sediment accumulation (a) are used to establish the volume of compacted sediment deposited annually
(v~):
Vc = aB b
(1)
where a is in m a l and Bb is in m 2. The value of Vc is then used to calculate the uncompacted volume of sediment deposited annually (Vu) using the relationship: Vu = Vc pc/Pu
(2)
where Pc and Pu are specific weights of compacted and uncompacted sediment with assumed values of 2700 and 1380 km m -3, respectively. The latter is an average value derived from 'open air' density estimates of 1490, 1314 and 1 2 5 0 k g m -3 for alluvial sand, silt and clay, respectively (FAO 1981, p. 173; see also Ingles & Grant 1975, p. 305). It should be noted that values of 9u will vary according to grain size and clay content. A value for a of 0.4 m m a -1 provides an estimate of c. 107 m 3 a -1 for Vu. An alternative method would be to use porosity evaluations to decompact the vertical sequence, although this is largely precluded by the virtual absence of primary porosity within these sediments as a result of their low grade metamorphic character (lower greenschist facies). However, in spite of the lack of porosity estimates and complexities relating to non-quantifiable volume loss during deformation, decompaction factors of 2 for sandstone and 3 for siltstone/ claystone are considered reasonable. In order to establish a crude mean decompaction factor for the basin-fill succession, the distribution of sandstone and siltstone must be accounted for. Sandstonedominated proximal and medial sequences are effectively "balanced' by more silt-prone distal and basinal deposits. The vertical alternation of proximal and distal sequences within the basin-fill therefore suggests that a crude mean decompaction factor of 2.5 is likely to be most applicable. This value compares well with that derived from Equation 1. The value of Vu derived in Equation 2 may be used to estimate the size of the drainage basin (Ad) through the relationship: Ad = 0.0168 VO'92
(3)
27
(R 2 = 0.82: standard errors; intercept (i) = 0.34, coefficient (c) = 0.05 log units). This relationship was derived from regression analysis of data from 89 drainage basins with values of Vu in the range 103-101~ 3 a -1 and Ad varying in the range 10~ 2. For a = 0.4 mm a t , Ad is estimated to be c. 50 000 km 2. The empirical relationships (2) and (3) suggest a direct link between drainage area and sediment yield. Both run-off and sediment yield from drainage basins are influenced by a number of additional, dependent and independent variables, which include bed-rock properties, drainage basin elevation/relief, main channel slope, basin length, rainfall intensity (duration and frequency) and vegetation (Langbein & Schumm 1958; Jansen & Painter 1975; Walling & Webb 1983). However, for a wide range of climatic regions, sediment yield (m 3 km -2 a -1) consistently decreases as a function of drainage area (Branson et al. 1981). This trend appears to be related to storage of eroded sediment within slope and channel systems (Trimble 1975). Climatic change and sediment source also have implications for sediment yield and the dimensions of the resulting alluvial fan (Bull 1991, p. 276). Empirical relationships between alluvial fan area, Af, and drainage basin are Ad, are expressed in the form Af = c A ~ (Bull 1964; Hooke 1968). The constants vary according to the lithology of the drainage basin. For a drainage basin with an area of 50 000 k m 2 comprising 60-70% sandstone, a fan area of c. 18 000 k m 2 is anticipated. For a drainage basin of the same size, but schistose in composition, a fan area of 43 000 krn 2 is predicted. The volumes of Holocene fan deposits correspond to estimated bedload sediment yields that range from 50 to 200 m 3 km 2 a ~ for watersheds underlain by granitic and metamorphic rocks. Significantly, estimates for the drainage basin that supplied the GPF depositional system suggest a sediment yield of c. 2 0 0 m 3 km -2 a -1 which is consistent with a proposed derivation from predominantly metamorphic rocks and older sediments within the Caledonides. The estimated drainage basin area (Ad) is used to establish mean annual discharge (Qd) through the relationship: Qd = 0.019A ~
(4)
(R 2 = 0.81: standard errors; intercept (i) = 0.11, coefficient (c) = 0.03 log units). Equation (4) is based on data from 259 drainage basins worldwide, with Ad in the range 10110 7 k m 2 and Qd in the range 10-1-105 m 3 s-1. This relationship is similar to that of Leopold et al. (1964), who further state that the relationship between discharge and drainage area can be
28
S.B. KELLY & S. P. SADLER
Table 2. Estimates of drainage basin area (Ad), length (Ld) , mean discharge (Qd) using Method I Sediment accumulation r a t e (mm a-1) a 0.1 0.2 0.3 0.4 0.5 0.6
Drainagebasin area (km2) Ad
Drainage basin length (km) Ld
Drainage basin mean annual discharge (m3 s-1) Qd
13 000 25 000 36 000 47 000 58 000 68 000
330 450 540 620 690 760
80 130 180 230 280 320
Various sediment accumulation rates (a) have been used, although the most likely rate is thought to be approximately c. 0.4 mm a-l.
expressed in the form Qf = aA b, where Qf is bankfull flood discharge and Ad is drainage area. The values of a and b depend on numerous factors including relief and climate. These relationships suggest estimated palaeodischarge volumes for the GPF drainage basin of c. 200 m 3 s-1 (Table 2). The length of the drainage basin (Ld) may be estimated using the relationship: L~ = 2.5A ~
(5)
(R2 = 0.83: standard errors; intercept (i) = 0.19, coefficient (c) = 0.03 log units). This relationship is based on data from 52 drainage basins with values of Aa in the range 104-107 km 2 and Ld in the range 102-104kin. Equation 5 is similar to that given by Hack (1957) and indicates that the relationship is fractalic (Leeder 1993). The estimates for Ad, Ld and Qa are given in Table 2 for varying sediment accumulation rates. The values for Ad and Ld are similar to those derived for Lower Old Red Sandstone drainage in the British Isles by Simon & Bluck (1982), who estimated a drainage basin of at least 5.2 x 104 km 2, with a 'valley distance' of c. 530 km. Method II
The second method of palaeohydraulic reconstruction uses the thickness of sets and cosets of cross-strata within sandbodies to assess discharge and channel depth. Allen (1968, Fig. 6.4) has shown that the mean height of dunes is proportional to flow depth. This has been supported by a number of field studies (Williams 1971; Wijbenga & Klassa 1983). Echo sounder profiles illustrated by Cant (1982, Fig. 31, p. 127) and Collinson (1986, Fig. 3.11, p. 28) indicate that bedform size will respond rapidly to discharge variations. A recent detailed field study by Gabel (1993) confirms the rapid response of dune height, length and migration rate to changes in discharge.
The size of the original bedforms is assessed by reference to preserved set height. The relationship between these parameters has been considered by several workers (Nummedal 1973; Paola & Borgman 1991). Harms &Fahnestock (1965) found that approximately half the dune height is preserved. Relative dune height [ratio of dune height (h) to water depth (d)] can vary from c. I (Simons 1971) to a very small value, with an average between 1/6 (Yalin 1964) and 1/3 (Nordin & Algert 1965). Allen (1968) derived the relationship h = 0.086d 1"19, where h is the mean bedform height (m) and d the mean water depth over the bedform (m). Although Allen's data indicate a _+50% variation in water depth for a given bedform height the relationship has been used by previous workers (e.g. Miall 1976). An alternative method of estimating channel depth is to assume that the bedforms achieved 'optimal' steepness values such that h/I= 0.6, where h is dune height and I is bedform wavelength (Haque & Mahmood 1986), and that the depth of the flow is I/7 (Jackson 1976). In order to reconstruct the dune height a preservation ratio of 0.24 is used (Paola & Borgman 1991). Combining these relationships yields d = 9.9Sp, where d is flow depth and Sp is preserved cross-strata set thickness. In order to estimate mean annual basin discharge (Qb) the following relationship is used: Qb = 32d18
(6)
(R2 = 0.68: standard errors; intercept (i) - 0.04, coefficient (c) = 0.01 log units). This relationship was derived from regression analysis of 150 rivers with Qb in the range of 0.5-18 000 m 3 s-1 and depths (d) in the range of 0.15-8.0 m. The power function given in Equation 6 can be compared to those of other workers (e.g. Knighton 1987) which indicate that the water surface width, w, and mean flow depth, d, of a stream change in proportion to some power of
SEQUENCES IN THE DEVONIAN MUNSTER BASIN water discharge, Q, such that w = aQ b and d = cQ a. For sand bed rivers, w - - a Q ~ and d = cQ ~ where likely values of a and c are 4.3 and 0.56, respectively. These relationships have been used to estimate the discharge of the rivers within the depositional basin. The minimum, mean and maximum set sizes observed in the basin-fill sequences indicate the range of discharges within the channels. The results are given in Table 3. Significantly, estimates of discharge in the depositional basin (Qb) derived from Method II are comparable to those for the drainage basin (Qa) provided by Method I though it is emphasized that Method II is independent of estimated sediment accumulation rate. Both methods suggest that the GPF depositional system could have been fed by a single river which had a mean annual discharge of c. 200 m 3 s-1. It is also possible to estimate gradient (slope) within the depositional basin, as various studies have demonstrated an empirical relationship between gradient and water discharge for different channel patterns (e.g. Leopold & Wolman 1957). Power functions derived by a number of workers indicate that gradient (G) and mean discharge (Q) are related by a power function of the form G = aQ c, in which the exponent c is of a constant (fixed) value and all variation, caused largely by the effects of sediment, is expressed in the coefficient a. For braided streams with discharges in the range 1.25-200 m 3 s-1 and a mean slope of 0.00096, Osterkamp (1978) established that a = -0.0019 and c = -0.31. These values are similar to those given by Leopold & Wolman (1957)(a= 0.0125 and c = - 0 . 4 4 ) and Lane (1957) ( a = 0 . 0 0 4 1 and c = -0.25). The differences between these relation-
Table 3. Channel depth (d) and discharge (Qb) estimates based on cross-strata thickness (Method II) Section within GPF
Depth (m) d
Discharge (m3 s-l) Qb
KE (mean) KE (min) KE (max) KF (mean) KF (min) KF (max)
2.8 1.9 4.2 2.6 1.9 3.6
210 100 420 180 100 330
Within each coset, the maximum set thickness observed has been used. The depth is then estimated using the two methods described in the text, the value d being the average of these two estimates. (Note: the difference between the two values for depth was never more than 17% and was, on average, only 5%). Section KE and KF refer to the lower and upper logged sections from Iveragh (see Fig. 4).
29
ships are due mainly to variations in the grain size of the sediment transported and the use of bankfull discharge rather than mean discharge in some of the earlier data sets. The relationship given by Osterkamp (1978) is for streams dominated by fine- to medium-grained sand, whereas the other relationships are based on data sets derived from both sand and gravel bed channels. Typical gradients on large fans dominated by braided streams are typically in the range 0.001-0.0003 (Stanistreet & McCarthy 1993). Using the relationship G = 0.0041Q -~ and assuming a mean Q of 200 m 3 s-1, a slope estimate of 0.001 is obtained for the GPF depositional system.
Basin equilibrium and response Many recent basin modelling exercises use a concept of diffusion which assumes that sediment influx is proportional to slope and transport in the downslope direction (Colman &Watson 1983; Kenyon & Turcotte 1985; Waltham 1992). However, it is noted that diffusive relationships between deposition and streamflows (e.g. Flemmings & Jordan 1989) imply a linear relationship between transport and slope and it has been suggested that this may represent too simplistic an approach (Leeder 1993). The following discussion is based on the model of basin filling developed by Paola et al. (1992). This model combines a linear diffusion equation for sediment dispersal with a simple mass balance model based on selective deposition for grain size partitioning and downstream fining. Paola et al. (1992) demonstrated from first principles that diffusivity is mainly controlled by the average rate of water supply and the type of streams in the system. The model allows the assessment of four independent governing variables whose influence on basin sedimentation is significant: input sediment flux, subsidence rate, gravel fraction in the sediment supply and diffusivity (which is mainly controlled by water supply). The stratigraphic response of the basin to all of these factors, except gravel fraction, depends on the ratio between the timescale over which the changes take place and a time constant that is unique to each basin or depositional system. The models of Paola et al. (1992) make use of the fact that natural systems take a finite length of time in order to respond fully to any imposed change and that the period of time necessary for recovery will reflect the magnitude and frequency of the imposed change and the magnitude of the geomorphic system (Howard 1982). Based on the scaling of the equations, Paola et al. (1992) defined the characteristic time constant for basin response, the equilibrium time
30
s.B. KELLY & S. P. SADLER
constant (T..), as To. = L2/v, where L is the basin length in th~ directl~)n of transport and v is the diffusivity. In physical terms, if a basin is allowed to evolve under constant rate and distribution of sediment input, the surface topography (stream profile) will reach steady state in a time approximating to Teq. For basin modelling purposes, the real importance of the equilibrium time constant is that the form of the basin response is determined by whether the timescale of the superimposed variations, T (e.g. variations in basin subsidence or variations in water/sediment flux due to climatic change), is substantially greater or less than Teq. Paola et al. (1992) state that 'slow' processes operate at timescales such that T >> Teq' whereas 'rapid' variations are those operating at timescales of T << T Estimates for the Te of the western Munster Baslqn are given in Tables ~ and 5. These use the palaeo-
PI |
0,;1
o.&
P3
South Beara Caha M o u n t a i n Formation
i
0,00
hydraulic estimates derived from Methods I and II as inputs to the model of Paola et al. (1992). The variation of Teq values in Table 4 (1.4-6.1 Ma) illustrates the influence of the sediment accumulation rate (a) in Method I. When these estimates are translated into stratigraphic thickness (Heq) the results have a relatively narrow range (600-860 m). Mean discharge values derived from Method II indicate a Teq of 2 "3 Ma ' which would correspond to a Heq value of 900 m (for a = 0.4 m m al). These results, based on two substantially different methods of palaeohydraulic reconstruction, therefore suggest a Teq of c. 2 Ma, which corresponds to a Heqvalue of 8 0 0 - 9 0 0 m. The consistent P1, P2 and P3 cycles detected in the GPF have wavelengths of c. 40-150 m, and it may therefore be concluded that these Milankovitch Band cyclicities are indeed the expression of 'rapid' variations, such that T < < Teq. Furthermore, by
o.d3
o.&
u,&
0,00
o,dl
0,02 o,d3 Cycles per metre
~
~;
o,dl
o,d2 0s Cycles per metre
Cycles per metre
P2
0,00
0,01
South Beara G u n Point F o r m a t i o n M a x . B e d t h i c k n e s s data
P2
South Beara Caha Mountain Formation S a n d s t o n e F a c i e s % data
1
0.04
0,05
PI S~t:Pi!i~aetr~7:r g adi ~n
nl
0,02 0,03 Cycles per metre
0,04
0,05
0,00
o,&
o,d5
P4 Section F Purple Sandstone Formation c o s e t data
~,L 9
Pl |
PI
,
0,00
0,01
0,(~2 0,(~3 Cycles per metre
0,(~4
0,65
0,00
0,01
,
,,,
South Beara G u n Point F o r m a t i o n C a l c r e t e f r e q u e n c y data
,
0,02 0,03 Cycles per metre
,
0,04
,
0,05
Fig. 6. Periodograms for the pseudo-time series generated from logs through the GPF and Caha Mountain Formation. Power spectra have been smoothed using a Hanning filter. 'Spot'-data (e.g. coset size and maximum bed thickness) were log transformed prior to Fourier analysis as the distribution of these sedimentary parameters is log-normal. Note the regular presence of the three main periodicities, P1, P2 and P3, which correspond to wavelengths of 31-41, 55-62 and 110-175 m, respectively. Using an estimated sediment accumulation rate of 0.4 mm a-1, the periodicities correspond well with elements of the Milankovitch Band eccentricity component.
SEQUENCES IN THE DEVONIAN MUNSTER BASIN
31
Table 4. Estimates for the diffusivity (v) and basin equilibrium time constant (Teq) using Method I Sediment accumulation rate (mm a-1) a 0.1 0.2 0.3 0.4 0.5 0.6
Diffusivity v
Equilibrium time constant ( 106 a) req
Equilibrium thickness (m) Heq
0.002 0.004 0.005 0.007 0.008 0.009
5.5 3.2 2.3 1.8 1.5 1.3
550 630 690 730 760 790
Note that the effect of varying the sediment accumulation rate (a) is limited in determining the 'equilibrium thickness' (neq).
comparison with the models of Paola et al. (1992, Fig. 4) it seems most likely that the cyclic variations observed within the GPF are the result of rapid variations in sediment flux or diffusivity (water flux). This is in accordance with their interpretation as the expression of orbitally-forced climatic oscillations. Significantly, the relationship T -- T,~ (Paola et al. 1992) may be applicable to the largerS~cale (c. 700 m) P4 cyclicity. This is also in accordance with the interpretation of P4 as the product of long-term orbital eccentricity variations. Significantly, the models of Paola et al. (1992) indicate that rapid changes in subsidence have a limited impact on the nature of the basin-fill. In order for changes in subsidence rate to control sedimentation patterns, sediment must be redistributed across the whole basin. This becomes increasingly difficult on timescales less than the equilibrium time (Paola et al. 1992). It is therefore unlikely that the rapid, high magnitude cyclic variations observed in the Munster Basin-fill are the result of subsidence variation. In contrast, the 1.5-2 km scale of the major clastic wedges that dominate the gross architecture of the Munster Basin (Fig. 3) indicates time intervals greater than that of the estimated Ton (i.e. T >> T,n). It therefore seems possible that the"se features w"ere generated by long-term ('slow') variations in basin sub-
sidence (cf. Paola et al. 1992; Heller & Paola 1992). Base level appears to have remained effectively constant in elevation during the deposition of the main basin-fill component considered here - the GPE However, transverse base level migration may have occurred in response to long-term (106a) changes in:(1) basin subsidence; (2) sediment supply. Prior to the deposition of the GPF relatively high rates of subsidence effectively 'trapped' vertically aggrading, sand-prone fan systems near the basin margin (cf. Blair & Bilodeau 1988; Steel & Ryseth 1990). However, the GPF (Stage 2 of Graham et al. 1992) appears to record a net decrease in subsidence rate that promoted the basinward progradation of the fluvial systems, as well as their vertical aggradation (Sadler 1992; Graham et al. 1992). Williams et al.(1989, p. 127) state that the GPF, '...exhibits an overall coarsening upwards'. However, work by the present authors (Sadler 1992; Sadler & Kelly 1993, p. 382) and Graham et al. (1992, Fig. 2, p. 657) indicates that the bulk of the GPF is characterized by a net upward fining trend (see Fig. 5), with only the basal portion of the formation coarsening upwards from the underlying fine-grained sediments of the Caha Mountain Formation and its equivalents. These trends probably reflect an initial decrease in basin subsidence
Table 5. Estimates for the diffusivity (v) and basin equilibrium time constant (Teq) using Method II Estimated mean annual discharge (m3 s-i)
Diffusivity
Equilibrium time constant (106 a)
Equilibrium thickness (m)
Qb
v
r~q
/4~q
100 200 400
0.003 0.005 0.011
4.6 2.3 1.1
1840 900 440
The equilibrium thickness, Heq, is estimated using a 0.4 m m a -1.
32
s.B. KELLY • S. P. SADLER
rates followed by a gradual increase. However, the uppermost part of the GPF and its equivalents (Stage 3 of Graham et al. 1992) records a return to a more sand-prone system with relatively incised channel deposits containing abundant reworked calcrete material. This appears to indicate a further, late-stage relative decrease in basin subsidence. The c. 1.5 km scale 'tectonic' cyclicity observed within the GPF may also reflect the evolution of the drainage basin as suggested by the model of Fraser & DeCelles (1992). Following a pulse of tectonic uplift, large sediment fluxes from the drainage basin would initially have promoted the upward coarsening of sequences in the depositional basin. However, as drainage basin nets in the source area expanded and valleys increased their capacity to store sediment eroded from interfluves, the quantity and calibre of the sediment load at the outlets are likely to have diminished. This would then have promoted the upward fining of fan sequences. It is therefore possible that the formation-scale cycles (probably representing time periods of 3.5-4 Ma) could represent 'slow' variations in sediment flux, such that T>> Teq (see Paola et al. 1992, Fig. 2a). Summary
Throughout much of the Late Devonian, the weste m province of the Munster Basin, southwest Ireland, was characterized by a large-scale, sandprone, terminal fan system (e.g. the GPF). The depositional system was effectively closed and was therefore extremely sensitive to any changes in input (sediment and/or water) at all timescales. Although water could escape by evaporation, all sediment was effectively 'trapped', leaving a nearcomplete stratigraphic record. Biostratigraphic data and the radiometric dating of stage durations suggest sediment accumulation rates of c.
0.4 mm a-l.This indicates that the duration of the sedimentary cyclicities identified throughout the basin-fill are within the Milankovitch Band (105 a). These features are considered to be a function of orbitally-forced climatic changes (predominantly related to orbital eccentricity) which had a profound influence on discharge patterns and the resulting alluvial architecture. The terminal fan covered an area of c. 13 000 km 2 and had a downslope extent of over 100 km. Empirical relationships derived from modern drainage basins allow the palaeohydrological reconstruction of this depositional system and its parent drainage basin. For the GPF, data suggest a drainage area of c. 50 000 km 2 with an overall length of c. 600 km. The mean annual discharge of the drainage basin was c. 200 m 3 s-1 and this is considered to have transported c. 0.005 km 3 of sediment annually. These estimates allow the diffusion (v) and the equilibrium time constant (Teq) of the basin to be assessed (0.006 and 2 Ma, respectively) using the model of Paola et al. (1992). These estimates confirm that the Milankovitch Band cyclicity detected within the basin-fill records a relatively 'rapid' cyclical variation in sediment/water flux and that the most likely explanation is climatic perturbation. These orbitally-forced cycles are superimposed on largerscale clastic wedges that dominate the basin-fill architecture. These are considered to be the product of 'slow' changes in base level related to variations in basin subsidence and/or sediment flux. The authors acknowledge funding from BP (S.B.K.) and the Geological Survey of Ireland (S.P.S.). Many thanks are due to P. Meere for assistance in structural matters and Henrik Olsen for fruitful discussions concerning terminal fan deposition. D. Lambourne generously and unknowingly provided the inspiration for the work on alluvial cyclicity for which the authors are very grateful.
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Devonian precessional and other signatures for establishing a Givetian timescale MICHAEL
R. H O U S E
Department of Geology, The University, Southampton, S 0 1 7 1BJ
Abstract: Palaeozoic radiometric dates are so isolated and crude as to offer little numerical precision in stratigraphy. In a Givetian sequence of pelagic micrites in the Montagne Noire of southern France there is a succession of small-scale microrhythms which appear to represent precessional signatures. Correlations can be made with other sections in the Tafilalt of southern Morocco and, for the late Givetian, in the Marble Cliff Beds of north Cornwall. Now that the IUGS has approved boundaries defining the upper and lower limits of the Givetian it is possible to estimate the number of putative precessional signatures for the whole Givetian and, using the calculation of periods for these in the Devonian, to estimate the duration of the Givetian and its conodont zones independantly from radiometric dates. Graphical correlation techniques are used to establish a time framework for the Givetian and this is used to estimate the duration of zones in relation to the preliminary precessional scale, and hence to estimate their duration in real time. Biological changes and sedimentary events can similarly be constrained against a scale for the Givetian. It is suggested that such techniques, using a stage for a base, is likely to be useful long before continuous orbital forcing timescales (OFT) can be established. It would be a start in eliminating present crude estimates of the length of stages based on the false presumption that zones within them are of equal duration. Given boundaries, defined internationally by IUGS at Global Stratotype Section and Points, such OFT scales would enable events and zonal durations to be given in relation to OPT graphic composite scales and on the basis of estimated precessional signatures.
The problem of giving numerical ages to system and stage boundaries through geological time is well known. For the Tertiary and late Mesozoic considerable progress has been made, but the Palaeozoic has lagged far behind and estimates are crude in the extreme. The author has published a short paper (House 1992) suggesting that for the Givetian stage of the Devonian improvements might be possible. Attention was drawn to a sequence of pelagic micrites at the Pic de Bissou, Languedoc, southern France, which suggested, by their homogeneity, that rather constant environments continued through much of the Givetian. Measurements of the microrhythms led to the suggestion that they were probably due to the precessional orbital forcing climatic signature. Because, from fossil and other evidence, the rate of the Earth's rotation, and the rate of rotation of the Moon around the Earth can be estimated, the revised calculation of precessional and obliquity orbital periods in the Devonian (Berger et al. 1989a, b) could be used to estimate the duration of the Givetian. Since 1991 the Subcommission on Devonian Stratigraphy has defined a Global Stratotype Section and Point (GSSP) for the base of the Givetian at a rather higher level than used for the orbital forcing calculations (House 1992), so a revision is necessary. Further work has also been
done to try to take the Pic de Bissous section up to the Middle/Upper Devonian boundary. In addition, an analysis of the late Middle Devonian rhythmic sequence of north Cornwall has been completed. Further work in Morocco, in the area of the new basal Givetian GSSP, has enabled a revised correlation against real-time precessional module to be presented.
Estimates from radiometric data of the duration of the Devonian period and Givetian stage Estimates of the duration of the Devonian based on radiometric data show considerable divergence as is indicated by figures published over the last decade (Table 1).
Other estimates of relative duration of the Givetian In an attempt to 'refine' the radiometric scale, subdivision of periods has been attempted in several ways. Friend & House (1964) estimated maximum sedimentation thickness worldwide per stage as a basis. More usual has been the assumption that stages might be of equal duration, and the period has been simply divided by the number of stages
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescales and Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 37-49.
37
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1s narrow
I
bench
L__
,8
=
C~---
' [~
i
~
r
,~
.........' '
,i UPPER TRACK CROSSES
,,,-
Nord
,~
,,i z
small bench
==/
,~,Marbri~re
!,
;" i,-'~'. / t
PIC
DE
" ~ ' ~1 .r,,~arbr i~re / /
VS-W VS-E
Sud
c. _OCALITY MAP ! | ROUVILLE/ ZONE
bench
-5
METRES
SECTIONS IN QUARR Y AND CLIFF INVERTED -0 Fig. 1. The Pic de Bissous (Vissous) area, 7 km southwest of Clermont l'Herault, Languedoc. (A) Middle Devonian section exposed in the Marbri6re Nord: note that the quarry section is inverted. The beds lettered are referred to in the text. The bed numbers are marked on the rock face. (B) The upper Middle Devonian section on the southwestern face of Pic de Bissous (VS-W), the section is inverted: selected numbers are shown which are marked on the rock face (C) Sketch map showing the location of sections on Pic de Bissous.
ESTABLISHING A GIVETIAN TIMESCALE
39
Fig. 2. The main face of Marbri6re Nord, Pic de Bissous, showing the lettering used in Fig. IA. The numbers marked on the rock face also enable alignment with Fig. 1A.
40
MICHAEL R. HOUSE Table 1. Estimates of the duration of the Devonian and Givetian
Author Harland et al. 1982 Palmer 1983 Odin 1985 McKerrow et al. 1985 Snelling 1985 Harland et al. 1989 Cowie & Bassett 1989 Menning 1989 Odin & Odin 1990 Fordham 1992
Devonian duration (Ma)
Givetian duration (Ma)
Givetian as a percentage of Devonian duration
40 48 40 58 50 46 55 46 50 44
6 6 5
12.5 12.5 12.5
(Harland et aL 1982). More recently, the number of biostratigraphic zones within a stage has been used, presumably on the assumption that this bears a relation to a rate of evolution for fossil groups (Harland et al. 1989). A different approach was adopted by Boucot (1975) who asked six specialists (Boucot, Dutro, House, Klapper, Oliver & Ormiston) with some knowledge of the Devonian worldwide to give their estimate of the relative duration of stages. The duration of the Givetian Stage was calculated as 15.4% of the Devonian. Later, Ziegler (1978) gave a similar estimate (14.9%) based on published views of continental specialists. In 1992 the author took the view that the only way to approach a first-order estimate for the length of the Devonian using radiometric dates was to average the various estimates then available, which gave a figure of 47.1 Ma. For estimating the length of the Givetian stage the view was taken that the Harland et al. (1989) figures were too unreliable, and the above estimates of Boucot (1975) and Ziegler (1978) were averaged, which gave a figure of 15.15%, and thus an estimated period for the duration of the Givetian of 7.1 Ma. Using the microrhythms measurable in the Pic de Bissous Marbri~re (House 1992), with corrections based on missing parts of the Givetian succession, it was estimated that there might be 353 such cycles in the Givetian as a whole, and that if these represented climatic variations due to the precessional signature, which Berger et al. (1989a, b) had shown to be 19.9 ka, this gave an estimate using orbital forcing methods of 7.0 Ma. A remarkable agreement with the crude estimate using radiometric data, and that of Boucot (1975) and Ziegler (1978). However, since 1992 the IUGS has accepted the recommendation of the Subcommission on Devonian Stratigraphy that the base of the Givetian should be drawn at a rather higher level than used in the orbital forcing calculation (Walliser 1991;
11
18.96
(10) 3.4
(20) 7.39
5 c. 8.7
10.0 19.8
House & Walliser 1993). In addition, the author has attempted to obtain figures for the latest Givetian, which were previously missing from the discussion, and further work has been done on the Moroccan sections.
Pic de Bissous, Marbri~re Nord On the northern slope of Pic de Bissous, 7 km southwest of Clermont l'Herault, in the southeastern Montagne Noire of Languedoc, is an old marble quarry in the Middle Devonian (1:25 000, Lod~ve sheet, 2643 ouest, Lambert coordinates, x = 682.36, y = 3145.150) (Fig. 1A-C). In recent years the quarry has been in use intermittently, but not for marble production. The bedding in the quarry is near horizontal but gently flexed and completely inverted. The sequence is one of rhythmically bedded pelagic micrites with regular master bedding planes which much facilitated quarrying. For analysis the main units were lettered from oldest to youngest (Fig. 1A), but since then other workers have numbered the succession (Fig. 2) and these numbers are also shown on Fig. 1A. In an earlier paper (House 1992) attention was drawn to the fine-scale rhythmicity in this succession and this is illustrated here (Fig. 2, note inversion) on the main face of the quarry section and elsewhere on exposed and weathered joint surfaces (Fig. 3A-C). The fine-scale rhythmicity, which may be referred to as couplets, was shown earlier (House 1992) to be bimodal with average thicknesses, based on 107 measurements, of 5.6 and 10.5 cm (Fig. 4A). Attention was then drawn to how separating bedding surfaces, when exposed, showed ramiform patterns filled with argillite and these are of T h a l a s s i n o i d e s - t y p e and probably represent burrows (House 1992, fig. 2C), but how far bioturbation may have contributed to the nodular micrite griotte structure of the couplet
Fig. 3. Photographs showing the microrhythmicity in the Givetian of Marbri~re Nord, Pic de Bissous. (A) A vertical joint face in the inverted youngest beds exposed in the quarry comprising Beds O-R below the lower wide bench of Fig. 1A. The scale is 2 m long. Note the grouping within Bed Q of couplets into bundles. (B) The sequence in a weathered joint surface in the lower part of the main face at the eastern end showing the succession above the upper broad bench of Fig. 1A. (C) A continuation of the section into older beds, the photograph overlaps with that in B at the prominent ledge separating Bed M from L.
42
MICHAEL R. HOUSE 20"
10-
A /
/I
Marbri~re
B VS-W
8
/ /,
15-
9
/
m ~ 10-
m 4" /
x~ 5E z 0
/
i
\
.a
/ o
~ Bed
~
1'5. . . . . .
Thickness
2'o
cm
~
\ ~
Bed
,~ Thickness
-
-
i
15
20
cm
Fig. 4. Graphs of microrhythm thickness in Middle Devonian rocks at the Pic de Bissous. (A) Based on measurements in the northern Marbri~re (after House 1992). (B) Based on measurements on the southwestern face of Pic de Bissous. units themselves is not clear. There is no evidence of bed thinning or major gaps in the sequence. The small-scale rhythmicity is not considered further here, but the implications for timescale subdivision is considered in a later section. The interpretation of the bimodality remains uncertain, although the earlier interpretation was that it might represent interpenetration of the precession cycles with those of obliquity (tilt). The grouping of the bimodal couplets into bundles was not discussed in the earlier paper since it is not apparent visually except in certain parts of the sequence but is noticeable in the youngest exposed beds, especially between Beds S and O (Fig. 3A). In Bed Q the bundles appear as emphasised hard units near the base and at the top of Bed Q with two intervening groups, making four altogether. When the thickness of couplets within the interval Bed S-O is plotted (Fig. 5), the systematic modulation of couplet size is noteworthy, and it is a common feature of microrhythmic sequences of this sort elsewhere. Even
~=~1 ~159~-1o- ---
0
)PI
Q
/ ~
I R /I
I
~\
~o'~JJlJtll/llllJt[ll/lliiliill{l,illl (.?)
Couplets given equal d u r a t i o n
where such prominence of hard beds does not show, characteristic signatures of couplets may be of the same frequency; four such are shown in the higher part of Bed L (Fig. 3C, top right) where there is a repeated thick bed with re-entrants above and below: less clearly these may also show in the lower part of Bed L in Fig. 3C. There are yet larger scale groupings of couplets in the Marbrifre Nord section which were not considered in the earlier paper. These show as master bedding planes often used by the quarrymen as working levels and benches (Figs 1 & 2). There is a crude'impression of units < 1 m in thickness (Fig. 1). Couplets have been measured or estimated within the lettered units as follows: G = 20; H = 14; I = 29; J = 41; K = 44; L = 22; M = 33; N = 13; O = 12; P = 3; Q = 13; R = 7), but these include both types of bimodal couplet, and since these differ by a factor of two they cannot be used simply to estimate duration. It seems wiser, as a crude estimate, to calculate the number of the smallest couplets they might theoretically contain using the 5.6 cm average. If this is done for the span of Beds F - L (total number of units, 18; total thickness, 12.16 m) then the average unit has 12 couplets. Using the precession estimate above (19.9 ka) this would suggest a period of 240 ka, approximately twice the lower Eccentricity period.
YOUNGI~-G
Fig. 5. Microrhythmicicity of Beds R-O at Marbrifre Nord. A plot of couplet thickness against successive couplets on the assumption that the couplets have equal duration. Note the modulation of couplet thickness.
Pic de Bissous
(Vissous)
On the southern side of the vertical face of the Pic de Bissous a Middle and Upper Devonian
ESTABLISHING A GIVETIAN TIMESCALE
sequence is accessible at the foot of the precipitous cliffs (Fig. 1B & C); the locality has been abbreviated as V S - W by Feist & Klapper (1985). Horizontal and inverted late Givetian strata are first encountered walking from the car park at the first overlook ( 1 : 2 5 000, Lod~ve sheet, 2643 ouest, Lambert coordinates, x = 682.25, y -- 3144.65). This is locality VS-W of Feist & Klapper (1985, p. 5). A normal and near-vertical fault then cuts the cliff and there is then another accessible section, also inverted and near horizontal, which embraces the Frasnian; that section continues intermittently below into the early Famennian griotte in Marbri~re Sud below. The Frasnian part is section VS-E of Feist and Klapper (1985, p. 5) and does not concern us further. Collections for conodonts through section at VS-W by Feist & Klapper (1985) showed it to pass from the Middle varcus Zone and into the Frasnian, the Lowermost/Lower a s y m m e t r i c u s Zone boundary being between their Beds 41 and 42, below which overgrown scree conceals the section. Bed 41 yielded goniatites of the Pharciceras Stufe (House & Dineley 1985, p. 12). This section was discussed in relation to the establishment of a GSSP for the Middle/Upper Devonian and Givetian/Frasnian boundary, although the GSSP chosen by lUGS to define the boundary is at Col du Puech de la Suque, near St Nazaire de Ladarez, 24 km to the southeast (Klapper et aI. 1987). Measurement of couplet thickness was not so simple as in the etched joint faces of the Marbri~re Nord. The results are plotted in Fig. 4B. There is some similarity in pattern when compared to the Marbri6re Nord, but there is not a clear correspondence of peaks apart from one near the 5.6 cm level. Since this is well documented in relation to the conodont zones it is later used for graphic correlation with a sequence from north Cornwall. It may be, however, that VS-W shows in the latest Givetian some of the environmental fluctuations more obvious as lithological changes in the Frasnian sequence at VS-E and is less stable that appears to be the case for the earlier Givetian. The VS-W data is later used to establish a graphic correlation with a sequence in north Cornwall and is also used to establish an estimated period for individual conodont zones.
Bou Tchrafine, Morocco This celebrated Devonian locality in the Moroccan Tafilalt lies 9 km southeast of Erfoud on the east side of the road to Taouz (1:100 000 Erfoud sheet NH-30-XX-2, Lambert coordinates x = 616.6, y = 87.8). The succession of the whole Devonian has been much studied, the Lower Devonian by
43
Alberti (1981), the Middle Devonian by Bultynck & Hollard (1980), Bultynck (1985, 1987, 1989), Walliser (1988, 1991) and Bultynck & Walliser (1990), and the Middle/Upper Devonian boundary beds by Bensaid et al. (1985). The goniatite faunas have been reviewed by House & Chlupfi~ (1987) and, more thoroughly, by Becker & House (1994). However, as with most localities in this area, deposits formed on the Tafilalt Platform are in a platform setting where non-sequences and diastems are to be expected; these, together with sharp lithological changes, are strongly developed in the late Givetian and Frasnian. It is not a wholly appropriate section for study of orbital forcing signatures. However, it was a contender section for the GSSP to define the Eifelian/Givetian boundary, and whilst the stratotype actually chosen was nearby, at Mech Irdane, 25 km south-southwest of Erfoud (Walliser 1991; House & Walliser 1993), correlation of the earliest Givetian beds is well established. The assignment of conodont zones to the Givetian of the Bou Tchrafine section is illustrated in Fig. 8.
Marble Cliffs Probably the finest section in Europe for conodont biostratigraphy across the Middle/Upper Devonian boundary is at Marble Cliffs, near Trevone, north Cornwall, UK (Grid Reference SW 891 764). Given favourable tides, the foot of the cliffs can be approached either through a tunnel at the foot of a blow hole, or, from the eastern end, and best using a rope, from near Porthmissen Bridge. The succession is shown on Fig. 6. The section has been much studied since the first work on the conodont biostratigraphy by Kirchgasser (1970), who published details of the succession and showed the sequence was inverted, supporting the unpublished views of R. G. Walker who first recognized the limestone turbidites and their inversion. A sedimentological analysis was published by Tucker (1970). Further conodont work was published by Mouravieff (1977) and in a field guide by Scrutton (1978). de Cartier de Marchienne (1976) carried the succession down to the contact with the dolerite sill of Roundhole Point, an area where access is dangerous. The resulting stratigraphical column is illustrated in Fig. 6, which is based on a figure published by House & Dineley (1985, p. 302) in which the conodont zonation had been updated by Professor Kirchgasser. Since that time, Klapper & Johnson (1990, p. 934) have suggested using the term hermanni Zone, in place of hermanni-cristatus Zone as shown on Fig. 6, and they also regard the Lower and Upper 'parts of the disparilis Zone as formal subzones. Sandberg et al. (1989, p. 201)
44
MICHAEL R. HOUSE
K20 99
97
METRES
136
9B
4 K19
135
] TUFF
134 133 --
K18
9 __132
131
130 128 126 124 123 122 121 120
K17
w 119 z 0 IM 118 K6 K5
K4
K3
K16
-~ 117 nIJJ 116 .o Q z LU 115 0 114 _---J n,
~
K15
111
(,/.} 110 K2
N~o9
K14~
K12
K137
LU 108 Q_ K1
o_ 106 ::3 105 104
KIO
103
THICKNESS UNCERTAIN
ffl::7
102 E W
I0~
9
139 138 137
Fig. 6. Section of part of the Marble Cliff beds, uppermost Middle Devonian, at Marble Cliffs, Trevone, north Cornwall, showing the succession of limestone turbidites. Partly based on sections of Kirchgasser (1970) and, for the stratigraphically lower part, Mouravieff (1977) and de Cartier de Marchienne (1976). Conodont zonation by W. T. Kirchgasser. Modified from House & Dineley (1985).
have recently replaced the Lowermost asymmetricus Zone by afalsiovalis Zone, and whilst the definition of the base remains the same there is a difference in the definition of the upper boundary, but that does not concern us. The section is inverted, and part of an isoclinal fold, but it is possible to measure a near-complete sequence of the limestone turbidites and the intervening slates. The turbidites are exotic and
thought to have arisen from neighbouring crinoidrich shoals (Tucker 1970). As in the case of the Pic de Bissous, sufficient measurements can be made for plots to be made of the frequency of the classes of bed size. These results are shown in Fig. 7A & B. There is the possibility that the rhythmicity seen at Marble Cliffs might be due to primary climatic changes which enhanced or slowed down crinoidal
ESTABLISHING A GIVETIAN TIMESCALE
o/'l
40
A
B
20-
LSTS
SHALES
15
/'\,
"1o
mlO.
\ 15 Bed
45
/ \ .\ /
3'0 Thickness
415 cm
E 5~ o
/
/
o
15 Bed
3'0 Thickness
4'5 cm
Fig. 7. Graphs of unit thicknesses for the uppermost Marble Cliff Beds. (A) Of shale beds between limestone units; (B) of limestone turbidite units.
formation on the shoals. If that were so it would be expected that regular rhythms might be interpretable in the Marble Cliff succession. Two different approaches are possible. The first would assume that the sedimentation rate of the background shale was fairly constant, then intervals between limestone turbidites might be expected to reflect orbital rhythmicity; this analysis should remove the turbidites from consideration. Secondly the limestone turbidites themselves might show patterns related to climatic maxima. For the analysis of shale intervals, Fig. 6A shows that there is indeed a clumping of bed size and there is an indication of bimodality in the lower unit thicknesses with peaks at about 5 cm and 20 cm. Some analogy with the Pic de Bissous seems apparent. However, the Marble Cliff succession is substantially thicker than that in the Pic de Bissous area. This comparison is shown by means of a correlation between the two successions which is given in Fig. 9. The linearity of the conodont zonal boundary plots in the two sections is striking, suggesting indeed that both sections had rather constant sedimentation rates. However, that of the Marble Cliffs is c. 9.3 times as great as in the Pic de Bissous section. Neither of the two high frequency bed rhythms of Marble Cliff can therefore be compared with the c. 5.6 cm rhythms of the Pic de Bissous and those rhythms would be expected at c. 0.52m intervals. Since there is not a significant signature at this frequency it seems unlikely that orbital forcing played a significant
role in controlling input of allochthonous crinoidal debris from the crinoid shoals and hence an interpretation has to be sought in tectonics and shoal instability, rather than the supposition that crinoid development on the shoals was climatically controlled and the rhythmicity resulting from orbitally forced climatic change. However, using the background sedimentation rate at Marble Cliff, and in conjunction with data from Pic de Bissous, confirmatory evidence is given to provide a framework for estimating the duration of conodont zones during the time represented by the sequence.
Estimation of the duration of the Givetian stage using couplets An estimate of the duration of the Givetian, based in part of the Marbribre Nord section and in part on graphic correlation with a section at Bou Tchrafine, Morocco, has previously been attempted (House 1992). Deductions were based on the assumption that the basic microrhythmicity (couplets at 5.6 cm) might be due to the second precession cycle which Berger et al. (1989a, b) has estimated to be 19.9 ka in the Middle Devonian, and this formed the basis for an estimation of Givetian duration. Values for the two eccentricity periods, not dependant upon changes in the Earth-Moon system, are not thought to have changed with time. The actual estimates of Berger et al. (1989a) are given in Table 2.
46
MICHAEL R. HOUSE Table 2. Calculationof orbitalfrequencies (from Berger et al. 1989a, b) Cycle Precession
Present-day
Mid-Devonian
19.0 ka 23.0 ka 41.0 ka 54.0 ka 123.0 ka 413.9 ka
16.8 ka 19.9 ka 32.1 ka 39.5 ka 123.0 ka 413.9 ka
(1) (2) (1) (2) (1) (2)
Obliquity Eccentricity
More recently the base of the Givetian has been approved by IUGS at a level in a section at Mech Irdane, 12 km southwest of Rissani, Morocco (Walliser 1991; House & Walliser 1993). The new stratotype level (GSSP) is higher than the rouvillei Zone which was used in the earlier calculation and is now based on the entry of the conodont Polygnathus hemiansatus The goniatite Maenioceras undulatum enters immediately
m [ Lower asymrnetricus Zone 2 5 ~ ct!spar/lis & L.most asyrnrn. Zone J31 U . ~ ~ e herrnanni Zone Middle
preceding Po. hemiansatus at Mech Irdane (Becker & House 1994) so that the rouvillei Zone must now be assigned to the Eifelian. A correction is needed for this modification, but this has to be done indirectly because Po. hemiansatus has not yet been located in the Marbrih'e Nord (Walliser 1990), and has been achieved using graphic correlation techniques. At Bou Tchrafine (Fig. 8) the base of the
FRASNIAN
G/VET,,,
varcus Subzone
/
Zone
28
LIJ
Z20
LL. <
-
n,, "1-
_
oI-- 1
Lower varcus 5ubzone
s!
_.~ 0 nn
hemiansatus Zone
10
15
ensensis Zone I
~e::.il, N' E/F, EL/A G/VETIAN
Mo-n8
I
!m nc /11 MO ,
3.:>" I
6
lb PIC
DE
1~
m
BISSOUS
Fig. 8. Correlation between the Givetian sequence between the Pic de Bissous and Bou Tchrafine, Tafilalt, Morocco, showing a timescale on the presumption that the principal microrhythm component is due to precessional orbital forcing.
ESTABLISHING A GIVETIAN TIMESCALE goniatite rouvillei Zone lies below Bed 9a, whilst the entry of Po. hemiansatus, defining the new base of the Givetian, is at Bed 15 bis (Bultynck & Walliser, in Walliser 1991, p. 54), i.e. a separation of 6.95 m. Using the crude correlation between Bou Tchrafine and Marbri~re Nord, and the tie levels of the rouvillei and terebratum Zones in both cases, this gives an estimate for the hemiansatus entry to top terebratum Zone at Marbri~re Nord of 13.8 m. Using the average of 5.6 cm for the lower couplet of Fig. 4A, and assuming that this represents the 19.9 ka period of Berger et al. (1989a), this thickness should represent 246 putative couplets and c. 4.9 Ma for the duration of this part of the succession. Younger levels to the base of the Frasnian at VS-W are 4.51 m thick and, using the same couplet unit, this would give 80.5 couplets and an estimated 1.6 Ma for the upper part of the Givetian. That gives a total estimate of 348 of the 5.6 cm couplets for the whole Givetian, and an estimate of c. 6.5 Ma for the duration of the Givetian. However, it should be noted that Feist & Klapper (1985, fig. 6) indicated uncertainty in the placing of the Upper varcus/ hermanni Zone boundary and further work is needed on the conodont subdivision of the varcus Zone. This new figure cannot really be compared either with the radiometric estimates given in an earlier section, or the averages from specialists quoted by
47
Boucot (1975) and Ziegler (1978) since the criteria for definition of both the base of the Givetian and the top of the Givetian have changed since those times.
Estimations of conodont zonal durations within the Givetian Since complete documentation is now available for conodont zonation of Bou Tchrafine (Walliser 1991) the graphic correlation diagram for the whole Givetian at Pic de Bissous and Bou Tchrafine (House 1992, fig. 4) can be modified to enable estimations of the duration of conodont zones within the Givetian (Fig. 8), but various factors limit the precision of the conodont subdivision of the varcus Zone. Calculation of the period of conodont zones for the early Givetian can only be done approximately because the Pic de Bissous conodont evidence is poor (Walliser 1990). Using the crude graphic correlation with the Marbri~re Nord, and accepting the ties provided by the rouvillei and terebratum Zone faunas, periods for the hemiansatus and varcus Zones are given below. The upper Givetian data for VS-W enables approximate designation of periods for the higher zones based on the assumptions regarding couplets and precessional period outlined previously. The results are given in Table 3.
9-
Table 3. Estimated duration of Givetian conodont zone
asymmetricus Zone
Lower
/"
8-
Lowermost E
7-
U3 6 > v
Upper
Lower
asyrnrnetricus Zone
Givetian conodont Zone f
disparilis S u b z o n e ~ /
disparilis /
Subzone / /
0 5_
/ herm/ r
el m "El
-ann/
4-
2
/ 654 ka
345 ka
246 ka
Zone
13-
3
0
1'0
20
30
4tO
M a r b l e Cliff shales m Zonal duration at Precession of 19.9 ka
Fig. 9. Correlation between the late Givetian of Pic de Bissous (VS-W) and the Marble Cliff succession.
Lowermost asymmetricus Zone Upper disparalis Subzone Lower disparilis Subzone hermanni Zone Upper varcus Subzone Middle varcus Subzone Lower varcus Subzone hemiansatus Zone
Estimated duration (Ma) 0.25 0.35 0.65 0.36 0.21 0.92 4.05 0.43
It is necessary for more sections in pelagic facies to be studied in order refine these figures, which should be regarded as preliminary estimates only at this stage. As has been explained, the use of rather condensed shelf developments, such as those of the Tafilalt, has not been driven by their excellence for the purpose, but only because they provide, at present, not only the standard GSSP for definition, but sections which have been well reported.
48
MICHAEL R. HOUSE
Conclusions A review is given of some estimates of the duration of the Devonian Givetian stage based on radiometric criteria and intuitive estimates of duration based on views of specialists familiar with the Devonian System worldwide. The figures vary greatly and the criteria for defining of the limits of the stage is rarely given. Recent definitions of the base of the Givetian and the top of the Givetian by the IUGS, and the establishment of a defining GSSP for both these boundaries, clarify the discussions. An estimate of the duration of the Givetian is given based on the presumption that a dominant microcyclicity in sections at the Pic de Bissous, southern France, is due to environmental changes orbitally forced by the second precessional cycle, and using estimates recently made by Berger et al. (1989a, b) of the period of those cycles in the Middle Devonian. The Pic de Bissous sections can be well correlated to the GSSPs. The estimated Givetian duration using these assumptions is 6.5 Ma. Calculation of the period of the conodont zones which are used to subdivide the Givetian is attempted using graphic correlation with other sections, especially Bou Tchrafine, Morocco, for the early Givetian, and Marble Cliffs, north Cornwall, for the late Givetian. These estimates are based on the same presumptions. Attention is
drawn to the need for the incorporation of data from other pelagic sequences and to the weaknesses in the Tafilalt environmental setting for cyclostratigraphic studies. Extension of the quarrying at Marbri~re Nord to higher levels may bridge the current uncertainty of correlation between this section and that of Pic de Bissous VS-W. The advantages of a borehole through the whole succession to enable geochemical signatures to be studied and to replace the subjective approach to microrhythm measurement adopted here represent investigations for the future. It is suggested that subdivision of stages using orbitally forced signatures may be the best approach to giving a numerical timescale to biostratigraphic scales and that these would be useful for the establishment of a tool for study of processes requiring a time base long before radiometric scales, in the Devonian at least, are likely to be helpful. When combined with graphic correlation techniques such OFF scales could give considerable time resolution. Whatever corrections may be needed to the estimates presented for zonal and stage duration of the Givetian one thing is quite clear: the zonal estimates given here are considerably nearer the truth than the na'ive assumption currently made by many that biostratigraphical zones are of equal duration.
References ALBERTI, G. K. B. 1981. Daten zur stratigraphischen Verbrietung der Nowakiidae (Dacryoconarida) im Devon yon N-Afrika (Marokko, Algerien). Senckenbergiana Lethaia, 62, 205-216. BECKER, R. T. & HOUSE, M. R. 1994. International Devonian goniatite zonation, Emsian to Givetian, with new records from Morocco. Courier Forschungsinstitut Senckenberg, 169, 79-135. BENSAID, M., BULTYNCK,P., SARTENAER,P., WALLISER, O. H. & ZIEGLER,W. 1985. The Givetian-Frasnian boundary in pre-Sahara Morocco. Courier Forschunginstitut Senckenberg, 75, 287-300. BERGER, A., LOUTRE, M. F. & DEHANT, V. 1989a. Influence of the changing lunar orbit on the astronomical frequencies of Pre-Quaternary insolation patterns. Paleoceanography, 4, 555-564. ,--, & 1989b. Pre-Quaternary Milankovitch frequencies. Nature, 342, 133. BOUCOT, A. 1975. Evolution and Extinction Rate Controls. Developments in Palaeontology and Stratigraphy, 1. Amsterdam, Elsevier. BULTYNCK, P. 1985. Lower Devonian (Emsian)-Middle Devonian (Eifelian and lowermost Givetian conodont successions from the Ma'der and the Tafilalt, southern Morocco. Courier Forschungsinstitut Senckenberg, 75, 261-286.
-
-
1987. Pelagic and neritic conodont successions from the Givetian of pre-Sahara Morocco and the Ardennes. Bulletin de l'Institut Royal des sciences naturelles de Belgique, Sciences de la Term, 57, 149-181. 1989. Conodonts from a potential Eifelian/Givetian Global Boundary Stratotype at Jbel Ou Driss, southern Ma'der, Morocco. Bulletin de l'Institut Royal des sciences naturelles de Belgique, Sciences de la Terre, 58, 95-103. & HOLLARD, H. 1980. Distribution compar6 de Conodontes et Goniatites d6voniennes des plaines du Dra, du Ma'der et du Tafilalt (Maroc). Aardkundig Medelingen, 1, 73 pp. & WALLISER, O. H. 1990. Proposal for the EifeliardGivetian Stratotype in Southern Morocco. Document submitted to the Subcommission on Devonian Stratigraphy, Frankfurt/Main, September 1990. COWlE,J. W. & BASSETT,M. G. 1989. International Union of Geological Sciences, 1989 Global Stratigraphic Chart. Episodes, 12 (2), Supplement. DE CARTIER, DE MARCHIENNE,R. 1976. Etude sddimentologique et micropaldontologiqe (conodonts) de turbidites de la coupe de marble cliff (Cornouailles, Grand-Bretagne. M6moire pr6sent6 en vue de -
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ESTABLISHING A GIVETIAN TIMESCALE l'obtention du grade Lincenci6 en Sciences Grologique et Minrralogiques, Universite Catholoque de Louvain. FEIST, R. (ed.) 1983. The Devonian of the Eastern Montagne Noire. Guidebook Field Trip, International Subcommission on Devonian Stratigraphy, Montagne Noire, Montpellier, September 1983. 1985. Devonian stratigraphy of the South-Eastern Montagne Noire (France). Courier Forschungsinstitut Senckenberg, 75, 331-352. • KLAPPER, G. 1985. Stratigraphy and conodonts in pelagic sequences across the Middle-Upper Devonian boundary, Montagne Noire, France, Palaeontographica, Abteilung A, 188, 1-18. FISCHER, A. G. 1986. Climatic rhythms recorded in strata. Annual Reviews, Earth and Planetary Sciences, 14, 351-376. FORDHAM, B. G. 1992. Chronometric calibration of midOrdovician to Tournaisian conodont zones: a compilation from recent graphic-correlation and isotope studies. Geological Magazine, 129, 709-721. FRIEND, P. F. & HOUSE, M. R. 1964. The Devonian period. In: HARLAND, W. B., SMITH, A. G. & WILCOX, B. (eds) The Palaeozoic Time-Scale. Quarterly Journal of the Geological Society of London, 120S, 233-236. HARLAND, W. B., ARMSTRONG, R. L., Cox, A. V., CRAIG, L. E., SMITH, A. G., SMITH, D. G. 1989. A Geological Time Scale 1989. Cambridge University Press, Cambridge. , Cox, A. V., LEWELLYN, P. G., PICKTON, C. A. G., SMITH, A. G. 8~ WALTERS, R. 1982. Geologic Time Scale. Cambridge University Press, Cambridge, 131 pp. HOUSE, M. R. 1992. Devonian microrhythms and a Givetian time scale. Proceedings of the Ussher Society, 7, 392-395. -& CHLUP,~(2, I. 1987. Goniatite faunas relevant to the definition of the Eifelian/Givetian boundary. Document submitted to the Subcommission on Devonian Stratigraphy, Calgary, 1987. & DINELEY,D. L. 1985. Devonian series boundaries in Britain. Courier Forschungsinstitut Senckenberg, 75, 301-310. -& WALLISER, O. H. (eds) 1993. Proposal for a Global Stratotype Section and Point (GSSP) for the Eifelian-Givetian boundary. Document submitted to the International Commission on Stratigraphy, International Union of Geological Sciences, Southampton.
-
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49
KIRCHGASSER, W. T. 1970. Conodonts from near the Middle-Upper Devonian boundary in North Cornwall. Palaeontology, 13, 335-354. KLAPPER, G. 8z JOHNSON,J. G. 1990. Revisions of Middle Devonian conodont zones. Journal of Paleontology, 64, 934-936. MCKERROW, W. S., LAMBERT,R. ST J. & COCKS, L. R. M. 1985. The Ordovician, Silurian and Devonian periods. In: SNELLING,N. J. )ed.) The Chronology of the Geological Record. Geological Society of London, Memoir, 10, 73-80. MENNING, M. 1989. A synopsis of numerical time scales 1917-1986. Episodes, 12, 3-5. MOURAVIEFF,N. A. 1977. Additional conodonts from near the Middle/Upper Devonian boundary in North Cornwall: a progress report. Proceedings of the Ussher Society, 4, 63-66. ODIN, G. S. 1985. Remarks on the numerical scale of Ordovician to Devonian times. In: SNELLING,N. J. (ed.) The Chronology of the Geological Record. Geological Society of London, Memoir, 10, 93-98. & ODIN, G. S. 1990. Echelle numrrique des tempts g6ologique. G(ochronique, 35, 12-21. SANDBERG, C. A., ZIEGLER, W. & BULTYNCK, P. 1989. New standard conodont zones and early Ancyrodella phylogeny across the Middle-Upper Devonian boundary. Courier Forschungsinstitut Senckernberg, 110, 195-230. SERUTTON, C. T. 1978. A field Guide to selected areas -
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of the Devonian of South-West England. Palaeontological Association. SNELLING, N. J. 1985. An interim time-scale. In: SNELLING, N. J. (ed.) The Chronology of the Geological Record. Geological Society of London, Memoir, 10, 261-265. WALLISER,O. H. 1988. Proposal for an Eifelian/Givetian boundary stratotype. Document submitted to the Subcommission on Devonian Stratigraphy, Frankfurt/Main, 1990. -1990. Marble Quarry at Pic de Bissous. Remarks on the stratigraphy. Document submitted to the Subcommission on Devonian Stratigraphy, Frankfurt/Main, 1990. (ed.) 1991. Morocco Field Meeting of the Subcommission on Devonian Stratigraphy, International Union of Geological Sciences, Nov. 28-Dec. 5, 1991, Guide Book. Grttingen. ZIEGLER, W. 1978. Devonian. In: COHEN, G. V., GLAESSNER, M. F. 8s HEDBERG, H. D. (eds) Contributions to the Geological Times Scale. American Association of Petroleum Geologists, Studies in Geology, 6, 337-339.
Orbital-climatic forcing of Namurian cyclic sedimentation from spectral analysis of the Limestone Coal Formation, Central Scotland G. E W E E D O N 1 & W. A. R E A D 2
1 Department of Geological Sciences, University of Luton, Park Square, Luton, Bedfordshire, LU1 3JU, UK 2 Department of Geology, University of Leicester LE1 7RH, UK Abstract: The Limestone Coal Formation of the Glasgow-Stirling area of the Midland Valley of Scotland was deposited in a composite basin divisible into the fluvially-influenced or proximal Kincardine Basin in the east and the distal, delta-dominated Kilsyth Trough. Three main factors influenced deposition and determined the geometry of the various lithologies: tectonism, glacioeustatic changes and localized irregular crevassing/avulsion processes. If the sea-level changes were induced by orbital-climatic control of global ice volumes, then the resulting sedimentary cyclicity might be expected to be regular (near-constant wavelength). However, the irregular processes would produce superimposed irregular sedimentary cyclicity. In this study spectral analysis of time series was used to test for regularity and to distinguish the various forms of sedimentological control acting within the basin. Six localities spanning proximal to distal settings, and exactly the same stratigraphic intervals in the upper part of the formation, were used. Digitization of rock types using different numeric codes for coal, claystone, siltstone, fine-, medium- and coarse-grained sandstone was used for time series generation. Blackman-Tukey Fourier power spectra revealed that in the lower half of the sections four localities contained evidence for regular cyclicity significant at the 90% level. In the upper half of the data only one section revealed regular cyclicity. However, following decompaction the same regular cyclicity was detected in the lower part of the data, but with a considerable concentration of variance at particular frequencies, the significance had risen to 95%. In the upper data decompaction revealed two regular cycles. The decompaction procedure appears to have substantially removed compactional distortions related to rock type. Very crude estimates of the periodicity of the regular cycles indicate that a high-frequency, possibly obliquity or short eccentricity cycle is involved. All the regular cycles detected in the decompacted series occur in dominantly distal fluvio-deltaic sections. This suggests that the main signal relates to successive progradation of deltas in distal settings under regular glacio-eustatic control. Recategorizing the data into simple non-sandstone/sandstone and coal/non-coal categories revealed no major changes in the spectral results. This supports the model of the cyclicity as arising from a simple grain-size oscillation related to deltaic progradation. The lack of regular cyclicity in the more proximal areas and in the upper half of the data probably reflects fluvial sedimentation producing more irregular grain-size variations related to channel avulsion, crevassing and meander migration plus, perhaps, increased hiatuses and the influx of coarse siliciclastic material forming thick wedges of multi-storey sandstones. The study illustrates the importance of applying decompaction and of using multiple sections for the study of regular cyclicity in cyclothem series.
It has been clear for some time that the variation in depositional environments associated with Carboniferous cyclothems can be explained by various combinations of factors including autocyclic (e.g. channel switching), tectonic and glacioeustatic controls (reviewed by Reigel 1991). Unfortunately, without a refined timescale comparable to the Plio-Pleistocene, high-resolution seismic profiles tied to borehole logs and detailed sedimentological studies, it is extremely difficult to disentangle these various factors for any particular
region. This paper attempts to show that via spectral analysis, it is possible to go some way towards this goal. In particular, the authors' believe that the effect on the rock record of highfrequency glacio-eustatic cycles under orbitalclimatic control can be demonstrated. A key result from this study is that the results vary across the basin under investigation in a way which is readily explicable in terms of the proximal to-distal location relative to clastic supply and fully marine settings.
From HOUSE,M. R. & GALE, A. S. (eds), 1995, OrbitalForcing Timescalesand Cyclostratigraphyy, Geological Society Special Publication No. 85, pp. 51-66.
51
52
G. P. WEEDON t~ W. A. READ
Solsgirth 0 k
10 20km ~ KINCARDIN~/ I Isopachs in metres Doll~l~ldl BAS,N IZ, Torwood~~ ~, North 7oo~ .,/~. "~Y""-/[_~oo O'rchardhead t
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~50,~ Fig. 1. Isopachs of the study interval and location of the six selected boreholes. Gartcosh No. 1 Borehole, 1966, (NS 7074 6837), IHF [106.44 m]; Glenboig No. 1 Borehole, 1967 (NS 7217 6796), IHF [118.32 m]; Torwood Borehole,1960 (NS 8376 8487), WAR [1533.30 m]; Doll Mill Borehole, 1955 (NS 8748 8805), WAR [179.04 M]; Orchardhead Borehole, 1956 (NS 9237 8412), WAR [162.24 m]; Solsgirth Borehole, 1941-61 (NS 9971 9483), EHF [203.52 m]. EHF, logged by E. H. Francis; IHF, logged by I. H. Forsyth; WAR, logged by W. A. Read. The thicknesses between Index Limestone to Knightswood Gas Coal are indicated in square brackets. All these boreholes were drilled for the National Coal Board and their logs are stored with the non-confidential borehole records of the Edinburgh Office of the British Geological Survey.
The dataset and sedimentology The Limestone Coal Formation from the Pendleian E l (Namurian) of the Midland Valley of Scotland (Read 1988), though largely obscured by Pleistocene glacial deposits, has been extensively cored. Read &Forsyth (1989, 1991) have investigated this unit from a stratigraphical/ sedimentological standpoint over the last 30 years. Their studies, and commercial exploitation of the coalfields, have demonstrated that the boreholes can be correlated to the scale of a few metres using coals, claystones rich in Lingula and marine bands. The cores of five of the six boreholes referred to in
this study were examined by these authors. The six borehole sections have the merit of being complete, well correlated with other localities and described to the nearest centimetre (Figs 1 & 2). The Formation is divided lithologically predominantly into coals, claystones, siltstones and sandstones. Typically, the coals are sharply overlain by claystones, which commonly contain a brackishwater fauna and grade up into siltstones. These pass into sheet sandstones up to 5 m thick, which are capped by rooty siltstones or claystones (palaeosols) overlain by coals. Some of the sandstones show a two-storey profile in which a finergrained, upward-coarsening, probably deltaic, lower portion is truncated by a coarser-grained,
Fig. 2. Graphic logs of the boreholes indicating lithological subdivisions and correlations. Note that an interval of basalt was encountered towards the base of the Orchardhead log.
53
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G.P. WEEDON & W. A. READ
upward-fining, probably fluvial, upper portion. Sporadic erosive-based, upward-fining, elongate, fluvial channel sandstones cut down right through sheet sandstones. Such channels may be up to 12 m deep and 1.5 km wide. Other thicker (up to 42 m) multi-storey fluvial channel sandstones occur in proximal (northeastern) environments (Read 1994a). Palaeogeographically the Limestone Coal Formation was deposited in a series of basinal structures within the Midland Valley of Scotland (Read 1988). Two of these, namely the rapidly subsiding Kincardine Basin and graben-like Kilsyth Trough, both in the north-central part of the Midland Valley, constitute the study area (Fig. 1). The surrounding areas were subject to reduced rates of subsidence as revealed by isopach studies, and some were associated with occasionally active volcanism and episodic tectonic uplift. Across the study area as a whole correlation and isopachytes indicate increased rates of subsidence and sedimentation towards the northeast, particularly in the Kincardine Basin. In general, sand entered the basin from the northeast and was then transported westwards down the trough (Read 1988). Sandstones associated with deltaic processes become more common towards the southwest within the Kilsyth Trough, as do indications of brackish- and marine-water influences. Thus a picture emerges of a basin filling with predominantly deltaic sediments in the southwest and more distal settings, while fluvial sediments become more important in the more proximal northeastern areas. It has recently been suggested that a series of transgressions and regressions, plus linked changes in fluvial base levels, frequently and profoundly affected the facies distribution (Read 1994a, b, 1995). Aside from the major marine intervals, such as the Index Limestone, that occur outside the stratigraphic interval studied here, the area was apparently transgressed by brackish water at maximum flooding during the highest rate of sealevel rise. Thus, Lingula-bearing claystones represent 'condensed sections' in terms of sequence stratigraphy. Progradation of delta fronts during highstand and initial sea-level fall led to the deposition of laterally extensive coarsening-up sheet sands. The isolated elongate sandstones represent the fills of channels incised during lowstands in this model. Coal-forming peats and most palaeosols accumulated during rises in relative sea level. In the more proximal areas, fluvial sediments, especially the multi-storey sands, constitute much of the succession. The study interval has recently been described in terms of sequence stratigraphy by Read (1994a, 1995).
Rationale for spectral analysis It is now well established that, in the late Pleistocene, changes in insolation distribution driven by the Milankovitch orbital cycles were responsible for controlling global climatic variations on timescales of 20-400 ka (Hays et al. 1976; Imbrie et al. 1984, 1992). In the Early Pleistocene and Late Pliocene the system was not dominated by the 100 ka cycles which are so well known from the Late Pleistocene (Ruddiman et al. 1989; Shackleton et al. 1990). Although clearly related to non-linear climatic amplification of insolation cycles, the origin of the Late Pleistocene 100 ka climatic cycles remain poorly understood. Consequently, it would be unwise, without further proof to assume, from the evidence for mid-late Carboniferous continental ice sheets (Veevers & Powell 1987; Frakes et al. 1992), that 100ka cyclicity necessarily dominated the glacio-eustatic variations of the time. Nevertheless, there is good evidence for the whole of the Phanerozoic that orbital-climatic cyclicity is recorded stratigraphically in many facies settings (Fischer et al. 1990). The analysis of time series, derived from prePliocene sequences, using power spectra now forms the basis for the detection of orbital-climatic records (Weedon 1993). The reason for this is that it provides a method for detecting objectively the regular sedimentary cyclicity which can be expected as a product from a system controlled by the quasi-periodic orbital-climatic cycles. Simply counting the numbers of 'sedimentary cycles' between particular stratigraphic horizons and estimating cycle periods can be erroneous (e.g. Algeo & Wilkinson 1988). Such an approach cannot distinguish regular from irregular oscillations nor distinguish the superimposition of several scales of regular cycles. The problem with spectral analysis of time series from the pre-Pliocene is that the timescale constraints are usually too poor for confident estimation of cycle periods. Hence, power spectra are currently used to test for the presence of regular sedimentary cycles as a function of stratigraphic thickness. Estimated average sedimentation rates are then used to constrain the periodicity. It is currently assumed that if the, albeit crude, dating indicates a periodicity in the range 20-400 ka then an orbital-climatic cycle is indirectly responsible. If it can be demonstrated that, in a certain sedimentary system, other processes are capable of generating regular cycles in the tens to hundreds of thousands of years band, repeated for millions of years, this approach will need to be modified. A major constraint of the use of a thickness rather than a timescale for analysis is that undetected missing cycles increase the noise level in
NAMURIAN SPECTRAL ANALYSIS, SCOTLAND the spectrum. Given the ever-present possibility of missing cycles the prudent approach is to use estimated cycle duration as the maximum possible figure. The effects on pre-Pliocene spectra of variations of sedimentation rate, compaction and bioturbation, as well as hiatuses, have been addressed elsewhere (Weedon 1991; Herbert 1994). Critically, power spectra can be used to detect regular cyclicity, but the various geological processes mentioned can all have the effect of preventing such detection. In most cases to date, spectral analysis of ancient sequences has been applied to pelagic, hemipelagic, lacustrine and evaporitic sequences (Fischer et al. 1990). However, in a few cases spectra of siliciclastic sequences have also revealed regular cyclicity (van Echelpoel & Weedon 1990; Van Buchem et al. 1994; Carrs & Neidell 1965). Recently Maynard & Leeder (1992) demonstrated that the occurrence of marine bands in Late Carboniferous series may be regular. In this case we attempt to examine the cyclothems themselves, as did Carrs &Neidell (1965) and Schwarzacher (1967). For this study it was decided to use the simplest approach possible for spectral analysis with the hope of learning more about the core data by experimentation. The first simplification was to use information from exactly the same stratigraphic interval in six carefully selected borehole sections spaced through the study area. These are, from west to east, the Gartcosh, Glenboig~ Torwood, Doll Mill, Orchardhead and Solsgirth boreholes. Their sites, full titles, etc., are shown in Fig. 1 and their corresponding lithological sections in Fig. 2. The analysis is restricted to the interval from the base of the Knightswood Gas Coal (or the eqivalent stratigraphical horizon) up to the base of the marine Index Limestone, within the upper part of the Limestone Coal Formation, where we can be reasonably certain of accurate correlations (Forsyth & Read 1962). The selected sections represent a spectrum of environments ranging from strongly fluvially-influenced proximal settings in the east, through settings with mixed fluvial and deltaic influences, to distal, predominantly deltaic, marineinfluenced settings in the west (Read 1994a, 1995). Ideally, the sedimentary environment of each bed needs to be deduced and numerical coding for time series analysis based on this. However, it became clear that too many ambiguous or unjustifiable categories could be generated by too detailed a set of classes. For instance, the claystone could be divided into that with rootlets, that containing Lingula and/or non-marine bivalves, and that with no special characteristics. Likewise, the sandstone beds might have been classified as thin laterallyextensive deltaic or fluviodeltaic sheet sandstones,
55
isolated channels and thick multi-storey fluvial sediments, and so on. The authors believe that such a genetic classification is too prone to errors of interpretation, particularly in situations with mixed criteria, although clear end-members can be found. Instead a simpler description was based on mean grain size. The six categories used are coal, claystone, siltstone, fine-, medium- and coarse-grained sandstone (Fig. 3). This categorization would mean that for an idealized cyclothem, solely reflecting glacio-eustatic controls, the sequence coal, claystone, siltstone, fine-, medium-, coarse-grained sandstone, siltstone and/or claystone (palaeosol), coal, would, in this scheme, produce a single simple oscillation when digitized (Fig. 3). A regular succession of such sequences would then be readily recognized during spectral analysis. In reality the successions are very much more variable, particularly in proximal environments, because of the effect of processes such as channel switching, crevassing meander migration, etc. (Read 1994a). Nevertheless, our hope was that coding in this manner might allow any regular cyclicity related to glacio-eustactic controls of delta progradation to emerge in the power spectra from the record of the other less regular processes.
Methods of time series analysis The time series generated by this simple coding procedure are illustrated in Fig. 4. Code values were obtained at 4 cm intervals by allocating numeric values to each rock type within the range 1-6. However, at Orchardhead (Figs 1 & 2) a basalt horizon is present. This was treated numerically as a series of missing values in the spectral computer program by coding it as -9999. The power spectra were based upon the standard Blackman-Tukey method (Priestley 1981). They were generated using Fourier rather than Walsh methods (Weedon 1991) since, although the data are stepped, there are more than two states and recent work has demonstrated that where more than two states are used difficulties can arise (van Echelpoel 1994). The computer program used restricts analysis to a maximum of 999 data values. Data were therefore interpolated so that fewer values described each record. For this a Gaussian interpolation program was used to generate values spaced at 20 cm based on data windows 55 cm long, with maximum weighting given to data nearest to the interpolation positions. This procedure slightly smooths the data and results in spectra with a slightly increased background continuum slope, but otherwise did not change the spectral shapes (cf. Weedon 1993). In each case the time series was detrended using a linear best-fit line. The spectra were estimated using the number
56
G. E WEEDON •
Distal
W. A. READ
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of lags equal to one-third of the number of times series values up to a maximum of 199 lags. The spectral estimates were produced from the truncated lagged autocovariance function and a Tukey-Hanning window (Priestley 1981; Weedon 1991). The spectra were normalized by dividing each power estimate by the sum of all estimates. This permits direct comparison of spectra from several time series where the average amplitude of each series differs. Approximate 90 and 95% onesided confidence intervals were applied to the spectra based on the Z2 distribution applied to the background spectral values (Priestley 1981). The background spectrum, from which the confidence intervals could be found, was located using a robust least-squares straight-line fit applied to the low frequency half of the spectrum as plotted on a loglinear basis. The advantage of the robust leastsquares method is that outliers due to unusually large spectral peaks and troughs are identified in an iterative procedure so that the fit is unbiased (Press et al. 1986). In this study two criteria were used to decide whether regular cyclicity could be detected with reasonable certainty. Firstly, a concentration of variance at a particular frequency is accepted as distinguishable from the background noise level if it emerges above the 90% (significant) or 95% (highly significant) approximate one-sided con-
fidence levels indicated on the spectrum. The frequency error, or bandwidth, of a spectrum is constant at all frequencies. This means that lowfrequency spectral peaks relate to a much larger range of possible wavelengths than peaks occurring at higher frequencies. For instance, a spectral peak denoting a wavelength of 1 m (1 cycle per metre) might have a bandwidth of 0.05 cycles per metre. Consequently the peak might relate to cycles between 0.95 and 1.05 cycles per metre or wavelengths of 1.05-0.95 m. The ratio of longest possible to shortest possible wavelength is 1.11. However, for a peak indicating a cycle of 10 m, with the same bandwidth, the corresponding range frequencies are 0.05-0.15 cycles per metres or wavelengths of 20.00-6.66 m (a ratio of 3.0). In this latter case it could not be claimed that the peak definitely denoted a regular cyclicity. Hence, the second criteria adopted here is that a significant spectral peak is taken to denote regular cyclicity if the maximum possible wavelength is less than twice the minimum. This criterion means that at least six cycles must be present in the complete time series. This criterion can be applied by using a frequency cut-off lying one and ahalf bandwidths from the left-hand side of the spectrum; it is denoted using a vertical dashed line on the spectra. As discussed earlier, various geological factors can affect power spectra where the time series
57
NAMURIAN SPECTRAL ANALYSIS, SCOTLAND
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I 25456
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parameter is expressed as a function of stratigraphic thickness. Particularly relevant in this case are the effects of hiatuses, variable sedimentation rate and variable compaction. The first two factors cannot be tackled without considerably increased information about the sedimentology, but a correction for compaction has been attempted. The degree of overburden involved is difficult to estimate, but a value of c. 3 km appears most appropriate from consideration of likely end- and post-Carboniferous tectonism, plus the thickness of younger units preserved in adjacent areas (Read 1988). Values of 2 and 4 km were also employed, and results for the 4 km trials are illustrated later. To keep the analysis as simple as possible the same compaction correction factors were applied to all occurrences of the same rock type. As the sections are only 200 m thick at maximum the slightly greater overburden at the base of the sections was ignored (maximum 5% excess). Following the methods and formulae used by Baldwin & B u t l e r (1985) we used values for correction factors listed in Table 1. As a reference, the solidity at 10 cm burial was compared with that at 3 km for normally compacting claystone and sandstones (regardless of grain size). For siltstone it was assumed, admittedly simplistically, that solidity at every level would lie half way between
that of claystone and sandstone at corresponding depths. Siltstone forms the smallest portions of the cores at all localities. The compaction factor for coal (16.0) is based on the work of Elliott (1985). Finally, the basalt horizon at Orchardhead is assumed not to have been compacted at all, so following decompaction this forms a relatively thin
Table 1. Solidity and compaction factor by rock type for 3 and 4 km overburden (Baldwin & Butler 1985; Elliott 1985 Rock type
Solidity
Compaction factor
3 km Claystone Siltstone Sandstone Coal Basalt
0.8961 0.8392 0.7822 -
5.071 2.444 1.533 16.00 0.000
4 km Claystone Siltstone Sandstone Coal Basalt
0.9377 0.8858 0.8338 -
5.307 2.580 1.635 16.00 0.000
58
G.P. WEEDON & W. A. READ
interval, treated, as before, as missing data values. Since the sections increased in thickness after decompaction, the interpolation interval for the new time series was changed to 80 cm (interpolation window = 202 cm).We emphasize that the decompaction procedure used is considered as a first-order correction and not a definitive solution. A separate petrographic and sedimentological study of the sandstone grain size and mineralogical composition and environments of deposition would be needed for a more accurate approach. Another geological factor which may affect the spectral results would be the effect of thick wedges of proximal fluvial multi-storey sandstones replacing more distal facies. From first principles the effect is akin to replacing part of a regular signal with an interval of high-frequency irregular oscillations or noise. The spectrum records the squared average amplitude (or variance) contribution of each wavelength component. Removal of some, but not all, regular cycles from a record would reduce the average amplitude at that frequency and slightly increase the average ampli-
tudes at the noise frequencies but not affect it otherwise. Thus, the effect would be to reduce the size of any spectral peaks, making them harder to distinguish from the background continuum.
Spectral results - original data The spectra for all locations based on the complete digitized logs are presented in Fig. 5. The two most northeasterly sections (Orchardhead and Solsgirth) have concentrations of variance almost entirely at very low frequencies (0.05 cycles per metre). On the other hand, the two most southwesterly locations of Glenboig and Gartcosh yielded the flattest spectra, or the most uniform distribution of variance with frequency. This pattern of overall spectral characteristics is clearly related to the relative proximal/distal positions of the cores and the associated facies associations. This is reflected visually in the character of the time series with generally thin fine-grained sandstones in the southwest and more frequent thicker, coarser sandstones in the northeast (Fig. 4). In terms of the detection of regular cyclicity, the
Gartcosh o.H
Glenboig
', I
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Torwood
ol
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Doll Mill
0.12 o.i,
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.
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I
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Cycles per metre
0.~2
~ .
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',,16.70m ',
i
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~.o
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0.~
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i
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H
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BY
. . . . ,, .... \\ 9 \~',
...
..... 0.02
" -.."..
o.o) o.oo
~-~.... 0.0
o.I
0.2
0.3
_
o.~, o.5 0,6 o.r Cyctes p.r .atre
Fig. 5. Power spectra for all locations for the interval Knightswood Gas Coal to Index Limestone.
NAMURIAN SPECTRAL ANALYSIS, SCOTLAND spectra indicate a pair of significant cycles at Gartcosh, one at Torwood and Doll Mill, and at Orchardhead one significant and one highly significant cycle. No regular cycles are apparent at Glenboig or Solsgirth where a highly significant peak lies to the left of the low frequency cut-off and thus cannot be demonstrated to represent regular cyclicity. Since it appeared that regular cyclicity was demonstrated at some localities it was decided to test the stability of the cycles by examining the top and bottom of the data sets. In order to maintain comparability between locations the sections were split at the correlative of the top of the Hartley Coal (Figs 2 & 4). Figures 6 & 7 demonstrate, for both the upper and lower parts of the time series, the same grouping of spectra into pairs of locations by the overall distribution of variance as observed before. However, the spectral results differ markedly above and below the Hartley Coal in terms of the frequency of the dominant spectral peaks. Below the Hartley Coal significant regular cyclicity has been detected at Gartcosh, Glenboig, Torwood and Doll Mill. On the other hand, above the Hartley Coal, a significant regular cycle is present only at Doll Mill. The spectra generated from the complete time series represent the average variance densities throughout the records, whereas the shorter series show that the first cycles detected are restricted to only certain parts of the overall record. This change in the frequency of the dominant peaks and the predominance of regular
Gartcosh I
o.to ~
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~
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I
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2 99 4 m
J
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0.5 0.6 o.v CycLes ~ e r me':re
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i ~176 .... ,',,I, -
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Figure 8 illustrates the graphic logs of the decompacted sections and Fig. 9 the time series. Typically, the decompaction exercise changed the series from the rather cuspate profiles of Fig. 4 to a more sinusoidal pattern. This is, of course, especially influenced by the much greater decompaction of the coals and claystones relative to the sandstones. Unlike the original data, in the decompacted series, Doll Mill represents the thickest section in the upper part of the Limestone Coal Formation. The spectra related to the decompacted sections are illustrated in Figs 10 & 11. Compared to the original data, there has been a dramatic concentration of variance towards particular frequencies. This is exemplified by the data from below the Hartley Coal from the Gartcosh, Glenboig and, especially, the Torwood boreholes (cf. Figs 6 & 10). Below the Hartley Coal a single, highly significant, regular cycle has been detected at Gartcosh, Glenboig and Torwood with a significant cycle at Doll Mill, i.e. the four most southwesterly locations. Above the Hartley Coal a significant cycle was detected at Gartcosh and two were
0, 05 9
r ,'..
o. 03. o. oz. o.o~
r e c o d e d data
o. o~ O.O8 o.OZ
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~ .... A ~',~,, 0.0,. I
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0,12 O,t:
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cyclicity in the lower part of the succession suggests a major stratigraphic/temporal change in the nature of the signal being investigated (as discussed later).
Glenboig
0.:2 o.~o
59
-o.o
o.I
o.2
o. 3
o,~
o.~
o.6
CycLes p e r
o,;'
me:~e
Fig. 6. Power spectra for Gartcosh, Glenboig and Torwood for the intervals Hartley Coal to Index Limestone (above) and Knightswood Gas Coal to Hartley Coal (below).
60
G . P . WEEDON t~ W. A. READ
Doll Mill
Orchardhead
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per metre
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Fig. 7. Power spectra for Doll Mill, Orchardhead and Solsgirth for the intervals Hartley Coal to Index Limestone (above) and Knightswood Gas Coal to Hartley Coal (below).
detected at Torwood. The higher frequency spectral peak might represent a second harmonic of the low-frequency cycle, hence indicating an asymmetric cycle shape (Weedon 1991).
Gartcosh
Glenboig
The spectra for the lower part of the series at Gartcosh and Torwood were regenerated using decompaction related to 2 and 4 km overburden with virtually no resulting change in spectral shape
Torwood
Doll Mill .
.
.
.
Orchardhead .
.
.
.
.
.
Solsgirth
.
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....... ....... .......
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Decompacted Fig. 8. Graphic logs illustrating the lithological subdivision of the boreholes after decompaction (cf. with Fig. 3).
61
NAMURIAN SPECTRAL ANALYSIS, SCOTLAND
Gartcosh
Decompacted Torwood Doll Mill
Glenboig
! 2 3 4 5 6
1 23/,
1 2 3 4 5 6
56
Orchardhead
Solsgirth 1 2 3 4 5 6
1 2 3 4 5 6
1 2 3 4 5 6
100
\\
200
\_ .J= 3OO ~9
-~_ ~00
500
/-
--2
Fig. 9. Time series of the boreholes after decompaction (cf. with Fig. 4; codes as for Fig. 3). Dashed correlation line represents the Hartley Coal (Fig. 2).
Gartcosh ~
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moore
per
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9 .
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, o.oo
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o,2o per
95gC. I. ~ol c.I.
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~
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---
84 95g C. L
---
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9
~ o.o2
-~--'-~ L - ~ = o, r CycLes pep
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0. N me:re
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.... o.o~ 0.10
0.15 Cycles per-
Fig. 10. Power spectra based on decompacted section time series from Gartcosh, Glenboig and Torwood for the intervals Hartley Coal to Index Limestone (above) and Knightswood Gas Coal to Hartley Coal (below).
0.20 me:re
62
G. P. WEEDON & W. A. READ
Doll Mill
Orchardhead
0.12 o. lo 0. 09 o. 00
C
Xxl',
~-H Bw '.
i
[
1
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0.12 0.11 0.10 ,~
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Cycles p e r m e c r e
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0.~0
metre
o. io c~ o. 09
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o. co
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0.12 o. tl
\
o.oi
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o.o~ o.ol
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i [
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o. o0 ~'
\ , \] ',
o.~o
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"'T. -- -0.05
Cycles per metre
O. I0
O. 15 Cycles
0.20 per
metre
o. oo 0.00
0.05
o. 1o
o, 15 Cycles
o.~o per
me~re
Fig. 11. Power spectra based on decompacted time series from Doll Mill, Orchardhead and Solsgirth for the intervals Hartley Coal to Index Limestone (above) and Knightswood Gas Coal to Hartley Coal (below).
(Fig. 12). The result for Torwood with 4 km decompaction produced a slightly longer, regular cycle reflecting the greater thickness of the decompacted section. These localities were also reanalysed for the original 3 km decompaction, but using different coding systems. Firstly, the codes for coal, mudstone, siltstone, fine-, medium- and coarse-grained sandstone were changed to 3, 3, 3, 4, 4, 4 (i.e. non-sandstone/sandstone categorization), next the codes were changed to 3, 4, 4, 4, 4, 4 (i.e. coal/non-coal categorization), the results are shown in Fig. 12. Discussion The dramatic concentration of variance towards particular frequencies in the decompaction spectra are interpreted as indicating the successful removal of compactional distortions of the primary environmental signal (namely, sediment type as a function of stratigraphic position). This approach cannot, of course, remove variations due to sedimentation-rate changes. This may mean that even more regular signals might ultimately be revealed. It is emphasized that the same decompaction procedure was applied to all locations simply on the basis of rock type. The major change in the nature of the signal being investigated centred at the Hartley Coal prompted re-examination of the primary data. Above the Hartley Coal thick coals become markedly fewer and less laterally persistent and
thick intervals of multi-storey, channelized sandstones and isolated erosive-based, fining-up channel sandstones become more common. The spectral analysis has apparently revealed a major change towards greater fluvial influence in the younger data. In the lower data it is clear that the more westerly locations preserve a single regular sedimentary cycle. The location of these time series indicates that this cyclicity is associated with more distal/deltaic facies. The lack of regular cyclicity in the east cannot be unambiguously interpreted, but may reflect large variation in accumulation rate, dominance of fluvially controlled variations in grain size unrelated to base level and/or increased numbers of hiatuses. Above the Hartley Coal only two locations possess evidence of regular cyclicity. These lie again, in the more distal setting. The reduction of locations recording regular cyclicity probably reflect the increased fluvial influence towards the latter part of the Limestone Coal Formation. This may indicate a long-term fall in average relative sea-level and/or rejuvenation of hinterland source areas due to tectonic uplift or climatic change. The changes in overburden corrections do not affect the spectral results if maximum burial lies between 2 and 4 km (Fig. 12). Another line of evidence supports our crude approach to decompaction. In Fig. 13 Shaw plots, based on the occurrence of correlated coals at pairs of locations, are presented. The plots are constructed with both sections in each case represented as one unit or
NAMURIAN SPECTRAL ANALYSIS, SCOTLAND
Gartcosh
63
Torwood
4km Decompaction 0.12 0.11 0.~0 0. 09
o. ii
'
o. . . .o. ,
o.o8
i
i ':i o, 12o.o8
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9SZ
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0.05
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O. IS 0.~0 Cycles per metre
g
o. o~
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o. io
o, t5 0,20 Cycles per metre
Non-sandstone/sandstone
0. lZ 0. r t , 0. to. 0.09. o.
-~
o.oo
o.l~ o. fo r~ I
8~'
08. -
o.o;' o.o6.
- - 95% C. I. ~1 c.u
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Coal/non-coal O.Q o. 11 ", i
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i L,
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-
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o. ls Cycles per
o.2o metre
o.oo
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Fig. 12. Power spectra for decompacted section time series from Gartcosh and Torwood using overburden = 4 km (top), code values 3, 3, 3, 4, 4, 4 [i.e. non-sandstone/sandstone categorization (middle)] and codes 3, 4, 4, 4, 4, 4 [i.e. coal/non-coal categorization (bottom)].
100% long. Plots for the original data (dashed) and decompacted data (continuous line) are indicated. In the plots it is assumed that the coals were formed essentially simultaneously at all locations, but it is not assumed that the coals formed at constant time intervals. If there was a variation in the relative accumulation rates and/or compaction comparing the pairs of sites, this is apparent from the deviation of the plots from straight lines at 45 ~. Both graphs compare sections with Gartosh (from a distal setting). Starting with the original data, the more proximal location of Orchardhead has a much more erratic line compared to Glenboig in the distal setting. This could result from more variable sedimentation rates and/or greater compactional distortions in the proximal facies. The effect of the decompaction is to produce smoother, more nearly straight lines. This straightening is interpreted as indicating that the decompaction method used was at least partly successful. However, deviations from
straight lines remains, suggesting that perhaps the decompaction is incomplete or that sedimentation rates varied independently at different locations. Again, the comparison of Gartosh with the more distal location (Glenboig) reveals a straighter line that the comparison with the more proximal location (Orchardhead). With the rock-type codes changed to non-sandstone/sandstone categories (Fig. 12) the main cyclicity detected earlier remains detectable at Gartosh and Torwood. On the other hand, with the categories as coal/non-coal the Torwood cyclicity was detected as before, but a major change in spectral character occurred at Gartosh. The latter observation may reflect the thinning out westwards of some coals and their replacement by siliciclastic palaeosols in more distal areas. However the consistent results for Torwood for all code categorizations supports our model of an idealized regular distal cyclothem with coal and sandstone as
64
G.P. WEEDON •
W. A. READ
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Fig. 13. Shaw plots comparing the relative stratigraphic positions of coals for sections at Glenboig (distal setting) and Orchardhead (proximal setting) with Gartcosh (distal setting). The plots are based on data from original measured stratigraphic positions and stratigraphic positions following decompaction. end members. The regular cyclicity in this dominantly deltaic setting is attributed to orbital control of glacio-eustatic forcing. The regular sedimentary cyclicity detected in the more distal/deltaic locations in this study can be used for comparison with an earlier debate related to processes originating within the basin (autocyclic) and outside the basin (allocyclic). Read & Forsyth (1989) argued that a purely autocyclic signal would involve more 'sedimentary cycles' in areas of greater net sedimentation rates. This would relate to the presence of a characteristic wavelength for cycles formed autocyclically. On the other hand, a purely allocyclic signal would generate equal numbers of cycles across the basin, presumably with longer wavelengths in areas of higher sedimentation rate. Their actual data of cycle counts within a certain stratigraphic interval indicated a combination of allocycles and autocycles. Thus, the total numbers of cycles increase in areas of higher sedimentation rate but extrapolation to sections of 'zero thickness' indicates a finite minimum number or a background of allocyclicity. Figure 14 plots the numbers of regular cycles implied in each decompacted section of the Knightswood Gas Coal to Hartley Coal sections. Only four values are available but the number of regular cycles decreases with increase in decompacted section thickness. As the regular cycles are believed to be allocyclic in origin this result was unexpected. However, in this case, the thick-
ness of the section increases towards the proximal area (Fig. 1), suggesting the possibility that the more proximal settings may include further undetected hiatuses. Figure 2 demonstrates that fewer horizons can be unambiguously correlated in the
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Fig. 14. Changes in the number of regular sedimentary cycles between the Knightswood Gas Coal and the Hartley Coal as a function of decompacted stratigraphic thickness (above) and theoretical numbers of 'sedimentary cycles' as a function of thickness (below), based on Read & Forsyth (1989).
NAMURIAN SPECTRAL ANALYSIS, SCOTLAND
more northeasterly areas, so the possibility of gaps cannot be totally ruled out. Additionally, the Shaw plots (Fig. 13) demonstrate a more erratic line of correlation with proximal areas. This might be a function of a less complete record which would cause apparent variations in accumulation rate relative to more distal locations. Finally, the question of estimating the periodicity of the regular cycles arises. The time that this interval represents can be estimated only considering one or other of two scenarios, neither of which is readily verified. Both assume that there is a complete sedimentary record of orbitallyforced, glacio-eustatic cycles at Gartosh where the maximum number of regular cycles are present. Scenario A postulates that major marine transgressions took place at approximately equal time intervals during the Pendleian and Arnsbergian. This time intervals is now thought to be c. 10 Ma (Claoue-Long, pers. comm. 1993). On this basis, Read 1994b) calculated the recurrence interval of major transgressions as c. 1 Ma. In Gartosh the maximum number of cycles in the study interval is 13 below the Hartley Coal and 11 above, or 24 cycles overall, so the periodicity of the cycles studied here would be between 38 and 43 ka. Scenario B postulates that the rate of basin subsidence stayed approximately constant throughout the Pendleian and Amsbergian. The present study interval represents 20-25% of the compacted succession thickness, and thus 2-2.5 Ma, giving an estimated periodicity of 83.3-104 ka. During the late Carboniferous the dominant periods of orbitalprecession were only 17 and 21 ka, and those of the obliquity cycles were 34 and 43 ka, whereas the eccentricity periodicities remained at their present values of c. 100 and 400 ka (Berger et al. 1989). Thus, scenario A and B would suggest forcing by the obliquity or short eccentricity cycles, respectively. Unfortunately, both methods used for these estimates of the periods are poorly constrained for many reasons. Consequently, at best, it can only be claimed that, if an orbital cycle was recorded, it lay in a high frequency band.
Conclusions The spectral results obtained here indicate that between the deposition of the Knightswood Gas Coal and the Hartley Coal a single regular
65
sedimentary cycle exists in the sections from the more distal part of the basin. This is interpreted as cyclicity, with end members of coal and sandstone, which recorded regular fluvio-deltaic progradation cycles responding to orbital-forcing of glacioeustatic variations. These cycles may possibly record the effects of either the obliquity or short eccentricity cycle, although the dating is extremely uncertain. The irregular autocyclic/fluvial components of the distal areas contribute to the background continuum of the spectra. In the more proximal areas fluvial sedimentation produced a more irregular sedimentary signal, probably related to irregular fluvial processes such as channel avulsion, crevassing and meander migration, plus the influx of coarse clastic material forming thick wedges of multi-storey sandstones. Between deposition of the Hartley Coal and Index Limestone only two of the more distal localities record regular cyclicity. Thick multi-storey sandstones are more common and laterally persistent coals are less frequent. This could indicate a change in facies related to long-term increased episodic hinterland rejuvenation and/or a drop in average relative sea level. Despite the large variety of factors affecting deposition of these Namurian 'coal measures' the spectral analysis is apparently capable of detecting regular components through the noise. However, interpretation of the results from a single site would have produced a misleading impression of the presence or absence of regular cyclicity across the whole basin. The decompaction exercise, although crude, seems to have substantially removed differential compactional distortion of the primary environmental signal. The differing numbers of regular cycles apparently preserved in different locations may relate to the frequency of hiatuses increasing towards fluvially dominated areas. This interpretation increases the difficulty of trying to reliably estimate the periodicity of any cyclicity encountered in cyclothem series.
During this work G.EW. was supported consecutively by BP Research (Stratigraphy Unit), Downing College, Cambridge, and the University of Luton. The staff of the British Geological survey, Edinburgh Office, are thanked for facilitating access to their borehole records. J. ClaoueLong is thanked for kindly providing pre-publication information on a revised radiometric timescale for the Carboniferous.
References ALGEO, T. J. & WILKINGSON, B. H. 1988. Periodicity of mesoscale Phanerozoic sedimentary cycles and the role of Milankovitch orbital modulation. Journal of Geology, 96, 313-322.
BALDWIN, B. & BUTLER, C. O. 1985. Compaction curves. American Association of Petroleum Geologists Bulletin, 69, 622-626.
BERGER,A., LOUTRE,M. E & DEHANT,V. 1989. Influence
66
G . P . WEEDON •
of changing lunar orbit on the astronomical frequencies of pre-Quaternary insolation patterns. Paleoceanography, 4, 555-564. CARRS, B. W. &NEIDELL, N. S. 1965. A geological cyclicity detected by means of polarity coincidence correlation. Nature, 212, 136-137. ELLIOT'r, R. E. 1985. Quantification of peat to coal compaction stages, based especially on phenomena in the East Pennine Coalfield, England. Proceedings of the Yorkshire Geological Society, 45, 163-172. FISCHER, A. G., DEBOER, P. L. & PREMOLISILVA,I. 1990. Cyclostratigraphy. In: GINBURG,R. N. & BEAUDOIN, B. (eds) Cretaceous Resources Events and Rhythms. Kluwer Academic Press, Dordrecht, 139-172. FORSYTH, I. H. & READ, W. A. 1962. The correlation of the Limestone Coal Group above the Kilsyth Coking Coal in the Glasgow-Stifling region. Bulletin of the Geological Survey of Great Britain, 19, 29-52. FRAKES, L. A., FRANCIS, J. E. & SKYTUS, J. I. 1992. Climate Modes of the Phanerozoic. Cambridge University Press, Cambridge. HARLAND, W. B., ARMSTRONG,R. L., COX, L. E., CRAIG, L. E., SMITH, A. G. &SMITH, D. G. 1990. A Geologic Time Scale 1989. Cambridge University Press, Cambridge. HAYS, J. D., IMBRIE, J. &SHACKLETON, N. J. 1976. Variations in the Earth's orbit: pacemaker of the ice ages. Science, 194, 1121-1132. HERBERT,Z. D. 1994. Reading orbital signals distorted by sedimentation: models and examples. In: DE BOER, E L. & SMITH, D. G. (eds) Orbital Forcing and Cyclic Sequences. International Association of Sedimentologists, Special Publication, 19, 483-508. IMBRIE,J., HAYS,J. D., MARTENSON,D. G. et al. 1984. The orbital theory of Pleistocene climate: support from a revised chronology of the marine 8180, record. In: BERGER, A., IMBRIE, J., HAYS, J., KUKLA, G. & SALTZMAN, B. (eds), Milankovitch and Climate, Part 1. Reidel, Dordrecht, 269-305. , BOYLE, E. A., CLEMENS, S. C. et al. 1992. On the structure and origin of major glaciation cycles, 1, linear responses to Milankovitch forcing. Paleoceanography, 7, 701-738. MAYNARD, J. R. & LEEDER, M. R. 1992. On the periodicity and magnitude of Late Carboniferous glacio-eustatic sea-level changes. Journal of the Geological Society, London, 149, 303-311, PRESS, W. H., FLANNERY, B. P., TEUKOLSKY, S. A. & VErrERLING, W. T. 1986. Numerical Recipes: The Art of Scientific Computing. Cambridge University Press, Cambridge. PRIESTLEY, M. B. 1981. Spectral Analysis and Time Series. Academic Press, London. READ, W. A. 1988. Controls on Silesian sedimentation in the Midland Valley of Scotland. In: BESLY,B. K. & KEELING, G (eds) Sedimentation in a Synorogenic Basin Complex. The Upper Carboniferous of Northwest Europe. Blackie, Glasgow, 222-241. - 1994a. High frequency, glacio-eustatic sequences in early Namurian coal-bearing fluviodeltaic deposits, central Scotland. In: DE BOER, P. L. & SMITH, D. G. (eds) Orbital Forcing and Cyclic Sequences. Special Publication of the International Association of Sedimentologists, 19, 413-428. 1994b. The frequencies of Scottish Pendleian
W. A. READ allocycles. Scottish Journal of Geology., 29, 91-94. Sequence stratigraphy and lithofacies geometry in an early Namurian coal-bearing succession in central Scotland. In: WHATELEY, M. K. & SPEARS, A. (eds) Coal - Exploration, Energy Policies and the Environment. Geological Society, London, Special Publication. 82, 285-297. - & FORSYTH,I. H. 1989. Allocycles and autocycles in the upper part of the Limestone Coal Group (Pendleian E l) in the Glasgow-Stifling region of the Midland Valley of Scotland. Geological Journal, 24, 121-137. - - & 1991. Allocycles in the upper part of the Limestone Coal Group (Pendleian E l) of the Glasgow-Stifling region viewed in the light of sequence stratigraphy. Geological Journal, 26, 85-89. REIrEL, W. 1991. Coal cyclothems and some models for their origin. In: EINSELE, G., RICKEN, W. & SEILACHER, A. (eds) Cycles and Events in Stratigraphy. Springer, Berlin, 733-750. RUDDIMAN W. F., RAYMO, M. E., MARTINSON, D. G., CLEMENT, B. M. • BACKMAN,J. 1989. Pleistocene evolution: northern hemisphere ice sheets and North Atlantic Ocean. Paleoceanography, 4, 353-412. SCHWARZACHER,W. 1967. Some experiments to simulate the Pennsylvanian rock sequence of Kansas. Kansas Geological Survey Computing Contributions, 1 8 , 5-14. SHACKLETON,N. J., BERGER,A. L. t% PELTIER,W. R. 1990. An alternative astronomical calibration of the Lower Pleistocene timescales based on ODP Site 677. Transactions of the Royal Society of Edinburgh: Earth Science, 81, 251-261. VAN BUCHEM, F. S. P., MCCAVE, [. N. & WEEDON, G. E 1994. Orbitally induced small-scale cyclicity in a siliciclastic epicontinental setting (Lower Lias, Yorkshire, UK). In: DE BOER, P. L. & SMm-I, D. G. (eds) Orbital Forcing and Cyclic Sequences. Special Publication of the International Association of Sedimentologists, 19, 345-366. VAN ECHELPOEL,E. 1994. Identification of regular sedimentary cycles using Walsh spectral analysis with results from the Boom Clay Formation, Belgium. In: DE BOER, P. L.& SMITH, D. G. (eds) Orbital Forcing and Cyclic Sequences. Special Publication of the International Association of Sedimentologists, 19, 63-76. &WEEDON, G. P. 1990. Milankovitch cyclicity and the Boom Clay Formation: an Oligocene siliciclastic shelf sequence in Belgium. Geological Magazine, 127, 599-604. VEEVERS, J. J. & POWELL, C. M. 1987. Late Palaeozoic glacial episodes in Gondwanaland reflected in transgressive-regressive depositional sequences in Euramerica. Geological Society of America Bulletin, 98, 475-487. WEEDON, G. P. 1991. The spectral analysis of stratigraphic time series. In: E~NSELE, G., RICKEN, W. & SEILACHER, A. (eds) Cycles and Events in Stratigraphy. Springer, Berlin, 840-854. - 1993. The recognition and stratigraphic implications of orbital-forcing of climate and sedimentary cycles. WRIGHT, V. P. (ed.) Sedimentary Review, Blackwell, Oxford, 31-50.
- - 1 9 9 5 .
Calibration, analysis and interpretation of depositional cycles in the Early Toarcian of Yorkshire, UK G W Y N E G. M O S E S
Phillips Petroleum Company United Kingdom, Phillips Quadrant, 35 Guildford Road, Woking, Surrey, GU22 7QT, UK Abstract: The Toarcian of Yorkshire, UK, has been analysed and interpreted in order to determine whether shale-nodule couplets or cycles represent a regular cyclicity, and, if so, what the cyclicity could represent. This was done by calculating rates of deposition by analysing shale laminae and calculating durations of hiatus from rates of growth of nodules which are believed to be formed during hiatuses. This has enabled a new ammonite geochronology to be calculated, giving average zonal durations of 700 ka and subzonal durations close to 300 ka. The cyclicity is not regular enough to represent Milankovitch cycles and a tectonic origin related to pulsed North Sea thermal doming is suggested as an alternative mechanism.
This paper attempts to suggest an alternative method for the calibration of cycles in a welldocumented section from the Toarcian of Yorkshire, UK. The section consists of a shale sequence with layers of calcareous nodules, which are early diagenetic in origin and are believed to represent hiatuses. The shale and nodule bands alternate and can be considered to represent couplets, or cycles, which can then be analysed. The most effective way of analysing cycles for periodicity is by time series analysis, but the measured thicknesses must first be converted to a time section, independent of cyclicity. Otherwise, if regular cyclicity is invoked to convert to a timescale, the data are already biased towards cyclicity before analysis can be done. Following the calibration, time series analysis, using Fourier transforms, was used to identify underlying cycle periods, which were then tested against possible causes such as orbital forcing and local or regional tectonic influences. Previous w o r k The Toarcian of Yorkshire has been the focus for stratigraphic study for over 100 a, and has been analysed in various ways as a classic section. The first modem detailed stratigraphy was given by Dean (1954), who described the beds at Blea Wyke near Ravenscar. This was followed by Howarth (1962) on the Jet Rock and Alum Shale Series, and Howarth (1973) on the Grey Shales. These three papers give detailed bed-by-bed descriptions including faunal lists, on which further work can be based. Cope et al. (1980) summarized the lithostratigraphy and biostratigraphy for the Upper Lias, and
this was followed by Powell (1984) and Knox (1984) who erected a formal lithostratigraphy for the Lias, independent of biostratigraphy. Hallam (1961) proposed the first cyclostratigraphy for the Lias, his units IX, X and XI covering the Toarcian. However, the duration of the cycles is too long to represent orbital-forcing cycles on the Milankovitch scale. Hallam (1978, 1981) also showed that worldwide eustatic sea-level changes, related to mid-ocean ridge uplift and subsidence, would be of too long a duration. The first proposal that orbital cycles could be used to subdivide a Jurassic succession using sedimentary microrhythms was by House (1985), who showed data from the Kimmeridgian of Kimmeridge Bay, UK, and the Lower Lias of Lyme Regis, UK, with well-developed microrhythms. This was built upon by Weedon (1989) who analysed the Lias of Switzerland for orbitally forced variations using Walsh power spectra and filtering. Using ratios of thicknesses of cycles he attributed three cycle periods to the 21 000, 41 000 and 100 000 a Milankovitch cycle periods. In this paper the stratigraphy of Dean (1954) and Howarth (1962, 1973) was confirmed by remeasuring the section and rationalizing discrepancies. Beds and nodule diameters were remeasured to a precision of 5 mm, compared to the previous interval of 1 inch. The biostratigraphy of the Toarcian of Yorkshire is well understood, with a complete sequence of ammonite zones and subzones. The main problem in analysing cycles is in establishing the duration of the cycles, as opposed to the thickness of the cycles in the rock record. In other words, cycles must be analysed in time, not thickness.
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescales and Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 67-74.
67
68 Thickness
G.P.G. - time conversion
With varying lithologies it is unlikely that depositional rates will be constant, so a method must be used to convert the measured section into a time series. This is commonly done in two ways (outlined below), and a third method is proposed in this paper. Conventional method
Geochronological dates are given for the beginning and end of the Toarcian, and it is then divided into equal time periods based on equal duration ammonite zones or subzones. There are two major flaws to this method. Firstly, it is unlikely that all ammonite species had a collective death wish at the end of the same time period, and secondly, there is little agreement about the duration of the Toarcian (Table 1). Thus, for example, the Variabilis zone or subzone could have a duration from 292 to 700 ka. Cyclostratigraphy method
This requires that mesoscale cycles are identified and attributed to external causes, such as Milankovitch cycles. Although this is a proven method within the Pleistocene, using it for a Mesozoic succession assumes that: (1) sedimentary and orbital cycles are related, and of fixed and predictable duration; (2) no other factors, such as local tectonics, influence deposition, and (3) that the succession is complete. The methods of recognizing these cycles vary from ratios of thicknesses of beds to statistical analyses of sequences. This method has been extensively utilized recently, and has given some convincing results, e.g. in the Lower Cretaceous. Proposed method
Quantitative methodology can be used to calculate short-term sedimentation rates and the duration of hiatuses. The durations of deposition and hiatus are Table 1. Toarcian chronostratigraphy Author
Start
End
Duration
Van Hinte 1976 Harland et al. 1982 Kennedy & Odin 1982 Odin 1984 Westerman 1984 Hallam et al. 1985 Kent & Gradstein 1985 Haq et al. 1987 Bayer & McGhee 1986 Harland et al. 1989
178 194 189 189 191.5 187 193 186 198 187
174 188 181 181 184 179 187 179 190 178
4 6 8 8 7.5 8 6 7 8 9
MOSES then summed and the total elapsed time calculated. This results in a more direct chronostratigraphy, against which other methods can be tested. In the Toarcian of Yorkshire this depends on shale laminae representing annual cycles (Hallam 1967), as well as the geochemistry of nodule diagenesis being understood so that rates of nodule growth can be correlated with durations of hiatuses. The reasons that the nodules are thought to represent hiatuses are that they are confined to discrete horizons, are formed at an early stage of diagenesis at the oxidizing/reducing boundary and take a significant period to form. Shales. There are two main types of shale present in the Toarcian of Yorkshire; the grey shales and the bituminous black shales. The bituminous black shales are part of the Jet Rock Member, which represents the Falciferum maximum flooding event. The effective (post-compaction) rate for the black shales is 0.3 m m a -1 and for the grey shales 0.53 m m a -1. However, further up the succession, in the Dispansum to Aalensis subzones, the depositional environment is high-energy reworked sands, so sedimentation rates cannot be determined. Nodules. There are two major types of nodules present, which vary with the enclosing lithology. In the grey shales the nodules tend to be small (< 100 mm), whereas those in the black shales are large (commonly > 150 mm). It has been suggested (Weedon 1993 pers. comm.) that the different types of nodules could be related to organic content in the enclosing shales. The formula for calculating the rates of growth and elapsed time for nodule formation was derived by Berner (1968), assuming no unidirectional flow in the substrate. The formula is: t = R2/2vD (Coo - CR), where t is the time for nodule formation, D the diffusion coefficient (cm 2 s -1) assumed to be 10 -5, R the radius of concretion, v the molar volume of concretion cement (cm 2 mol-1), Coo the CaCO 3 concentration at infinite distance from concretion, C R the concentration at surface of concretion, and (Coo- C R) the degree of super saturation (mol cm-3). Raiswell (1988) applied this formula to a nodule band in the bituminous shales assuming a normal seawater concentration of 400 ppm CaCO 3, giving the degree of supersaturation as 10 -5 mol cm -3. Berner (1968) used 10 ppm and 10 "7 for calcareous concretions in the grey shales. These values resolve at t = 4.53 R 2 in the bituminous shales and 453R 2 in the grey shales. However, it must be stressed that these numbers should be treated with caution, as the variability in parameters is probably greater than recognized, e.g. organic content may effect the diffusion coefficient.
CYCLES IN THE TOARCIAN OF YORKSHIRE
resampling, and a 500 a interval was chosen as it would sample every depositional and hiatal episode. The data was analysed using Statistica from Statsoft. The analysis was carried out in two parts, firstly the whole section was analysed over the entire duration, followed by detailed analysis of the oldest three subzones, Paltum to Tenuicostatum and Variabilis to Striatulum subzones.
Table 2. Ammonite zonal geochronology Zone
Duration (ka)
Subzone
Thouarense
Fallaciosum Striatulum Variabilis Crassum/Braunianus Fibulatum Commune Falciferum Exeratum Semicelatum Tenuicostatum Clevelandicum Paltum
Variabilis Bifrons Falciferum Tenuicostatum
516.6 166.2 333.7 466.1 122.3 468.7 688.6 63.6 305.8 182.0 20.2 169.2 3503.4 700.7 292.0
Total Average zone Average subzone
Paltum to Tenuicostatum This part of the sequence was chosen because it consists of a well-defined series of shales and nodules, with 14 cycles unevenly spaced represented in c. 370 ka (Fig. 1). The data was then run through a fast Fourier transform (FFT) to give a raw frequency distribution. The frequency spectrum is difficult to interpret, due to the block nature of the time series producing a spike at very low frequencies (Fig. 2). In order to remove this effect, a 15% taper was introduced to the data, but the resultant spectrum is still very spiky (Fig. 3). Finally, a Harming transform was applied which resulted in a smooth and interpretable spectrum (Fig. 4). Dominant frequencies are 20, 36, 55 and 80 ka, with 12, 28, 95 and 120 ka also present.
By summing the periods of deposition and hiatus, and correlating them with ammonite biostratigraphy, a chronostratigraphy can be constructed for the lower part of the Toarcian (Table 2). Data analysis
One of the main problems with statistical analysis of a geological data set is that normally a geological section is measured at changes in lithology, but time series analysis needs the data points to be evenly spaced. This means that the data set needs
9
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Variabilis to Striatulum The same sequence of processing was applied to a younger section containing 10 cycles and a
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70
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within only five subzones, or 20% of the calculated section. Hence, it is unlikely that there is a regular cyclicity recorded in this sequence.
different set of dominant frequencies obtained, namely 10, 17, 26, 37 and 42 ka (Fig. 5), with 53, 69, 86, 101 and 130 ka also identifiable. This shows that 16 different frequencies have been obtained,
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Entire section Due to software limitations, the entire section was further resampled to a 5000 a interval, and longperiod cyclicities calculated (Fig. 6). The signifi-
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cant cycle lengths obtained were 200, 320, 400, 500 and 630 ka, with a tailing off of many other frequencies which do not appear to be statistically significant.
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time (Me)
Fig. 6. Paltum to Fallaciosum; FFT with 15% taper and Harming applied.
Milankovitch origin It is noticeable on the smaller scale that the cycle lengths fall within the Milankovitch band, from 10 to 130 ka. This poses the questions of whether these are valid cycles and what is the causal mechanism. The first point is that the methodology is untested on other sequences, so, by virtue of the measurement interval and analysis, a regular cyclicity may be generated. A further point is raised by Algeo & Wilkinson (1988) who state, 'For many cyclic sequences, calculation of a Milankovitch range period may be a virtual certainty, regardless of the actual genetic mechanism of cycle formation'. Pisias & Mix (1988) discuss the aliasing of time series in the geological record, and conclude that variance in sampling and measurement can significantly affect the cycle length that can be obtained through analysis. Weedon &Jenkins (1990) showed that similar sediments from the Pliensbachian of the Wessex Basin show a similar range of durations, with shale-marl couplets at 8-34 ka, and couplet bundles at between 54 and 215 ka. The couplets have been assigned to the precession or obliquity cycles, but the bundles to large, irregular climatic variations.
Tectonic origin In order to test a tectonic origin for the cycles, the
depositional environment of the sediments and the regional tectonic framework should be taken into account. The Toarcian of Yorkshire was deposited in a shallow epeiric sea within the photic zone, and was very susceptible to sea-level changes. The regional tectonic framework at this time was controlled by regional rift tectonics, including the early stages of thermal doming in the North Sea (Underhill & Partington 1993). The thermal uplift caused the 'Central North Sea Dome' which is centred on the Rattray or Forties volcanic centre (Fig. 7). The dome is calculated to have a radius of up to 600 km, whereas the Yorkshire coast is c. 400 km from the centre, although relief would have been minor. The reason that there are cycles in the studied sequence is that uplift is not believed to have occurred at a constant rate, but was pulsed. Supporting evidence for this comes from studies of the East African Rift (Asfaw et al. 1992), where even under conditions of uniform strain, rifting is episodic in nature. There is no evidence to suggest that strain during Jurassic rifting, and therefore uplift of thermal domes and pluming, is constant, so it is probable that the distal effects of uplift would be episodic. By analogy, it is possible that Volcanic centres in the Porcupine Basin could affect the Lower Jurassic sediments in Dorset and Somerset, UK, in a similar way.
CYCLES IN THE TOARCIAN OF YORKSHIRE
ENGLAND TOARCIAN
NORTH SEA
73
NORWAY
. . . . .
SIMPLIFIED PALAEOSTRUCTURE - INTRA AALENIAN (Not to scale)
RATTRAY
#
VOLCANIC CENTRE
RIASSIC OR BASEMENT
200~m
m
I . a p
Fig. 7. Effect of North Sea Thermal Doming. Above Simplified palaeostructure. Below Effect of thermal doming (after Underhill & Partington 1993).
74
G. P. G. MOSES
Conclusion Cycles of couplets exist in the Toarcian of Yorkshire, but no regular period can be identified, although calculated frequencies lie within the range of Milankovitch. A possible alternative m e c h a n i s m is episodic uplift of the Central North Sea Dome.
The author is grateful to Phillips Petroleum Company for granting permission to publish this paper. However, the company does not accept responsibility for any of the views put forward, which were developed from a thesis for an MSc in Stratigraphy at Birkbeck College, London University.
References ALGEO, Z. J. • WILKINSON, B. H. 1988. Periodicity of mesoscale Phanerozoic sedimentary cycles and the role of Milankovitch orbital modulation. Journal of Geology, 96, 313-322. ASFAW,L. M., BILHAN,R., JACKSON,M. & MOHR, P. 1992. Recent inactivity in African rift. Nature, 357, 447. BAYER, U. & MCGHEE, G. R. 1986. Cyclic patterns in the Paleozoic and Mesozoic: implications for time scale calibrations. Paleoceanography, 1, 383-402. BERNER, R. A. 1968. Rate of concretion growth. Geochimica et Cosmochimica Acta, 32, 477-483. COPE, J. C. W., GETTY,T. A., HOWARTH,M. K., MORTON, N. &TORRENS, H. S. 1980. A Correlation of
Jurassic Rocks in the British Isles. Part One: Introduction and Lower Jurassic. Geological Society, London, Special Report, 14. DEAN, W. T. 1954. Notes on part of the Upper Lias succession at Blea Wyke, Yorkshire. Proceedings of the Yorkshire Geological Society, 29, 161-179. HALLAM, A. 1961. Cyclothems, Transgressions and Faunal Change in the Lias of North-West Europe. Transactions of the Edinburgh Geological Society, 18, 124-174. - - 1 9 6 7 . An environmental study of the Upper Domerian and Lower Toarcian in Great Britain. Philosophical Transactions of the Royal Society of London, B252, 393-445. - 1978, Eustatic cycles in the Jurassic. Palaeogeography, Palaeoclimatology & Palaeoecology, 23, 1-23. - - 1 9 8 1 . A revised sea level curve for the early Jurassic. Journal of the Geological Society, London, 138, 735-743. , HANCOCK,J. M., LABRECQUE,J. L., LOWRIE, W. & CHANNELL, J. E. Z. 1985. Jurassic and Cretaceous geochronology and Jurassic to Paleogene magnetostratigraphy. In:SNELLING, N. J. (ed.) The Chronology of the Geological Record. Geological Society, London, Memoir, 10, 118-140. HAQ, B. U., HARDENBOLL, J. & VAIL, P. R. 1987. Chronology of fluctuating sea levels since the Triassic. Science, 235, 1156-1167. HARLAND,W. B., Cox, A. V., LLEWELLYN,P. G., PICKTON, C. A., SMITH, A. G. & WALTERS, R. 1982. A Geological Time Scale. Cambridge University Press, Cambridge. , ARMSTRONG,R. L., Cox, A. V., CRAIG,L. E., SMITH, A. G. & SMITH, D. G. 1989. A Geological Time Scale. Cambridge University Press, Cambridge. HOUSE, M. R. 1985. A new approach to an absolute timescale from measurements of orbital cycles and sedimentary microrhythms. Nature, 315, 721-725. HOWARTH,M. K. 1962. The Jet Rock Series and the Alum
Shale Series of the Yorkshire Coast. Proceedings of the Yorkshire Geological Society, 33, 381-422. - 1973. The stratigraphy and ammonite fauna of the Upper Liassic Grey Shales of the Yorkshire Coast. Bulletin of the British Museum (Natural History), Geology, 24, No. 4. KENNEDY, W. J. & ODIN, G. S. 1982. The Jurassic and Cretaceous Time scale in 1981. In: ODIN, G. S. (ed.) Numerical Dating in Stratigraphy. Wiley, New York, 557-592. KENT, D. V. & GRADSTEIN,E M. 1985. A Cretaceous and Jurassic geochronology. Bulletin of the Geological Society of America, 96, 1491-1427. KNOX, R. W. O'B. 1984. Lithostratigraphy and depositional history of the Late Toarcian sequence at Ravenscar, Yorkshire. Proceedings of the Yorkshire Geological Society, 45, 99-108. ODIN, G. S. 1984. Geochronology of the Jurassic time: status in 1984. In: MICHELSON,O. & ZEISS, A. (eds) International Symposium on Jurassic Stratigraphy. Symposium volume 3, Geological Survey of Denmark, 768-776. PISIAS,N. G. & Mix, A. C. 1988. Aliasing of the Geologic record and the search for long-period Milankovitch cycles. Paleoceanography, 3, 613-619. POWELL, J. n. 1984. Lithostratigraphical nomenclature of the Lias Group in the Yorkshire Basin. Proceedings of the Yorkshire Geological Society, 45, 51-57. RAISWELL, R. 1988. The microbiological formation of carbonate concretions in the Upper Lias of NE England. Chemical Geology, 18, 227-244. UNDERHILL, J. R. & PARTINGTON,M. A. 1993. Jurassic thermal doming and deflation in the North Sea: implications of the sequence stratigraphic evidence. In: PARKER, J. R. (ed.) Petroleum Geology of Northwest Europe:Proceedings of the 4th Conference. Geological Society, London, 337-345. VAN HINTE, J. E. 1976. A Jurassic time-scale. Bulletin of the American Association of Petroleum Geologists, 60, 489-497. WEEDON, G. P. 1989. The detection and illustration of regular sedimentary cycles using Walsh power spectra and filtering, with examples from the Lias of Switzerland. Journal of the Geological Society, London, 146, 133-144. --& JENKYNS, H. C. 1990. Regular and irregular climatic cycles and the Belemnite Marls (Pliensbachian, Lower Jurassic, Wessex Basin). Journal of the Geological Society, London, 147, 915-918. WESTERMAN, G. E. G. 1984. Gauging the duration of the stages: a new approach for the Jurassic. Episodes, 7, 26-28.
High-resolution palynofacies investigation of Kimmeridgian sedimentary cycles H E L E N K. W A T E R H O U S E
Department of Geology, The University, Southampton, S17 1B J, UK Abstract: Palynofacies analysis is used as a tool to investigate in detail the palaeoenvironmental variations through several sedimentary cycles in the Kimmeridge Clay of Kimmeridge Bay, Dorset, UK. Evidence is given of palaeoenvironmental variations within cycles corresponding to those expected for obliquity orbital forcing. In addition, a second cyclical palaeoenvironmental variation, probably precessional forced, is seen in the palynofacies data. Further small-scale variations in palynofacies characteristics, which are not evident in the sedimentology, are also identified and allow cycles to be divided into a number of distinct palaeoenvironmental units. It is proposed that the obliquity cycle had its greatest effects on the marine environment, while the precessional cycle mainly affected the terrestrial environment. The abundance of useful palaeoenvironmental and palaeoclimatic information obtainable through high-resolution sampling in conjunction with a tool such as palynofacies analysis, provides evidence for and information about orbital forcing additional to that of most orbital forcing studies as it allows variations within cycles to be investigated.
High-resolution studies of cyclical sedimentary sequences have been few, and detailed intra-cycle investigations even fewer. Much research into orbitally forced sedimentary cycles has concentrated on the interpretation of palaeoenvironments and palaeoclimates, and the calculation of the duration of cycles through thick cyclic sequences. Palynofacies analysis is the study of all the organic particles recovered from palynological preparations. In the present work, palynofacies analysis is used to characterize and to interpret lithological and organic particle cyclicity in terms of possibly orbitally-forced climatic changes. An attempt is made to interpret any variations through individual cycles, in relation to the known effects of the orbital forcing cycles on climate, using high-resolution sampling techniques in order to identify changes within individual cycles. Cycle durations are also estimated using time series analysis and extrapolation methods. Only after attempts have been made using the methods separately are they compared. Attempts to calculate absolute palynofacies particle abundances, a measure little used in pre-Quaternary studies, and palynofacies, were made. These were then able to be employed in order to make influence about the marine and terrestrial ecological realms separately. This work also provides a means of evaluating the effectiveness of palynofacies as a tool in such studies. Orbital forcing has been invoked as a cause for variations in a wide variety of palaeoenvironmental situations. It has also been investigated in terms of a new approach to the establishment of an absolute timescale (e.g. House 1985) and an estimation of stratigraphic errors (e.g. Schwarzacher 1989).
Mechanisms for linking changes in orbital parameters and solar insolation to climate, ocean circulation and pelagic sedimentation are more obscure than those for the response of global ice volume which have received greater attention (e.g. Hays et al. 1976). Other studies on orbitallyforced cyclicity, which have also employed highresolution sampling techniques, are those of Ditchfield & Marshall (1989), Cottle (1990) and Leary et al. (1989). The term 'palynofacies' was originally defined (Combaz 1964) to refer to the general aspect of palynological preparations, and is the study of all of the particulate organic matter present in a palynological preparation. As a micropalaeontological study it is very suitable for small-scale investigations. Unlike many branches of micropalaeontology, however, the analysis of palynofacies can be made as simple or as detailed as necessary by using an adaptable classification scheme. In addition, if a method of estimating absolute particle abundances is employed, an attempt can be made to assess variations in the marine and terrestrial components independently without the effects of data closure encountered when using percentage data, thereby providing information about the reactions of both the marine and terrestrial environments to any environmental variations. The Kimmeridge Clay Formation of Dorset, UK, was chosen for this investigation because of its well-documented sedimentary cyclicity and biostratigraphy. The section studied is located immediately west of Gaulter's Gap at Kimmeridge Bay (Ordnance Survey Landranger sheet number 195,
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescalesand Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 75-114.
75
76
H. K. WATERHOUSE
grid reference 907793). Approximately 12 m of section were examined in detail, which are situated directly above and below the Washing Ledge Stone Band within the Washing Ledge Shales of Arkell (1947), and in the autissiodorensis Zone of the uppermost Lower Kimmeridge Clay Formation. Details of the complete succession are given by Cox & Gallois (1981). This part of the Kimmeridge Clay consists essentially of alternating bituminous shales and dark grey mudstones, interrupted by the Washing Ledge Stone Band, a dolomitic limestone bed. Although 12 m of section were studied, detailed results and discussion of variations within cycles are limited mainly to the three uppermost lithological cycles in the section. Figure 1 gives a detailed lithological log of the whole sampled section. Cycles referred to in this work have been named after the bituminous shale bed which comprises their lower part. Thus, the WL4 cycle consists of the WL4 shale bed plus the mudstone bed above it. The Kimmeridge Clay is generally considered to have accumulated in a relatively shallow marine environment on a broad continental shelf (Gallois 1976) between 30 and 40 ~ N. Estimates of the depositional water depths for the Kimmeridge Clay of Dorset vary considerably, for example, from < 10 to 30 m (Hallam 1967, 1975), up to a few tens of metres (Aigner 1980), 50--100m (Myers & Wignall 1987) and up to 100 m (Gallois 1976). The whole sequence is thought to have accumulated below wave base because it shows only rare evidence of current activity (Tyson et al. 1979) or storm beds (Wignall 1991). The Kimmeridge Clay contains substantial amounts of organic material. Discussion of other workers' observations on its organic-matter accumulation is given in the discussion. Dunn (1974) suggested the possibility of orbital cycles in the Kimmeridge Clay through using Fourier analysis of geochemical data. He studied a 20 m section, situated slightly higher in the succession than that used in the present study, and recognized no apparent cyclicity in the lithological variation over the interval. Fourier analysis of his trace element data indicated cyclicity and, assuming an estimated duration of 10 ka for the deposition of each metre of sediment, Dunn proposed several superimposed periodicities with durations of 200, 100, 40, 14 and 11 ka. Several of these periods were coincident with the duration of astronomical phenomena which could cause climatic fluctuations. His study revealed a 1.43 m thick cycle, which is in close accord with the average cycle of 1.45 m observed over a much larger interval of the Kimmeridge Clay by Downie (1955), and with the lithological cycles observed in this study (see Fig. 1). Dunn estimated a duration
of c. 14 ka for these cycles, in contrast to that of c. 40ka later given by House (1985). House obtained estimates ranging from 36.9 to 48.1 ka, depending on which of the many estimates for the duration of the Jurassic was used. He concluded that the obliquity cycle may be assumed to be the dominant one. Herbin et al. (1991) revealed a cyclic distribution in the organic content of the sediments of the Kimmeridge Clay Formation and found the cycles to represent 30 ka, by dividing an ammonite zone by the number of cycles contained in it. Several idealized cyclothems have been proposed for the Kimmeridge Clay (e.g. Tyson et al 1979; Cox &Gallois 1981), with some upper Lower Kimmeridge Clay and Upper Kimmeridge Clay sediments showing a sequence from the base of a cycle upwards of oil shale or bituminous mudstone to dark grey mudstone to calcareous mudstone, with the carbonate content increasing and the kerogen content decreasing upwards (Cox & Gallois 1981). The cycles of the present work, however, lack the upper carbonate part or the full development of oil shale at the base. Using the stratigraphic timescale of Harland et al. (1989), the actual age of the section used in this study was calculated to be c. 153 Ma. The expected periods of orbital cycles at that time have been estimated using data given in Berger et al. (1989a,b) who calculated the variation back in time of the periods of the precession and obliquity cycles, which are known to have altered during the evolution of the Earth-Moon system. Their estimated cycle durations were 37.8 ka for the obliquity cycle and 21.9 and 18.3 ka for the two main precessional cycles. These alterations are all, however, relatively small when compared to the amount of stratigraphic information which may be lost during and after deposition and diagenesis, and the possible introduction of errors during processing and analysis.
Methods Palynofacies classification The classification of organic particles has always been rather subjective. Classifications often have a particular objective in mind and are therefore usually tailored for that purpose. For example, particles have been divided by their modification and thermal alteration, their depositional environments, botanical classification, degree of terrigenous supply and thereby distance from land, degree of degradation, and allochthonous (transported) and autochthonous (in situ) fractions. For this study, a method of classification was devised in order to obtain as much palaeoenvironmental and palaeoclimatic information as possible
@
KEY.
B15 B5
-
C
~.~
,.,
~
@
@
"~
B5, G4 etc.
] r
D10
"'
Indicates presence of ammonites Indicates every fifth sample point.
WL4
[
C25
o N
@
@
C5 C15
Indicates presence of bivalves.
J WL5
~
1
i
1 .... rr3)
F__7_
Medium to dark olive grey, hard laminated bituminous shale (laminae c. 2-5ram). Medium to dark grey, hard massive, organic-rich (but less so than above) mudstone Sometimes poorly laminated near contacts with above.
D20 r~)
Sample number
WL3
Transitional between the above two lithologies.
E7 E12
I
F3 F8
As next above but softer and lighter in colour and homogenous
1 WL2
I I I ~[~
o~
Limestone
F13 F18 (fib
F23
WL1
(i)
F28 O
F33 o
[ j [ J [2 1 I I I I I i I
G4
WLSB
G9
Fig. 1. Graphic lithological log of the sampled section of Kimmeridge Clay at Kimmeridge Bay in Dorset, UK. Bed numbers follow those of House (1985), indicating bituminous mudstone beds (WL) above the Washing Ledge Stone Band and (F) above the Flats Stone Band.
r ~
G14 --~" ~
~
~
t
Fll
G19
1
G24 G29
~
F10
,,
I
1m
Ga4
~ F9
78
H. K. WATERHOUSE
from palynofacies analysis. Table 1 shows the classification scheme used. It is essentially a morphological classification but it also incorporates the broad areas of provenance of particles. In order to investigate a large number of cycles, systematic taxonomy was considered too time consuming for this study. However, in the classification, palynomorphs were divided into broad systematic groups and, in some cases, morphological groups (see Table 1), The three main particle groups were marine palynomorphs, terrestrial palynomorphs
Table 1. Palynofacies classification used in this study. Particles are divided into three main groups based primarily on provenances, and then further subdivided by morphology. A second count is undertaken on the structured (terrestrial) debris group. Amorphous organic matter (AOM) is not included in the classification Marine palynomorphs
Chorate dinocysts Proximate dinocysts Otbers acritarchs foraminiferal test linings tasmantids prasinophytes others Terrestrial palynomorphs
Bisaccate pollen Non-saccate pollen Spores
medium and heavy spores light spores
Degraded palynomorphs Structured (terrestrial) debris
Cuticle Black debris
Brown debris
laths - length > 2x width equidimensional sharp debris length < 2x width equidimensional rounded debris length < 2• width Second count (structured debris)
Black laths Black equidimensional debris Dark laths and tubes Dark equidimensional laths Medium brown laths and tubes Medium brown equidimensional debris Medium brown laths and tubes with lateral thickenings - stripy Medium brown equidimensional debris with thickenings - stripy Light brown-yellow laths and tubes Light brown-yellow equidimensional debris Other structured debris
and terrestrial debris. These categories were subdivided into the categories shown in Table 1. 'Degraded' particles (actually only partially degraded) were also counted. These consisted of particles which had obviously once been palynomorphs but were too indistinct to assign to any group, even to distinguish between marine and terrestrial palynomorphs. However, they were more distinct than amorphous organic matter (AOM), A number of categories were also used in a second particle count (see Table 1). This was undertaken to investigate whether a more detailed analysis of the large variety of structured terrestrial debris could provide any additional palaeoenvironmental information. These second count results are largely not presented separately in the figures as they generally did not provide any additional information for palaeoenvironmental interpretation. They were useful, however, when assessing the reliability of some of the data. A second count provided a method of checking the reliability of the counting process by comparing the two count results. The counts were carried out following the same practices for both the first and second particle counts as described below. Much palynofacies work is semi-quantitative. It may consist of visual estimations of the character of a sample or, at best, percentage particle abundance counts. Lorente (1990) introduced quantitative data generated through digital image analysis of kerogen concentrates. In addition to using percentage particle abundances, Tyson (1989) illustrated absolute abundance trends by combining percentage particle abundances with total organic carbon (TOC) values. He found these parameters extremely valuable as they allowed a much better appreciation of real stratigraphic and regional variations in terrestrial carbon flux and marine organic matter preservation which could significantly clarify palynofacies interpretations and palaeogeographic reconstructions. In this work, absolute abundances per unit rock weight of palynofacies particles were estimated. The particular value of using absolute particle abundances is that any particle type may be studied in isolation from the complete suite, or compared between samples. This avoids the data closure effects and the ambiguity inherent in using percentage data and helps to enable any variations in the different provenance areas of the particle categories to be investigated separately. In addition, further or more detailed counts may be made at a later date and added to the original count without upsetting the results. In this study, percentage data are not used but ratios between selected pairs of particle types are presented. The methods used to calculate the absolute abundances are explained below. The term 'absolute'
PALYNOFACIES OF KIMMERIDGIAN CYCLES abundance, as employed here, refers to the abundance of organic particles per unit weight of rock. It should be noted that this will not necessarily equate to absolute abundance per unit weight of rock per unit of time, which is the correct usage of the term absolute abundance, as employed by Quaternary workers. This is generally not possible, however, to determine in ancient sediments such as those of this work. It must be remembered, therefore, when interpreting the absolute abundance data, that variations may be due to the effects of mineral dilution as well as absolute organic particle abundance variations. This is why the employment of ratios in addition can be useful. This is discussed further in the interpretations.
Sampling The 12 m section of the Kimmeridge Clay was initially sampled as a continuous 130 cm block through one lithological cycle in order to choose the best sampling interval. This block was sampled first at 5 cm intervals. Subsequent analysis of several complete cycles at both 5 and 10 cm intervals indicated that the resolution of the results was generally not altered by spacing samples at intervals of 10 cm. However, a sample spacing of 20 cm intervals did appear to alter the accuracy of the results. Therefore, 10 cm was chosen as the standard sample interval. About 15 g of rock was normally collected for each sample and consisted of no more than 1 cm thickness of rock, in order not to destroy the fine resolution of the sampling.
Processing Samples of c. 2-7 g were processed using standard palynological processing techniques (e.g. Traverse 1988). Table 2 gives details of the techniques used in this study. Additional points to note are as follows. (1) No oxidation was carried out during processing as all the particles in a sample were of interest in this study and oxidation may selectively break down some particles. (2) It is important to note that AOM was not included in the palynofacies counts. AOM was often so abundant that it obscured many of the other particles on a slide. In addition, it was not particulate in the same sense as much of the other organic matter and was therefore more difficult to quantify. For these two reasons it was not practical to attempt to include AOM in the normal particle counts. It was therefore removed from the samples using an ultrasonic probe followed by sieving. Estimates of the AOM content can, however, be inferred from the TOC data. (3) The final organic residue was mounted on
79
Table 2. Palynofaciesprocessing techniques (see textfor further explanation) Procedure 1. Wash 2. Dry 3. Crush 4. 5. 6. 7. 8. 9.
Weigh out 2-7 g Dissolve in HC1 Rinse 2-3 times Dissolve in HF Rinse twice Add Lycopodium spike 10. Wash until neutral 11. Sieve at 20 lain 12. Boil in HC1 13. Sieve at 20 lam 14. Examine 15. Ultrasonic (if necessary) 16. Sieve at 20 pm 17. Mount in Elvacite 2044
Purpose To remove surface contamination To increase surface area for action of acids To remove carbonates To remove silicates For calculation of absolute particle abundances To remove any net-formed fluorides To remove AOM
glass microscope slides as strew mounts, ensuring that samples were well mixed prior to pipetting of the portion to be analysed. The residues were mounted in Elvacite 2044 under rectangular glass coverslips. Some samples were also mounted in Metset resin and polished in order that they could be examined under both transmitted and reflected light (Hillier & Marshall 1988). (4) A number of samples were also subjected to elemental analysis. This provides measures of the percentage TOC content of a rock. Whole rock samples through two sedimentary cycles (WL3 and WL4) were powdered and analysed in an elemental analyser twice, once before and once after removal of CaCO 3 by HC1 digestion. This provided two measures of carbon content; % total inorganic and % TOC contents. From these measures the % CaCO 3 content may be calculated, the remaining percentage also providing an estimate of the combined clay and quartz portion of a rock together with any other remaining minerals, such as pyrite. L y c o p o d i u m spiking All samples were weighed as a first step in processing and a known number of exotic spores, of the genus Lycopodium, were introduced to the samples during the processing. This technique, of spiking samples with foreign spores, allows the
80
H. K. WATERHOUSE
calculation of absolute abundances of particles, per unit weight of rock, in a sample and was first used by Quaternary palynologists (Davis 1967; Davis et al. 1973), The Lycopodium spores can be readily identified in a sample by their morphology and colour and were counted simultaneously to, but separately from, the palynofacies particle count. The absolute abundances mg -1 of each particle category can be determined from these counts using a few simple equations. The calculations are based on the knowledge that the number of Lycopodium counted is in the same proportion to the total number of Lycopodium added as the (rock) weight of the sample portion counted is to the weight of the total rock sample processed. Thus:
Lycopodium counted Lycopodium added
weight of sample counted weight of total sample (1)
Therefore, since we need to know the weight of the portion of the sample that was counted, we can derive the following equation from Equation 1: weight of sample counted (mg) =
Lycopodium counted Lycopodium added x total sample weight (mg).
(2)
Once the weight of the sample portion analysed is known, the number of particles in any category can be calculated using the following equation: particle count particles per m g = weight of sample analysed (mg) (3) As the weights are in mg (a convenient form of the more standard numbers g-1 _ does not imply greater accuracy), this equation provides the absolute number of particles in any particular category per mg of original rock. The absolute particle abundances in any particle group may therefore be compared between and within samples. Lycopodium spores may be obtained in the form of tablets (available from the Laboratory of Quaternary Biology, Tornavagen 13, S-223 63 Lund, Sweden), each containing a known number of Lycopodium spores held together by sodium bicarbonate and water soluble organics, which are washed away by the water and HC1 during subsequent processing. The use of tablets, a method developed by Stockmarr (1971), provides an accurate and uncomplicated method of introducing exotic spores into a sample. To avoid bias between different batches of samples, the tablets were
always added after the second rinse subsequent to HF digestion. There were a number of other reasons for this choice of timing. The organic matter is not completely released from the mineral matrix of the rock sample until after HF digestion. The introduced and indigenous microscopic organic particles must behave in as similar a manner as possible during the processing in order to avoid bias in the absolute particle abundances. The Lycopodium spores should therefore not be added until all the organic matter has been released from the rock matrix. Great care must, however, always be taken when decanting samples prior to the addition of tablets in order not to lose any of the sample, which might bias the results by causing the particle abundances contained in the sample to be under-represented relative to the Lycopodium spores. Excess sample may be poured off subsequent to the addition of Lycopodium spores providing it is well mixed first so that particles are poured off in the same proportions as Lycopodium spores. Finally, to add the tablets after the HF has been partially diluted is safer for the operator. The number of Lycopodium spores added to a sample varied because the number of tablets added was not always constant. Initially, seven tablets were added to each sample. However, due to the large number of samples to be processed and a halt in production of Lycopodium tablets at their source, it was necessary to economize on the use of the tablets. The amount of Lycopodium required per unit of rock for ease of identification of Lycopodium during counting varied according to the concentration of organic matter in the rock. In organic-rich sediments, one tablet per g of rock usually provided a good enough ratio of introduced spores to organic residue to allow both particle types to be easily counted on a slide. Other, less organic-rich, rock types, such as limestone, needed fewer but more rock needed to be processed. Therefore, as a general rule, one tablet was added per g of rock. The significance of the number of Lycopodium spores added was tested. This could only be done, however, after processing by calculating the standard deviation on the number of Lycopodium counted as a percentage of the total particle count in each sample. The likely errors caused by too few Lycopodium being added to produce a significant ratio for calculation of absolute abundances of particles could then be calculated. This would allow the calculated absolute abundances in adjacent samples to be identified as being either statistically separable, i.e. any variation from one sample to the next is a real variation, or not statistically separable, i.e. variations seen could be due to statistical error or chance and therefore not reliable enough to use in the palaeoenvironmental interpretations.
PALYNOFACIES OF KIMMERIDGIAN CYCLES For the purposes of interpretation, the statistical separability of adjacent samples was taken as the main test of the significance of the count results in the less abundant groups as well as of the Lycopodium in all groups, as far as the error on the number of Lycopodium added and counted was concerned. It was found that for some of the initial counts of 300 particles, the Lycopodium spore percentage in the total count occasionally did not allow adjacent samples to be statistically separated. It was decided, therefore, to undertake a third count, of 1000 particles, on the samples in the hope of ensuring that the most statistically significant results that were realistically obtainable would be achieved. Further counts of 1000 particles were therefore undertaken, dividing particles into Lycopodium spores, spores, bisaccate pollen and 'others'. The 'others' group contained all of the particles counted in the original count except the spores, pollen and Lycopodium. It was considered unnecessary to distinguish these groups from one another as it would have made the process very time consuming and, in the first count, those groups which were subsequently used for interpretation were all well represented through the section. In other words, they showed large enough percentage values in a sample to be considered significant. In addition, all of their major variations in abundance through a section (i.e. those used in the interpretations) were found to be sound when tested for statistical separability of samples according to their count values as a percentage of the total count. The original count results for these groups were therefore considered reliable for subsequent use subject to testing of the Lycopodiumratios. The actual percentage of any group could not necessarily be increased by counting a greater number of particles, but to obtain the same (low) percentage in a larger particle count gives a smaller standard deviation and therefore less likely error. This was particularly important for the particle groups which occurred in low abundances - the spores and bisaccate pollen. These groups, despite their low abundances, seemed to show interesting patterns of variation which, it was hoped, could be used for interpretation. By increasing the total particle count, therefore, the significance of their values would increase even if their actual percentage values in a count did not. For example, in a count of 1000 particles, 1% is statistically distinguishable from 2% but this is not the case for smaller total particle counts. The graphs displaying the absolute abundances through the section (except in Fig. 2) show clearly which pairs of samples cannot be statistically distinguished from their neighbours and whose variation from each other was therefore not used when interpretations
81
were being made. For the ratio graphs, values were considered significant if their combined total was more than 50. Samples whose ratios were not considered to be completely statistically significant are indicated on the diagrams.
Counting. Counts were undertaken on the strew mounts of the samples in transmitted white light using an Olympus BHT 112 microscope and a mechanical stage controlled by an electronic point counter. Usually, 300 particles were identified and counted per sample, excluding Lycopodium spores which were counted simultaneously but separately, and AOM. For each field of view only one particle, the most central, was counted. Lycopodium spores were counted in exactly the same way as all other particles, that is, if a Lycopodium spore was the most central particle it would be the one counted, but the count of Lycopodium would be registered separately from the others. This technique was used for all three particle counts undertaken (the reasons for undertaking each count are explained elsewhere) - the initial (first) count on all particle groups, the second count on the woody particle groups, and a third count of 1000 particles which was undertaken for the purposes of statistical testing. However, in the third count, Lycopodium were not registered separately because with such a large total they would not prevent a large enough count of the other less abundant particles from being achieved, as they might in a smaller total particle count. They were, however, separated afterwards for all of the subsequent calculations. With hindsight it was realized that an easier method of ensuring the significance of results would have been to set the first total particle count to a value such that a pre-set number (that was considered to be significant) of Lycopodium spores, and/or of the least abundant particle group of interest, was reached in each count. The new ratios of Lycopodium counted to total Lycopodium added, calculated using the third particle count results, were used to recalculate the original (first) absolute abundance values for all particle groups, except spores and bisaccate pollen which were recounted anyway in the third count. This was done using the same calculations as for the first count, and the particle count data for most particle groups from the first count, but included an adjustment of the particle count data to correspond to the new (more significant) Lycopodium count values. This would 'boost' up the significance of the original counts (on 300 particles) as far as the role of Lycopodium in the calculations was concerned. This would add further to the statistical significances of the counts of these particles. When the counts for these groups were recalculated using the new Lycopodium ratios the overall patterns of
82
H. K. WATERHOUSE Marine palynomorphs
Terrestrial palynomorphs
wLs
Terrestrial debris
WL4
WL3
WL2
WL1
WLSB
F11 135 FIO lm
m 140
F9
Dmudst] shale
DIst
I I
particles per mg rock
Fig. 2. Absolute palynofacies particle abundance variation in the three main particle groups through the Kimmeridge Bay section. Bed numbers are the same as in Fig. 1. Extra values on bars indicate abundance values exceeding the scale of the graphs.
variation through a cycle varied very little from those displayed by the results of the original count. All figures, except Fig. 2, display the new absolute abundances, calculated using the recounted Lycopodium ratios, and those which are still not statistically separable from their neighbours are clearly marked. In addition, the spores and pollen graphs on these figures were also plotted using the recounted (third) spores and pollen values. A much smaller number of qualitative analyses were also undertaken, using polished thin sections under oil immersion, with a Zeiss UMSP 50 microscope in reflected light mode in order to investigate the different types of wood particles. Results and discussion of these analyses are given elsewhere in the text.
Sources of error in processing
Number of particles counted The number of particles counted per palynofacies sample varies greatly between workers from 300 (Nehr-Hansen 1989), to at least 500 (Tyson 1984) to over 1000 (Lorente 1990). The particle counts of 300 used in this study are at the lower end of
this spectrum of examples. However, tests indicated that a count of 300 was sufficient to ensure accuracy. Rarefaction curves were plotted for some particle count results, with count values registered at random points through the count. On most graphs the count lines became horizontal (additional particle counts did not alter the results) after only c. 150 particles had been counted and, after 250 particles, almost all counts had stabilized. The wood particles in many samples were also recounted as described above. This second count provided larger particle counts per sample and therefore more reliable results for those particle categories recounted. The final results of the second count were, however, not used in the palaeoenvironmental interpretations as they did not provide any additional information. In addition, the other particles were also recounted into two groups; marine palynomorphs and terrestrial palynomorphs in order to check the counting process by comparing results with the first count. These suggested that a sufficient number of particles were counted to obtain reliable count results and that the methods used to undertake the particle counts were efficient. These two tests
PALYNOFACIES OF KIMMERIDGIAN CYCLES add evidence that the number of particles counted was sufficient to ensure statistical reliability, in addition to the tests carried out subsequent to the third particle count, as described above.
The L y c o p o d i u m tablets Two batches of Lycopodium tablets were used in this study. The first batch were used in the processing of all samples except those prefixed with a G (see Fig. 1). The second batch were used for the G samples. The first batch of tablets were calibrated at the laboratory in Lund, Sweden, using the following procedure. One hundred samples of five tablets, each taken from different places in the batch, were prepared by dissolving tablets in 0.9% NaC1 solution in 100 ml flasks. Twenty counts of 0.5 ml were made on each sample with an electronic coulter counter. The results given were: x = 67 498, s = + 1543, v = + 2.3%, therefore, for one tablet, x = 13 500 Lycopodium spores. The second batch of tablets were calibrated in the Department of Geography at the University of Southampton. Four samples of five tablets were added to 200 ml of Isoton II, particle-free 1% NaC1 solution. Fifteen samples of 2 ml were taken from each and the number of particles counted with an electronic coulter counter. A background check of the electrolyte was made before the Lycopodium spores were added. The results were as follows:
Sample 1
Sample 2
Sample Sample 3 4
Average of 15 counts of 2 ml 935 925 Background count 30
33
924
910
31
26
therefore,
Lycopodium per tablet = average count per 2 ml
• 100.
83
It was noted, however, that these counts were only carried out on four samples. Additional attempts to verify this calibration were unsuccessful due to problems with equipment. Since the detailed interpretations of the results were carried out using only the data from the uppermost three lithological cycles in the section, and therefore the first batch of Lycopodium tablets, the likely error due to variation in Lycopodium spores per tablet can be taken as 2.3%. The implications of this when combined with other errors are discussed below.
Effects of lithologicaI variation During the initial stages of processing, a slight difference in the weights of different lithologies was noted which, it was thought, may possibly have an effect on the absolute particle count results which were calculated using rock weights. The weights and volumes of a series of samples from two cycles were therefore measured and the weight-to-volume ratios calculated. The results indicated that, although there was a slight difference in the ratios between the two lithologies, the mean difference was only 3.25%. Compared to other possible sources of error, this was considered to be of little importance, especially as it would only affect adjacent samples between which there is a lithology change. It is, however, still considered when making interpretations.
Combined errors A calculation of the combined errors was not made as different errors do not affect all samples equally. For example, lithological variation will not affect samples located within a single lithological bed, and errors on the number of Lycopodium spores per tablet vary within the section depending on which batch of tablets was used. Therefore, any errors which were considered to have a possible effect on the palaeoenvironmental interpretations are clearly labelled on all of the graphs used for these interpretations. The division of the typical palynofacies cycle for interpretation (see Fig. 8) is based only on palynofacies variations and characteristics which have been proved to be statistically significant.
number of tablets So, the number of Lycopodium per tablet for each sample are as follows:
N = 18 470
Sample 1
Sample Sample Sample 2 3 4
18 700
18 500
s = + 207
18 480 v = + 1.12%
18 200
Time series analysis In this work the time series, a collection of observations made sequentially in time (Preistley 1981; Chatfield 1980), consists of equally, vertically, spaced observations about the organic content of sediments. The imperfect calibration of time in the rock record produces 'time series' with measurements obtained at inconstant intervals in time. This
84
H. K. WATERHOUSE
w167
~t".-.-.....,tt'3
er5
~, t..r
~
.........
2
m
~d
w
~176
PALYNOFACIES OF KIMMERIDGIAN CYCLES
ow167
oo
o
~~~~~ooOOoooo
~8
9
O
i
E "6 ~ - ~ ~
~
~
~
~
~
-~ ~ ~ ' . ~ ~= .~ ~ . ~
-= ~ o-,.., ~ >..,
85
86
H. K. WATERHOUSE
therefore demands that great caution be exercised when using time series techniques to analyse geological sequences, as thickness measurements may not necessarily be equated with time. Power Spectral Analysis can be used to detect regular cycles within a time series. It can detect and separate a number of different frequencies, or regular cycles, in a time series consisting of several cycles with different frequencies, amplitudes and starting points superimposed upon one another, as orbitally forced cycles often are. The results of the power spectral analyses of the palynofacies results are displayed in Figs 5 & 6.
Results The raw palynofacies count data is given in Table 3. Absolute palynofacies particle abundances are displayed, in Fig. 2, as bar charts of the three main particle categories (marine palynomorphs, terrestrial palynomorphs and terrestrial debris) for the complete sampled section at Kimmeridge Bay. In addition, the individual particle categories are displayed as barcharts, in Fig. 3, for the top three sedimentary cycles only. The ratios are not
chorate
proximate prasino saccate pollen
spore5
dinocysts dinocysts phy'tes
brown wood
cuticle
displayed in full - only those which were considered to contain useful palaeoenvironmental information are shown. The results of time series analyses are presented as periodograms in Figs 5 & 6. It should be noted that none of the time series analyses were carded out on smoothed data sets. Smoothing the time series would eliminate high frequency cycles, or 'noise', and perhaps produce clearer periodograms, but precision may be lost. It was possible that, in this high-resolution study, organic particle cycles of quite short duration may have been present and these would be lost if the data were smoothed. The bar charts show a cyclic pattern of variation in the absolute abundances of all the organic particle groups. In general, there is one peak, sometimes a double peak, and one trough in the abundance graphs per lithological cycle, the peak occurring within the bituminous-rich shale bed and the trough spanning the remainder of the cycle. The cycles in absolute organic particle abundance therefore appear to have the same duration as the lithological cycles, and to be more or less in phase with them. The periodograms support this observation. Towards the base of the lithological section the
eq. r'ded eq. shp black wd black wd
hii
.=
i
N 9
iN Im
m mm I
9
ir
9
I
I
9 I
9
9
m
,
|
I
Hk= . -IDI,'-!
,,m, t
aN
"" i I I
I
r
m
9 B
m m
i m p
R m i m m
i
TOC 9 I
9 |
i
m
I I
BB ira=
g l Lm
!ii!i !!!!
,
degraded total p'morphs particles
i
I !
black laths
CaC03 clay. qz etc.
r
BB m m Bi m /i | i i l i I
I i
L particles per mg
% of rock wt adiacenl samples not significantly separable due to Lycopodium variation p
adjacent samples not significantly separable due to low count values of these categories in the total count (see text)
Fig. 3. Variation in absolute palynofacies particle abundances in individual palynofacies categories, and of mineral component percentages of the rock, through the topmost three cycles of the Kimmeridge Bay section. Bed numbers follow those of Fig. 1. Where the statistical significance of variations between adjacent palynofacies samples is not sufficient to separate them this is indicated on the graphs.
87
PALYNOFACIES OF KIMMERIDGIAN CYCLES cyclic pattern becomes less distinct, with increased variation between peaks. In addition, there is a trend towards higher absolute particle abundances in the terrestrially derived groups in the lower part of the section. In Figs 3 & 4 the variations within individual cycles can be seen more clearly. The chorate and proximate dinoflagellate cysts display very similar patterns; a mostly smooth cyclicity corresponding to the lithological cyclicity, with occasional unusually high abundances in places. Terrestrial palynomorphs tend to show more abrupt changes. The bisaccate pollen in particular show this very clearly, with a strong peak on the periodogram indicating a wavelength of 122 cm for the cyclicity. The spores have a very similar distribution to that
marine p'morphs :
terrestrial debris
dinocysts :
prasinophytes
of the bisaccate pollen. Both pollen and spores display very similar abundance patterns and display a cyclicity more or less in phase with that of the dinocysts, but perhaps with the peak in abundance occurring slightly higher in the bituminous shale. Cuticle shows a rather different cyclicity to that of the marine and terrestrial palynomorphs. In contrast to other groups it has low, or decreasing, abundances during the bituminous-rich shale beds. In addition, it appears to display two peaks per lithological cycle. The first peak occurs immediately before, or at, the base of the bituminous-rich shale, and the second about half way through the lithological cycle, immediately above the bituminous-rich shale. The peaks vary slightly in size but there does not appear to be one
.
f
spores and r'ded:sharp black laths : pollen: equid, equid, terrest, debris black wood black wood
spores :
saccate pollen
.
.
.
.
.
.
.
.
i~.....'
pollen and spores :
dinocysts
~.-.
--ii..........
m
!:
I
I
.... i::
I
h=e I
.
-
sample in which the total value of all the constituents of the ratio is less than 50. the ratio thereby possibly lacking significance at that point
Fig. 4. Variation in the ratios between selected palynofacies particle groups through the topmost three cycles of the sampled section at Kimmeridge Bay. Bed numbers follow those of Fig. 1. The ratios are plotted on a logarithmic scale and adjacent samples which were not found to be statistically separable are indicated.
~
H. K. WATERHOUSE
dominant peak within each cycle. This is confirmed by examination of the periodogram which has its dominant peaks at 76 and 60 cm, which represent a cycle, or cycles, of approximately half the duration of the 120 cm cycle of other particle groups. The equidimensional sharp and the lath-shaped black wood, the brown wood, the degraded palynomorphs and the total particle abundances all display a clear cyclicity in phase with the lithological cyclicity and with that of the terrestrial palynomorphs. The periodograms record strong cycles at c. 119 cm. The equidimensional rounded black wood, like the cuticle, has two peaks in abundance per lithological cycle. The periodogram records a cycle at 64 cm but nothing near 120 cm, indicating that both of the two peaks within each cycle are equally important. Many of the ratios in Fig. 4 display cyclic variation with the most prominent cyclicity being at c. 120 cm. These groups do vary from one another, however, in their types of distribution. Several groups also have a clear cyclicity but are not in phase with the lithological cyclicity or the main palynofacies cyclicity. Figure 3 contains the results of the elemental analysis evaluations for the WL3 and WL4 cycles. It can be seen that the % TOC displays a cyclic distribution, with the pattern being very similar in both cycles. It is low in the mudstone, increasing gradually through the bituminous shale above to a definite peak at the top of it. It then decreases rapidly at the base of the mudstone above. The percentage of clay, quartz and other minerals displays a variation similar to that of the % CaCO 3, both being almost exactly the opposite to that of the % TOC.
Discussion Explanations for the deposition of such large amounts of organic matter as those found in the Kimmeridge Clay Formation have been advanced by a number of workers, often using analogies with Recent environments such as the Black Sea. There have been two main schools of thought, one favouring a mainly preservational control on organic matter accumulation and the other favouring productivity. Calvert & Pedersen (1990) have suggested that the settling flux of organic carbon, closely linked to the rate of primary production, was the main control on accumulation of organic matter irrespective of bottom-water oxygen values. Pedersen & Calvert (1990) noted that sediments accumulating in the modem Black Sea are not particularly enriched in organic matter despite the presence of an anoxic water column. They suggested that oceanic anoxic events, such as in the Cretaceous Atlantic for example, were brought about by sluggish circulation and that increases
in primary production, reflecting changes in the ocean-atmosphere system, constitute a more tenable explanation for the occurrence of modem and Quaternary carbon-rich sediments and Cretaceous black shales. Lallier-Verg~s et al. (1993) proposed that tile organic sediments of the Kimmeridge Clay Formation were deposited under an oxygenated water column with a redox boundary that oscillated above and below the water-sediment interface, mainly due to variations in the organic flux or productivity. Huc et al. (1992), in a study of the Dorset Kimmeridge Clay based on petrography and geochemistry of organic matter, proposed that variations in the amount and nature of organic matter were mainly controlled by changes in biomass. They also showed that changes in depositional conditions occurred together with changing productivity. These changes were mainly reflected in the modification of burrowing features. Tribovillard et al. (1994) favour phytoplankton productivity as the main driving force of variations in organic matter concentration, with redox conditions of the depositional environment acting as a positive feedback effect. Gallois (1976) invoked a delicate balance between palaeogeography, a series of subsiding basins linked by relatively narrow straits between land areas of low relief, and nutrient supply as particularly favourable to the formation of algal blooms. Another theory suggests that phytoplankton blooms were a symptom, rather than a cause, of widespread anaerobic bottom conditions and that preservational factors rather than productivity were the major control (Tyson et al. 1979). Tyson (1987) proposed a widespread depletion of oxygen with deposition of organic-matter-rich mudstone cycles due to limited, probably silled, connections to the open ocean, and high biological productivity. Wignall (1989) and Myers &Wignall (1987) suggested, in this context, that a climate-induced temperature stratification (thermocline) was the predominant control on the different lithologies. This situation may have been storm-limited, with water depth being the principal control on the distribution of bituminous shales (Wignall 1989). Many other authors invoke water stratification as a characteristic parameter of the depositional conditions of the Kimmeridge Clay Formation. Miller (1988) suggested that stratification was due to salinity variations. Pratt (1984), studying the Cretaceous Greenhom Formation of Colorado, found that sedimentary and organic geochemical data indicated a close association between palaeoclimate, salinity of surface water, strength of bottom-water currents, and amount and composition of organic-matter preservation in the sediment. She suggested that
PALYNOFACIES OF KIMMERIDGIANCYCLES palaeoclimatic and palaeoceanographic factors that influenced mixing and current strength in the water column profoundly affected the amount and type of organic matter preserved. She proposed that dry periods were characterized by a well-mixed water column and low preservation of organic matter, and wet periods with high river discharge led to salinity stratification of the water column and quiescent oxygen depleted bottom water and high preservation of organic matter. Bottom-water oxygenation was therefore primarily controlled by the strength and frequency of benthic currents rather than by the rate of oxygen consumption in the benthic environment. Oschmann (1988) proposed the North Atlantic Water Passage model, which is a seasonallycontrolled modified upwelling model with temperature stratification. His interpretation was not related to any cycles of possible orbital origin but referred to shorter-scale variations and the overall palaeoecological situation. Cyclic variation in the lithology of the Kimmeridge Clay Formation has often been suggested as being due to vertical movement of the oxic-anoxic interface (Irwin 1979; Tyson et al. 1979; Myers & Wignall 1987). Scotchman (1991), using Tyson's (1989) method of calculating kerogen abundances, found that the second highest phyTOC values (an estimate of the absolute abundance of terrestrially derived structured organic debris) from a number of Kimmeridge Clay locations were recorded at Kimmeridge Bay in Dorset. This suggested to him that increased terrigenous kerogen input was associated with high sedimentation rates. He also found that proximal shelf samples had the lowest phyTOC values, and suggested a sedimentation rate control on kerogen facies in the Kimmeridge Clay. In addition, he suggested for the Kimmeridge Bay location that dilution of a relatively constant flux of amorphous marine organic matter by variable amounts of structured terrestrial-derived kerogen occurred towards basin centres, where higher sedimentation rates increased the relative proportions of the terrestrial components. He also demonstrated that both sedimentation rate and differential organic-matter preservation were major controls on the kerogen facies of the Kimmeridge Clay Formation. Myers &Wignall (1987) have suggested that the sedimentation rate exerts a control on the organic matter and carbonate content of the mudstones of the Kimmeridge Clay Formation. Downie (1957), studying the Kimmeridge Clay of Dorset, found a close correlation between the quantity of kerogen and the proportion of terrestrial palynomorphs. This suggested that bulk kerogen is also terrigenous and he proposed that the Kimmeridge Oil Shale was formed by an increase
89
in swamp vegetation and therefore an increase in transported swamp-generated organic matter. Tribovillard et al. (1994) have stated that dilution effects by inorganic components (detrital clay and quartz and biogenic coccolith carbonate) of the sediment cannot account for the TOC cyclicity in the Kimmeridge Clay in Yorkshire. They studied the clay minerals in two cycles and obtained results which indicated a homogeneous terrigenous detrital supply. No qualitative variation was detected in the clay assemblage and minor quantitative variations were seen within and between the cycles. Muller (1959) noted that there was apparently no direct relationship between the rate of sedimentation and the total pollen content of the modern sediments of the Orinoco Delta, which was not surprising as the sources and supply conditions were different. He stated that climatic changes may be expressed by the bulk changes in composition of vegetation and the amount of river discharge. Both Tribovillard et al. (1994) and Herbin et al. (1991) report that the main source of carbonates in the cycles of the Yorkshire Kimmeridge Clay was coccoliths. NChr-Hansen (1989) used visual and chemical kerogen analysis to investigate the Lower Kimmeridge Clay. He recognized three distinct kerogen facies (palynofacies); one dominated by AOM, a second consisting mainly of dinocysts, acritarchs, spores, pollen, cuticles and foraminiferal test linings, and a third facies containing predominantly woody particles. He interpreted these associations to be the result of vertical fluctuations in the oxic-anoxic boundary of the sediment-water interface and only to a lesser extent of variations in the organic input. This was in agreement with earlier models. Williams & Douglas (1983) conducted an examination of sequences of shale, clay and limestone from the Kimmeridge Clay of Dorset. They sampled from three points within the section used in the present study and, using a number of techniques, found no major qualitative differences in the organic-matter content of the sedimentary units. However, major variations in the quantity of organic matter were observed, shales containing greater proportions than either the clays or limestones. They suggested a combination of mineralogical dilution of kerogen and sedimentary preservation effects as the causes. Ebukanson & Kinghorn (1985) investigated the influence of environmental factors on the distribution of kerogen types contained in the different lithologies of the Kimmeridge Clay of southern England. They found no simple linear relationship between kerogen types and the associated lithologies. However, there was a relationship between kerogen type and organic richness of rock samples. The laminated mudstones were found to
90
I-I. K. WATERHOUSE
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12 8 4
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0 0.2 cycles per sample interval
0.3
PALYNOFACIES OF KIMMERIDGIAN CYCLES be very organic rich and were thought to represent oxygen-deficient bottom conditions with low benthic activity. The non-laminated mudstones, marls and limestones were deduced to have been deposited in oxic environments and were generally poorer in organic content. They suggested that fluctuating oxic-anoxic boundary levels were a major factor controlling the distribution of kerogen facies, largely by influencing kerogen preservation. Bertrand & Laltier-Vergrs (1993) have studied one cycle from the Kimmeridge Clay of north England in high-resolution detail. They found variations in TOC of from 1.8 to 9.5 Wt%, which showed a pattern very similar to the TOC variations in the cycles of the present work (see Fig. 3). The peak in TOC was slightly higher in their cycle, possibly due to different cycle boundaries being taken, or slightly different sedimentation rates or compaction within cycles in the two areas. They also measured the sulphate reduction index (SRI), which represents the intensity of sulphate reduction relative to TOC. This estimates the fraction of organic carbon that was oxidized by sulphate reduction as a result of anaerobic bacterial degradation by sulphate reducers. The resulting reduced sulphur is often concentrated in organic-rich marine sediments and in the Kimmeridge Clay of northern England most of it is diagenetically trapped and can therefore be measured. Except at the beginning of the cycle, the SRI displayed a similar pattern to the TOC variation in their cycle. As the TOC variations could not be explained by variation in dilution by detrital (clay and quartz) and biogenic (coccolith produced) minerals, these patterns were interpreted as due to increased plankton productivity causing increased sedimentary accumulation of refractory organic matter accompanied by an increase in anaerobic degradation of metabolizable organic matter. The main method of accumulation was selective preservation of biologically-resistant organic matter from plankton. These interpretations were found to be in agreement with Calvert & Pedersen (1992) who proposed a dominant control
91
by primary production and a minor control by the redox conditions in deep water. Bertrand & LallierVergbs (1993) state that variations in redox conditions play a role only in changing the pathway of degradation of metabolizable organic carbon from aerobic to anaerobic. Lallier-Vergrs et al. (1993), studying the Kimmeridge Clay of Dorset, found that in samples with a TOC content of 2-10%, sulphate reduction (mainly expressed by presence of pyrite) consumed all available organic matter that could be metabolized. The samples were characterized by genetically variable organic matter, i.e. with different abilities to be metabolized. In samples with greater than 10% TOC the easily reducible iron was no longer available. Organic matter which could be metabolized had exhausted the available dissolved sulphate in the pore waters. Excess organic matter which could be metabolized was probably degraded by methanogenesis. These sediments corresponded to high monospecific phytoplankton input. Lallier-Vergrs et al. (1993) concluded that variability in intensity of sulphate reduction depended on the type of planktonic organic matter, its transportation to the sediment, and oxidation degradation during transport. The p a l y n o f a c i e s counts
The parameters used in this work in making palaeoenvironmental interpretations of the data collected, and their basic interpretations, are displayed in Table 4, with references. For clarity, the table contains interpretations only for parameters which were recognized as providing potentially useful palaeoenvironmental information for this work. Other parameters are not discussed, even if they were recorded during the counting, but selected references are given here. Further discussion of a few additional points now follows. Downie (1957) investigated the conditions of deposition of the Kimmeridge Clay and examined microplankton from Dorset. Gitmez & Sargeant
Fig. 5. Periodograms for the variations in palynofacies absolute particle abundances in the groups displayed in Fig. 3, but using count data from the complete (12 m) sampled section at Kimmeridge Bay. Power spectral analysis consists of the transformation of a time series into power spectra, and their subsequent interpretation. The values resulting from the transform operation are plotted as the squares of the amplitudes of each different frequency component, or cycle, to produce a periodogram. A component of greater amplitude is presumed to be more important in the time series, therefore its amplitude is squared to produce a more prominent peak on the periodogram. The height of each peak on the periodogram therefore represents the 'power' or importance of each frequency component, or cycle, in the series, while its width indicates the range of possible cycle lengths. Peaks to the right-hand side of the periodograms represent high frequency 'noise' and peaks to the far left, long-term trends. The power is plotted linearly against frequency, which is shown as the number of cycles per sample interval (10 cm). It should be noted that the values for cycle lengths given on the peaks on the periodograms represent centimetres of rock section, not years. Clear cycles at c. 120 cm are recognized in most groups but in some particle groups the shorter, 60-70 cm cycle is dominant.
92
H.K. WATERHOUSE saccate pollen:spores
20241I" 6
10
122!!
, spore s § pollen 9terrest, debris 125
8 6 4
2 0
0.1
0,2
0.3
0.4
0.5
o__ 0.1 0.8'
~
300
0.6
9
200
62
100[ / o~ I 0
n-~~ di 30
0.~
wo.o_ds
,
0.2
0.3
0.2 ..... 0.4
s :Prasin_.__oPbyt 9S
4 0,5
.
" "
125
25 20 15 10 5 0i
04
0.4
\/'"\"/~'~40 0.1
0.3
black laths i e quid; black
rounded:sharp_ equid black wood 400 ~ - T z b . . . . . . . . . . . .
~
02
0
0
0.1
0.2
0.3
0.4
0.5
marine p'morphs:terrest, debris
10 8 6
. . .38. . . . . . 0
0.1
0.2
0.3
0.4
0.5
0
0.1
0.2
0.3
0.4
.i ~J 0,5
cycles/sampling interval
Fig. 6. Periodograms showing cyclic variations through the top three cycles of the Kimmeridge Bay section in the ratios between selected particle categories displayed in Fig. 4. A clear cycle at c. 125 cm is displayed in all of the periodograms.
(1972) have given an extensive account of the microplankton assemblages of the Kimmeridge Clay of England, Scotland and France. Ioannides et al. (1976) studied microplankton of the Kimmeridge Clay of Dorset a few tens of metres higher in the section than the sediments of the present work. They found little change in the dinocyst assemblages with increasing organic content. This uniformity was found in ammonites and other nektonic and planktonic organisms. The benthonics did, however, show marked differences. The composition of dinocyst assemblages was found to be fairly constant, but the relative proportions of spores and pollen to dinocysts and acritarchs changed markedly. Wall et al. (1977) give an account of the environmental and climatic distribution of dinocysts in modern sediments. Acritarchs have been used as indicators of water turbulence and depth (Wall 1965). Dinocysts have also been used as indicators of salinity (Wall et al.
1977) and water temperature and, together with reworked palynomorphs, are of use in indicating proximity to shoreline (Williams & Sargeant 1967). Prasinophytes have also been used as indicators of low salinity (Prauss & Riegel 1989). The distribution of spores and pollen has been much studied from modern models (e.g. Muller 1959) and the information used to refer back to older sediments. A spore or pollen grain ready for fossilization is a more or less hollow, spherical body consisting of the sporopolleninous outer wall of the grain (Traverse 1988), but the variation in abundances at source, and morphology of different types, produces much variation in their subsequent dispersal and occurrence in sediments. Erdtman (1943) suggested that many pollen grains will float for a short period of time before settling. He stated that the settling velocity of empty pollen exines depends on weight (which will increase as they fill up with water) and on the size, shape and ornamen-
PALYNOFACIES OF KIMMERIDGIAN CYCLES tation. Muller (1959), in his study of Recent palynomorph distribution of the Orinoco Delta, noted a gradual increase offshore of total pollen grains down to 3000 grains per g, followed by a rather steep decline to 500 grains per gm (between 30 and 50 miles offshore) due to a deflection of the pollen carrying terrestrial output from the rivers by a marine current. In the present study c. 10 000 pollen and spore grains were counted g-1 of sediment in their peaks of abundance, with 1000 grains per g or less in the troughs. Total pollen content of modern sediments is influenced by marine currents, distance from shore, rate of sedimentation, coarseness of sediments, presence of carbonates and organic matter, and wind transportation. Muller (1959), in two of his cores, estimated that a minimum of 10% of pollen was wind blown and that for saccate pollen wind transportation is important, especially in certain seasons, with patterns of distribution being modified by sea currents. Melia (1984) noted that palynomorph distributions are closely related to both source vegetation and to atmospheric and oceanic transport mechanisms. Pollen and spores per g of bottom sediment ranged from > 2000 off the coast to < 50 in deep ocean basins. Distances of transport may have exceeded 5000 km. Groot (1966) noted that there appeared to be a relationship between the number of pollen grains and the quantity of mineral matter in suspension in the estuary of the Delaware river. He suggested that the two types of particles are transported together and in a similar fashion, i.e. primarily by water currents. Pollen grains per g of suspended sediment (clayey silts and silty clays) varied between 16 000 and 25 000. Traverse & Ginsburg (1966), studying the surface sediments of the Great Bahama Bank, stated that pine pollen (bisaccate) is large and buoyant and, once delivered to the water, remains suspended for long periods of time and is deposited as a rather sensitive indicator to hydrographic features, especially turbulence. On swamps and tidal fiats they found nearly 5000 pine pollen g-l, and in all other samples 0-558 per g. Hoffmeister (1954) reported 7500 grains of pollen per g in a sample and a ratio of 1 : 4 of large to small pollen, indicating proximity to a shoreline, in ancient basins of deposition, with regularly decreasing amounts of pollen per g and a decreasing ratio of large to small grains moving away from the shore. But, Traverse & Ginsburg (1966) noted an inverse relationship between pine pollen in the water and that in the sediment. This they explained as being due to sediment patterns not source. Pollen highs were regularly found where water turbulence was minimal. Cross et al. (1966), studying the source and distribution of palynomorphs in the gulf of California, found a few thousand to 80 000 spores
93
and pollen grains per g of dried bottom sediment; a few hundred to over 20000 pine pollen; 2000-25 000 cuticle fragments; up to 3 000 000 tracheids (5-50 lam) and up to 17 000 dinocysts per g of sediment Low concentrations of palynomorphs were found to occur along some shores, probably due to dilution by terrigenous sediments, but they built up offshore. Some palynomorphs decreased in number offshore while others (e.g. pine) became selectively increased in relative frequency. In this study no taxonomic determination of the spores and pollen was undertaken and, as a result, it would not be feasible to relate their abundances in the samples to changes in the source vegetation. This would be difficult anyway in such a distal setting due to differential sorting during transportation. Ratios such as the bisaccate pollen:spore ratio, the l i g h t : h e a v y spore ratio, and the terrestrial palynomorphs : wood ratio, in which one particle type is less buoyant than the other whilst having similar provenances, may all record variations in energy conditions of the transporting medium and, thus, variations in marine current circulation or perhaps run-off. The terrestrial debris found in palynological preparations is composed of broken-down parts of terrestrial plants. It may therefore provide a record of several aspects of the terrestrial environment and of the conditions of transportation of particles from there to the environment of deposition. Two characteristics of these particles were noted during analysis - the absolute abundances, and the condition of the particles as classified by size and shape. As the particles have different morphologies, they are likely to have different hydrodynamic properties and, therefore, different distributions. Various measures of environmental variation which can be inferred from these particles are listed in Table 4. The black laths are likely to have a greater buoyancy than other particles as they would have a lower mass to surface area ratio, and therefore be hydrodynamically equivalent to smaller solid particles. These particles would remain in suspension for relatively longer periods of time than other particles and therefore be an indicator of low energy conditions or a distal setting. This is the case as indicated in the literature (e.g. Tyson 1989; Cope 1981), In addition, most black laths were noted to be angular rather than rounded which is also suggestive of low energy conditions with very little abrasion occurring and/or that they consisted of a more resistant material (Cope 1981). Van Buchem et al. (1994) noted that opaque particles found in their samples can be rounded, blocky or splinter-like. SEM examination revealed to them that most of these particles showed anatomical details of the kind found in carbonized wood fragments of terrestrial plants (Cope 1981).
94
H. K. WATERHOUSE
~.~
~
"1~ ~'~ ~- ~
~
-
~
o
D'- ~
"~
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~
~
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~
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~=
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PALYNOFACIES OF KIMMERIDGIAN CYCLES
~
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95
96
I-I. K. WATERHOUSE
Van der Swan (1990) has noted that opaque particles are considered to be more buoyant than other particles. Cope (1981) noted that, although not common, the Kimmeridge Clay of Dorset contains occasional structured intertinite microclasts, indicating the presence of charcoal in this formation. Inertinite phytoclasts accumulate as small, splinter-like clasts, some showing rounding due to abrasion. Charcoal may survive unaltered in a variety of geological environments and the particles are of sufficiently low density that only very weak currents would be necessary to move them into deep-basinal sedimentary accumulations. A variation in the amounts of charcoal flux into the Pacific Ocean has been attributed to changing vegetation patterns with time (Herring 1977). Berlin & Brosse (1992), studying the Yorkshire Kimmeridge Clay Formation, noted that where TOC is < 4%, terrestrial organic debris consisted of irregularly and angularly shaped elements, and where the TOC was 4-8%, the angular particles decreased. It was suggested that when the energy was lower (and TOC subsequently greater) the particles spent more time in transit and therefore became more rounded, as mineral grains do. Woody material has been noted as being more resistant to abrasion and corrosion, especially when partly or wholly carbonized (Venkatachala 1981). Therefore, a relative increase in carbonized wood may indicate a longer time spent in transit and therefore less energetic transporting media. Wood is also more likely than cuticle to be reworked (Cope 1981). Muller (1959) noted a selective concentration of wood in shallow marine sediments when moving seawards, and Farr (1989) also noted a downstream increase in the proportion of woody fragments relative to terrestrial palynomorphs in estuarine samples, possibly because they are more resistant and can remain in suspension longer than most other types of plant debris.
Other types of variation Figure 7 gives a diagrammatic summary of types of variation which may cause palynofacies or sedimentary cyclicity. They may be divided into four types - biological (discussed above), climatic (types relevant to this study are discussed), physical (may essentially be discounted for this study), and chemical. The main type of chemical processes likely to cause variation are diagenetic. These have been the cause of controversy over the origin of some sedimentary cycles, although they are now generally only considered to be of importance as a secondary enhancer of already present cycles (Hallam 1986), and most controversy has centred on other stratigraphic tsections, particularly those alternating between limestone and shales.
Other types of variation which may be relevant are those which affect the palynofacies particle abundances on which the interpretations are made. Variation in sedimentation rate may dilute organic particles by increasing the inorganic output. However, this would affect different palynofacies groups in different ways and may therefore be recognizable. For example, if terrestrially derived inorganic particles (e.g. clay, quartz) increased the terrestrially derived organic particles would be expected to show similar variations (e.g. Scotchman 1991), while the marine ones would not, as they would be diluted by the terrestrial output. As mentioned above, Myers & Wignall (1987) suggested that sedimentation rate exerted a control on the organic matter content of the mudstones of the Kimmeridge Clay. Tribovillard et al. (1994) obtained results which indicated a homogeneous terrigenous detrital supply through two cycles of the Kimmeridge Clay in Yorkshire and stated that dilution effects by inorganic particles could not account for observed cyclicity in TOC. In contrast, if marine derived inorganic particles (e.g. CaCO 3) increased then this would dilute all organic particles to the same degree. However, the CaCO 3 content of the Kimmeridge Clay sediments studied in this work is low. As each sample consisted of about the same thickness of sediment, the amount of time represented by a sample, and therefore the amounts of organic particles deposited in that time, may vary if sediment accumulation rate or compaction varies through a cycle. It is thought that, even if sedimentation rates were different in the two lithologies, these factors are unlikely to have caused the more subtle variations in palynofacies characteristics which have been recognized within the apparently fairly uniform lithological beds. Bioturbation causes mixing of sediments and their contents and may, therefore, mix samples if they are situated closely together. The distance of mixing can vary greatly, depending mainly on the degree of benthic oxygenation. In well-bioturbated sediments palynofacies patterns present have been found to be smoothed, but not completely destroyed, by bioturbation (Waterhouse 1992). Hart (1987) suggested that intense bioturbation would make it difficult to produce a detailed set of faunal data from visible chalk cycles, as sampling at less than 50 cm would not provide a sufficiently restricted time-slice. Cottle (1990) had little doubt that bioturbation affected the nature of cyclicity recorded, but he proved that it is still possible to record cyclical changes in foraminiferal abundance at intervals of 10 cm and to demonstrate the existence of cycles of the order of 30--40 cm. In this work, however, low oxygen conditions were generally prevalent through much of the section
PALYNOFACIES OF KIMMERIDGIANCYCLES
BIOLOGICAL VARIATION
PHYSICAL VARIATION BENTHIC OXYGENATION ~pp( /~..~o..~..,~..! ~"
......9
PLANKTON ......../ " BLOOMS ;,\
97
=~./
PLATE MARGIN ACTIVITY / =oa~v=chan~ I
\=~.-
~,~r.::~
/
~
~bl~/
VOLCANICITY
~,,.~t m
TERRESTRIAL VEGETATION ~ _ ~ A ,.
........ '~
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DIAGENESIS
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i
TECTONISM
1> ..<
0u m
tet~r~re
\
CHEMICAL ~
j
EVAPORATION ~
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NUTRIENTS
\ "" SEA LEVEL
"2- . . . . . . . . . .
t~ature/
SEDIMENT ~ INPUT
|
|
CLIMATIC VARIATION
Fig. 7. Summary diagram of the possible causes of sedimentary and palynofacies cyclicity. These may be divided into four main types; biological, climatic, physical and chemical variations.
and bioturbation was recognizable only as a few burrows which were avoided during sampling. Benthic oxygen conditions are known to have an important influence on the accumulation of organic matter in sediments due to the variations in preservation potential at the site of deposition. In low-oxygen conditions all organic particles have a high preservation potential, although different types of particle can have different preservation potentials. Variation in benthic oxygen conditions may be interpreted from TOC, total particle abundances and abundances of degraded palynomorphs, as well as other measures not used in this work. A discussion of the productivity vs. preservation argument about organic matter accumulation
has been given above. Benthic oxygen conditions will therefore not be discussed further here.
Estimation of cycle durations Time series analysis (Chatfield 1980; Preistley 1981) was employed to estimate the durations of the palynofacies cycles. It was noted that the two main cyclic variations recognized in the palynofacies data were at c. 120 and 63 cm. The ratio between these two cycles is remarkably close to that between the obliquity and precessional orbital cycles (Berger et al. 1989b). It may, therefore, be reasonable to conclude that the effects of these two
98
H. K. WATERHOUSE
i
vegetation type precession == o
obliquity marine bottom currents and circulation
-
Q
benthic oxygen
~Ill
ph'plankton prod.
0
I11 |
O
|fll
"
1111 ~>
~Itl
~, g !
w
J}
clay, qz etc. i
%CaC03 Q
Total Organic Carbon (%rock wt) degraded palynomorphs
|
J;
total particles lithology black laths : equid, black wood , ~ ~ . -
........
....... .... .
~
.
,
!: .... .
rounded : sharp equid, black wood pollen + spores : wood 8
:~
spores : saccate pollen
pollen + spores : dinocysts dinocysts : prasinophytes
6
marine p'morphs : terrest, debris
o brown wood+ black laths
~m"'" '~
,:~
~
,
equid, sharp black wood "~ r equid, rounded " black wood .JD IB cuticle ,=
d
pollen and spores marine palynomorphs ;
segment
J
o
i
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E i
i
PALYNOFACIES OF KIMMERIDGIAN CYCLES cycles were the dominant influences on the palynofacies variations. It is advisable, however, to use more than one method to estimate cycle duration as all methods have flaws. Therefore, two methods of extrapolation were also used. For the first method the duration of the sampled section was estimated by the proportion of an ammonite zone it represents and was divided by the number of cycles contained in it to give an estimate of the average duration of the cycles (e.g. House 1985). In the second, the duration of the section was divided by its height in metres to give the average duration of a metre of sediment. The cycle lengths, in metres, were then converted into their duration in years. The autissiodorensis Zone at Kimmeridge Bay is known to be c. 60 m in thickness (Cox & Gallois 1981). If the common method of assigning a duration of 1 Ma to a zone (Harland et al. 1989) is used, the duration of the total 12 m section sampled represents c. 20% of the ammonite zone, and therefore c. 200 ka. The number of cycles (both lithological and palynofacies) in the section is either 8, 9 or 10, depending on whether the WL3 shale band is taken to represent one or two cycles (the lithology suggests one but the playnofacies results suggest two, due to the two peaks in abundance during the cycle; see Figs 1 & 2) and whether the shale band within the Washing Ledge Stone Band represents a distinct cycle. All three values were therefore used and the results obtained were 25, 22.2 and 20 ka, respectively, per cycle, the mean of the three values being 22.4 ka. Using the data published by House (1985) and Cox & Gallois (1981), an average duration of 22.2 ka is obtained for each of the 45 lithological cycles in the whole autissiodorensis Zone. This is almost identical to the estimate obtained for the palynofacies cycle durations of this study by this method. If the assumed duration of 1 Ma for the autissio-
99
dorensis Zone is divided by 60 (the thickness of the zone in metres), a duration of 16.667 ka is obtained for the deposition of each metre of sediment. The time series analysis results and the absolute abundances for the palynofacies cycles picked out the main cycles at Kimmeridge Bay to be at c. 125 and 63 cm. Durations of 20.8 and 10.5 ka may therefore be obtained for the durations of the 125 and 63 cm palynofacies cycles, respectively. Therefore, using both extrapolation methods, the estimates of the durations of the main Kimmeridge Bay cycles are: lithological cycles (using House's 1985 data) = 22.222 ka; lithological and palynofacies cycles of this study = 22.407 ka; palynofacies cycles = 10.500 ka (63 cm cycle), 20.834 ka (125 cm cycle). A number of assumptions are necessarily made when using estimates such as these. Varying sedimentation rates, differential compaction and hiatuses in any part of a cycle, sampled section or ammonite zone, may introduce bias into the results (Weedon 1989; Schwarzacher 1989). Sub-dividing the existing radiometric timescale into zones or sub-zones of equal duration, and assuming that the evolutionary changes used to identify ammonite zonations proceed at a uniform rate, are both unwise practices (House 1985). The results of such extrapolation methods, which necessarily make many assumptions, should not, therefore, be used in isolation to make interpretations. Of the extrapolation results, it would be expected that those of House (1985) would be the most accurate since they were based on the thickest lithological sections. His results clearly suggested the obliquity cycle as the main controlling orbital cycle. In addition, Weedon (1986) concludes that the obliquity cycle was the most likely control on the longest sedimentary cycles of his study. He also recognized a shorter cycle using time series
Fig. 8. Summary of the palaeoenvironmental variations through a typical palynofacies cycle. All of the most important parameters through one 'typical' palynofacies cycle are represented. At each of the sample positions in the cycle the value represented is the mean value, for that position in the cycle, of the values for the three cycles displayed in Fig. 3 (the WL3, WL4 and WL5 cycles). The main particle groups, some ratios and the mineral component percentages are displayed. The cycle can be divided, using only the palynofacies characteristics, into five segments or subsegments (whose characteristics are more similar to each other than to other segments). The most important palaeoenvironmental interpretations are represented diagrammatically in terms of phytoplankton productivity (more dinocysts indicate higher productivity); benthic oxygenation (increased oxygen content of sediments indicated by wavy lines, decrease indicated by straighter lines); ocean current intensity (increase in intensity indicated by solid lines and single circles, decreased intensity indicated by broken lines, and very low intensity, with Kimmefidge Bay area cut off from neighbouring areas, represented by two separate circles); variation in the obliquity of the ecliptic - only estimated maximum and minimum points shown; variation in the precession of the equinoxes - only estimated maximum and minimum points of the cycle shown; variation between two vegetation types (e.g. scrub- and forest-type vegetation). All variations between adjacent samples which may be statistically due to variations other than palaeoenvironmental effects are indicated. Where the value in the box is 1 this variation may be considered as probably statistically significant. Where the value is 2 or 3 the variations should not be used for palaeoenvironmental interpretation. Division of the cycle into segments of different palaeoenvironmental characteristics was not based on any variation which may possibly have been insignificant.
100
H.K. WATERHOUSE
analysis at about half the length of the longer one, which he interpreted as the precessional cycle. The detailed high-resolution analysis of this study also recognized this shorter cycle without the need for time series analysis on a long section - it is recognizable by eye in the bar charts of some categories (especially the equidimensional rounded black wood). The results of House (1985) suggest that the shorter section studied in this and in Weedon's (1986) study were not as complete as they may have appeared. However, the cycles used for the detailed palaeoecological interpretations here can be picked from the complete section and the best (apparently most complete) ones chosen. In addition, averaged cycles (see Fig. 8) were used for the interpretations, which are likely to make the results more reliable by decreasing the importance of any anomalies.
Interpretations The main interpretations of the playnofacies cycles are presented diagrammatically in Fig. 8, which summarizes all the important variations in palynofacies and related data (TOC, CaCO 3, clay content) and their interpretations through a typical Kimmeridge Clay cycle. The cycle shown was constructed by taking the means of each value at each sample position from the three cycles WL3, WL4 and WL5. Using only palynofacies data, the cycle may be divided into five segments or sub-segments, each one having different palynofacies characteristics. This indicates the greater sensitivity of palynofacies data over sedimentological data for recognizing small-scale environmental variations. Using only the sedimentological characteristics the cycle can be divided into only two distinct segments. There is a little variation within the lithological beds but it is not repeated in other cycles, as the palynofacies data often is. Where the detailed palynofacies data is presented in the figures any variations which might not be significant (i.e. might be due to variations other than palaeoenvironmental ones) are clearly marked. No interpretation or division of the cycle is based on possibly nonsignificant variations. Detailed palaeoenvironmental and orbital interpretations of the variations displayed in Fig. 8 are given in Table 5, which is self-explanatory so is discussed further only as a summary and in the expansion of a few points which warrant more detailed discussion.
An orbital interpretation The results of the estimations of other workers of the durations of the sedimentary cycles suggest that the most likely orbital explanation of the cycles is
that they were controlled by the obliquity cycle. It is also possible that the precessional cycle was the main control, but the presence of secondary cycles at half their wavelength within all of the main cycles of this study, and other workers' estimations (e.g. Weedon 1986), suggest the presence of precessional effects in addition to obliquity. One of the aims of this study was to see whether the known effects of any of the main orbital cycles would be recognized in the detailed palaeoenvironmental interpretations of the palynofacies data irrespective of any recognizable in the sedimentological variation, as an independent method of determining the dominant controls. A summary of the ways in which orbital variations could influence palynofacies assemblages is presented diagrammatically in Fig. 9. The main palaeoclimatic effect of the 100 ka eccentricity cycle is that it causes simple variations in the strength of the solar beam, but this is usually linked with the effects of the precessional cycle. The resulting pattern of cycles is usually seen as clusters of 3-8 precessional cycles with amplitudes which vary in response to changes in eccentricity. No such patterns were seen in the palynofacies data of this study and these should have been recognizable in the longer 12 m section if they were present. Rather, the variations appeared to be much more regular, simple alternations between two extreme points. The climatic effects of the precessional cycle are complex - the structure of the seasonal cycle is altered and is reversed in the two hemispheres. The result is to shift the climatic belts between hemispheres. It was not possible to look for evidence of variation in pollen and spore taxa through a cycle, as might be expected if climatic belts shifted, because in this distal setting the true diversity of spore and pollen taxa could not be obtained due to selective transportation of particles. In addition, the precessional cycle is known to have its greatest effect in low latitudes. During Kimmeridgian times, southern England was situated at c. 40~ (Hallam & Sellwood 1976), The obliquity cycle is known to oscillate with a steady beat and to yield simple, regular (sedimentary) alternations (Fischer 1986). At times of high obliquity, seasonality is greater and the pole-to-equator insolation gradient is lower, thereby producing less intense atmospheric and marine currents. At times of low obliquity the situation is reversed. Most of the palynofacies and related data display fairly regular alternations between two extreme conditions. The amplitudes of variations may alter through the section but in a different manner to the 'bundles' which would be expected if the amplitude of cycles were controlled by variations in eccentricity, as precessional-
101
PALYNOFACIES OF KIMMERIDGIAN CYCLES
ABUNDANCE polen ab
TYPE ~ooresl~olen
spore lib
VEGETATION
t~
WILDFIRES wood ( ~ ~ )
i J
/ A M O U N T AND DISTRIBUTION
I RAINFALLI
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INSOLATION
RATE
all al:)$ ~',s
1
DISTRIBUTION
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RUNOFF
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TERRESTRIAL INPUT
r,pore+ l:x)km ab.
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Fig. 9. Summary of the many possible effects of insolation changes on climate (in boxes), the resulting environmental impacts (in upper case text and no boxes), and the ways in which these may be reflected in the resulting palynofacies assemblages (lower case text). controlled cycles would be. Variations in amplitude are probably due to long-term trends in sedimentological or palaeoenvironmental factors. Several components of the palynofacies data can be interpreted in terms of variations in intensity of marine current activity, such as energy of marine currents and/or water column stratification, and related anoxia and high plankton productivity (see Fig. 8 and Tables 4 & 5). Since these indicators display a fairly regular oscillation between two extremes in each cycle, and marine current intensity is known to be a major effect of obliquity orbital forcing, it is reasonable to suggest that these palynofacies data support the various proposals
(of this and previous workers) that these cycles represent the effects of the obliquity orbital cycle. It is proposed that the main cycles (c. 120 cm) in palynofacies variation and in the sedimentology of this work were controlled by obliquity orbital forcing of the climate which controlled the marine current system. It is further suggested that the secondary cycles recorded by some palynofacies categories represent the effects of precessional orbital forcing of the climate. The ratio between the two cycles would certainly suggest this. The effects of the obliquity cycle are known to be most pronounced at high latitudes and those of the precessional cycle most
102
H.K. WATERHOUSE
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106
H.K. WATERHOUSE
pronounced at low latitudes. Perhaps, therefore, the position of southern England during the Kimmeridgian was such that it was within an overlap of the latitudes where the controls of both the precessional and obliquity cycles had effect. This would explain the presence of the secondary (precessional) cycles within each major (obliquity) cycle. One of the precessional peaks in palynofacies abundance is situated (segment la) close to, but slightly below, the major peak in palynofacies abundance (the obliquity 'high' of segment lb) of the longer obliquity cycle, and is therefore partly disguised by it. It may be noted, when the complete sampled section is studied (Fig. 2), that the major peaks in abundance often appear as double peaks. When studied in more detail these can be seen as the peaks in absolute abundance of almost all palynofacies categories in segment lb and of marine palynomorphs, equidimensional rounded black wood and cuticle in segment la of the typical cycle. It is thought that the distance between these two 'peaklets' varies through the longer complete sampled section (see Fig. 2). This would be expected to occur if precessional variation were superimposed on obliquity variation whose duration is not exactly twice that of the precessional variation. For example, if the terrestrial palynomorph abundances in the top part of the complete Kimmeridge Bay section (Fig. 2) are studied, it can be seen that the distance between the two 'peaklets' of each major peak seems to decrease moving down-section to a point within the WL3 shale bed, where there appear to be two major peaks situated more closely together than the rest, within one bituminous shale bed with no mudstone bed between them (note that the WL3 cycle analysed in detail in this work does not include the lower part of the WL3 bed which was considered to possibly be a separate cycle). Each peak in the WL3 bed also possibly has two 'peaklets' within it. It is suggested that this pattern represents the gradual shift of a superimposed precessional cycle (the lower 'peaklets' in these particular cycles, plus a second peak in equidimensional rounded wood which cannot be seen clearly in Fig. 2) along the wavelength of an obliquity cycle. At times the obliquity and one of the precessional peaks in palynofacies abundance would coincide (as they nearly do in the three cycles studied here in detail) and at others they would be completely separate, as has possibly occurred in the lower half of the sampled section at Kimmeridge Bay (Fig. 2). It was for this reason that the study of the Kimmeridge Bay section was continued downwards - to investigate whether it was possible to recognize this expected shift. However, as the patterns lower down the section became more complicated it was
rather difficult to interpret them in any meaningful way using the current sampling interval, although it certainly does show more than one major peak in palynofacies particle abundance per lithological cycle. Van Buchem et al. (1992), studying chemical cyclicity of Lower Lias mudstones of Yorkshire, UK, found that at some times the eccentricity signal was best preserved, while at others the obliquity and/or precessional signal was dominant. Van Buchem et al. (1994) further suggested that, in the Lower Lias mudstones of Yorkshire, orbitally forced climatic changes influenced the depositional system at two levels: (1) short-term variations affecting the storm frequency and perhaps magnitude; and (2) longer-term variations affecting weathering and clay production in the source area. It is proposed here that two orbital cycles produced different palaeoenvironmental effects in the Kimmeridge Clay of Dorset. The precessional cycle is suggested to have had its greatest effects on the terrestrial environment and terrestrially-derived variations in palynofacies (the second peak in abundance in each cycle occurred only in terrestrially derived particle groups), while the obliquity cycle had its greatest effect on the marine environment (the current intensity, plankton productivity and preservational conditions). A discussion of the possible nature of these effects now follows. A discussion of the variation in dinocyst abundance is first required. It may be noted that the first peak in dinocyst abundance, in segment l a, is not accompanied by peaks in the other particle groups (except equidimensional rounded black woods and cuticle). It is therefore likely that this peak is productivity, rather than preservation, controlled. This distinction could not be made without the use of absolute abundance techniques. Indeed, Berlin & Brosse (1992), in a high-resolution petrographical and geochemical study on cycles from the Exodus Zone of the Yorkshire Kimmeridge Clay Formation, proposed a higher primary productivity and higher sedimentation rate at the base of their cycle which, according to the patterns of their results, would probably correspond to the segments 3 and la of the present work. It is also possible that the peak in dinocyst abundance could be due to good benthic preservational conditions in the marine environment but low terrestrial output into the marine basin. It should be noted, however, that the dinocysts display two peaks in absolute abundance situated closely together, one at the same time as the first equidimensional rounded black wood precessional peak (la) as well as one during the obliquity controlled (lb) peak, while other particle groups display one peak. Perhaps this indicates both an obliquity and a
PALYNOFACIES OF KIMMERIDGIANCYCLES (different) precessional control on the dinocysts. However, they do not display the expected second precessional peak in segment 2a. There are several possible explanations for this pattern. Dinocyst production could be precessional controlled, with the peak in segment l a being a productivity peak and the (obliquity) peak in segment lb being preservation controlled, as the high absolute abundances of other particle groups would suggest. The lack of a productivity peak in segment 2a could therefore be due, in some way, to the probably rather stronger influence of the obliquity cycle producing a breakdown of stratification and benthic anoxia by increasing the intensity of ocean-current circulation. This would have prevented the otherwise expected (productivity generated) peak in dinocyst abundance from being preserved due to the low preservation potential, with the higher preservation potential in segment la (due to a slowing of ocean circulation and build-up of stratification and anoxia) allowing the first productivity generated peak to be preserved. It is proposed that this is the most likely explanation for the patterns. There is still, however, a peak in equidimensional rounded black wood recorded in segment 2a despite the apparently low preservation potential. If this explanation for the dinocyst peaks in abundance were correct, it would suggest that the abundance of equidimensional rounded black wood was controlled by either the energy of the transporting medium, which is thought to be greatest in segment 2a, or some factor acting on the terrestrial environment which affects neither the other terrestrially derived particles nor the dinocysts, A further discussion of the equidimensional rounded black wood variation patterns is given below. Alternatively, the dinocyst peak in segment la could be part of the obliquity cycle to which most other palynofacies particle groups responded slightly later. It would still, however, be productivity controlled. In this case, the earlier peak in dinocyst abundance than in other particles could be due to the obliquity generated conditions in that part of the cycle particularly favouring dinocyst production. For example, the conditions prevalent during the build-up to the low-energy, low benthic oxygen probably stratified water-column conditions of segment lb could be ideal for high dinocyst production, possibly including blooms. This production would then decrease as the water column became increasingly stratified, with maximum stratification (and minimum oxygen and minimum current intensity) in segment lb. The second dinocyst abundance peak (in lb) would be preservation controlled, as are most other particle groups, despite possibly much lower plankton
107
productivity. This would explain the lack of a second (precessional) peak in dinocyst abundance in segment 2a. A third explanation could be that two preservation-controlled dinocyst peaks were preserved within the bituminous shale, but that the supply of terrestrial debris was much lower during the first peak, therefore resulting in the lack of a peak in terrestrial particles in la due to low supply, despite high preservation potential. However, there is little conclusive evidence from other palynofacies indicators (e.g. the ratios), or from previous workers' observations, to support such a large variation in terrestrial supply in such a small segment of the cycle and in such a distal setting. It is suggested, therefore, that the first dinocyst peak in abundance was productivity controlled. Tribovillard et al. (1994) state that TOC acts as an indicator of productivity of organic-walled plankton. Perhaps this would explain the fairly high TOC values which correspond to the first dinocyst abundance peak, with the main peak in TOC values being an indicator of preservational conditions, as the other high particle abundances are. However, this still requires an explanation for the presence of a productivity peak at this part of the cycle and it needs to be determined why this peak occurs during the build-up of stratification rather than during its breakdown. Breakdown of anoxia is likely to be a more common trigger of high productivity due to remixing of stored nutrients back into the water column, although it does not necessarily require large amounts of stored nutrients or widespread anoxia. Irwin (1979) has suggested vertical mixing of nutrients as a method for coccolith propagation. Calvert & Pedersen (1990) suggested primary production as a main control on accumulation of organic matter irrespective of bottomwater oxygen values. Whichever interpretation is preferred, it is still likely that the two peaks in dinocyst abundance per cycle have different origins, the first being productivity controlled and the second preservational. At the beginning of their cycle in TOC variation, which showed similar patterns to the TOC variations of the cycles of this work, Bertrand & LallierVerges (1993) noted high SRI values where the TOC values were lowest. This difference in the two measures, which was in contrast with the rest of the cycle where they showed similar patterns, they interpreted as most probably being due to early sulphate reduction in anaerobic microenvironments, such as fecal pellets, aggregates, or living or dead organisms, close to the photic zone. This occurred in the same part of the cycle as the first dinocyst abundance peak of this work, for which an explanation of high dinocyst productivity is difficult to determine.
108
H.K. WATERHOUSE
If the TOC peak of Bertrand & Lallier-Verg~s (1993) is taken to represent the same point within a cycle as the TOC peaks of the present work (lb), then both peaks (according to Bertrand & LallierVerges 1993) would be the result of increased planktonic productivity, with the lower part of the cycle (below the peak in la) having the lowest plankton productivity but a high degree of sulphate reduction in microenvironments in the upper part of the water column. The results of the present study, however, suggest the opposite pattern for the bituminous shale part of the cycle. The other parameters requiring an explanation are the distributions of equidimensional rounded black wood and cuticle, which are very different from that of other palynofacies data. Their main features are a low absolute abundance during segment lb, where all other particles peak, a peak in segment 2a of equidimensional rounded black wood when all other particles have a low abundance, and a peak in segments 3/la where all other particle groups, except dinocysts, have low abundances. The low equidimensional rounded black wood and cuticle abundances in segment lb occurs during the time of lowest marine energy conditions, lowest oxygen and probably stratification of the water column. The low abundances of these groups must be due to a low supply to the site of deposition, as all other organic particles (both terrestrial and marine) are preserved in high absolute abundances. It must, therefore, be determined why these components have a lower supply than other terrestrially derived debris. The low abundance of cuticle can be explained by the low marine energy conditions. Cuticle is the strongest indicator of proximity of all the terrestrial particles analysed. In such a distal setting it would be low in abundance anyway, but where the energy of the transporting medium is especially low cuticle would be expected to be absent. An explanation for a low supply of equidimensional rounded black wood may be obtained from the paragraphs below where possible reasons for its peaks in abundance are discussed. One explanation of why there is no peak in dinocyst abundance in segment 2a when equidimensional rounded black wood peaks (if both were precession controlled) could be that the two particle types have very different origins, the dinocysts settling through the water column and probably being degraded on their way through it, while the equidimensional rounded black wood has a more horizontal supply, perhaps carried mainly by bottom marine currents once in the marine basin. It is also possible that the equidimensional rounded black wood is more resistant to degradation than
the dinocysts and is therefore not degraded in the well oxygenated conditions of segment 2a. There are very few clues in the data obtained to provide assistance with explaining the presence of the two peaks in equidimensional rounded black wood in segments 2a and 3/la. However, a number of speculative explanations can be suggested. Berlin & Brosse (1992) suggested that terrestrial organic debris particles became more rounded when energy was lower (and TOC therefore greater). An increase in equidimensional rounded black wood with respect to other, angular wood particles would be expected to represent a decrease in energy of the transporting medium allowing more time to be spent in transit for abrasion of black wood particles (from sharp to more rounded) and oxidation from brown to black wood to occur. The peaks in absolute abundance of equidimensional rounded black wood could not be due to any transportational effects as they show the opposite pattems to those expected, such as high abundance during segment 2a where energy is thought to be much higher than in lb where their abundance is low. It is possible that the equidimensional sharp wood and black laths consist entirely of charcoal and the equidimensional rounded black wood represents the only true (unburnt) wood, which was subsequently oxidized in transit. The equidimensional rounded black wood would therefore have a different origin, within the terrestrial environment, from the other types of black wood and therefore possibly be a reflection of vegetation change (Cope 1981). However, the brown wood displays a similar pattern of variation to the equidimensional sharp black wood and black laths (charcoal), and not the equidimensional rounded black wood as would be expected in this case. Therefore, an explanation for the similar abundance variations but different origins of those particles would be needed which could also explain the similar origins but different abundance variations of brown and equidimensional rounded black wood. An explanation which does not require different origins for the different equidimensional black wood types and similar origins for the equidimensional rounded black wood and brown wood is that the equidimensional rounded black wood has undergone a different method of transportation (including storage - see below) on land, or intertidally, before reaching the marine basin, from the other types of wood, whose abundances could then be a reflection purely of transportational processes or preservation in the marine basin. There is no indication from the other recorded data that the difference in abundance patterns of the woods
PALYNOFACIES OF KIMMERIDGIANCYCLES could be due to differences in marine transportation processes. Equidimensional rounded black wood could therefore be brown wood that has undergone oxidation or equidimensional sharp black wood that has undergone abrasion during transit (mostly whilst still on land). It would therefore have had to have spent a longer time in transit than the fresh brown wood or the sharp black wood and its shape would certainly indicate this. This could be due to extended periods of storage on land or intertidally, and not necessarily slower transit times (for which there is no evidence). This scenario, and those described below, could apply equally if the equidimensional rounded black wood was derived from oxidation of brown wood or physical abrasion of sharp equidimensional black wood particles (perhaps charcoal), or both. Scotchman (1991) noted that in almost all Kimmeridge Clay samples that he studied, vitrinite (wood) particles were generally small, not very abundant and occasionally showed signs of oxidation. Their reflectances suggested that many were reworked from older sediments. This he found to be more common in basin margin areas rather than in basinal areas, such as Kimmeridge Bay. Nevertheless, it is possible that, during certain parts of the cycle (3/la and 2a), a situation occurred where older sediments were able to be reworked, or the storage areas (on land or intertidally) of the oxidizing brown wood or abrading equidimensional sharp black wood to be plundered, and their particulate organic contents incorporated into the resulting marine deposits. A possible climatic explanation of this could be that the shift of climatic belts, which is a known effect of precessional forcing, could result in reworking of previously little eroded areas. Kutzbach & Otto-Bliesner (1982) stated that this shift would alter the boundaries between wet and dry zones, and the balance between monsoonal and zonal wind circulation. This could result in reworking of sediments in a number of ways. A change in direction of the prevailing winds could expose previously sheltered coastal areas, possibly storing equidimensional rounded black wood, to wave erosion. An increased frequency of storms would have a similar effect. Increased run-off would also expose terrestrial sediments to renewed erosion. A difference in type of vegetation may alter the supply of wood. For example, a forest-type vegetation would increase the supply of wood on land (although not necessarily its output to the marine basin). Another type of vegetation, such as desert or scrub, could induce greater terrestrial erosion (by rain and wind) and perhaps, therefore, remove previously stored equidimensional rounded black wood to the marine basin. An established vege-
109
tation of this type would also then keep it low in the marine basin by preventing its accumulation by oxidation and abrasion on land. A taxonomic study of the pollen grains may provide additional information, as would a study of the reflectivities of the different types of wood. If the difference in origin between equidimensional rounded black wood and other types of black wood were vegetational (or proxy-vegetational due to vegetation affecting erosion; Herring 1977), this could also fit into a precessional-controlled model for the patterns of variation seen. For example, a slight shift in climate belts could alter the type of vegetation. In the climatically sensitive 'zones', between 20 and 40 ~ latitude, even small changes in climatic conditions would be likely to change to position of the climate belts and, as such, the depositional environment (Van Buchem et al. 1994). The high sensitivity of the climate system has been illustrated by Kutzbach & Guetter (1984) and Kutzbach &Otto-Bliesner (1982), who demonstrated the importance of low latitude solar insolation changes as a modulator of monsoon circulation during the last 9 ka. It is proposed, therefore, that the most likely explanation for the apparently anomalous abundance patterns of the equidimensional rounded black wood is not any marine process or a direct result of vegetation change but that in some way a precessional-controlled shift in climate belts, and therefore vegetation, alters the supply of equidimensional rounded black wood (whether derived from equidimensional sharp black wood by abrasion or brown wood by oxidation) by alternatively allowing storage of the particles somewhere outside the marine basin and plundering these storage areas, therefore increasing the supply to the marine site of deposition. Any of the possibilities discussed above could also alter the horizontal supply of sediments and nutrients to the site of deposition, as well as possibly affecting the vertical mixing in the marine basin (e.g. salinity stratification due to increased run-off). This would, in turn, affect the plankton productivity which is also thought to respond to precessional orbital forcing. A variation in marine current circulation accompanying the shift in climate belts (but less important than the more dominant obliquity marine current control) could also explain the presence of a plankton peak here. Quantitative results were obtained from the reflected light work (on five samples - C9, C l l , C13, D2 and D20) and used to try to determine whether there was a difference between equidimensional rounded black wood and equidimensional sharp black wood, apart from their shape. Initial results indicated that equidimensional sharp black
110
H.K. WATERHOUSE
wood consisted of c. 15% vitrinite ('fresh wood'), 70% fusinite/semifusinite ('charcoal') and 12% inertinite ('oxidized wood') (Stach et al. 1982), while equidimensional rounded black wood consisted of c. 25% vitrinite, 65% fusinite/semifusinite and 7% inertinite. In other words, they both consisted of similar percentages of fusinite/semifusinite but equidimensional rounded black wood contained a greater proportion of vitrinite and equidimensional sharp black wood a greater proportion of inertinite. If equidimensional sharp black wood had spent less time in transit than equidimensional rounded black wood, it would be expected to contain a higher proportion of vitrinite ('fresh wood') than the equidimensional rounded black wood. That it does not could suggest several things. It could contain a similar or greater absolute amount of vitrinite but the larger amount of inertinite lowers the relative amount of vitrinite. Altematively, to fit into the storage theory, the two types could have had similar origins but have taken different journeys to the site of deposition, with equidimensional sharp black wood undergoing more chemical degradation during transit and the equidimensional rounded black wood more physical abrasion. Or, perhaps, vitrinite from different plants had different resistances to degradation and the proportions of wood types within each group reflects vegetation changes. A final explanation could be that the equidimensional sharp black wood contains less vitrinite because the vitrinite is still brown in colour, and therefore classified in a different group, because it has been little oxidized due to less time spent in transit. The question of the different compositions of the individual wood groups warrants further investigation. It should be remembered that the reflectance results discussed here were obtained from a small number of particles in a small number of samples. It is not unusual, when using percentage data, for one component to vary at the expense of another, as do the TOC and the clay, quartz and other mineral components of the samples of this work. What is interesting in the data of this work is that the CaCO 3 displays a similar variation to that of the clay and other minerals which comprise the major (80-98%) portion of the rock. When one component comprises such a large proportion of the total it would seem more likely for smaller components, such as the TOC and CaCO 3 here, to vary in the opposite way to the major one and in a similar manner to each other. The patterns seen here suggest that the conditions which favour an increased percentage of clay and other minerals to occur in a sample also favour the deposition of CaCO 3. This is reasonable, as CaCO 3 is more readily precipitated in oxygen-rich (and therefore more energetic) waters, and clay and other
minerals, most of which would be terrestrially derived in these samples, are likely to be more abundant when terrestrial run-off is higher or marine currents more energetic and therefore more able to transport particles. It has been shown that the percentage of terrestrially derived inorganic particles (clay, quartz, etc.) varies through a cycle, decreasing during the top of the bituminous shale. This could be a reflection of two things. It could represent an actual decrease in terrestrially derived mineral matter, which may give a false high reading of absolute organic particle abundances due to less dilution by mineral matter. This would only be true if the relative decrease in terrestrially-derived mineral matter was also an absolute decrease, which cannot be determined from the existing data. It could also be due to an increase in TOC in the marine basin during the deposition of the bituminous shales, due to either an increase in organic matter preservation (lower marine energy) and/or an increase in plankton productivity. Despite the fact that the TOC percentage is much lower overall than that of the clay and other minerals, it is suggested that during the bituminous shale deposition it had the same major control as the factors affecting the mineral matter percentages. In this case, a decrease in marine-bottom current energy would be the cause of increased preservation and/or productivity of organic matter and therefore TOC and, hence, relatively less clay and other terrestrial minerals. There are two factors which support this theory. This was a relatively distal setting so variations in terrestrial run-off, and therefore dilution by terrestrial particles, would have less effect than marine currents would have. Van Buchem et al. (1994), studying Lower Lias mudstones, proposed that terrestrial run-off could be excluded as a cause for TOC variation as the clay mineral distribution showed no correlation with TOC variations on a layer scale. In addition, the terrestrial organic particles (except the equidimensional rounded black wood and cuticle) show the opposite distribution to the clay and other minerals through the whole cycle. This suggests that their abundance is preservation-controlled because both of these components would otherwise be expected to vary in a similar manner since they are transported by the same agents. This also lends weight to the argument that a high absolute abundance of marine palynomorphs in the lower part of the bituminous shale was productivity controlled. The terrestrial organic particles and TOC have lower values in the lower (la) than in the upper (lb) part of the bituminous shale but the per cent of clay and other minerals is higher. If this was due to higher terrestrial output it should both dilute the marine component and coincide with an increased terrestrial organic debris component, neither of
PALYNOFACIES OF KIMMERIDGIAN CYCLES which occur. Therefore it is probably only an effect of a decreased TOC component. A summary of the interpreted palaeoenvironmental changes through a typical Kimmeridge Bay palynofacies cycle is as follows. Precessional (positive ?) orbital forcing shifts climatic and vegetational zones, and allows areas on land or intertidally storing wood, which is being altered (from equidimensional sharp black wood or brown wood) to equidimensional rounded black wood, to be eroded and their contents carried into the marine basin. This is accompanied by a peak in dinocyst productivity in the marine basin (1 a). The high point of an obliquity cycle causes marine currents to slow sufficiently to allow a build-up of anoxia in the particular palaeogeographic position of the Kimmeridge Bay area at that time. This allows all organic particles to be preserved in high abundances. There is a low supply of equidimensional rounded black wood and cuticle at this time due to the (negative ?) precessional forcing (lb). Precessional (positive ?) forcing again shifts climate belts and allows storage areas of equidimensional rounded black wood to be plundered and the particles to be transported to the marine basin. Decreasing obliquity values keep ocean currents circulating normally and keep preservation potential of organic matter low, except where there is a particularly high supply of resistant material (equidimensional rounded black wood) (2a, b). The cycle then begins again (3). It must be remembered, however, that if the proposals advanced here that the palynofacies particle abundances show the effects of the precessional cycle 'moving along' an underlying obliquity cycle, the effects of the two cycles will not always coincide in the pattern as those described for the 'typical' palynofacies cycle here, as this is based mainly on the top three cycles of the Kimmeridge Bay section studied. In conclusion, it is proposed that the 'major' palynofacies cycles in the Kimmeridge Bay section represent the climatic effects of orbital forcing by the obliquity cycle and the 'secondary' palynofacies cycles, those of forcing by the precessional
11 |
cycle, and that the position of southern England during Kimmeridgian times was such that the Kimmeridge Bay area was under the influence of both the obliquity and precessional orbital variations, but perhaps with the orbital influence being slightly stronger. It is also proposed that the obliquity cycle exerted a stronger control on the marine realm by influencing the ocean-current circulation and, therefore, plankton productivity and preservation conditions, and that the precession cycle affected both the marine and terrestrial realms but exerted its control mainly on the terrestrial environment. By using quantitative palynofacies analysis as a tool and employing high-resolution sampling techniques, detailed palaeoenvironmental variations, more detailed than those recognizable in the sedimentological characteristics, can be identified through a lithological section. The most important of these were the precessional-forced peaks in abundance and the variation within the apparently uniform bituminous shale beds. In addition, the use of a method of estimating absolute palynofacies particle abundances has allowed particles with different provenances, but subsequently found superimposed upon one another in the distal setting of the Kimmeridge Bay section, to be studied separately. This provides a possible means of identifying separate variations, that may have different palaeoenvironmental causes, within and between the marine and terrestrial environments. Quantitative palynofacies analysis, in conjunction with high-resolution sampling techniques, has therefore proved to be a powerful tool and a sensitive indicator in the palaeoenvironmental and palaeoclimatic investigation of the cycles in this study, and future work on the detailed effects of orbital forcing of climate would benefit from the use of such techniques. This work was supported by a NERC research studentship. Thanks are also due to the Smedmore estate for permission to collect samples, C. Mar-Molinero for advice on time series analysis, to J. E. A. Marshall and M. R. House, and to R. V. Tyson and another, anonymous, referee for many useful comments.
References AIGNER,T. 1980. Biofabrics and stratinomy of the Lower Kimmeridge Clay (U. Jurassic, Dorset, England). Neues Jahrbuch fiir Geologie und Pal~iontologie, Abhandlungen, 159, 324-338. ARKELL,W. J. 1947. The Geology of the Country around Weymouth, Swanage, Corfe and Lulworth. Memoirs of the Geological Survey of Great Britain. BERGER, A. L., LOUTRE, M. F. & DEHANT, V. 1989a. Influence of the changing lunar orbit on the astronomical frequencies of pre-Quaternary insolation patterns. Paleoceanography, 4, 555-564.
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Modelling Late Jurassic Milankovitch climate variations E J. V A L D E S , B. W. S E L L W O O D 1 & G. D. P R I C E 1
Department of Meteorology, University of Reading, 2 Earley Gate, Whiteknights, Reading RG6 2A U, UK. 1Postgraduate Research Institute for Sedimentology, The University, Whiteknights, Reading RG6 2AB, UK. Abstract: Although largely circumstantial in character, evidence for orbitally-forced
(Milankovitch) climate changes in Jurassic microrhythmic successions, such as those of the Lias and Kimmeridgian, is becoming more persuasive. We present here the results of experiments, using a general circulation model, testing ways in which orbitally-induced variations in solar energy might be translated into a Jurassic climate response. In particular, we address the problem of the 100 ka (eccentricity-forced) cycle. This is generally considered to have only a small direct effect on solar input. It would be expected to have little impact on an ice-free Earth (commonly assumed for the Jurassic). Nonetheless, this weak signal is claimed to have been recognized in many Jurassic successions. Our results simulate, for the Late Jurassic, the possible effects at the 'minimum' and 'maximum' extremes of seasonal forcing (i.e. comparable with those affecting the Earth at 115 and 9 ka BP, respectively). Model predictions are critically evaluated against the geological database. At times of 'minimum seasonal forcing' there is a significant expansion in the area of the Northern hemisphere monsoon. In the tropics, changes in precipitation predominate over changes in temperature, whereas at high southern latitudes there are very large seasonal variations in temperature, and heavy winter snows. During these times the model is close to predicting a modest, but significant, Jurassic ice-cap in the Antarctic. Ice build-up is particularly likely over uplands. Such an ice-cap disappears at times of 'maximum seasonal forcing'. Waxing and waning of ice may thus provide the elusive mechanism for metre-scale sea-level changes. It is argued that apparently similar microrhythms (e.g. limestone-shale) might be the sedimentary response to different climatic signals in different climate zones, cause and effect being exactly the opposite in some circumstances.
Successions rhythmically bedded on a decimetre to metre scale (often interpreted to be cyclic) are a common feature of Jurassic successions worldwide, 'pervading every Jurassic facies' according to House (1986). Since the pioneering days of Kltipfel (1917) and Brinkmann (1929) a vast literature has blossomed, concerned with both the description and interpretation of rhythms, microrhythms and cycles, interpretations often reflecting the fashion of the time. In the early parts of this century epeirogenetic interpretations were popular. Subsequently diagenetic and autocyclic mechanisms (reviewed in Hallam 1964, 1986) were favoured, followed in the last couple of decades by eustatic and generalized climatic controls (Sellwood 1972; Anderton et al. 1979). At the moment interpretations involving refined climatic controls (orbitally-forced Milankovitch models) are in vogue (e.g. House 1985, 1986; Weedon 1985, 1993; Weedon & Jenkyns 1990), even though the evidence in support of such interpretations are often no better than circumstantial [see discussion in Hallam (1986) and Wignall (1989)]. Evidence for eccentricity,
obliquity and precession cycles has also been claimed on the basis of natural gamma-ray logs used at outcrop and in boreholes (e.g. Van Buchem et al. 1992). The objective of this paper is not to re-review the nature and origin of rhythmic cyclic successions in the Jurassic but to shed light on the way in which Milankovitch climatic signals might effect preservable responses in the sedimentary record of the Jurassic Earth. Because of growing interest in the mechanisms of climate change, controls of eustasy and the potential use of Milankovitch cyclicity in high resolution stratigraphy, it is important to understand the ways in which Milankovitch signals might effect changes in climatic patterns on the Earth. The approach adopted here is the employment of a version of the general circulation model (GCM) of the UK Universities Global Atmospheric Modelling Programme (UGAMP). As will be seen, some of the climatic responses predicted are not, perhaps, those that might intuitively be expected. The results of some of modelling experiments underline the principle that apparently similar
From HOUSE,M. R. & GALE,A. S. (eds), 1995, OrbitalForcing Timescalesand Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 115-132.
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cyclic responses may be produced for different reasons in different climate regimes. Although there has been much work on Jurassic cyclicity, and much recent speculation about possible climatic controls, there has been much less effort in trying to understand the nature of climate change itself, yet this is of fundamental importance. The effects of Milankovitch forcing are well recognized in Quaternary and some Pliocene successions, particularly in the deeper marine record (e.g. Shackleton 1993; Thunell et al. 1991). Climate change is forced by variations in the radiative energy arriving at the top of the atmosphere. This is controlled by three orbital parameters, namely the eccentricity of the Earth's orbit (which has periods of c. 100 and 400 ka), the obliquity of the rotation axis (period of c. 40 ka) and the precession of the equinoxes (period of c. 20 ka). They change the seasonality of incoming solar radiation. The annual mean solar radiation changes only by very small amounts, yet the climate itself changes both seasonally and over an annual average. Further, the 100 ka eccentricity forcing has a small direct effect on the solar input. It's effect mainly controls the size of the precession signal. Thus, it is surprising that the 100 ka signal is so large in the Pleistocene, and is frequently claimed to have been recognized in the prePliocene (e.g reviews in Crowley & North 1991). There are now many explanations for the dominance of the 100 ka signal (for review see Imbrie et al. 1992). All of them imply some feedbacks between various components of the climate system, of which the most important, given late Pleistocene boundary conditions, is the growth of the major ice sheets. The standard theory is that the initial glacial response to orbital forcing will occur at high Northern hemisphere latitudes. If the orbital parameters are such that Northern summers receive less radiation than normal, then summers will be relatively cool and winter snows will not be entirely melted. Ice albedo feedback will then result in further cooling and, hence, further growth of ice. Such an orbital configuration occurred at 115 ka BP. A similar argument, but in reverse, explains deglaciation, for which the orbital configuration at 9 ka Be is a prime example. If an expansion in the area of ice-fields is the main feedback process which amplifies the 100 ka Milankovitch period, then we would expect to see a much weaker eccentricity climate signal during times when the Earth was generally warmer ('equable'), such as the Mesozoic (e.g. Crowley & North 1991; Hallam 1993). Alternatively, ice may not be the major feedback process during such periods. Oceans, continental configuration and changes in the carbon dioxide content of the
atmosphere (e.g. Berner 1992) may also be of crucial importance. Further, Saltzman & Maasch (1988) and Maasch & Saltzman (1990) have suggested that the 100 ka period signal is the result of a self-sustaining oscillation of the climate system. They propose that this signal would exist even if there were no external forcing at the eccentricity period. If this is true, then it introduces some serious problems for interpretation of pre-Pliocene sedimentary rhythms because such a self-sustaining oscillator could possibly change frequency in a climate regime very different from that of the Peleistocene.
Numerical climate modelling of Milankovitch cycles In order to understand the quantitative links that exist between changes in the incoming solar radiation and climatic parameters, such as temperature and rainfall, it is necessary to develop numerical models of the climate system. There are currently two main types of models, both of which have their strengths and weaknesses. The first type of model, generally referred to as energy balance models (EBM), can be used to simulate the climatic change over a full glacial/interglacial cycle. The most sophisticated versions include sub-models for the atmosphere, ocean, ice sheets, and lithosphere (e.g. Berger et al. 1993). However, even with the most powerful computers, such complicated models can only be run for Milankovitch time-scale periods by making a number of severe approximations. The most common of these is to calculate the latitudinal/ height variations only. No longitudinal variations are included. Further, some potentially important processes, such as clouds and monsoons, cannot be properly represented. Berger et al. (1993) have used such a model to simulate the last glacial/interglacial cycle with considerable success. For the last glaciation the evolution of predicted temperatures and ice volumes are remarkably similar to the observed variations. In particular, it correctly reproduces the longer timescale (100 000 a) climate variation. Further diagnosis of their model shows that the amplification of the longer timescale variations are the result of feedbacks related to ice cover and carbon dioxide. The clear implication is that the 100 000 a response would be considerably weaker on an Earth without significant ice. The strength of this type of model is that it calculates the transient response. The weakness is that the model has many simplifications and assumptions. It would be potentially difficult to
LATE JURASSIC CLIMATE VARIATIONS modify such a model for a different climate regime, and a different palaeogeography. Further, the assumption of a zonally symmetrical climate may be particularly unreliable for the Triassic and Jurassic, when the supercontinent of Pangaea, and its related mega-monsoons, significantly disrupted zonal circulation patterns (Kutzbach & Gallimore 1989). The alternative type of numerical climate model is the GCM. This model can be used to simulate the full regional and global climate, but cannot be run for 100 000 a. Instead, the model can only be run for a 'snapshot' at particular times during a Milankovitch cycle. It predicts all climate variables. In particular, it is possible to examine the relative changes of temperature and the hydrological cycle. This information can be used to aid the interpretation of sedimentary deposits. The GCM approach has had considerable success in explaining the evolution of climate over the last 20 000 a (e.g. COHMAP 1988), as well as increasing our understanding of regional climate change for specific periods (e.g. Kutzbach & Wright 1985; Kutzbach & Guetter 1986). However, there are also a number of problems. Rind & Peteet (1985) showed that GCM simulations, using CLIMAP sea-surface temperature reconstructions for the last glacial maximum, were in contradiction with estimates of land temperatures at high altitudes. There is currently much debate as to whether it is the data, or the models, that are incorrect. There is also a problem in understanding the onset of glaciation. Rind et al. (1989) showed that the GCM of the Goddard Institute for Space Studies (GISS) would not start to grow an ice sheet if the orbital parameters alone were modified to those for 115 ka BP The cool Northern hemisphere summers still melted all of the winter snows. Winter snow lasted through the summer only if sea-surface temperatures were changed to those at the last glacial maximum (probably an unreasonable assumption). Such a serious disagreement, between models and data, may just be an artifact of the particular model used. Oglesby (1990) showed that the NCAR CCM1 model had no problems at maintaining winter snow, and he speculated that the GISS model may be a 'hot' model that has difficulty maintaining snow cover, whereas the NCAR model may be a 'cold' model that has difficulty removing snow cover. The UGAMP model, employed by us, appears to lie in between these two extremes (D. Buwen, pers. comm.). In the northeast of Canada, Alaska and Siberia, winter snow cover does last through the summer. GCMs and EBMs have also been used to
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examine 'snapshots' of climate of the more distant past. Oglesby & Park (1989) and Park & Oglesby (1990, 1991, 1992) have examined the role of orbital changes on Cretaceous sedimentary rhythms. They show that the precession effects dominate their model, especially the hydrological balance in the tropics. They suggest that orbitally induced changes of the 'precipitation minus evaporation' budget can result in cyclic variations of anoxic waters in shallow basins and in fluctuations in terrigenous input to shelf environments. Obliquity changes seemed to be of less importance, even at high latitudes. Crowley et al. (1992a, b) showed that the 100 ka period could be amplified at times of extreme continentality, as in the Triassic and Early Jurassic. They showed that the temperature maximum of low-latitude land exhibited a significant 100 ka oscillation as a result of the interaction between the twice yearly passage of the Sun across the Equator, and the seasonal timing of perihelion. The magnitude of the temperature response implies significant changes in the hydrological cycle of the monsoon and are of sufficient magnitude to leave an imprint on the geological record. Their model suggested that high latitudes did not experience the same amplification, although feedbacks within the climate system could influence high latitudes. Crowley et al. (1993) have also investigated Milankovitch variations for Carboniferous climates. Using a GCM they showed that there were large changes in climate between cold and warm summer orbits. In particular, snow cover extended over large parts of Gondwanaland during cool summer orbits, but almost disappeared during warm summer orbits. The extent of the ice was in reasonable agreement with the geological data concerning ice cover. Further, they suggested that the interpretation of warm winters in Gondwanaland (Yemane 1993) could be the result of preferential preservation of the deposits representing the warm Milankovitch phases. They suggested that GCMs should be run for a number of orbital parameters in order to fully document climate for the past. These studies, therefore, suggest that the 100 ka variations should be strong during periods of glaciation (especially at high latitudes), or during periods when the Earth has supercontinents (a strong climate signal occurring at low latitudes only). For other periods the 100 ka period is less well understood and should provide a weaker, and less well-represented, signal. In this paper results are presented from a GCM investigating the sensitivity to Milankovitch for the late Jurassic (Kimmeridgian Stage). This is a time period in which the supercontinent of Pangaea is breaking
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E J. VALDES ET AL.
up but, as in the Cretaceous, there are generally 'equable' conditions.
Model description The model presented here is similar to that used in Valdes (1993) and Valdes & Sellwood (1992). It is a version of the UGAMP model. This model is spectral, using a triangular truncation at total wave number 31 (a horizontal grid of 3.75 ~ x 3.75~ and 19 levels in the vertical. The parameterization schemes are the same as those used in Valdes (1993). They include the Betts-Miller convective adjustment scheme (Betts & Miller 1986; Slingo et al. 1994), and Morcrette radiation (Morcrette 1990). The former scheme resulted in significantly moister tropics than was found in Valdes & Sellwood (1992), which used Kuo convection (Kuo 1974). The model is adapted to Kimmeridgian boundary conditions by changing the coastlines (from Alan Smith 1991, pers. comm.), mountains (both the mean height and the variance), surface roughness and albedo, sea-surface temperature (SST), and carbon dioxide concentration. All other parameters, including the Earth's orbital characteristics, were kept at their present-day values. The details of these changes are explained in Valdes & Sellwood (1992). In Valdes (1993) and Valdes & Sellwood (1992), we prescribed a zonally uniform SST of 2 7 ~ cos (latitude) and it did not change with the seasons. The temperatures were consistent with the energy balance of the model, provided that carbon dioxide concentrations were 3-4-times present day values. Such figures are consistent with those calculated by Berner (1992). By fixing these temperatures, we effectively implied a change of poleward transport of oceanic heat flux. It was found that the annual mean oceanic heat transport was significantly weaker than that of the present day. In order to examine the model's response to Milankovitch variations it is inappropriate to keep the SST constant because this will artificially damper the climate changes. The best approach would be to use a fully dynamic, coupled ocean-atmosphere model. However, such models are still being developed for present-day simulations and are also extremely expensive computationally. Instead, a similar approach to that used for many future climate change scenarios was adopted (Gates 1990), namely a mixed-layer ocean. The ocean is modelled as a single thermodynamic slab, 50 m thick, with a prescribed oceanic heat flux. Sea ice is modelled using the scheme of Slingo (1985). Thus, the SST can change due to changes in the energy entering the slab
(a sum of latent, sensible and radiative fluxes), but not due to changes in the oceanic circulation. The oceanic heat flux is the annual mean value from the simulation in Valdes (1993). Thus, it has both longitudinal and latitudinal structure, but no seasonal variations. This approach is similar to that used by Barron et al. (1993) for the Cretaceous, except that they used an oceanic heat transport which was a simple multiple of the present day value. Since there is currently no reliable way of determining palaeoocean heat flux, both approaches represent a gross simplification of the true climate system. In the following sections the results of three simulations, representing different extremes, are presented. The first experiment will be referred to as the control. It uses current orbital parameters (obliquity -- 23.44 ~ eccentricity -- 0.0167, longitude for perihelion relative to vernal equinox -282.04~ The only difference between this simulation and that employed in Valdes (1993) is the use of a mixed layer ocean. The other two experiments were chosen so as to demonstrate the possible effects, on the Jurassic Earth, of extremes in the radiative forcing corresponding with the changes that have affected the Earth's orbital parameters over the last 120 000 a. At 115 ka bp the orbital configuration (obliquity = 22.41 ~ eccentricity = 0.04142, longitude of perihelion- 291.02 ~ corresponded to a cold Northern-hemisphere summer. This allowed winter snow to last through the summer, resulting in the growth of the major ice sheets (a time of 'minimal seasonal forcing'). The opposite occurred 9 k a ago (obliquity=24.238 ~ eccentricity= 0.01928, longitude of perihelion = 131.26~ Warm summers throughout the mid-Holocene resulted in the melting of the huge Devensian ice sheets that had extended over North America and Europe (time of 'maximum seasonal forcing'). Both of these configurations primarily altered the seasonality of incoming solar radiation. At most latitudes, the annual mean changes were relatively small. Figure 1 shows the changes in incoming solar radiation for the two periods. For many months and latitudes, the changes exceed 40 W -2. By comparison, a doubling of carbon dioxide, or a 1% change in the solar constant produce a change in the net radiative forcing of c. 4 W m -2. However, both of these processes operate throughout the year. The control was run for a simulated-time 'snapshot' of 10 a, whereas the sensitivity experiments were run for 5-6 a, starting from the end of the 10 a control integration. We also performed a further experiment using the control simulation but with initial land ice covering all land polewards of 45 ~ latitude. All results presented here are averages over the last 3-4 a (simulated) years of the
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Fig. 1. The change of incoming solar radiation at the top of the atmosphere, caused by Milankovitch orbital variations, for (a) 'minimum seasonal forcing', and (b) 'maximum seasonal forcing'. The horizontal axis shows day number from the beginning of the year and the vertical axis shows latitude from the South Pole (-90 ~) to the North Pole (+90~ The contour interval is 5 Wm. integrations. Since (September 1993) (simulated) has been all the main features
the paper was submitted an integration for 2 5 - a completed. This confirms that are robust.
Results of the control simulation The results for the control simulation (using present-day orbital parameters) are similar to, but not identical with, those shown in Valdes (1993). The inclusion of a seasonally and longitudinally varying SST results in important changes in the tropical circulation. The seasonal mean surface temperatures for December-February average (DJF) and June-August average (JJA) shows the development of a warm pool of water in the Western Tethys. During the Northern-hemisphere summer this region is particularly notable, there being an extensive region of water in excess of 28~ just North of the equator. The temperatures are 1-2~ warmer than those noted in the simulation of Valdes (1993), which used prescribed SSTs. These relatively warm ocean temperatures have an important effect on the general circulation, and especially the hydrological cycle. The amount of water vapour in the atmosphere is a non-linear function of temperature. Thus, an increase of temperature from 27~ to 28~ is more important than a similar change between 10 and ll~ for example. The total precipitation for DJF (Fig. 3a) and JJA (Fig. 3b) closely follows the maximum in the SST. There is a band of heavy precipitation over the
Northern Tethyan ocean in JJA that only impinges on land in a small belt over Northern Africa and the Eurasian Peninsula. This predicted area of precipitation over Africa is in better agreement with geological observations of bauxite, coal and evaporite distributions (Fig. 4a and b) than were the original simulations in Valdes (1993) and Valdes & Sellwood (1992). The formation of evaporites and other hydrologically sensitive deposits is probably more clearly shown by examining the surface-soil moisture for DJF (Fig. 5a) and JJA (Fig. 5b). This is a more straightforward diagnostic than the 'precipitation minus evaporation' field since, in the latter, the evaporation should be a measure of the potential evaporation, not the actual value. Soil moisture takes this into account. Regions which are seasonally dry are regions where, potentially, either laterites or evaporites may form (Sellwood & Price 1993). Such regions occur over many of the areas bordering the tropical Tethys sea. This situation was not observed when the model was run with prescribed annual mean SSTs (Valdes 1993).
Modelling the Kimmeridgian simulating 'minimum seasonal forcing' This orbital configuration produces less summer and more winter solar radiation. It produces relatively cool summers and warm winters, as shown in Fig. 6a and b. These show the difference in seasonal mean surface temperature between the 'minimum seasonal forcing' simulation (based on
120
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Fig. 2. The simulated December-February (DJF) (a) and June-August (JJA) (b) mean surface temperatures from the control simulation of the GCM (using present-day orbital parameters). The contour interval is 4~ and the zero contour is dotted. Sub-freezing temperatures are shown by dashed contour and light hatching. The coastlines are shown by the thin solid line. The projection is a simple longitude/latitude style, extending to both poles. Tick marks on the edge of the box are every 30 ~.
121
LATE JURASSIC CLIMATE VARIATIONS
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Fig. 3. The simulated December-February (DJF) (a) and June-August (JJA) (b) mean total precipitation from the control simulation of the GCM (using present-day orbital parameters). The contours are 0.5, 1, 2, 4, 8, 16 mm day-l, and amounts greater than 4 mm day-1 are hatched.
122
P . J . VALDES E T AL.
ETe .. KEY
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(b) Fig. 4. (a) Distribution of climatically-significant facies for the Kimmeridgian and including ammonite provinciality data from Ross et al. (1992). Data derived from many sources (on Reading University Climate Modelling Database) and including references cited in text. Other key references are: Arkell (1956), Jansa (1972), Beydoun (1988), Fltigel and Fltigel-Khaler (1992). Data are plotted on 'Atlas' base map of A. Smith (Cambridge University). (b). Distribution of Late Jurassic humid and arid belts on the 'Atlas' palaeogeography (climate zones from Hallam 1985).
the Pleistocene 115 ka orbital configuration) and the control. There are changes in temperature of c. 5~ in the tropics and at high latitudes. Those regions between 15 and 45 ~ latitudes show only very modest changes. The low-latitude changes are associated with
changes in the monsoonal circulation. This is particularly dramatic in the Northem African area. In the control simulation (Fig. 3 ) there was heavy precipitation during Northern-hemisphere summer over a relatively small region. The changed solar insolation results in a southward shift, and expan-
123
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Fig. 5. The simulated December-February (DJF) (a) and June-August (JJA) (b) mean surface soil moisture from the control simulation of the GCM (using present-day orbital parameters). The contours are every 2 mm (which represents 10% of saturation).
124
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Fig. 6. The simulated December-February (DJF) (a) and June-August (JJA) (b) mean surface temperature difference between the 'minimum seasonal forcing' simulation and the control. The contour interval is 1~ the zero contour is not shown, and negative anomalies are shown by dashed lines. Regions with negative anomalies less than -2~ are hatched.
LATE JURASSIC CLIMATE VARIATIONS
sion, of the inter-tropical convergence zone (ITCZ). There is a consequent increase in the land area affected by the summer monsoon (Fig. 7). A similar shift occurs over the ocean in the JJA precipitation. However, over land there is less difference. There is only a small enhancement of the precipitation over Eastern Africa, an effect which decreases further west. The low level flow is strongly influenced by our reconstruction of orography (see Valdes & Sellwood 1992). This has a large plateau over central Africa (as a result of tiffing). This upland reduces the extent to which flow can penetrate into the heart of the continent. In the tropics the changes in precipitation predominate over changes in temperature (and associated changes in evaporation). Thus, the soil moisture barely reflects the changes in precipitation. There are large changes in Northern.Africa, and on the eastern coast (Kenya), but much smaller changes elsewhere. The situation at high latitudes in the Southern hemisphere is especially interesting. The large land mass results in large seasonal variations in temperature. Winter minima are < -20~ whereas in summer the temperatures can exceed 20~ In the control simulation, although there are heavy winter snows over much of the continent, the warmth in the summer is sufficient to melt all of this snow. Thus, there is no possibility that, in the control case, permanent snowfields could start to accumulate. At first sight this is not the case for the 'minimum seasonal forcing' simulation. Figure 8 shows the snow depth for DJF. There is a large area of snow cover in the Southern hemisphere (summer), and this would imply that snow could accumulate. However, a more detailed analysis shows that, although close to the threshold for progressively incremental snow cover, the climate has not quite deteriorated that far. A monthly analysis shows that snow cover does entirely disappear in February (the warmest month for the Southern hemisphere). For all other months some snow is present. The model, thus, stand at a knife edge in predicting a Jurassic ice-cap. In the Northern hemisphere the situation is much simpler. Winter snow melts completely in the summer. There is at least a three month summer period during which there is no snow cover. We performed a further integration to investigate to what extent the model could sustain an ice sheet. The control simulationwas repeated, but with an initial condition which included 5 m of snow cover upon all land poleward of 45 ~ latitude. However, the ice melted rapidly in both hemispheres. Within 3 a (simulated) the ice sheets had contracted to much smaller areas. There was no significant region in which an ice sheet grow thicker. This
125
suggests that any ice sheet produced in a cold phase of Milankovitch would quickly melt when the orbital parameters return to near those of present day values. None the less, the interesting possibility exists for sufficient ice both to accumulate and disappear on a Milankovitch timescale, and in volumes sufficient to produce metre-scale changes in sea level over upland areas in the cooler parts of the southern continent.
Modelling the Kimmeridgian simulating 'maximum seasonal forcing' The orbital parameters for this period correspond to an enhancement of incoming summer solar radiation. To a first approximation, this orbit is the opposite to that which produces the 'minimum seasonal forcing'. Thus, it is expected that the summer and autumn temperatures will be warmer than the control, and spring and winter ones to be cooler. Figure 9 shows the difference in surface temperature between this simulation and the control. High latitude winter temperatures are colder by up to 7~ In the Southem hemisphere the cooling is only apparent over Australia. Over Antarctica there is a warming of up to 5~ We believe that this is related to the orography. The barrier effect of the mountains disrupts the zonal flow and deflects tropical air on to Gondwanaland. In the tropics the temperature response is simpler. Land temperatures are high during JJA and low in DJE The magnitude of the changes are 4-5~ The sense of the response is as would be expected. The decrease of land temperature in the Northern-hemisphere winter has the effect of increasing the equator-to-pole temperature gradient. This results in an increase in mid-latitude depressions, and hence in the rainfall. This results in a substantial increase in the soil moisture content over Southern Europe (Fig. 10). In the Southern-hemisphere winter a similar band of increased precipitation and soil moisture occurs. However, this is related to a shift in the convective precipitation. The ITCZ moves c. 5 ~ equatorward. This expands the region which experiences wet seasons. In the tropics the changes in soil moisture are modest. However the largest changes occur in regions similar to those in the 'minimum seasonal forcing' simulation, namely in Northern Africa and Eastern Africa.
Discussion The model simulations suggest that the tropics will experience significant Milankovitch variations,
126
P.J.
VALDES ET AL.
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Fig. 7. The simulated December-February (DJF) (a) and June-August (JJA) (b) mean total precipitation difference between the 'minimum seasonal forcing' simulation and the control. The contours are at 1, 2, 4, 8, 16 mm day-1. The zero contour is not shown, and negative anomalies are shown by dashed lines.
127
LATE JURASSIC CLIMATE VARIATIONS
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Fig. 8. The simulated December-February (DJF) snow depth in the 'minimum seasonal forcing' experiment. The contour interval is 10 cm of water equivalent. The large area of snow cover in the Southern-hemisphere winter was not present in the control simulation. particularly over the Northern and Eastern coasts of Africa in areas where unfortunately, in areas where precise stratigraphic information is currently lacking. The variations are mainly seen in the hydrological cycle (most notably in the convective rainfall), with much smaller changes in temperature. Convective rainfall is associated with thunderstorm activity and related phenomena. Such storms may trigger wildfires, particularly at the end of dry spells, and the heavy downpours associated with them frequently result in immediate run-off (Sellwood & Price 1993). Such processes might produce signals detectable in the sedimentary record. In mid-latitudes, Southem Europe also experienced large changes in soil moisture; however, these were associated with changes in rainfall connected to mid-latitude depressions. This type of rainfall is a slower type of rainfall and is much less likely to produce excessive amounts of run-off. Instead, the model predictions are reflected in the general absence of evaporites within peri-tidal successions, and a general predominance of kaolinitic clay suites in Atlantic-Tethyan sediments (Chamley 1989), and the localized development of bauxites (Fig. 4). In high latitudes the changes are largest in terms of temperature not moisture. In the Northern hemisphere the changes in snow cover are minimal.
In particular, winter snow melts for a substantial part of the summer. In the Southern hemisphere, however, snow cover only clears for one month in each year. Small changes in SST, orography or carbon dioxide concentration could result in the growth of a significant ice sheet. Such a sheet would be expected to wax and wane in concert with Milankovitch periodicity. A further implication, consequent upon our model simulations, is the possibility that similarlooking rhythms might be generated in different climate regimes, but for different reasons. As Thunnell et al. (1991) have shown, for a Pliocene succession in Southern Italy, limestone-marl rhythms originated in response to Milankovitchdriven changes in surface-water productivity. In this case, more productive conditions (giving pelagic limestone accumulation) were caused by cooler, more windy, phases which promoted increased up-welling. Warmer, less productive phases led to marl accumulation. In other climate systems warmer conditions might well lead to increased rates of carbonate precipitation, especially over adjacent platform areas, and give rise to more rapid rates of peri-platform lime-mud accumulation. Thus, superficially similar rhythms might be generated by climatic changes that were exactly out of phase. The changes noted above could be related to
128
P. J. VALDES ET AL.
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Fig. 9. The simulated December-February (DJF) (a) and June-August (JJA) (b) mean surface temperature difference between the 'maximum seasonal forcing' simulation and the control. The contour interval is I~ the zero contour is not shown, and negative anomalies are shown by dashed lines. Regions with negative anomalies < -2~ are hatched.
129
LATE JURASSIC CLIMATE VARIATIONS
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Z J. VALDES ET AL.
some of the limitations inherent within the model itself. These fall into four categories: lack of time domain; lack of ocean feedbacks; use of constant carbon dioxide concentration; and inaccuracy in dealing with ice accumulationn. As stated earlier, general circulation model studies will always be constrained to the investigation of one short time slice. Their strength is in their ability to predict regional climate change. As we have shown, these have the potential to predominate over global change. Because sedimentary deposits are a record, albeit imperfect, of both climate and climate change in a localized region (e.g a single sedimentary basin), GCM simulations are especially useful in heightening the awareness of geologists to some of the climatic processes that may cause sedimentary cycles. However, the subtlety of some climate signals means that only weak sedimentary responses may be generated. Such responses may only be preserved in special types of sedimentary regime (e.g. basins prone to anoxia, lakes, restricted epeiric shelves, etc.). There may be a poor chance of both recording and preserving such signals in more dynamic sedimentary regimes, such as flashy alluvial systems and basins affected by active tectonics. The apparently widespread preservation of such expectedly weak signals may say a great deal about the more general state of the Earth's depositional systems during the Jurassic, or it may suggest that our expectations require revision. The neglect of ocean feedbacks is clearly a limitation of our modelling work. The ocean model used has been described as a 'no surprise' ocean (Hansen et al. 1988). It does not include processes such as deep water formation, which is thought to be important in late Quaternary climate change (Broecker & Denton 1989). However, it is difficult to deduce too much from the results for the Quaternary. Models of ocean circulation in the Cretaceous (Barron & Peterson 1990) have suggested that deep water may form in low-latitude regions. Such a radically different ocean circulation
suggests that few lessons can be learnt from the Recent. Our simulations used a constant level of carbon dioxide. Ice core data show that carbon dioxide (and methane) concentrations varied over the last glacial cycle. Such changes were approximately in phase with the temperature changes and acted as a positive feedback process. Our Jurassic simulations would potentially show larger changes, especially in temperature, if carbon dioxide changes were included. This is especially true for the 'minimum seasonal forcing' simulation, where a relatively small decrease of carbon dioxide would result in the formation of a large Southern-hemisphere ice sheet. The land area covered by snow in the 'minimum seasonal forcing' run was c. 6-9 million km 2. If such an ice sheet did develop, then it could easily result in metre-scale variations of sea level. Our recent measurements of Cretaceous sea-surface temperatures (Sellwood et al. 1994) suggest that global mean temperature may have been significantly cooler than previously believed, with minimum mean equatorial values close to those of the present day and polar temperatures close to 0~ In the light of these findings, the climatic role of atmospheric carbon dioxide in determining Cretaceous climate becomes unclear, thus also casting a similar doubt upon its role in the Jurassic. The possible existence of an ice sheet for parts of the late Jurassic is controversial (Frakes & Francis 1988). However, it is possible that an ice sheet could have existed over parts of what is now Antarctica, particularly where elevations exceeded c. 1 km altitude, without having yet been detected. There are no late Jurassic deposits yet known in the heart of that continent. The existence of such an ice sheet would, however, also explain the presence of an apparently strong 100ka signal in the sedimentary record. This work, as part of the Mesozoic climate modelling project, has been funded by the NERC under grant GR3/7939, which we gratefully acknowledge. Figure 1 was produced by N. Hall.
References ANDERTON, R., BRIDGES, P.H., LEEDER, M.R. & SELLWOOD,B.W. 1979. A Dynamic Stratigraphy of the British Isles. George Allen and Unwin, London. ARKELL,W.J. 1956. Jurassic Geology of the World. Oliver and Boyd Ltd., Edinburgh. BARRON,E. J., & PETERSON,W. H. 1990. Mid-Cretaceous ocean circulation: Results from model sensitivity studies. Paleoceanography, 5, 319-337. , FAWCETT,P. J., POLLARD,D. & THOMPSON,S. 1993. Model simulations of Cretaceous climates: the role of geography and carbon dioxide. Philosophical Transactions of the Royal Society of London, B341, 307-316. BERGER, A., TRICOT, C., GALLI~e,H. & LOUTRE, M. E
1993. Water vapour, CO2 and insolation over the last glacial-interglacial cycles. Philosophical Transactions of the Royal Society of London, B341, 253-261. BERNER, R. A. 1992. Palaeo-CO2 and climate. Nature, 358, 114. BETrS, A. K. & MILLER, M. J. 1986. A new convective adjustment scheme. I: Observational and theoretical basis. Quarterly Journal of the Royal Meteorological Society of London, 112, 677-692. BEYDOUN, Z. R. 1988. The Middle East: Regional Geology and Petroleum Resources. Scientific Press Ltd, Beaconsfield. BRINKMANN, R. 1929. Statistich-biostratigraphische
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der Gesellschaft Wissenschaft zur Gb'ttingen, Maths-Physics, 13, 373 pp. BROECKER, W. S. & DENTON, G. H. 1989. The role of ocean-atmosphere reorganizations in glacial cycles. Geochimica et Cosmochimica Acta, 53, 2465-2501. CHAMLEY,H. 1989. Clay sedimentology. Springer-Verlag, Berlin. COHMAP MEMBERS, 1988. Climatic changes of the last 18,000 years: Observations and model simulations. Science, 241, 1043-1052. CROWLEY, T. J. & NORTH, G. R. 1991. Paleoclimatology. Clarendon Press, Oxford. KIN, K.-Y., MENGEL, J. G. & SHORt, D. A. 1992a. Modelling 100,000-year climate fluctuations in pre-Pleistocene time series. Science, 255, 705-707. , BAUM, S. K., & HYDE, W. T. 1992b. Milankovitch Fluctuations on Supercontinents. Geophysical Research Letters, 19, 793-796. YIP, K.-J.J. & BAUM, S. K. 1993. Milankovitch cycles and Carboniferous Climate. Geophysical Research Letters, 20, 1175-1178. FLOGEI, E. & FLOGEI-KHALER,E. 1992. Phanerozoic reef evolution: basic questions and data base. Facies, 26, 167-278. FRAKES, L. A. & FRANCIS, J. E. 1988. A guide to Phanerozoic cold polar climates from high-latitude ice-rafting in the Cretaceous. Nature, 333, 547-549. GATES, W. L., ROWNTREE, P.R .& ZENG, Q.-C. 1990. Validation of climate models. In: HOUGHTON,J. T., JENKINS, G. J. & EPHRAUMS, J. J. (eds) Climate ,
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Change: The IPCC Scientific Assessment, 93-130. Cambridge University Press, Cambridge. HALLAM, A. 1964. Origin of the limestone-shale rhythm in the Blue Lias of England: a composite theory. Journal of Geology, 72, 157-169. 1985. A review of Mesozoic climates. Journal of the Geological Society, London., 142, 433--445. 1986. Origin of minor limestone-shale cycles: climatically induced or diagenetic? Geology, 14, 609-612. 1993. Jurassic climates as inferred from the sedimentary and fossil record. Philosophical Transactions of the Royal Society, London, 341, 287-296. HANSEN, J., FUNG, i., LACIS, A., RIND, D., LEBEDEFF, S., RUEDY, R., & RUSSELL, G. 1988. Global climate changes as forecast by the Goddard Institute for Space Sciences three dimensional model. Journal of Geophysical Research, 93, 9341-9364. House, M. R. 1985. A new approach to an absolute timescale from measurements of orbital cycles and sedimentary microrhythms. Nature, 93, 721-725. 1986. Are Jurassic sedimentary microrythms due to orbital forcing? Proceedings of the Ussher Society, 6, 299-311. IMBRIE, J., BOYLE, E. A., CLEMENS,S. C., ET AL. 1992. On the structure and origin of major glaciation cycles. 1. Linear responses to Milankovitch Forcing. Paleoceanography, 7, 701-738. JANSA,L. 1972. Depositional history of the coal-bearing Upper Jurassic-Lower Cretaceous Kootenay Formation, Southern Rocky Mountains, Canada. -
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Geological Society of America Bulletin, 83, 3199-3222. KLOPFE1, W. 1917. Ober die Sedimente der Flachsee im Lothringer Jura. Geologische Rundschau, 7, 97-109. KUTZBACH, J. E. & WRIGHT, H. E. 1985. Simulations of the climate of 18,000 yr BP: Results for the North American/North Atlantic/European sector and comparison with the geological record. Quaternary Science Review, 4, 147-187. & GALLIMORE, R. G. 1989. Pangean climates: Megamonsoons of the megacontinent. Journal of Geophysical Research, 94, 3341-3357. -& GUETTER, P. J. 1986. The influence of changing orbital parameters and surface boundary conditions on climate simulations for the past 18,000 years. Journal of Atmospheric Science, 43, 1726-1759. Kuo, H. L. 1974. Further studies of the parameterization of the influence of cumulus convection on largescale flow. Journal of Atmospheric Science, 31, 1232-1240. MAASCH, K. A. & SALTZMAN, B. 1990. A low-order dynamical model of global climatic variability during the full Pleistocene. Journal of Geophysical Research, 95, 1955-63. MORCRETTE, J.-J. 1990. Impact of changes to the Radiation Transfer Parameterizations plus cloud optical properties in the ECMWF model. Monthly Weather Review, 118, 847-873. OGLESBY, R. J. 1990. Sensitivity of glaciation to initial snow cover, CO 2, snow albedo, and oceanic roughness in the NCAR CCM. Climate Dynamics, 4, 219-236. & PARK, J. 1989. The effect of precessional insolation changes on Cretaceous climate and cyclic sedimentation. Journal of Geophysical Research, 94, 14 793-14 816. PARK, J. & OGLESBY, R. J. 1990. A comparison of precession and obliquity effects in a Cretaceous paleoclimate simulation. Geophysical Research Letters, 17, 1929-1932. & 1991. Milankovitch rhythms in the Cretaceous: A GCM modelling study. Global Planetary Change, 4, 329-355. & -1992. The effect of orbital cycles on Late and Middle Cretaceous climate: a comparative GCM modelling study. In: Orbital Forcing and Cyclic Sequences. IAS Spectral Publication, ??????. RIND, D. & PETEET,D. 1985. Terrestrial conditions at the last glacial maximum and CLIMAP sea-surface temperature estimates: Are they consistent? Quaternary Research, 24, 1-22. - - - , KUKLA, G. & PETEET, D. 1989. Can Milankovitch orbital variations initiate the growth of ice sheets in a general circulation model? Journal of Geophysical Research, 94, 1285 ! - 12871. Ross, C. A., MOORE, G. T. & HAYASHIDA,D. N. 1992. Late Jurassic paleoclimate simulation - paleoecological implications for ammonoid provinciality. Palaios, 7, 487-507. SALTZMAN, B. & MAASCH, K. A. 1988. Carbon cycle instability as a cause of the late Pleistocene ice age oscillations: Modelling the asymmetric response. Global Biogeochemical Cycles, 2, 177-185. SELLWOOD, B. W. 1972. Regional environmental changes -
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across a lower Jurassic stage-boundary in Britain. Palaeontology, 15, 125-157. & PRICE, G. D. 1993. Sedimentary facies as indicators of Mesozoic palaeoclimate. Philo-
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~ & VALDES, P. J. 1994. Cooler estimates of Cretaceous tempratures. Nature, 370, 453-455. SHACKLETON,N. J. 1993. The climate system in the recent geological past. Philosophical Transactions of the Royal Society of London, 341,209-213. SLIN~O, A. 1985. Handbook of the Meteorological Office 1 l-layer Atmospheric General Circulation Model. Meteorological Office 20 DCTN 29. Meteorological Office, Bracknell. ~, BLACKBURN,M, BETTS, A., El" AL. 1994. Mean climate and transience in the tropics of the UGAMP GCM. Part I: Sensitivity to convective parameterization. Quarterly Journal of the Royal Meteorological Society, 120, 881-922. THUNNELL, R, RIO, D., SPROVIERI,R & RAFFI, I. 1991. Limestone-marl couplets: origin of Early Pliocene Trubi Marls in Calabria, southern Italy. Journal of Sedimentary Petrology, 61, 1109-1122. VALDES, P. J. 1993. Atmospheric general circulation models of the Jurassic. Philosophical
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& SELLWOOD,B. W. 1992. A palaeoclimate model for the Kimmeridgian. Palaeogeography, Palaeoclimatology, Palaeoecology, 95, 47-72. VAN BUCHEM,E S. P., MELNYY, D. H. & MCCAVE,I. N. 1992. Chemical cyclicity and correlation of Lower Lias mudstones using gamma ray logs, Yorkshire, UK. Journal of the Geological Society, London, 149, 991-1002. WEEDON, G. P. 1985, Hemipelagic shelf sedimentation and climatic cycles: the basal Jurassic (Blue Lias) of South Britain. Earth and Planetary Science Letters, 76, 321-335. 1993. The recognition and stratigraphic implications of orbital-forcing of climate and sedimentary cycles. In: WRIGHT, V. P. (ed.), Sedimentology Review~I, Blackwell Scientific Publications, Oxford, 31-50. & JENKYNS, H. C. 1990. Regular and irregular climatic cycles and the Belemnite Marls (Pliensbachian, Lower Jurassic, Wessex Basin). Journal of the Geological Society, London, 147, 915-918. WIGNALL, P. 1989. Sedimentary dynamics of the Kimmeridge Clay: tempests and earthquakes. Journal of the Geological Society of London, 146, 273-284. YEMANE, K. 1993. Contribution of Late Precambrian paleogeography in maintaining a temperate climate in Gondwana. Nature, 361, 51-54.
Constraints for using high-frequency sedimentary cycles in cyclostratigraphy E COTILLON
Universitd de Lyon, Centre des Sciences de la Terre, 27-43 Boulevard du 11 Novembre, 69622 Villeurbanne Cedex, France Abstract. High-frequency cycles can be used for setting accurate correlations within basins,
between sites of the world-ocean or between two different continents. Besides, since Gilbert (1895, Journal of Geology, 3, 121-127), the cycles forming marl-limestone couplets have been regarded as chronostratigraphic tools for estimating durations of some stages, formations and biostratigraphical units. But this use is dangerous if applied without caution. Firstly, the period of decimetric couplets can vary along sedimentary successions because it is negatively correlated to the sedimentation rate. Indeed, in a dilated and quickly deposited series, many low-order cycles, unapparent in slowly deposited sediments because of bioturbation, occur as centrimetric to decimetric units and can be gathered in groups differing from the 100 ka bundles. Examples from the Lower and Middle Cretaceous show that 21 ka cycles are well-expressed with sedimentation rates of 5-30 m Ma-1. Secondly, the recording of orbital cycles depends on the facies and on the characteristics of sedimentary fluxes. Thirdly, corrections must be done for the variation of some orbital frequencies with time. In summary, facies suitable for cyclostratigraphic work show moderate sedimentation rate, homogeneous lithologies with a balance of limestone and marl and are related to high sea-level periods which coincide generally with regular fluxes.
The search and application of one or several metronomes for putting rhythms into the Earth story is an old dream of stratigraphers. The knowledge of metronomes and of their characteristics (solar heat variations driven by astronomical cycles) does not exclude the true difficulty of this approach, which consists of the deciphering the sedimentary cycles induced by the metronomes, then in the using of cycles as tools for correlations and time-measurement. Before this can be done, cycles must display the following properties: (1) they should be effectively linked to astronomical forcings of given periodicities (allocyclicity); (2) they should have as large an extension in time and space as possible; (3) they should have a frequency well fitted to the most useful range of datings, i.e. spanning from some thousands to 1-2 Ma; (4) they should be easily recognized in a geological series, i.e. they should be preserved from bioturbation and from mechanical and chemical diagenesis effects. In this study the lithological features driven by very short (lunar, seasonal, annual and so on) global cycles are excluded. Only the so-called high-frequency cycles, corresponding to decimetric sedimentary units, are taken into account. They
form alternating series of marl-limestone couplets which are suitable for cyclostratigraphy, especially when they have formed in deep environments sheltered from major autocyclic controls. As early as 1895, Gilbert was convinced that alternating successions resulted from a global cyclic control and so could be used for measuring time. Indeed, he was impressed by their regularity in the Upper Cretaceous series and by their extent through large areas of the Western Interior Basin of the USA. This global control was thought to originate from the mechanisms which themselves control the Earth's motion in space. This followed the discovery by Newton, d'Alembert and Le Verrier of orbital parameter cycles, and the use by Croll (1875) of the 21 000a precessional cycles for explaining the glacial fluctuations through Pleistocene time. Following Gilbert, the 21 000 a period was recorded in the alternating succession of American Upper Cretaceous. As the latter was composed of c. 1000 cycles, it was estimated that the series was nearly 21 Ma long, a value not very far from those yielded by present radio-chronological methods of 24-35 Ma. This revolutionary approach was reemployed much later, after the Second World War when, following Milankovitch (1941), highfrequency sedimentary cycles were again associated with climatic or eustatic fluctuations and, in turn, to orbital cycle periods.
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescalesand Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 133-141.
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P. COTILLON
The global control of cycles is demonstrated through two properties: firstly, their wide geographical distribution which allows their use as correlation tools; secondly, their dependence on climatic forcings of orbital origin, proved among other techniques by harmonic analysis of sedimentary cycles, giving to them a chronostratigraphic value.
High-frequency cycles as correlation tools Many correlations of various extent and accuracy are based on high frequency cycles. Basin- and transcontinental-scale correlations using the space continuity of cycles have been successfully attempted in the Upper Cretaceous of the Western Interior, USA (Hattin 1971), in the Lower Cretaceous of Southeastern France (Cotillon et al. 1980), in the Cenomanian/Turonian transition between central Italy and Colorado (De Boer 1983). The thickness variation of cycles in the Lower Cretaceous has confirmed correlations between Southern France and two DSDP sites: the one off Florida (Site 534) and the second in the Gulf of Mexico (Site 535) (Cotillon & Rio 1984). Lastly, the number of cycles per drilling core, which is proportional to the sedimentation rate (Cotillon 1985), has been used to correlate the Hauterivian and Albian deposits of a dozen of the DSDP sites distributed in the Atlantic and Pacific Oceans. These correlations are based on variables evolving cyclically, for example, the carbonate content or the sedimentation rate, all related to a carbonate productivity more or less climatically controlled; they are possible only if the successions have accumulated in comparable environments and with close average sedimentation rates.
High-frequency cycles as chronostratigraphic tools If high-frequency cycles result effectively from forcings of given periodicities then they can be employed for time measurement. However, they correspond to complex signals made of several components which, in addition, are not recorded with the same intensity according at all locations. So, the precession signal and its modulation by eccentricity are easily recognizable in low-latitude deposits, while the obliquity signal is clearly expressed only in high latitudes (Van Woerkom 1953). In addition, the sedimentary record of astronomical signals can be modified by local features such as tectonics, bioturbation (Pestiaux & Berger 1984), bottom currents, climate and so on. Therefore, the raw signal must be deciphered by a specific process such as harmonic analysis, taking
into account the successive values of a variable liable to reflect climatic and, further, astronomical signals. This variable may be the CaCO 3 content, measured according to a given step unit, or the thickness of marl-limestone couplets. Whatever the course and the basic assumptions of harmonic analyses applied, most of them result in estimating nearly 21 000 a duration to marllimestone couplets. From this feature, and following Gilbert's (1895) example, the duration of some chrono or biostratigraphical units have been recalculated. This approach occurs in papers of Weedon (1989) for the Lias, Rio et al. (1989) and Ten Kate & Sprenger (1989) for the Lower Cretaceous, De Boer (1983) and De Boer & Wonders (1984) for the Middle Cretaceous of Umbria, Hart (1987) and Gale (1989) for the Cenomanian and Hilgen (1991) for the Mediterranean Plio-Pleistocene. On this basis, the bundles made of four to five cycles, noticed in some successions by many authors (e.g. Herbert et al. 1986, Goldhammer et al. 1987; Schwarzacher & Fischer 1982; Weedon 1989) have been considered to be the record of eccentricity cycles. However, in addition to the fact that some hiatuses may occur in seemingly continuous series, no evidence exists in favour of high-frequency cycles provided with constant periods along series of important duration and where two parameters, facies and sedimentation rate, are liable to change with time. In fact, according to some authors (e.g. Pestiaux & Berger 1984, Schwarzacher & Fischer 1982, Pestiaux et al. 1988; Clerc-Renaud 1988; Weedon 1989) it is unrealistic to suppose that sedimentation rate is constant over long time periods, particularly when the fithology of a succession is heterogeneous. The relations between cyclicity, facies and sedimentation rate discussed by the authors help to indicate constraints which must be taken into account with the use of cycles in stratigraphy.
Constraint of sedimentation rate Tie with cyclicity
Sedimentation rate and cycle frequency are liable to vary in the course of time. Recent researches (Cotillon & Rio 1984, Cotillon 1985, 1987) have shown some correlation between successions from several tethyan regions: southeastern basin of France, Lombardian Basin (northern Italy), Djebel Oust Basin (North Tunisia), DSDP Sites 534 (Central Atlantic) and 535 (Gulf of Mexico). In these areas, investigations have shown the occurence of a positive correlation between the sedimentation rate and the number of high-
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F i g . 1. Diagram illustrating the relation between cycles and sedimentation rate. (a) Sedimentation rate vs. number of marl-limestone couplets deposited per 105 a in Lower Cretaceous sections drilled at DSDP sites and investigated in southeastern France, Northern Italy and Northern Tunisia (after Cotillon 1987). (b) Illustration of (a) by a field feature: correlation between pelagic (A) and hemipelagic (B) series. Bed-interbed cycles in the Vocontian basin (A), may correspond to a bundle of beds and interbeds in the hemipelagic zone (B). (after Ferry & Monier 1987, simplified).
frequency marl-limestone cycles counted through given intervals (Fig. la). The relation is linear up to 40 m Ma -1 in sedimentation rate. Only cycles > 5 cm thick, namely preserved from a bioturbation of medium intensity (e.g. Guinasso & Shink 1975; Peng et al. 1979, Nittrouer et al. 1984, De Master et al. 1985) are taken into account. It follows that in two alternating successions deposited during the same time interval, but represented by unequal sedimentary columns testifying to unequal sedimentation rates the mean duration of lithologically expressed cycles > 5 cm thick is dissimilar. It means that a bed-by-bed correlation between the two series, i.e. a 'first degree correlation' is not possible. The relation between sedimentation rate and the mean duration of cycles was already suspected by Arthur et al. (1984), but especially by ClercRenaud (1988) in the Upper Jurassic - Lower Cretaceous succession of Fontcalent (NE Spain), when she highlights, in the power spectra of cycles, frequency shifts linked to sedimentation rate variations. The same relation has been directly demonstrated in the southeastern basin of France by Ferry & Monier (1987). On the north side of the Mont Ventoux Chain, located on the palaeoslope separating the Vocontian Basin from the Provence Platform during the Lower Cretaceous, a block tilting created an active subsidence resulting in a high sedimentation rate reaching 470 m Ma -I for
the Lower Barremian. Now the latter is composed of a hemipelagic succession showing a major alternation where each cycle is a 10-15 m thick bundle of couplets with a more marly lower part (Fig. lb). This succession can be correlated very accurately with the Vocontian pelagic Lower Barremian where limestone beds and marly interbeds alternate, giving 0.5-1 m thick cycles, while the sedimentation rate is c. 30 m Ma -1. Each pelagic decimetric couplet fits exactly with a hemipelagic metric bundle. Note that the ratio between cycle thicknesses and sedimentation rates are nearly the same and close to 16. The non-homothetic transformation of a single cycle into a composite one when sedimentation rate increases implies the occurence of lower-order cycles within the former. But these thin cycles are not lithologically expressed owing to bioturbation which acted systematically upon fresh-deposited Vocontian sediments. In fact, examples from the southeastern basin of France have shown that little cycles disappear and become amalgamated within an homogeneous bed when bioturbation increases; the reverse is true when the latter decreases temporarily, owing to suboxic conditions; centrimetric cycles are then individualized as in the Upper Valanginian of the Angles succession. The hypothesis of potential for minor order cycles to occur within decimetric marl-limestone couplets is verified in the Lower Cretaceous series
136
r,. COTILLON
of some Atlantic DSDP sites. For instance, at Sites 534 and 535, where the decimetric alternation of clear limestone and dark marl exhibits a specific laminated structure, developed essentially in marls, testifying to a lack of bioturbation. Laminae are arranged in at least four orders; the smallest, often destroyed or deformed by compaction, correspond to 8-25 p m thick cycles regarded as varves (Cotillon & Rio 1984, Cotillon i991). The major feature observed here is the occurence of a continuum of superposed cycle orders from the varves to the decimetric marl-limestone alternation. It is tempting to link this range of lithological cycles to those of climatic origin with periods spreading from 1 to 21 000 a or more (e.g. Pestiaux et aI. 1988). An increase of sedimentation rate, affecting an alternating and laminated succession of an Atlantictype, leads to a homothetic transformation increasing the thickness of all cycles, the number of which remains constant (Fig. 2). The same transformation applied to an unlaminated alternating series, because of bioturbation, will get some cycles
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Signification o f bundles The previous features lead to the consideration again of the signification of cycle bundles. When these bundles succeed one another regularly, with a constant average number of beds between 4 and 5 [4.03 in the Majolica Formation; 4.88 in the Scaglia Bianca (Schwarzacher & Fischer 1982), 5 in the Middle Triassic of Italian Dolomites (Goldhammer et al. 1987)], this signifies that two superposed periodicities occur with a ratio of c. 5, like those of precession and eccentricity. But, in other cases, the signification of bundles is different. The average number of cycles per bundle may exceed 5 [5.9 and 7.5 in the Irish Carboniferous (Schwarzacher 1964, 1975); 7 in the Scaglia Cinerea (Lower Eocene at Gubbio, Italy)]. These features are close to that of the Lower Barremian of Mont Ventoux where the bundles are composed of 15-25 unit-cycles. When bundles of 3-6 cycles appear, either isolated, or grouped by two, three or four, within regular alternating successions it can be assumed these bundles correspond to temporary accelerations of sedimentation rate. They increase abnormally the number of cycles in a given series. From this follows this question: what are the most appropriate sedimentation rates in a series which allows the use of cycles as chronologic tools ?
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Fig. 2. Effect of sedimentation rate variation on cyclicity. It is assumed that, < 5 cm thick, a cycle is destroyed by bioturbation. (a) In the unbioturbated succession, all the cyclicity orders are individualized from the marl-limestone couplet to the lamination (not represented here); one cycle > 5 cm thick is taken into account; with an increase of sedimentation rate all the cycles grow thicker so that five of them outpace 5 cm. (b) In the bioturbated succession only one order of cyclicity (the marl-limestone couplet) is represented; the others, < 5 cm, have been destroyed by bioturbation; with an increase of sedimentation rate, the five potential cycles reaching, then outpacing 5 cm are individualized and may be taken into account.
The first example is taken from the Vocontian Basin (SE France), thought to represent sediment at c. 1000 m depth on average and corresponding to open and well-oxygenated environments. In the Lower Cretaceous stratotypic succession of Angles, marl-limestone decimetric couplets have been counted from the Berriasian to the Barremian. The total number of cycles is 1080 (Rio et al. 1989) for a mean total duration of 22.8 Ma (average value from five different geochronological scales). Consequently, the mean duration of a cycle is close to 21 000 a. The succession being 635 m thick, the average sedimentation rate is 28.8 m Ma -1. The Angles Succession shows a general composition balanced between marls and limestones but with a prominent carbonate fraction at the two extremities: the Berriasian and the BarremianBedoulian (Cotillon 1971). Bed bundles are relatively few, this fact agreeing with a rather low
CONSTRAINTS IN CYCLOSTRATIGRAPHY sedimentation rate. Besides, important sections including the Upper Berriasian, the Lower Valanginian and the Middle Hauterivian, exhibit a very regular pattern with nearly equal proportions between beds and interbeds. A second example is from the Appennines of Umbria (Central Italy) where the Middle Cretaceous testifies to a deeper and less oxidized environment than in the Vocontian Basin. A cored succession at Piobbico is characterized by a weak sedimentation rate of 5 m Ma -1 (Herbert & Fischer
Transmission .Orn
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1986; Premoli Silva et al. 1989); it shows 87 m of a very homogeneous series made of couplets, 10 cm thick on an average, grouped in bundles and super-bundles. These three orders of cycles are assumed to correspond to the precession and eccentricity cycles (Fig. 3). The small thickness of cycles and their good preservation are explained by dysaerobic to anaerobic depositional periods illustrated by layers of laminated black shales enriched with organic carbon. These periods of poor bottom circulation have prevented a major bioturbation. The examples of Angles and the Appennine Successions could be regarded as limit possibilities of cyclostratigraphy based on the precession cycles. The upper limit is located at Angles, where some tendancies to bundle individualization occurs, signifying episodic periods of sedimentation rate acceleration. The lower one, in Umbria, shows that in spite of a lithological contraction of cycle orders and of the thickness of marl-limestone couplets, the record of precession cycles is preserved thanks to a relatively closed environment having prevented major bioturbation. So, between 5 and 30 m Ma -1, the sedimentation rate may allow the individualization of 21 000 a cycles and their use for geochronological objectives. However, in the Piobicco series, cyclicity displays the same signification, from a periodicity point a view, whatever the level, owing to the homogeneous facies and regular cycles. On the contrary, the vertical facies succession at Angles is unhomogeneous, with carbonate formation in the Lower Berriasian, the Lower and Upper Hauterivian, the Upper Barremian and marly intervals with sparse carbonate beds in the Middle and Upper Valanginian, the Middle Hauterivian. From this follows this question: are the orbital signals similarly recorded in the two kinds of facies?
Constraints of facies
A
B Fig. 3. Part of the cored Upper Albian Piobbico series (Umbria, Central Italy). Analytical process, based on the brightness (A) and CaCO3 content (B) of deposits, clearly displays three superposed cyclicities corresponding to" fundamental marl-limestone couplets; bundles of couplets (lettered); and super-bundles (dashed lines). After Herbert & Fischer 1986.
A first answer comes from the investigations of Clerc-Renaud (1988) on the Upper Jurassic-Lower Cretaceous alternating series at Fontcalent (Eastern Spain). The spectral analysis of this succession has pointed out frequent shifts occurring particularly at the transition between lithological units. A second answer lies in the Angles series (Fig. 4). Harmonic analysis of cycles based on their constant duration, i.e. 21 000 a (Rio et al. 1989) has shown that most carbonate intervals of the series (Berriasian, Lower Valanginian and Barremian) record nearly identically all the cycle periods except the shortest; besides, the longest of them (up to 1 Ma or more) are well recorded. Conversely,
138
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the most marly intervals, for instance, the Upper Valanginian, have recorded the shortest periods, essentially from 20 to 100 Ka. Very similar features have been observed by Clerc-Renaud (1988) in Upper Jurassic - L o w e r Cretaceous succession at Fontcalent. Two possible explanations are suggested. A theory based on the modality o f sedimentation.
In the Angles Series, carbonate-rich formations, such as the Hauterivian and Berriasian, exhibit a rather high rate of sedimentation (30.6 and 30 m Ma -], respectively, against 28.8 m Ma -1 for the total Berriasian-Barremian interval). This feature agrees with previous results (Cotillon 1985), showing an increase of sedimentation rate from marls to limestones within basic couplets of alternating deposits. But the Berriasian at Angles is partially composed of thick carbonate units formed by amalgamated cycles separated only by millimetric clayey seams, or dry diastems, or even without physical boundary. These features imply an irregular and discontinuous sedimentation suited for the unusual recording of long cycles. This last trend is more developed in the Barremian succession (Fig. 4) which is also carbonate-rich but where the sedimentation rate is reduced (15 m Ma-1); so, a more important effect of hiatus between
cycles, i.e. on the bed contacts, may be suspected as well as an important loss of material during the burial diagenesis. Consequently, a bed-scale rapid deposition may be included in a slowly deposited formation. This might explain the seemingly inconsistent results obtained with the same section, through a harmonic analysis differing from that used by Rio et al. because the latter is based on the variation in CaCO 3 content (Giraud et al. this volume). Conversely, the sedimentation is slower, more regular and continuous in marly intervals such as the Valanginian. This fact is illustrated by a transitional contact between carbonate beds and marly interbeds. This character, and the less active bioturbation in marls favour the recording of short cycles (see also Arthur et al. 1984). On the contrary, long cycles often remain undetected in marls where beds are sparse and variations of CaCO 3 content are not always lithologically expressed. This explanation must be considered cautiously given the results obtained using a different method (Giraud et al. this volume). A theory based on outer controls o f sedimentation.
Terrigeneous sources would react more quickly than carbonate facies (namely platforms and plankton) to the climatic oscillations related to orbital cycles. Clerc-Renaud (1988) reached the same conclusion but also took into account the distance separating sources from depositional areas. In summary, cyclostratigraphy must be employed essentially in series where carbonate and marls are nearly equally represented and form regular altemations. It must be avoided for facies which are too restricted, either in marls or limestones.
Constraint
of sedimentary
fluxes
The characteristics of material fluxes controlling the deposition of alternating successions may influence the recording of orbital cycles. A previous study based on image analysis of oceanic cores (Cotillon 1992) has shown that, during low sea level, deposits with abundant nutrients of continental and marine origin result in especially high carbonate fluxes, marked by a great variability (Fig. 5). These features could contribute to explaining the weak record of short harmonic periods as observed in the Berriasian and Barremian carbonate intervals of Angles Succession, which essentially correspond to low-stand deposits. Conversely, during high sea-level periods, because of the withdrawal and interference of sources, and of the screen effect caused by drowned
CONSTRAINTS IN CYCLOSTRATIGRAPHY
.
.
139
B
Fig. 5. Variability of material fluxes held to characterize the response of sources forced by external influences: climate, tectonics sea-level variations. (A) Low sea-level. Continent and ocean are close. From the sources there is a good transmission (by amplitude and frequency) of signals to the oceanic domain via the material fluxes. The vicinity of sources makes them not very interferent causing a high flux variability. Abundant fluxes and important sedimentation rates can be expected. 1, Silicate terrigenous, carbonate (clastic) and organic (land-derived) fluxes; 2, planktonic flux; 3, detrital flux; 4, activated up-wellings. (B) High sea-level. Continent and ocean are distant. Poor transmission of signals to the ocean because they are partly trapped by epicontinental domain. The remoteness of sources makes them interfere, lowering the flux variability. Reduced fluxes and sedimentation rates can be expected. 1, Silicate terrigenous and organic (land-derived) fluxes; 2, partial flux diversion; 3, weakened residual flux; 4, planktonic flux; 5, detrital flux.
continental margins, fluxes decrease and become regular while sedimentation moves towards a more marly character.
Constraint of the age of deposits Lastly, cyclostratigraphy must take into account the variation of precession and obliquity periods with time, owing to the continuous evolution of the Earth-Moon system. So, at the beginning of the Phanerozoic, these periods were 17 000 and 28 000 a long, respectively, while they are 21 000 and 41 000 a today (Algeo & Wilkinson 1988; Berger et al. 1989). So, the precession cycle, which is the more convenient cyclostratigraphic unit because it is the more clearly recorded in deposits, must not be quantified as 21 000 a by referring to the Present. De Boer & Wonders (1984) have suggested this unit be named 'the Gilbert', which is not an intangible quantified value.
Conclusion Given the outer limits of forcing of orbital origin undergone by sedimentation, cyclostratigraphy, based on high-frequency cycles, seems to be the ideal way for accurate datings whatever the age, environment and diagenetic transformation of deposits. But this method must be used cautiously in successions showing some peculiar characteristics. (1) Successions must be as regular and continuous as possible, then the following depositional environments have to be excluded: those of too strongly dynamic waters which generate erosional features; those of coastal areas where sea-level variations induce sedimentary cycles but also emersions and their associated gaps; and localities where significant detrital or reefal sedimentation disturbs the record of cycles. Conversely, sheltered and relatively deep environnements, down the continental slope, are privileged. In them, planktonic productivity and
140
P. COTILLON
detrital fluxes have the best possibility to express orbital messages with maximum faithfullness. (2) Successions should show very weak diagenetic alterations. (3) In open and moderatly deep basins, where deposits are bioturbated, the sedimentation rate must fall within the 20-30 m Ma -1 range. In deeper and more restricted basins, with or without a weak bioturbation, cycles are preserved in spite of a sedimentation rate as weak as 5 m Ma -1 or less. Accurate correlation through high-frequency cycles requires close sedimentation rates in the compared series. (4) The most favourable facies for using cyclo-
stratigraphy is a balanced alternation between marl and limestone, deposited during high sea level periods related to the maximum regularity of material fluxes and lithological cycles. Such are the conditions for the best record of an orbital periodicity spectrum. Formations displaying major carbonate enrichment should be avoided because their deposition is too irregular and mineral transfer too important during the diagenesis. (5) Lastly, the duration of reference orbital cycles must be defined according to the age of deposits, taking into account their changes through geological time.
References ALGEO, T. J. & WILKINSON, B. H. 1988. Periodicity of mesoscale phanerozoic sedimentary cycles and the role of Milankovitch orbital modulation. Journal of Geology, 96, 313-322. ARTHUR, M. A., DEAN, W. E., BOTrJER, D. & SCHOLLE, P. A. 1984. Rhythmic bedding in Mesozoic Cenozoic pelagic carbonate sequences : the primary and diagenetic origin of Milankovitch like cycles. In : BERGER,A. L., IMBRIE,J., HAYS,J., KUKLA,G. (~ SALZMAN,B. (eds), Milankovitch and climate, 1. Reidel, Dordrecht, 191-222. BERGER,A., LOUTRE,i . F. & DEHANT,V. 1989. Influence of the changing linear orbit on the astronomical frequencies of pre-Quatemary insolation patterns. Paleoceanography, 4, 555-564. CLERC-RENAUD, T. 1988. Recherche de pdriodicitds relatives & la variation des param~tres orbitaux dans la sddimentation alternante du Jurassique supdrieur - Crdtacd infdrieur. Th~se de Doctorat de l'Universit6 Paris VI. COTILLON,e., 1971. Le crdtacd infdrieur de l'arc subalpin de Castellane. Stratigraphie et sddimentologie. Mrmoires du Bureau de Recherche Gdologique et Mini~re, 68. 1985. Les variations h diffrrentes 6chelles du taux d'accumulation srdimentaire dans les srries prlagiques alternantes du Crrtac6 infrrieur, consrquence de phrnom~nes globaux. Essai d'rvaluation. Bulletin Socidtd Gdologique de France, Srr. 8, I, 1, 59-68. 1987. Bed scale cyclicity of pelagic Cretaceous successions as a result of world-wide control. Marine Geology, 78, 108-123. 1991. Varves, bed and bundles in pelagic sequences and their correlation (Mesozoic of SE France and Atlantic). In: EINSELE,G., RICKEN,W. & SEILACHER, A. (eds) Cycles and Events in Stratigraphy, Springer, Berlin, 820-839. 1992. Search for eustacy record in deep tethyan deposits through the study of sedimentary flux variations. Application to the Upper TithonianLower Aptian series at DSDP Sites 534 (Central Atlantic). Palaeogeography, Palaeoclimatology, Palaeoecology, 91, 263-275. ~, FERRY, S., GAILLLARD,C., JAUTEE, E., LATREILLE, G. & RIO, M. 1980. Fluctuations des param~tres du
milieu marin dans le domaine vocontien (France SE) au Crrtac6 infrrieur : mise en 6vidence par l'6tude des formations marno-calcaires alternantes. Bulletin Socidtd Gdologique de France, SEt. 7, XXII, 1,733-742. & Rio, M. 1984. Cyclic sedimentation in the Cretaceous of D.S.D.P. sites 535 and 540 (Gulf of Mexico), 534 (Central Atlantic) and in the Vocontian basin (France). Initial Reports of the Deep Sea Drilling Programme, 77, 339-378. CRUEL, J. 1875. Climate and Time in their Geological Relations: A Theory of Secular Change in the Earth's Climate. Appleton, New York. DE BOER, P. L. 1983. Aspects of Middle Cretaceous pelagic sedimentation in Southern Europe. Geologica Ultraiectina, 3, 112 pp. & WONDERS, A. A. H. 1984. Astronomically induced rhythmic bedding in cretaceous pelagic sediments near Moria (Italy). In: BERGER, A. L., IMBRIE, J., HAYS, J., KUKLA, G. & SALZMAN,B. (eds) Milankovitch and climate, I, Reidel, Dordrecht, 177-190. DE MASTER, D. J., MCKEE, B. A., NITIROUER, C. A., BREWSTER, D. C. & BISCAYE,P. E., 1985. Rate of sediment reworking at the HEBBLE Site based on measurements of Th-234, CS-137 and Pb-210. Marine Geology, 66, 1/4, 133-148. FERRY, S. • MONIER, P. 1987. Correspondances entre alternances marno-calcaires de bassin et de plateforme (Crrtac6 du Sud-Est de la France). Bulletin Socidtd Gdologique de France, Serie 8, III, 5, 961-964. GALE, A. S. 1989. A Milankovitch scale for Cenomanian time. Terra Nova, 1,420--425. GILBERT, G. K. 1895. Sedimentary measurement of geologic time. Journal of Geology, 3, 121-127. GOLDHAMMER,R. K., DUNN,P. A. & HARDIE,L.A. 1987. High-frequency glacio-eustatic sea level oscillations with Milankovitch characteristics recorded in middle Triassic platform carbonates in Northern Italy. American Journal of Science, 287, 853-892. GUtNASSO, N. L. & SHINK, D. R. 1975. Quantitative estimate of biological mixing rates in abyssal sediments. Journal of Geophysical Research, 80, 21, 3032-3043. HART, M. B. 1987. Orbitally induced cycles in the
CONSTRAINTS IN CYCLOSTRATIGRAPHY chalk facies of the United Kingdom. Cretaceous
Research, 8, 335-348. HATTIN, D. E. 1971. Widespread, synchronously deposited burrow-mottled limestone beds in Greenhorn limestone (Upper Cretaceous) of Kansas and Central Colorado. Bulletin of the
American Association of Petroleum Geologists,
55, 412-431. HERBERT, T. D. & FISCHER, A. G. 1986. Milankovitch climatic origin of Mid-Cretaceous black-shale rhythms in Central Italy. Nature, 321, 739-743. --, STALeARD, R. E & FISCHER, A. G. 1986. Anoxic events, productivity rhythms and the orbital signature in a mid-cretaceous deep-sea sequence from central Italy. Paleoceanography, 1, 495-506. HIeGEN, F. J. 1991. Astronomical forcing and geochronological application of sedimentary cycles in the Mediterranean Pliocene/Pleistocene. Geologica Ultraiectina, 93. MILANKOVITCH, M. 1941. Kanon der Erdbestrahlung
and seine Anwendung auf des Eiszeitenproglen. Academic Royale Serbe, Belgrade, special edition, 133. NIT~OUSER, C. A. DE MASTER, D. J., MCKEE, B. A., CUTSHALe, N. H. & LARSEN, I. L., 1984. The effect of sediment mixing on Pb-210 accumulation rates for the Washington continental shelf. Marine Geology, 54, 201-222. PENG, T. H., BROECKLER,W. S. & BERGER, W. H., 1979. Rate of benthic mixing in deep-sea sediments as determined by radioactive tracers. Quaternary. Research, 11, 141-149. PESTIAUX, P. & BERGER, A. 1984. Impacts of deep-sea processes on paleoclimatic spectra. In: BERGER, A. L., IMBRIE, J., HAYS, J., KUKLA, G. & SALZMAN, B. (eds), Milankovitch and Climate,l, Reidel, Dordrecht, 493-510.
141
--,
VANDER MERSCH, I., BERGER, A. & DUPLESSY,J. C. 1988. Paleoclimatic variability at frequencies ranging from l cycle per 10,000 years to 1 cycle per 1,000 years: evidence for nonlinear behaviour of the climate system. Climatic Change, 12, 9-37. PREMOLI SILVA, I., RIPEPE, M. & TORNAGHI, E 1989. Planktonic foraminiferal distribution record productivity cycles : evidence from the AptianAlbian Piobbico core (Central Italy). Terra Nova, 1,443-448. RIO, M., FERRY, S. & COT~eeoN,E 1989. Periodicit6s dans les s6ries p61agiques alternantes et variations de l'orbite terrestre. Exemple du Cr6tac6 inf6rieur dans le Sud-Est de la France. Comptes Rendus Acad(mie des Sciences, Paris, 3119, (II), 73-79. SCHWARZACHER,W. 1964. An application of statistical time series analysis of a limestone-shale sequence. Journal of Geology, 72, 195-213. 1975. Sedimentation models and quantitative stratigraphy. Developments in Sedimentology, 19, Elsevier, Amsterdam. -& FISCHER, A. G. 1982. Limestone-shale bedding and pertubations of the Earth's orbit. In: EINSELE,G. & SEILACHER, A. (eds), Cyclic and Event Stratification, Springer, Berlin, 720-95. TEN KATE, W. G. & SPRENGER, A. 1989. On the periodicity in a calcilutite-marl succession (SE Spain). Cretaceous Research, 10, 1-31. VAN WOERKOM,A. J. J. 1953. The astronomical theory of climatic changes. In: SHAPLEY,H. (ed.), Climatic
Change. Harvard University Press, Cambridge, MA, 147-157. WEEDON, G. P. 1989. The detection and illustration of regular sedimentary cycles using Walsh power spectra and filtering, with examples from the Lias of Switzerland. Journal of the Geological Society, London, 146, 133-144.
Periodicities of carbonate cycles in the Valanginian of the Vocontian Trough: a strong obliquity control E G I R A U D 1, L. B E A U F O R T 2 & E C O T I L L O N 1
1Centre des Sciences de la Terre, Universit( de Lyon, 27-43 Boulevard du 11 Novembre 69622 Villeurbanne Cedex, France 2Laboratoire de G(ologie du Quaternaire, CNRS Luminy case 907, 13288 Marseille Cedex 09, France Abstract: A high-resolution study of variations in carbonate productivity, quantified by CaCO 3
content and by colour intensity (grey level) for a Valanginian rhythmic succession is presented in the Vocontian Trough. Spectral analysis techniques reveal the presence of strong CaCO3 cycles linked to cyclic variations in the Earth's orbit. The results of the analysis suggest that obliquity cycle is the most clearly defined signal whereas the presence of CaCO3 cycles related to the precession signal is less distinct. The transition from limestone-dominant alternations in the lower Valanginian to marl-dominant alternations in the upper Valanginian is characterized both by an increase of sedimentation rate and a change from the precession as the dominant forcing in the lower, calcareous part of the Valanginian, to obliquity as the dominant forcing for the upper marly part of the Valanginian. An orbitally calibrated chronology is presented for the Valanginian based on the identification of the carbonate cycles. The application of band-pass filterings to the original carbonate record allows the extraction of 91 carbonate cycles related to precession and 137 cycles related to obliquity. The proposed duration for the Valanginian stage is 7.04 Ma.
In pre-Quaternary alternating pelagic successions, the thickness of a marl-limestone couplet is often the main lithological signal used for the search for orbital cycles in the sedimentary record (De Boer 1982; Dean et al. 1981; Fischer & Schwarzacher 1984; Schwarzacher & Fischer 1982; Rio et al. 1989; Weedon 1989). However, the use of CaCO 3 content rather than bed-interbed thicknesses provides some advantages. The data do not suffer from the ambiguity of the qualitative designation of a couplet. In a more continuous record the calcium carbonate content in pelagic sediments is more valuable for studies of orbital forcing (Mayer 1991); information, i.e. higher resolution, of variations concerning periods are accurate and smaller than the duration of a bed-interbed couplet. A convincing demonstration of the predominant astronomical forcing in the development of the carbonate system is furnished by the works of Hilgen (1991) in the Mediterranean Pliocene and Pleistocene, and Mayer (1991) in the Central Pacific Plio-Pleistocene. The Vocontian lower Cretaceous alternations representing variations in carbonate productivity are chosen as the framework for a study of carbonate content fluctuations. The rhythmic bedding (from bed-interbed thickness) has been previously linked with cyclic variations in the Earth's orbit (Rio et al. 1989), but the large dis-
persion of frequencies, with only a narrow part in the Milankovitch frequency band, was not conclusive. A new quantitative study on those cyclic sequence (from bed-interbed thickness), through application of the Walsh spectral method (Huang et al. 1993), reveals for the ValanginianHauterivian interval the importance of obliquity in controlling the rhythmic sedimentation, whereas the precession cycle is less important. These results are interesting, especially as they are not in agreement with other examples of astronomical forcing in pelagic successions. On the one hand, De Boer (1991) concluded, from a compilation of literature examples, that sedimentary systems located at c. 30 ~ (the paleolatitude of the Vocontian Basin for the studied time period) are particularly sensitive to the precession of the Earth's axis. On the other hand, Fischer (1986), in a review on pelagichemipelagic series, remarks that many long sequences (such as the Vocontian alternations) are compound with segments reflecting both the precession and eccentricity cycles with, in other parts, the obliquity cycle. Therefore, the main goal of this present study is: (1) To demonstrate the connection between cyclic variations in CaCO 3 content and astronomical parameters. (2) To determine accurately the periodicities of carbonate cycles and then to compare them with the results of Huang et al.
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescalesand Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 143-164.
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F. GIRAUD ET AL.
144
(1993). This is performed by spectral analysis of the resulting carbonate record, at two different scales but focuses on the Valanginian stage using a high resolution analysis from a drilling core of an upper Valanginian set of beds with determination of variations in colour intensity concurrently with the quantification of carbonate content; and also a study of the carbonate content record through the whole Valanginian. On this an estimate of the duration of the Valanginian on the basis of the carbonate cyclicity is proposed. G e n e r a l i t i e s on s t u d i e d a r e a
Geological setting The Angles Section (southern Subalpine Ranges) is chosen to illustrate the Vocontian alternations (Fig. 1). This selection is guided by the following features: (1) Angles is the hypostratotype of the Valanginian stage (Busnardo & Thieuloy 1979); (2) it has a typical sequence of pelagic limestonemarl couplets; (3) the access and exposure are good; (4) tectonic disturbances are minor; and (5) no biostratigraphic hiatuses are observed. The Vocontian Trough is defined sensu stricto (Cotillon 1984) as a relatively deep pelagic sedimentation area with cyclic deposits where sedimentation is mainly characterized by vertical fluxes of fine pelagic biocarbonate and clay particles. From the late Jurassic, and throughout the early Cretaceous, this basin was open towards the Tethys
and surrounded by slopes and platforms with hemipelagic facies and temporarily by shallow carbonate banks. At that time, the tectonic regime was extensional. Platform margins were split into tilted blocks by NW-SE and WSW-ENE faults generating high subsidence contrasts (Cotillon 1985a). The sedimentation was influenced by this tectonic regime as demonstrated by the occurrence of slumps into the basin and by rapid lateral changes of thickness and facies. From the end of the Jurassic until the Aptian, the Vocontian Trough was located at a palaeolatitude of 25-30 ~ (Savostin et al. 1986). In terms of basin paleogeography the Valanginian stage is divided in two periods. The first begins with a transgressive episode emphasizing the transition from a system of carbonate platforms initiated during the Beniasian towards a system of deep-water carbonates with deposition of alternating cycles in the basin where fine terrigenous materials are mixed with biogenic carbonate in various proportions. The second episode is marked by a deepening trend. The basin is largely open to the Tethys and there are also communications with the Boreal basins by way of the Jura, Franche-Comt6 and Alsace platforms (Cotillon 1984).
Lithostratigraphy & biostratigraphy (Fig. 2) Lithology. At Angles the Valanginian is 240 m thick. The alternations consist of a repetition of
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Fig. 1. Paleogeography of the Subalpine Basin from the Late Jurassic to the Early Cretaceous.
VALANGINIAN OF THE VOCONTIAN TROUGH
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146
E GIRAUD ET AL.
decimetric to metric binary cycles rhythms with beige calcareous beds (c. 85% CaCO3) and dark grey marly interbeds (c. 65% CaCO3) (Cotillon et al. 1979). In the upper Valanginian alternation with a prominent marly part occurs with dark marly interbeds and a concomittant decrease of the average carbonate content from 79% for limestone beds to 59% for marls; the cycles appear as composite elements or bundles: for instance, the studied set of 10 beds located at the top of the Saynoceras verrucosum Zone. This set corresponds to a key horizon averaging 12 m thick, composed of 10 closely spaced principal beds and numerous minor clayey carbonate layers (Fig. 3). As shown by Cotillon et al. (1980), this horizon exists all over the Vocontian Basin and can be identified in 30 sections. The whole succession is intensively bioturbated, with limestone beds and marly interbeds both affected (Gaillard 1984). Some slumping disturbs the lower Valanginian but the continuity of the succession remain unaffected.
Biostratigraphy. Ammonites represent almost the totality of the nectonic macrofauna. Six zones and 12 subzones have been recognized at Angles and were adopted subsequently for the southeast of France (Busnardo & Thieuloy 1979). The associated microfauna (calpionellids, foraminifera and ostracods) is poor and provides no precise information concerning the biozonal subdivision. For the foraminifera, a zonation based on the appearance of some species without presumption of their extinction has been proposed (Moullade 1979, 1984). The Valanginian nannoflora is well preserved and is composed of coccoliths and nannoconids. Two biozones are identified (Cretarhabdus crenulatus and Calcicalathina oblongata) (Manivit 1979).
Origin of the rhythmic variations The primary cyclic variations that cause the sedimentary pattern of the Lower Cretaceous Vocontian alternations can generally be attributed to periodic fluctuations in (1) carbonate dissolution, (2) biogenic carbonate production; or (3) terrigenous dilution (ROCC Group 1986). Dissolution as the primary factor can be excluded, since the study of radiolarians and carbonate debris have shown no indications of this (Darmedru et al. 1982) and as the water depth of the basin never exceeded 1500 m (Ferry 1976); [for the Valanginian, it was estimated at 600 m (Donze 1979)]. The scarcity of intercalations of distal siliciclastics or carbonate turbidites indicates that dilution by terrigenous clastics could not have played a decisive role either. The following observations suggest that this cyclic
sedimentation records mainly variations in carbonate production (Cotillon 1985b): (1) nanoplanktonic debris (nannoconids and coccolithophorids) represents the bulk of the micrite both in the limestone beds and in the carbonate component of the marly interbeds; (2) the carbonate component is more important than those of the argillaceous component.
Methods Carbonate and colour analysis Samples for CaCO 3 analysis were taken at constant spacing without regard to the lithology encountered (243 samples spaced at 5 cm for the cored set). The high resolution sampling of the core was performed not only to obtain a detailed carbonate record but also to choose a good sampling interval for the whole Valanginian section. The objective was to find a sampling interval which met two contradictory conditions: it must be large enough to avoid too much field sampling and must be close enough to preserve a maximum the amount of information. From the 243 data points available, carbonate curves have been drawn for different sampling steps (between 10 and 100 cm). The field observation shows that the curve realized with a constant 50 cm space is satisfactory (the principal beds composing the set are identified) whereas with a 60 cm space some information is lost. Then, 448 samples equally spaced at 50 cm were taken for the whole series. The levels with slumps were not sampled. The bulk carbonate content was measured through the gas-volumetricaly method; the uncertainty in the CaCO 3 measurements is 1%. Variations in the colour intensity within the core have been determined by numerical image processing, leading to 256 grey levels being distinguished (for the method see Giraud et al. 1991), with the same spacing as for the carbonate content.
Spectral analysis Two different spectral analytic methods have been applied to the carbonate content and colour records: (1) the standard Blackman & Tukey method, which computes the fast Fourier transform of the autocorrelation function; (2) the maximum entropy spectral analysis, which computes a set of autocorrelations and determines the power spectral density of the most random time series. Detailed explanations of their principles are given in Hinnov & Goldhammer (1991), Weedon (1991) and Ten Kate & Sprenger (1992).
VALANGINIAN OF THE VOCONTIAN TROUGH
CaC03%
uack
30 40 50 60 70 80 90 100 I
I
I
I
I
147
J
,
i
grey levels
wh,e
30 40 50 60 70 80 90 100110 I
I
I
I
I
I
I
I
I
12 )ed 10
bed 9 10 beds 8b 8a
bed 7
bed 6
5b 5a ~4b 4a
13
bed 1 0 m limestones mady limestones fissile calcareous marls marls Fig. 3. The Angles core set (Upper Valanginian). CaCO 3 and grey intensity records.
148
F. GIRAUD ET AL.
The advantage of the first method is the accuracy of the estimation of spectral amplitude and the possibility to test the statistical properties of these estimates; whereas the second is chosen for its good resolution in the low frequency domain and between neighbouring peaks. The lack of valid chronostratigraphic data for the studied series initially obliged us to calculate all periods in metres rather than in time units. The results of the data processing will be converted into time later, using the average sedimentation rate inferred for the section from the duration of the stage and the thickness. As we work on carbonate percentages and grey levels and not on cycle thickness, commonly applied hypothesis which confers a value of 21 ka as the duration of the elementary bed-interbed couplet is not used.
4000
-
3500
-
3~0
-
2500
-
119
| Maximum Entropy
20001500
-
Blaokrnan - Tukey
ooq j/
1000
~1:i
0
iI1~
~'/
-
71
"" ......
0
0.02
~
"" '
5000
0.1
ti 19
3000
Blackrnan - Tukey
$ a.
0.08
(cycles/cm)
|
4000
The A n g l e s cored set
-" ' . . . . . . . . .
0.06
frequency
Results Carbonate content and grey level variations were measured at a 5 cm spacings. The same pattern of variation exists in the two records (Fig. 3). They are dominated by high-frequency oscillations showing carbonate maxima and lighter grey intensity centred in limestone beds and by lower carbonate values and darker grey intensity centred in marly interbeds. The high-frequency fluctuations thus depict the bedding couplets. Carbonate content and grey scale power spectra are in phase, showing a single strong peak with a wavelength of 119 cm (Fig. 4). According to five recent timescales (Table 1) the duration of the Valanginian stage varies from 5.7 Harland et al. (1990), to 5.9 Ma from Huang et al. (1993), to 9 Ma for Hallam et al. (1985), with intermediate values of 7 and 8 Ma for Kent & Gradstein (1985) and Odin & Odin (1990) respectively. The mean sedimentation rate calculated from these estimates varies between 25 and 40 m Ma -1. The duration of the 119 cm peak periodicities thus varies between 29.7 and 47.6 ka. According to a sedimentation rate
"""
0.04
2000
1000
-
Maximum Entropy
"
~~._.., 0
i
0
0.02
........,A~.
i
1
0.04
0.06
I
I
0.08
0.1
frequency(cycles/cm)
Fig. 4. Spectra of (a) CaCO3 distribution; (b) grey level in the Valanginian core. Cycle lengths (cm) are indicated on top of the peaks.
of 3 2 . 1 m M a -1 using Kent & Gradstein's timescale, the periodicity is 37 ka which is close to the obliquity frequency calculated by Berger et aL (1989) for the early Cretaceous (Fig. 5). To examine the relative importance of the obliquity signal in this series, we used a band-pass filtering centred on a frequency of 1/119 (Figs 6 &
Table 1. Duration of the Valanginian period based on recent geological timescales and the mean sedimentation rate for the Angles section. (The Valanginian thickness without slumped intervals is 225 m Hallam et al. 1985 Duration of the Valanginian (Ma) Inferred sedimentation rate (m Ma-1) at Angles
Kent & Gradstein 1985
Harland et al. 1990
Odin & Odin 1990
Huang et al. 1993
9
7
5.7
8
5.9
25
32.1
39.5
28.1
38.1
VALANGINIAN OF THE VOCONTIAN TROUGH .
.
.
.
.
.
.
54
000
50
000
|
'~176 80
41
149
70
000
,
t '~ 5
:
"
.
|
.~,o.
3
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,
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,
|
,
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|
~
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~
t. 5 3ia
40
ECESS,O.
'
; ;
-~ 0
~
30
23ooo
| - 10 200 400 600 800 1000 1200 1400 depth (cm)
0
,,ooo
-3oo
'
-2'00
'
-lOO
o
CaC03 % ] ........... band-pass filtering
TIME (MYR BP)
tFig. 5. Estimated values of the periods of the palaeoclimatic precession and obliquity from Berger et al. (1989).
| lO0
2 ~,k:: : :: ': i~ :
90 80
7a). Compared with the original profiles the filtered series is significant, especially for the carbonate content, and shows amplitude variations which coincide with the main excursions of high carbonate content (Fig. 7a). These high carbonate values correspond to the principal beds recognized in the lithological log. If the 119 cm cycle corresponds to an obliquity forcing, the 71 cm cycle indicated in the CaCO 3 periodogram should represent evidence of pre-
I
grey levels t band-pass filtering
...........
(I11 ~9 cm)
1101
/
1~176 80 go 70
fi ~: .... ':': !( !i ! " '~ . !~! ~ ):' : . ' :
,
/~f~i;!::
60
10
V ",." ' i"
: '!
A,,=.~.3 !s = g -
i
40 30 0
I
I
i
I
200
400
600
800
|-10 I 1000 1200 1400
0
_--~.r
3 =~
v~.o
~
!i:'
!
"
ii
.=.
1
cr
==
~'~
-'~, o. m 3
8
70
0
o
6o
~
50
-1 ~,
40
'i =" ~ i:~
30 0
i -2 200 400 600 800 1000 1200 1400 depth (cm)
Fig. 7. Comparison between original and filtered carbonate content series. (A) Band-pass filtering centred on 119 cm; (B) band-pass filtering centred on 71 cm.
cession cycle, using the Kent & Gradstein timescale. This cycle is easier to follow on a filtered carbonate series (Fig. 7b), but the amplitude of this signal is shorter than those of the 119 cm cycle. These CaCO 3 cycles at 71 cm could correspond to the minor marly-carbonate layers which double the elementary beds (Fig. 3). Thus, the filtered carbonate and grey level series compared with the original series allow us to distinguish the primary bed and interbed couplets from minor couplets and to propose an important obliquity forcing. We shall now study the carbonate content record for the whole Valanginian.
The whole Valanginian succession
depth (cm)
Fig. 6. Comparison between original and filtered (1/119 cm) grey level series.
On the original carbonate curve (which is not illustred here because of its length) the visual
150
F. GIRAUD ET AL.
Two kinds of data are available: (1) thicknesses of the four recognized facies in the section: limestone, marly limestone, fissile marly carbonate, and marls; and (2) carbonate content data on samples taken at a constant 50 cm spacing. Carbonate records for the different facies are presented
observation does not reveal any cycles. In order to get a usable carbonate record we take lithological information into account: bed and interbed thickness data combined with carbonate content will generate a carbonate curve that is more representative of the initial lithology.
CaC03(%) total C a C 0 3 record
2,0001
marls record
35
45
55
65
75
85
95
I
I
I
I
I
I
J
4045 5055 5065 7075 '
'
J
'
'
I
I
I
fissile calcareous marls record 455o556o65 70758o i
i
i
i
i
i
i
i
20 000
Io 000
5000
0J
Fig. 8. CaCO3 records (with 50 cm spacing) of the Angles Valanginian section.
marly limestones record 7075808590 l
I
I
i
i
limestones record 80859095 = t i |
!
VALANGINIAN OF THE VOCONTIAN TROUGH separately in Fig. 8, showing that the marl record contributes the most information to the total CaCO 3 record. Both signals will allow the construction of a more representative fit for the carbonate curve related to the initial lithology. The principal steps of this procedure are summarized in Fig. 9. A part of the final carbonate curve on which spectral analysis will be applied is shown in Fig. 10. Similar procedures of spectral analysis as used for the cored set were employed but only the Blackman-Tukey spectral density is presented here because of its better statistical validation; spectra obtained with the maximum entropy method show
initial lithology marl fissile calcareous marl marly limestone limestone
1
marl's carb. profile
/
Inltlal / ~ s carbonate l J - ~ curve ~ "
1
i digitized Ilthology resampled I with a constant 10 cm spacing J
l
the same peaks. The periodogram of CaCO 3 distribution for the Valanginian shows two populations of peaks with three dominant (169, 141 and 101 cm) and a group of minor peaks between 80 and 65 cm (Fig. 11 a). The Valanginian section is in two different parts: the early Valanginian has limestone-dominant alternations, in contrast the late Valanginian presents marl-dominant alternations. Figure l i b shows power spectra derived from the carbonate content for the early and late Valanginian separately. The populations of peaks identified for the whole Valanginian are present, but the early
I initial carbonate curve I with samples collected at a 50 cm spacing
digitized Iithology marl = 1 fissile calcareous marl = 2 marly limestone = 3 limestone = 4
(a)
151
fissile calcareous marl's carb. profile marly limestone's carb. profile limestone's carb. profile
carbonate profiles resampled I with a constant 10 cm
spacing I
final carbonate curve I codes are replaced by correspondingl calcium carbonate values i
~oI L.,-,2~ ~o i ~ " ' = " | 40 -111~'*---- s.~ou,~ ;
20
,
Gurve
,
10
0
(b)
0
I
I
I
I
l
5O
100
150
200
25O
thickness(cm)
Fig. 9. (a) Different steps in processing the Valanginian time series data. The 10 cm spacing has been chosen after represent the thickness frequency plot of Valanginian beds and interbeds shown in b). (b) Bed thickness frequency of Valanginian beds and interbeds. (The 10 cm thickness is taken as the intermediate value between the two principal modes).
152
F. GIRAUD ET AL.
initial carbonate curve
stratigraphic log
digitized carbonate curve
240
240'
239
238
238
237
236
236 235
v
E
234
E
234
91o
233
ill) "10 232
232
1
C-
231 230
230 229 -
228
228 -
227
226
226
I 40 50
= I ' 'l ' 'l 60 70 80 90
[ I limestones FTTq marlylimestones fissile calcareous marls marls
I40
50
60
70
i 80 90
CaC03 %
Fig. 10. A section through part of the T. callidiscus ammonite Zone.
Valanginian shows a weak peak at 164 cm, one larger peak at 100 cm and three groups of peaks at 78-65 cm and 55-43 cm; while the late Valanginian is characterized by larger peaks at 178, 143 and 119 cm: the last one corresponds to the
principal peak identified in the cored interval. From these results, it is clear that the more calcareous early Valanginian is characterized by shorter cycles and the marl-dominated late Valanginian by longer cycles. Two kinds of hypothesis can be proposed to
VALANGINIAN
OF
THE
| 1000
169
400
49
200
0
0
I
I
0.01
0.02
I
I
I
0.03 0.04 frequency (cycles/cm)
0.05
powerof carbonate content for the early Valanglnlan
(~)
......... powerof carbonate content for the late Valanglnlan
1500 qt 178
12oo
143 /119
ql
I
//
,~i z
J~
/ /
!i;/1~ 100
78
55 /~
300
4
i.':~ 0
/
~
"~",'.L,'" i
0
49
0.01
i
i
0.02
0.03
i 0.04
i 0.05
frequency (cyclesJcm)
Fig. 11. Blackman-Tukey spectra of CaCO3 distribution for (a) the whole Valanginian, (b) early and late Valanginian. Cycle lengths (cm) are indicated on top of the peaks.
explain the differences between limestone and marl periods: (1) a variation of the sedimentation rate with no change in the forcing period(s); (2) a change in the record of orbital forcing with or without a change in sedimentation rate. To answer to this question, a detailed study of the spectrum evolution is necessary (Ten Kate & Sprenger 1992). This evolution is investigated by examining spectra of successive segments of the Valanginian series and by comparing the location and the height of peaks. This assumes that, within each segment, sedimentation rate is more likely to
VOCONTIAN
TROUGH
153
be constant than through the entire record (Herbert & Mayer 1991). Spectral peaks are therefore less likely to be smeared by variations in sedimentation rate. The statistically more reliable procedure in the investigation of the evolution of the spectra is to use a moving window of standard length and to generate spectra for overlapping subsections along the length of the time series. Another method consists of the division of the original series into subsections of unequal length based on abrupt change in carbonate distribution. This second method is statistically more doubtful because of unequal lengths of the subsections, but it allows respect for one of the theoretical requirements in performing spectral analysis: the stationarity of the data series, especially as data series are geochemical data such as calcium carbonate. For the first procedure, the original Valanginian series was subdivided into 21 segments, each with a length of 28 m and having a mutual overlap of 65% (see Appendix); the non-stationarity of the time series is indicated by the common presence of double peaks in some sections, corresponding to lithological changes. Therefore, we have chosen to apply the second method based on the abrupt change in carbonate distribution. The Valanginian was divided into 14 segments of unequal length, between 800 and 2800 cm (Fig. 12). Independent spectral analyses were performed on each segment. Comparison of the 14 subspectra (Fig. 12) yields the following conclusions. For each segment, the spectrum of CaCO 3 distribution shows one dominant peak, but the location and the shape of the highest peak is never stable along the succession. This suggests that the rate of accumulation has varied. A first group of peaks detected in the early Valanginian spectrum (80-61 and 53-41 cm) is present in numerous segments and consists of two principal peaks. The changing periodicities found for these significant peaks are probably due to variations in accumulation rate or to minor stratigraphic hiatuses which affect the base of the Valanginian series. This group of peaks is predominant and is the only one present throughout the sections a-d. The power of the peaks also fluctuates, with a maximum in the sections b and d, which present the highest peaks. Section e, with a double peak at 80-100 cm, seems to be transitional to the higher sections, which are characterized by an increase in the cycle length. From section f-n other groups of peaks (159, 154-145 and 115-106cm) are prominent. These different peak lengths are interpreted as being the result of a changing sedimentation rate. As in the first sections, the power of the dominant peaks varies with a maximum in sections g and k. The modulation of the power of a peak is consistent
154
F. GIRAUD ET AL.
350
52
section a (1190 m) (0-1190 cm)
300
2000-
80 100.
section e (1460 m) (5240-6700cm)
1500
250 200
~1000
~.150 100
62
500
..,.;
50 I
0
]
I
I
--7
1400
section b (840 m) (1190-2330cm)
78
1200 1000
I
I
0.01 0.02 0.03 0.04 0.05 (frequency/cm)
0
I
I
0.01 0.02 0.03 0.04 0.05 (frequency/cm)
2000
106 t
section f (1370 m) (6700-8070cm)
1500
47
434
8. 600 400 200
...
0 i 0
._./
500 ~
",.. I
I
0
I
0.01 0.02 0.03 0.04 0.05 (frequency/cm)
350
47
300 250
section c (860 m) (2330-3190cm)
0
0.01 0.02 0.03 0.04 0.05 (frequency/cm)
2000
section g (2000 m) (8o7o-lO4oo cm)
154 1500
62
3 ]= 200 8. 15o
~
. 1000
,.,.. :',',
100 500
50 0
I
0
I
I
I
1
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I
0.01 0.02 0.03 0.0~ 0.05 (frequency/cm)
67
1500
section d (1280 m) (3190-5240cm)
500
36
o
0
400 350
0.01
0.02 0.03 0.04 0.05 (frequency/cm)
section h (2200 m) (10400-12600crn)
113
300 250
1000
8.
~
~
0 ,
0
:. 48
, ""';, " ~ , 0.01 0.02 0.03 0.04 0.05 (frequency/cm)
g3= 200 150 100 50 0
5
!
0
]
I
I
I
0.01 0.02 0.03 0.04 0.05 (frequency/cm)
VALANGINIAN OF THE VOCONTIAN TROUGH 1000 q |
8oo-t
115. ~[
section 1(2800 m)
il
(12eoo-1.~.4oD~ )
400
155
159
700
section 1(800 m)
60o
(2o03o-~,o~o~m,
300
30
.
200
200
100
0 0
O, 0
0.01 0.02 0.03 0.04 0.05 (frequency/crn)
1oolItlI/lit, i
~[
section j (2000 m)
, , ~ , 0.01 0.02 0.03 0.04 0.05 (frequency/cm)
14oo 4
section m (1550 m)
1000
| 800 e~ 600
e~ 650 .~, 325
7~t/5/i 0
1600 1400 1200
il 61
36
23
400
0.01 0.02 0.03 0.04 0.05 (frequency/cm)
145 section k (2450 m) (17400-20030cm)
1000
'I i\ 5 6 ~
0
2500
0.01
154
3_7
0.02 0.03 0.04 0.05 (frequency/cm)
section n (1600 m) (22380-23980cm)
2000 1500
. 800 600 400 l t ~ { t 78 63 200 0 0 0.01 0.02 0.03 0.04 0.05 (frequency/crn)
1000 500
0
0.01 0.02 0.03 0.04 0.05 (frequency/cm)
Fig. 12. Blackman-Tukey spectra (solid line) of 14 successive sections of the Angles Valanginian series. Maximum entropy spectra (dotted line) are presented to precise peaks resolution. Time base is the length of the time series. The important frequency peaks are marked with their wavelength. In brackets are the length of time series (cm) including slumped intervals (cf. Fig. 2). with orbital control of the underlying process (Ten Kate & Sprenger 1992). With the different subperiodograms it is possible to study the relationship between carbonate content and dominant spectral peaks. For each short
section the average percentage of carbonate is calculated, which allows the construction of scatter plot of main spectral peaks vs CaCO 3 (Fig. 13). The general trend of the plot does not allow us to promote one or the other of the two hypotheses
F. GIRAUD ET AL.
156
]
b. c, d. e, f, g: early Wdaagini~m
kA mA
140 -
g ~
[ ~ I, Jl I~ I, m, n: klle Valanginian
I hAA
120 -
IA
[
fA
100 -
t
eA
80 -
bA dA
6O
cA 40 55
aA
I
f
i
!
i
60
65
70
75
80
CaC03 %
Fig. 13. Relationship between average CaCO 3 percentage and principal spectral peak calculated for each short windows of the Valanginian. proposed: variations of sedimentation rate or change in the record of orbital forcing, but, on the contrary, both propositions should be considered. The relation between CaCO 3 percentage and dominant cycle length is nearly linear with a valid correlation factor (r = - 0.78). The correlation is especially good for segments a-e, where the length of peaks increases when the. CaCO 3 percentage falls. So, the proposed variations of sedimentation rate are correlated with the CaCO 3 content: the sedimentation rate increases when the sedimentation becomes marl dominated. The scatter plot
also shows a discrimination of groups of points, one concentrated towards short peaks and high percentages of CaCO 3 (sections a-e), another located towards long peaks and low percentages (sections g, and k-n) and an intermediate group between these two populations corresponding to sections (f and h-j). Nevertheless, there are too few points to support only the interpretation of a change of dominant orbital forcing (precession to obliquity?) from limestone-dominant alternations to marly sediments. Thus, the transition from limestone-dominant alternations to marly-dominant alternations could be characterized by both: (1) a change of the record of orbital forcing; (2) an increase of sedimentation rate. The search for Milankovitch periodicities in the carbonate record requires a last step: the transformation of observed cyclicities expressed in metres into time. This conversion is realized in two steps. The main carbonate cycles identified in the sections are transformed into time. From the different durations attributed to the Valanginian and the thickness of the Angles Section, average sedimentation rates have been calculated (Table 1). From these calculations the duration of the carbonate cycles can be estimated. The periods attributed to the carbonate cycles are shown in Table 2. The durations obtained do not match
Table 2. Carbonate cycle durations calculated for each section using different average sedimentation rate and orbital peaks inferred for the Angles Valanginian series. P, precession cycle; Q, obliquity cycle; E, eccentricity cycle
Sections a b c d e f g h i
Length of main peaks (cm) 52 41 78 47 47 62 67 80 100 106 434 154 113 115 189
j k 1 m n
106 145 159 137 474 154 77
Corresponding time period (ka) from different estimations of the sedimentation (m Ma-1) rate for the Valanginian stage at Angles (Table 1) 25
28.1
32.1
38.1
39.5
Inferred orbital cycle
20.8 16.4 31.2 18.8 18.8 24.8 26.8 32.0 40.0 42.4 173.6 61.6 45.2 46.0 75.6 42.4 58.0 63.6 54.8 189.6 61.6 30.8
18.5 14.6 27.7 16.7 16.7 22.1 23.8 28.4 35.6 37.7 154.4 54.8 40.2 40.9 67.2 37.7 51.6 56.6 48.7 168.7 54.8 27.4
16.2 12.8 24.3 14.6 14.6 19.3 20.8 24.9 31.1 33.0 135.2 47.9 35.2 35.8
13.6 10.7 20.5 12.3 12.3 16.3 17.6 21.0 26.2 27.8 113.9 40.4 29.6 30.2
13.1 10.4 19.7 11.9 11.9 15.7 16.9 20.2 25.3 26.8 109.8 39.0 28.6 29.1
P P P P P P P P O O E O O O
58.8
49.6
47.8
O
33.0 45.2 49.5 42.7 147.6 47.9 24.0
27.8 38.1 41.7 35.9 124.4 40.4 20.2
26.8 36.7 40.2 34.7 120.0 39.0 19.5
O O O O E O P
VALANGINIAN OF THE VOCONTIAN TROUGH exactly with the orbital periodicities, but they are within the range of the Milankovitch cycle frequencies. So, the family of peaks 41-80 cm match the precession cycles, whereas the cycles including 100-159 cm correspond to the obliquity cycles. The highest peaks at 434 and 474 cm match the shorter cycle of eccentricity; in pre-Quaternary series, the presence of 100 ka cycle results from the climate system's non-linear response to the precessional forcing (Herbert & Fischer 1986). According to the calculated values of the precession and obliquity periods for the Early Cretaceous, which are 20.2 (average value) and 38 ka (Berger et al. 1989), respectively, we can normalize the carbonate-cycle duration. Knowing the duration of the principal carbonate cycles recognized in a section, and its thickness, we can deduce the duration of the different Valanginian sections (Table 3). This allows the reconstruction of the Valanginian carbonate series as a function of time. A new spectral analysis is applied for the whole Valanginian to calculate periodicities in time units. The periodogram illustrated in Fig. 14, corresponding to a duration of 6.8 Ma for the Valanginian, confirms the occurrence of two kinds of peaks: dominant peaks between c. 43 and 35 ka, and less prominent peaks between 23 and 17 ka. Thus, the main signal recorded from the CaCO 3 cycles seems to be related to the obliquity. These conclusions confirm the results obtained by Huang et al. (1993) from bed-interbed c'yclicities.
157
Another approach in order to determine if the frequency peaks of the spectra are related to Milankovitch orbital elements is the comparison of ratios between the main orbital periods for the Early Cretaceous period and those calculated for the dominant carbonate cycles. The ratio comparison method does not work when there is only one major peak in the spectra, such is often the case in this study. Cyclostrati g r a p h y
We have explained the use of CaCO 3 cycles to make an orbitally calibrated chronology for the Valanginian. This estimation of Valanginian duration can be made more precise, using filtered components compared with the original carbonate record. In the study of the Upper Valanginian core, the great potential of the filtered series to identify cycles has been demonstrated. For the whole series we work on the last data set, in which the carbonate content is expressed as a function of time. As the precession seems to be the dominant forcing in the lower part of the Valanginian (sections a-e) (Table 3), and obliquity the primary cause of the dominant rhythmicity for the rest of the Valanginian (sections f-n), bandpass filterings centred on the principal frequencies 1/20.2 and 1/38 ka, respectively, are applied to the carbonate series. On these two curves all CaCO 3 cycles will be identified and numbered; knowing the number of cycles and making the above
Table 3. Mean sedimentation rate and estimated duration of the Valanginian sections
Sections
Section thickness (cm)
a
1190
b
840
c d e
860 1280 1460
f g h i
1370 2000 2200 2280
j k 1 m n
2000 2450 800 1550 1600
Length of main cycles (cm)
Duration (ka)
52 41 78 47 47 67 80 100 106 154 113 115 189 106 145 159 137 154 77
20.2 20.2 20.2 20.2 20.2 20.2 20.2 38.0 38.0 38.0 38.0 38.0 38.0 38.0 38.0 38.0 38.0 38.0 20.2
Sedimentation rate (cm ka-1) 2.57 2.03 3.86 2.33 2.33 3.31 3.96 2.63 2.79 4.05 2.97 3.03 4.97 2.79 3.81 4.18 3.6 4.05 3.81
Average sedimentation rate (cm ka-1)
Section duration (ka)
2.3
517.4
3.09
271.8
2.7 3.31 3.29
318.5 386.7 443.8
2.79 4.05 2.97 4.00
491.0 493.3 740.7 700.0
2.79 3.81 4.18 3.6 3.93
716.8 643.0 191.4 429.4 407.1
158
F. GIRAUD ET AL.
1200
38.7 /
1000
135.6
800
42.7~
600
23
400 2O0 0 0
0.004
0.008 0.012 frequency (&-yr}
0.016
0.02
Fig. 14. Blackman-Tukey spectrum of CaCO3 distribution for the Valanginian.
assumption about their identification and, hence, their duration, a new estimation of the duration of the Valanginian will be proposed. An example of band-pass filtering centered on the dominant frequencies of precession and obliquity and applied to different parts of the carbonate record is illustrated in Fig. 15. The filtered components are often in phase with the original series with small shifts corresponding to the slight changes in sedimentation rate observed between more or less calcareous sediments. As shown on Fig. 15, all carbonate cycles have been identified and numbered. Concerning carbonate cycles related to the precession cycles, 91 cycles have been identified, whereas 137 cycles related to the obliquity signal have been recognized. This gives a duration of 7.04 Ma for the Valanginian stage, which agrees best with the Valanginian timescale proposed by Kent & Gradstein (1985), which is in the middle of the age 5 . 7 - 9 M a covered by the other estimates (Table 1).
Discussion and conclusion What is the explanation of the change of spectral responses between limestone-dominant alternations in the lower Valanginian and marl dominant alternations in the upper Valanginian ? It is important to state that the presence of cycles related to the obliquity signal is not an artefact inherent to the section or/and to the spectral method used: (1) no hiatuses are identified in biostratigraphic studies; Atrops & Reboulet (1993) show that the Valanginian hypostratotype can be correlated bed-for-bed with another pelagic section at La Charce (Dr6me) which argues for the relative continuity of the sedimentary record; (2) the identical results obtained by Huang et al. (1993) from Walsh spectral analysis of marl-limestone
sequences validate the choice of spectral techniques employed. It is widely accepted that low-latitude regions, such as Vocontian Trough, are sensitive to the precessional and eccentricity parameters. Variations in tilt of the Earth's rotation axis affect seasonality and insolation gradient with cyclic latitudinal shifts of climatic zones generally being important closer to the poles (Berger 1978). If so, could it be possible that cycles related to obliquity in the Valanginian Vocontian series reflect a distortion of the spectral signal because of important variations of sedimentation rate? In the upper Valanginian the average length of a carbonate cycle is 143 cm. If this cycle is relative to the precession forcing this implies a sedimentation rate of c. 70 m Ma -1. The Valanginian alternating sediments are characterized by the presence of burrowing structures and trace fossils very well preserved both in limestone and marls. These structures attest to a depositional environment subject to slow and regular sedimentation (Gaillard 1984). Considering, on the one hand, these observations and, on the other hand, the estimation of sedimentation rate for pelagic carbonates, between 10 and 5 0 m M a -1 (Einsele 1992), the value of 70 m Ma -1 seems to be incompatible with our data. Consequently, the hypothesis of distortion of the spectral signal by variations of sedimentation rate is not very convincing. Carbonate cycles which represent variations of carbonate productivity are not directly linked to changes in insolation but represent a complex response of the ocean-atmosphere system to insolation forcing (Herbert & Mayer 1991). Carbonate productivity varies in response to nutrient supply which, in turn, depends on climatic and oceanographic factors such as wind driven upwellings (Thomsen 1989) and sea-level changes. The cyclic change observed in biological content between limestone beds and marly interbeds suggests a rhythmic succession of different marine environments (Table 4). These observations have caused different authors (Cotillon et al. 1980; Darmedru 1983, 1984; Ferry 1991) to interpret the variability in carbonate productivity, at the scale of marl-limestone couplets, in terms of 'minimum and maximum' climatic conditions [with regard to a Cretaceous time of global warmth (Barron 1983)] changing in response to external forcing. To explain the presence of a strong obliquity cycle in Valanginian altemations, we must consider palaeogeographic and palaeoceanographic conditions during this time interval. The obliquity becomes the predominant forcing in the marldominated sediments; this change of sedimentation corresponds, approximately, to the transition from the early to the late Valanginian. The change
VALANGINIAN OF THE VOCONTIAN TROUGH 7-10 5,6
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9
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Fig. 15. Comparison of the filtered components (precession 20.2 ka and obliquity 38 ka) (dotted line) in parts of the original carbonate series (solid line). The numbers figured on the right-hand side of the two plots correspond to the carbonate cycles recognized. towards marls in the Vocontian Trough is the result of a widespread transgression evident for the early late Valanginian over much of Europe (Tyson & Funnell 1990). This transgressive episode allowed
better connections between the Boreal and Tethyan regions, favouring migration of marine nannofloras and faunas (Mutterlose 1992, Tyson & Funnell 1990). The connection between these basins is
160
F. GIRAUD ET AL.
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Table 4. Distribution of the main sedimentary components in limestone-marl alternation Limestones Microfauna (Darmedru 1983)
Nanoflora (Darmedru 1982)
Specialized and poorly diversified assemblages Planktonic species: Benthic species: spherical radiolarians calcareous and arenaceous planktonic foraminifers foraminifer Nanoconids
Coccoliths
Smectite
Illite, kaolinite, chlorite and silty quartz
Fe, Mg and Sr
K, Li and Zn in calcareous phases; Zr, Nb, Rb and Ni in terrigenous phases
Major concentration in 180 and 12C
Major concentration in 160 and 13C
Wood debris
High concentration in organic matter
Mineralogy (Cotillon et al. 1980) Geochemistry major and minor elts (Jouchoux 1981)
Oxygen and carbon isotopes (Cotillon & Rio 1984) Organic matter (Darmedru 1982)
Marls
indicated by the occurrence on several occasions of boreal faunas (ammonites) in the Vocontian Trough through the late Valanginian (Besse et al. 1986; Reboulet et al. 1992). According to Michael (1979) and Kutek et al. (1989), a sea-way was constantly open from the Boreal Basins towards the Tethys existed throughout the Valanginian. However, during the early Valanginian no exchange of Tethyan and Boreal nanofloras and faunas have been recorded; this time period is characterized in both areas by a strong faunal endemism and provinciality. Mutterlose (1992) suggests that the spreading of organisms during the late Valanginian was controlled by other factors, such as climate, and consequently water temperature. According to Kemper (1987), colder
phases allowed Boreal species to expand to the south. For some intervals, Reboulet et al. (1992) observed out-of-phase migration of the Mesogean and Boreal faunas and argued for a dominance of the climatic factor in the Vocontian Trough at these levels. So, during the early late Valanginian climatic and oceanographic changes affected the Vocontian Basin. Because these major changes favoured the connection with the Boreal Basins, we can suppose that the Vocontian Trough predominantly records the climatic influences of the higher latitude realm. Thus, the strong occurrence of obliquity signal in southern areas during this time period could be explained.
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Modifications stdimentaires, renouvellement des faunes et inversions magnttiques dans le Valanginien de l'hypostratotype d'Angles. Bulletin des Centres de Recherches Exploration-Production Elf Aquitaine, 10, 365-368. BUSNARDO, R. & THIEULOY, J. P. 1979. Les zones d'ammonites du Valanginien. In: BUSNARDO,R., THIEULOY, J. P. & MOULLADE,M. (eds), Hypostratotype mdsog~en de l'~tage Valanginien (S-E de la France), Centre National de la Recherche Scientifique, Paris, 6, 58-68. COTILLON, P. 1984. Pa16ogtographie. In: DEBRANDPASSARD, S. ET AL. (eds), Synthkse gdologique du Sud-Est de la France. Mtmoire du Bureau de Recherches Gtologiques et Minitres, 125, 328-330. 1985a. Hauts-fonds de la marge Nord-Est
VALANGINIAN OF THE VOCONTIAN TROUGH provenqale au Crrtac6 infrrieur. Un exemple de contr61es tectonique et bathymrtrique. Bulletin de la Section des Sciences, IX, 25-8. 1985b. Les variations ~ diffrrentes 6chelles du taux d'accumulation s6dimentaire dans les srries prlagiques alternantes du Crrtac6 infrrieur, cons6quences de phrnom~nes globaux. Essai d'rvaluation. Bulletin de la Soci(t( g(ologique de France, 8, 59-68. --, FERRY, S. & RIO, M. 1979. Le Valanginien d'Angles: minrraux argileux, calcimrtrie. In: BUSNARDO, R. THIEULOY, J. P. & MOULLADE, M. (eds), Hypostratotype m(sog(en de l'(tage Valanginien (S-E de la France), Centre National de la Recherche Scientifique, Pads, 6, 30-32. - - , GAILLARD,C., JAUTEE, E., LATREILLE, G. & RIo, M. 1980. Fluctuation des param~tres du milieu madn dans le domaine vocontien (France Sud-Est) au Crrtac6 infOrieur: mise en 6vidence par l'rtude des formations marno-calcaires alternantes. Bulletin de la Soci(t( g(ologique de France, 7, 735-744. & RIO, M. 1984. Cyclic sedimentation in the Cretaceous of Deep Sea Drilling Project sites 535 and 540 (Gulf of Mexico) 534 (Central Atlantic) and in the vocontian basin (France). Initial Reports of the Deep Sea Drilling Project, Washington (US Government Printing Office), 77, 339-376. DARMEDRU,C. 1982. La microfaune dans les alternances marne-calcaire p(lagiques du Cr~tac( inf(rieur vocontien (Sud-Est de la France). Mise en (vidence d'oscillations climatiques. Universit6 Lyon I, Lyon, Th~se 3e cycle. 1983. La microfaune dans les alternances marnescalcaires pOlagiques du Crrtac6 infrrieur dans le Bassin vocontien (SE de la France). Mise en 6vidence d'oscillations climatiques. Comptes Rendus de l'Acad(mie des Sciences, 296, 715-718. 1984. Variations du taux de sOdimentation et oscillations climatiques lots du drprt des alternances marne-calcaire p6lagiques. Exemple du Valanginien suprrieur-Vocontien (Sud-Est de la France). Bulletin de la Soci(t( g(ologique de France, 7, 63-70. --, COTILLON,P. & RIO, M. 1982. Rythmes climatiques et biologiques en milieu matin pOlagique. Leurs relations dans les drprts crrtacrs alternants du bassin vocontien (Sud-Est de la France). Bulletin de la Soci(td g(ologique de France, 7, 627-640. DE BOER, P. L. 1982. Cyclicity and the storage of organic matter in Middle Cretaceous pelagic sediments. In: EINSELE, G. eT AL. (eds), Cyclic and Event Stratification. Springer-Verlag, Berlin, 456-475. 1991. Pelagic black shale-carbonate rhythms: orbital forcing and oceanographic response. In" EINSELE,G. ET AL. (eds), Cycles and Events in Stratigraphy. Springer-Verlag, Berlin, 63-78. DEAN, W. E., GARDNER,J. V. & CEPEK, P. 1981. Tertiary carbonate dissolution cycles on the Sierra Leone Rise, Eastern Equatorial Atlantic Ocean. Marine Geology, 39, 81-101. DONZE, P. 1979. Les ostracodes. In: BUSNARDO, R., THIEULOY, J. P. & MOULLADE, M. (eds), Hypostratotype m~sog(en de l'~tage Valanginien (S-E de
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Detailed relationships between platform and pelagic carbonates (Barremian, SE France) DIDIER QUESNE & SERGE FERRY Universit~ de Lyon, Centre des Sciences de la Terre, 43 Brulevard du 11 Novembre, 69622 Villeurbanne, Cedex, France
Abstract: Field observations suggest that prograding bioclastic wedges of depositional sequences in outer-platform settings correlate with packages of basinal limestone beds thought to represent third-order lowstand systems tracts. The rule is apparently the same for higherfrequency cycles within these sequences, since basinal beds seem to correlate with bioclastic clinoforms while basinal marly interbeds root between clinoforms. If confirmed in other sequences, it would suggest that limestone beds, or bundles of beds, in this basin have the same 'lowstand' significance 'highstand' vs intervening marls from second- to sixth-order cycles, whatever the source of the carbonate (periplatform oozes or planktonic rain). A major remaining challenge is to understand the mechanisms by which high-frequency sea-level changes, as documented in outer-platform bioclastic wedges, correlate so well with the changes in nannoflora (nannoconus-rich beds vs coccolith-rich marly interbeds) thought to be climatic in the pelagic Vocontian Trough series. Is it glacio-eustasy and how does it work to make the deepwater Cretaceous limestone-marl alternation?
Alternating limestone-marl deposits of the Vocontian Trough have previously been interpreted as distal turbidites (Beaudoin 1977), as deposits which come from diagenetic differentiation of an originally homogeneous sediment (Ricken 1986) or as deposits which have registered high-frequency climatic variations during the Mesozoic (Cotillon et al. 1980; Darmedru 1983). Mineralogic and biologic differences seem to be in agreement with the idea of Cotillon et al. (1980). The analysis of other similar cyclic series has permitted some authors (Fischer 1986)to suggest the permanence, during geological times, of the orbital forcing of climatic variations (Milankovitch theory) and to use this high-frequency cyclicity as a high precision chronometer (cyclostratigraphy). In another domain, the concepts of sequence stratigraphy (e.g. Vail et al. 1977; Posamentier et al. 1988) have proposed mechanisms to explain the detailed geometry of continental margin deposits. However, this model is still unclear about the relationships of the coastal deposits with the deep deposits, especially the pelagic ones. Yet it is these deposits that record the climatic variations which are implicitly invoked (Vail et al. 1991) to explain the eustatic part of the sea-level variations on the third-order cycles (1-3 Ma). It is therefore important to study basins where cyclostratigraphy and sequence stratigraphy can be linked, i.e. where it is possible to correlate, in detail, deep deposits (with climatic cycles) and coastal deposits (under the dependance of sea-level variations).
This work is focused on the Vocontian basin (SE of France) and its margins because of good biostratigraphic control which permits precise lithologic correlation between the pelagic area, the hemipelagic zone (corresponding to the slope) and the outer facies of the carbonate platform (bioclastic progradation clinoforms). The studied Barremian hemipelagic series are located on the south edge of the Vercors (north edge of the trough) and on the north side of the Ventoux mountain (south edge of the trough). The platform series are located on the Ardrche platform and on the south edge of the Vercors (Fig. 1).
Observations Bed-to-bed correlations between the pelagic basin and the hemipelagic slope As regard basin-slope correlations, Ferry & Monier (1987) have shown, by counting high-frequency cycles in several successive depositional sequences, for which the age is controlled by ammonites (Hauterivian, Sayni Zone and Lower Barremian, Hugii Tethyan Zone), that the bed-interbed couplet can expand by 3-15 times, according to the sedimentation rate, into a bundle of beds in the hemipelagic marly-limestones, when the accumulation rate is high. In contrast, when the accumulation rate is low, bundles of pelagic beds, that is fifth- to fourth-order cycles (Ferry 1991) may be condensed into only one bed in the hemipelagic area.
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescales and Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 165-176.
165
166
D. QUESNE & S. FERRY
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%
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'%.
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] Urgonian Platform
~
Archiane
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Vocontian Trough (pelagic facies) ( ~
studied areas
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200 km I
Fig. 1. Location of studied areas in southeastern France. (Site 2: side of the Archiane valley).
Is it possible to make high-resolution correlations upslope, i.e. into the true limestones of the carbonate platform?
Correlations between the slope (hemipelagic facies) and the outer platform (bioclastic clinoforms) We have made precise observations on several sites on the west and north edges of the Vocontian Trough (Fig. 1), at the cliff of Chames, near VallonPont-d'Arc (Lower Barremian, Hugii Zone of the Ard~che platform), at the Rochers de Combeau (base of the Upper Barremian), at the Montagnette and in the Archiane valley (top of the Lower Barremian), on the south edge of the Vercors. These observations concern four third-order successive depositional sequences.
a,
Rochers de Combeau; b, the Montagnette; c, east
The cliff of Chames to Pont-de-Laval. The cliff is formed of two massive superimposed prograding bioclastic limestones overlying the hemipelagic marly-limestone of the Upper Hauterivian. Each is c. 60 m thick. They are separated by a thin marly bed which becomes a seam towards the southwest, i.e. up the outer platform. Towards the north (Pontde-Laval, Fig. 2) the two bioclastic masses pass into two calcareous hemipelagic units, which have been dated to the first Tethyan ammonite zone of the Barremian (Schroeder et al. 1989). The Mouniers Marls terminate this first formation and are correlated with the marls of the basal compressissima Zone. The marly episode which separates the two calcareous masses is our local reference level and is continuous with the marls located between the two massive bioclastic limestone of the cliff of Chames. The correlations that it is possible to make from Chames (outer platform) to Angles (deep basin) (Fig. 2) integrate the
BARRREMIAN CARBONATES SE FRANCE
167
Platform~;:~' 9 (;ent3~ c ] ,'~on
Pont-de-Laval
Pont-de-LavaJ
tp. ~
, ~
I~
Chames
Vocontian
Trough
,~ , 9
~::~
C,harnes Ventoux mountain slope "
ktar~ciltc"' ' \
hemipe/agic slope i
=
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l
hemipelagmudslones.l icmassive ~
~
~ ....
~L~':~:~ -~-.: i
~.~
~om
~
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Fig. 2. High-resolution correlations between the outer platform and the basin.
previous observations of Ferry & Monier (1987). The two third-order successive cycles of the hugii Zone are in perfect continuity from the platform to the basin. Apparently, the two massive bioclastic units and the two limy hemipelagic masses correspond with each other. The intercalated marls
indicate relative highstands in the deep basin series, as well as between the two bioclastic masses at Chames. Our detailed observations concern the first bioclastic unit. At its base it is possible to see the successive clinoforms of the prograding wedge
168
D. QUESNE & S. FERRY
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BARRREMIAN CARBONATES SE FRANCE disappear into the hemipelagic marly-limestone, without lying on a downlap surface (Fig. 3). The successive sections made from the southwest to the northeast confirm this observation. The underlying hemipelagic beds can be followed at outcrop, especially due to some characteristic ones that can be recognized from one section to the next. Thus, an increase in the number of hemipelagic limestone beds basinward is observed (Fig. 4). This is proof of a lateral facies shift. Each bioclastic clinoform, which settled on the previous one, gives rise lateraly to an additional bed or a little hemipelagic bundle of beds in successive sections towards the northeast. The marly interbeds which are developed basinward between the beds seem to disappear upward between the clinoforms. These facts are in agreement with the observations made by Bosellini & Stefani (1991) in the Triassic of the Dolomites.
Rochers de Combeau (Fig. 5a). These are built by a vertical succession of bioclastic units, which are 10-15 m in thickness and can be traced over several hundreds of metres from west to east, up to the Men6e fault which limits the observations basinward. The massive bioclastic units (without
Platform (SW) .......:.....
~ 7 ~ m - ........
9.iiiil..iii.iiii~ii~i.
,~ ..
visible internal structures) stack and constitute a regressive/transgressive cycle (R/T) (Everts et al. 1992). Prograding clinoforms cannot be seen here. The interest of this site is that the accumulation rate was very high, which allows observation of the appearance, basinward, of the 'marly' interbed that separates two bioclastic units (Fig. 6). This interbed is made up of thin bioclastic beds (storm levels?), more and more separated basinward by marly seams. The Men6e fault prevents us from seeing the appearance of a real marly interbed. The massive units divide into thin beds, while the grain size fines basinward. If these thin beds are storm levels too, the alternation between massive bioclastic units and marly seams reveals R/T cycles within the major cycle. The calcareous unit seems to represent a relative lowstand in the high-frequency cycle and the marly interbed a highstand.
The Montagnette (Fig. 5b). The outcrop allows tracing of the facies evolution in a clearly prograding geometry (Arnaud-Vanneau & Arnaud 1991) over 300--400 m. The clinoforms, c. 10 m thick, pass laterally over < 300 m into small bundles of hemipelagic limy beds, separated by homogeneous marls. The thickness of these small
Basin (NE) ~
~i~iii~i~i~!i S -
169
I.
- 5 0 0 m ........ - ~
t.,....-........,
_.q ~i~:~ ~_]
cliff (bioclastic clinoforms) hemipelagic calcareous beds argilaceous limestone
Fig. 4. Correlations between sections of the cliff of Chames (Ard6che). Note the increasing number of hemipelagic limestones basinward, which shows the lateral shift of the bioclastic clinoforms to the hemipelagic limestones.
170
D. QUESNE & S. FERRY \V ( platform )
E tbasin)
,/
Mcndc fault
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I
bicx:laslic cliff
~ I00 m studicd area
a -- Rochers de Combeau
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NW (plallbrm)
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i
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c -- East side of the Archiane valley
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Fig. 5. Viewsof the outcrops studied in the south Vercors.
bundles of beds decreases basinward, from 3 to 4 m to c. 1 m; then they seem to thin and disappear into a homogeneous marly facies. Archiane valley (Fig. 5c). The lower bioclastic mass constitutes a prograding body which has received different sequential interpretations (Jacquin et al. 1991; Everts et al. 1992). From our observations, the existence of a downlap surface at its base and an onlap surface at its top are questionable. In agreement with Everts et al. (1992), the
body seems to be a R/T cycle from which only the lower part is visibly constituted by prograding clinoforms. This prograding part is overlain by rather aggrading units (mainly visible on the west side of the valley) which are overlapped by some thin units that prograded rapidly. We cannot see the passage from the prograding and aggrading units to the hemipelagic marly limestones on the west side of the valley, but the existence of a downlap surface, on which the bioclastic material could have prograded, seems improbable. On the east side of the
BARRREMIAN CARBONATES SE FRANCE Plattbrm (WNW)
171 Basin ( ES E) -"------~-
~400 to 500 m
h.V
seal1/ thin clastic beds ....:..:. ~ . . : ~ ..............
.. ======================
bioclastic limestone with joints
Men6e fault
.......... ..................:.,.:.;..:.;..:.?~.'..,.:;,.:.;.,:;.,.:.~-2"~ ~,~'~ ~,:
~.?.?..,
"marl)," interbed expanding basinward Fig. 6. Diagram showing the lateral change in the basal bioclastic units of the Rochers de Combeau.
valley we can observe and sample, as at Chames, the lateral change of two clinoforms in the lower part into two sets of bundles of hemipelagic beds (Fig. 5c). Here too, the interbeds, which develop basinward between the bundles, pass into seams between the bioclastic units. These, as at Combeau, are at first massive in the direction of the outer platform and then divide into beds basinward. Moreover, we can clearly see here that they are followed by a hemipelagic bundle, continuous over 3 km, and in which the thickness decreases very slowly. In all the above cases, thin sections made from successive samples show the gradual evolution of the microfacies. It goes progressively from foraminiferal bioclastic sand (grainstone/packstone) to bioturbated peri-platform mud with spicules. But it is impossible to analyse further the facies in these carbonates, for instance, to know
if these clinoforms are built by the accumulation of storm levels that become more and more bioturbated basinward. All the observations made in these four thirdorder depositional sequences show that the bioclastic formations, representing lowstand relative sea level in these R f f cycles, pass laterally into thick hemipelagic calcareous beds, then into bundles of more calcareous beds within the vocontian alternations (Fig. 7). This observation seems to be true for the elementary cycles that make up the third-order cycles. It will sometimes be possible to correlate in detail the hemipelagic m a r l y - l i m e s t o n e of the slope with individual basinal rhythms. In other cases, even if the correlations between the bioclastic clinoforms and the bundles of hemipelagic beds are correct, an individual clinoform cannot correlate with a particular pelagic bed. In different cases, one clinoform will correspond to a fifth-order bundle of beds
Outer platform
correlations between bioclasticclinoformsand limestone-marl hemipelagiccycles
floodin~marls
I
~ I(Im
correlations between bundlesof hemipelagicbeds and pelagic beds
--
'~mfs
~ 1 2 m ~ mrs
Fig. 7. Diagram showing the relationship between outer platform and basinal third-order cycles (mfs, major flooding surfaces),
172
D. QUESNE & S. FERRY
in the pelagic alternations, or to a single bed (sixthorder cycle). It is also possible that, as at the Montagnette, hemipelagic beds at the top of the clinoforms diffuse into rather homogeneous marly-limestones. But there is no question of the equivalence in age between the outer platform bioclastic formations and the bundles of calcareous basinal beds, as suggested by the lateral shift of facies. The marly episodes are correlated from the basin to the platform. They indicate a relative highstand between the calcareous units of the thirdorder cycles.
Platform
Hg ~
Basin
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i
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Discussion
Medium-frequency cycles (third-order cycles) The platform carbonate formations are generally interpreted as highstand prograding wedges, above all if they are overlain by an emersion discontinuity. According to Magniez-Jannin & Jacquin (1990) the bulk of calcareous alternations, representing third-order lowstands in the pelagic series, could be linked with the top discontinuity of the platform limestones (Fig. 8). Then platform limestones and basin limestones should be diachronous. Our observations suggest that this is incorrect. Between the marly flooding levels the bioclastic formations of the four studied Barremian cycles seem to be coeval with lowstand calcareous formations in the basin. Bioclastic limestones on the outer platform and calcareous beds in the basin may represent the upper and lower parts of relative lowstand systems tracts in simple R/T cycles. The question then is to know whether these bioclastic formations can be linked, and how (Fig. 9), with the Urgonian formations sensu stricto (the coral and rudistid facies) located higher in the depositional system, and which could represent the highstand prograding systems tracts of each thirdorder cycle (Fig. 9, 1st solution). If, on the contrary, the Urgonian formations pass laterally into the outer platform bioclastic formations they could be an integral part of the shelf margin wedges, i.e. of relative lowstand systems tracts, because coeval with them (Fig. 9, 2nd solution). The establishment of Urgonian platform facies, or of the platform limestone in general, often appears abruptly over the flooding marls; this change corresponds, according to us, with a downward shift of the depositional system which permits us to put here a Type II sequence boundary. It is very hard to answer to these questions in the first Barremian cycles, where the outcrops are not sufficiently continuous to see the relationship directly. Also, the biostratigraphy is not sufficiently precise to effect detailed correlations on long transects with exposure gaps. The available
Platform
Basin
-BFig. 8. (A) Correlations between platform shallowingup sequences and basinal sequences according to Magniez-Jannin & Jacquin (1990). SB, sequence boundary; LST, lowstand systems tract; TST, transgressive systems tract; HST, highstand systems tract; TS, transgressive surface; mfs, major flooding surface; Hg, hard ground; BS, black shales. (B) Our conceptions (Ferry & Rubino 1990; Ferry 1991).
field observations in the more Urgonian cycles of the Upper Ban-emian and Aptian of the subalpine chains and the Ardrchoise platform (Ferry & Rubino 1990) suggest that bioclastic wedges shifted basinward, corresponding to the discontinuity which overlies the platform carbonates; the Urgonian sequence seems to be coeval with with lowstand calcareous beds of the basin. Current work in the jurassic of the Jura (F. Cochet pers.
BARRREMIAN CARBONATES SE FRANCE Field observations
Inner platform (Urgonian)
173
on the platform edge
Outer platform (bioclastic progmdation)
Slope (hemipelagic marly-limestones)
:.,2,
q
'":"-".~........ .
,~
.
flooding marls
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Possible correspondance
Inner platform (Urgonian)
Inner platform (Urgonian)
-
~
with the Urgonian
Outer platform (bioclastic pro~radation)
Outer platform (bioclastic progradation)
..,,.:.
facies
I st solution
Slope (hemioela~ic
Slope (hemipelagic marly-limestones)
2 nd solution
~
4~.! ~
~ j~':
.-
Fig. 9. Correspondences between the inner platform, the outer platform and the slope (for explanation see text).
comm.) shows, meanwhile, that a bioclastic wedge and/or oolitic limestone of forced regression s e n s u Posamentier e t al. (1992) may sometimes appear below the extended platform, but it is biostratigraphically linked to what could be defined as a
transgressive system tract (passage of a calcareous alternation to a marly alternation) in the basin series. In a previous paper (Ferry 1991) we explained how, by calling upon different eustatic mechanisms which are out of phase, we can resolve
174
D. QUESNE & S. FERRY
this paradox and consider the Urgonian formations as relative lowstand wedges in spite of the discontinuity (maybe of emersion?) which overlaps them. This explanation is based on the idea of transgressive coolings which permit the superimposition of momentary high-frequency glacio-eustatic variations on the slow third-order transgressions. This hypothesis is in part based on comparison of the Vocontian rhythms with the high-frequency limestone-marl cycles of the Atlantic Quaternary. In both cases, nannofossil-rich calcareous beds could have an interglacial (i.e. warm period) value, and thus of relative highstand. The marly interbeds could have the opposite value. But our detailed correlations seem to invalidate this first idea.
High-frequency cycles and signification o f Vocontian alternations In the Atlantic Quaternary deposits the last glaciation (lowstand) is represented by a marly or argillaceous 'interbed' and the interglacial
Bahamas platform
Holocene by a calcareous bed (cf. Ferry et al. 1985). The rule is the same for the 50 climatic oscillations which preceded it (Ferry 1991), as the underlying limestone-marl (Leg DSDP 94) can be correlated with the standard isotopic scale (Opdyke in Harland et al. 1989). In the Bahamas (Fig. 10) the Holocene (highstand) is represented by c. 2 m of lagoonal limestone or oolitic dunes and reefs on the platform edge. This lagoonal Holocene is linked with the Holocene calcareous nannofossil-rich mud level of the Atlantic. The lowstand Wurmian ice age is represented by an emersion surface on the platform and by a marly interbed in the Atlantic Ocean. In the Mesozoic depositional sequences the supposed glacial-interglacial cycle has a sixthorder value in duration, i.e. a bed-interbed couplet. However, these pelagic bed-interbed couplets cannot be compared in terms of highstand or lowstand, even if they represent climatic cycles in both cases. On the contrary, the interbeds between the bioclastic clinoforms of the Vocontian margins, which seem to be linked with the pelagic interbeds according to our correlations, could, on the contrary,
Central Atlantic ocean
karstification
-120 m
Glacial Wtirm (lowstand) -25(X) m
actual m a lcvcl "lagt•
~~~~'~~ lagoonal HST(H~ previous sea level (- 120 m)
Interglacial Holocene (highstand)
ca[caret)us
~
mu.t /
= w a r m HST
(HF3 =coldLST (ttl~ ~ -2N~)m
Fig. 10. Holocene interpretation in the Bahamas (after Rasmussen et al. 1990) and central Atlantic Ocean during the last glacial cycle. LS, lowstand; HS, highstand; HF, high frequency (not to scale).
BARRREMIAN CARBONATES SE FRANCE represent flooding phases. So, there seems to exist an autosimilarity through the whole hierachy of the cycles in the Vocontian Trough. Calcareous beds could be equivalent to bioclastic progradation, from the high-frequency cycles (sixth order), up to those of third, or even second, order. At the opposite extreme, marl could represent flooding at every hierachical level of the cycle. But, we cannot consider that in the basin a more calcareous bed represents bioclastic progradation because it is made mainly of nannofossils (Noel 1968) (Nannoconus prevailing over coccoliths). So, these correlations link together two mechanisms of carbonate production which have no evident relationships. On the one hand the planktonic fall dominates in the basin, and on the other, the biodetrital sand and mud production characterizes the platform and the slope (or hemipelagic muds where nannofossils are rare). This raises the problem of the true nature of the high-frequency cyclic forcings which act on these mechanisms. Even if they are different according to the environments, and therefore not connected, our work shows that they act in rhythm. So, what are the mechanisms which generate the Vocontian cycles where the nannofossils are the same from the bed to the interbed? These correlations are important because one of the most famous problems in sequence stratigraphy is to understand the cause of high-frequency cycles in the sedimentary series. These cycles are often attributed to an orbital forcing (Milankovitch cycles) (Fischer 1986). The problem is to determine if these climatic cycles correspond to highfrequency sea-level variations too, as are those registered in carbonates of inner platforms (Goldhammer et al. 1987). A glacio-eustatic mechanism, although weak during the Mesozoic, could be the simplest solution to link the different
175
forcings. If the Vocontian alternations are really the result of climatic cycles and if, nevertheless, they cannot be regarded as equivalent to the Quatemary pelagic altemations, what are the mechanisms which create the bed-interbed succession in this basin?
Conclusion Our observations on four third-order depositional sequences made in the Barremian of the Vocontian Trough and its margins (Vercors, Ard~che and Mont Ventoux slopes) show that each bioclastic prograding wedge on the outer platform shifts laterally to one group of pelagic beds. It is sometimes possible to make detailed correlations between hemipelagic lime-marls of the slope and basin pelagic alternations in the basin. The marly flooding levels, which are correlated from the basin to the platform, signify highstands. So, the bioclastic progradational clinoforms of the outer platform and the calcareous bundles of the basin (which are coeval) signify lowstands. Can the correlation be continued up to the inner platform (rudistid Urgonian facies)? No observation has yet verified that. The bioclastic formations of the outer platform register high-frequency sea-level oscillations the pelagic alternations register high-frequency climatic altemations. We now need to know what mechanism is able to link these two parts, upslope and downslope, of the depositional system. Could it be glacio-eustacy, as for the Quaternary limestone-marls of the Atlantic, which presents lots of similarities with the Vocontian altemations? If so, we have to understand how it could act, because the Atlantic limestone bed is interglacial, so it has a relative highstand value, whereas the Vocontian bed may prove to have the opposite significance.
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Th6se Facult6 des Sciences de Caen, no. 85 004, 478 pp. BOSELLINI, A. & STEFANI, M. 1991. The Rosengarten: a platform-to-basin carbonate section (middle Triassic, Dolomites, Italy). Dolomieu Conference on Carbonate Platforms and Dolomitization. Guidebook excursion C, 1-22. COTILLON, P., FERRY, S., GAILLARD, C., JAUTI~E, E.,
LATREILLE, G. &RIO, M. 1980. Fluctuation des param6tres du milieu marin dans le domaine vocontien (France Sud-Est) au Cr6tac6 inf6rieur: mise en 6vidence par l'6tude des formations mamo-calcaires altemantes. Bulletin de la Socigtd gdologique de France, Paris, ser. 7, 22, 735-744. DARMEDRU,C. 1983. La microfaune dans les alternances mames-calcaires p61agiques du Cr6tac6 inf6rieur dans le Bassin vocontien (SE de la France). Mise en 6vidence d'oscillations climatiques. Comptes Rendus de l'Acaddmie des Sciences, Paris, 296, 715-718. EVERTS,A., SCHLAGER,W. • STAFLEU,J. 1992. Sequence analysis of the Vercors carbonate platform (Lower Cretaceous, SE France): problems with the existing sequence stratigraphic models. PlaOCorm Margins International Symposium,
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Book of Abstracts,
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FERRY, S. 1991. Une alternative au mod61e de stratigraphie s6quentielle d'Exxon: la modulation tectono-climatique des cycles orbitaux. Ggologie Alpine., M(moire H.S., 18, 47-99. --, PASTOURET, DE BEAULIEU, &MANDIER, 1985. Synchronisme remarquable entre les cycles calcaire-marne de vases quaternaires atlantiques et les alternances tourbe-limon d;anciens lacs p6riglaciaires europ6ens. Comptes Rendus de l'Acad(mie des Sciences, Paris, ser. II, 300, 573-578. MONmR, P. 1987. Correspondances entre alternances marno-calcaires de bassinet de plate-forme (Cr6tac6 du SE de la France). Bulletin de la Soci~t~ g~ologique de France, Pads, ser. 8, 3, 961-964. & RUBINO, J. L. 1990. Mesozoic Eustacy Record on Western Tethyan Margins. Post-meeting field-trip in the Vocontian Trough. Publication Association des Sddimentologistes Fran~ais, Pads, 12, 141. FISCHER, A. G. 1986. Climatic rhythms recorded in strata. Annual Revue of earth Planetary Sciences, 14, 351-376. GOLDHAMMER,R. K., DUNN, P. A. & HARDIE,L. A. 1987. High frequency glacio-eustatic sea level oscillations with Milankovitch characteristics recorded in middle Triassic platform carbonates in Northern Italy. American Journal of Science, 287, 853-892. HARDIE, L. A., WILSON, E. N. & GOLDHAMMER, R. K. 1991. Cyclostratigraphy and dolomitization of the Middle Triassic Latemar buildup, the Dolomites, northern Italy. Dolomieu Conference on Carbonate Platforms and Dolomitization. Guidebook excursion F, 2-37. HARLAND, T., ARMSTRONG, R. L., COX, A. V., CRAIG, L. E., SMITH, A. G. & SMITH, D. G. 1989. A Geologic Time Scale, 1. Cambridge University Press, Cambridge, XVI, 263 pp. JACQUIN, T., ARNAUD-VANNEAU, A., ARNAUD, H., RAVENNE,C. & VAIL, P. R. 1991. Systems tracts and depositional sequences in a carbonate setting: a study of continuous outcrops from platform to basin at the scale of seismic lines. Marine and Petroleum Geology, 8, 122-139. MAGNIEZ-JANNIN, F. &JACQUIN, T. 1990. Validit6 du d6coupage s6quentiel sensu Vail en domaine de bassin: arguments fournis par les foraminif6res dans le Cr6tac6 inf6rieur du sud-est de la France.
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-
&
Comptes Rendus de l'Acad~mie des Sciences, Paris, ser. II, 310, 263-269. NOEL, D. 1968. Nature et gen6se des alternances de marnes et de calcaire du Barr6mien sup6rieur d'Angles (Fosse vocontienne, Basses Alpes). Comptes Rendus de l'Acad(mie des Sciences, Paris, (D), 266, 1223-1225. POSAMENTIER, H. W., JERVEY, M. T. & VAIL, P. R. 1988. Eustatic Controls on Clastic Deposition. Conceptual Framework. SEPM, Special Publication, 42. - - - , ALLEN, G. P., JAMES, D. P. & TESSON, M. 1992. Forced regressions in a sequence framework: concepts, examples, and exploration significance. American Association of Petroleum Geologists Bulletin, 76, 1687-1769. RASMUSSEN, K. A., HADDAD, R. I. & NEUMANN, A. C. 1990. Stable-isotope record of organic carbon from an evolving carbonate banktop, Bight of Abaco, Bahamas. Geology, 18, 790-794. RICKEN, W. 1986. Diagenetic bedding: a model for marl-limestone alternations. Lecture Notes in Earth Sciences, 6, 210. RIO, M., FERRY, S. & COTILLON,P. 1989. P6riodicit6s dans les s6ries p61agiques alternantes et variations de l'orbite terrestre. Exemple du Cr6tac6 inf6rieur dans le Sud-Est de la France. Comptes Rendus de l'Acad~mie des Sciences, Paris, 309, 73-79. Schroeder, R., BUSNARDO,R., CLAVEL,B. & CHAROLLAIS, J. 1989. Position des couches h Valserina briinnimanni Schoeder et Conrad (Orbitolinid6s) dans la biozonation du Barr6mien. Comptes Rendus de l'Acad~mie des Sciences, Paris, ser. II, 309, 2093-2100. VAIL, P. R., MITCHUM,R. M. Jr., TODD, R. G., WIDMERI,J. W., THOMPSON, S., SANGREE, J. B., BUBB, J. N. & HATLELIb, W. G. 1977. Seismic stratigraphy and global changes of sea level, in seismic stratigraphy Application to hydrocarbon exploration. American Association of Petroleum Geologists Memoir, 26, 49-212. --, AUDEMARD, E, BOWMAN, S. A., EISNER, P. N. & PEREZ-CRUZ, G. 1991. The stratigraphic signatures of tectonics, eustacy and sedimentology. An overview. In: EINSELEet al. (Eds) Cycle and Events in Stratigraphy. 617-659. -
Cyclostratigraphy and correlation of the Cenomanian Stage in Western Europe A. S. G A L E
Department of Palaeontology, Natural History Museum, Cromwell Road, London SW7 5DB, UK and Department of Geology, Imperial College, Prince Consort Road, London SW7 2BP, UK Abstract: Basinal successions of Cenomanian age in Western Europe are dominated by pelagic
and hemipelagic sediments, and display conspicuous primary bedding cyclicity (couplets) which is attributed to climatically-controlled variations in carbonate productivity. Within single basins individual rhythmic couplets are widely identifiable and enable basin-wide decimetre-scale correlation. A high resolution ammonite biostratigraphy (aided by inoceramid bivalves) is used to correlate successions in separate basins, and allows a comparison of couplet numbers to be made. Couplet numbers are similar across Western Europe, in spite of an order of magnitude thickness variation, and the carbonate:clay ratios of beds and groups of beds are persistent and diagnostic of particular levels. Additionally, beds with distinctive ichnofabrics (e.g. abundant dark Chondrites) are regionally extensive in the basins of northern Europe. It has proved possible to construct a composite cyclostratigraphy for the Cenomanian, graduated by 212 precession units (mode at 21 ka), which indicate a duration for the stage of 4.45 Ma. Corroboration of this cyclochronology comes from radiometric dating; the Middle Cenomanian basal A. rhotomagense Zone to the Upper Cenomanian basal N. juddii Zone contain 107 precession couplets in Western Europe, giving a duration 2.24 Ma. The equivalent biostratigraphic interval in the Western Interior Basin of the USA recently yielded Ar-Ar dates of 2.2 Ma from sanidines in bentonites The dominance of the precession cycle in the mid-Cretaceous is in keeping with results from climate sensitivity modelling.
In the MiJnsterland and Lower Saxony Basins (NW Germany), the Anglo-Paris Basin (southern England, northern France), the Cleveland Basin (NE England) and the Vocontian Basin (SE France) (Fig. 1) the Cenomanian Stage is represented by bioturbated hemipelagic chalks and marls, which are typically rhythmically bedded on a decimetre to metre scale. Within individual basins it has proved possible to correlate these rhythmic couplets by means of distinctive marker beds, fossil occurrences and acmes, and trace fossil concentrations, as demonstrated for the Plenus Marls in the AngloParis Basin by Jefferies (1962, 1963) and more recently in part of the Middle Cenomanian of the same region (Gale 1990). This paper describes attempts at bed-scale correlations between five European basins, using a combination of biostratigraphy, marker-bed stratigraphy, and carbon isotope stratigraphy. Biostratigraphy underpins many of these correlations; the Cenomanian Stage can be finely subdivided by the use of ammonites, and accessory information is provided from the ranges and acmes of inoceramid bivalves. Cyclostratigraphy is beset by many of the general problems of fine-scale stratigraphic correlation. These include the following. (1) The recognition of minor hiatuses in or
condensation of part of the succession, which are generally suggested by sedimentary features such as the occurrence of glauconite, phosphate, intraclasts, winnowed beds, marked omission surfaces and hardgrounds. The fact that a consistent succession of couplets can be traced within and between basins is itself an argument for completeness, because it is unlikely that erosion would evenly strip a thin layer of sediment over such a wide area. In the present study one example of a basin-wide hiatus missing 3-5 couplets was discovered in the early T. costatus Subzone of the Anglo-Paris Basin. (2) The loss of rhythmicity for part of the succession is a frequently encountered difficulty in attempting to establish a cyclostratigraphy. Loss of rhythmicity is caused by swamping of clay by carbonate, or vice versa, and consequent loss of sensitivity to the climatic signal (Fischer et al. 1985 Fig. 2). (3) The development, in very expanded successions, of very fine-scale, often irregular, rhythmicity with a shorter periodicity than the Milankovitch Band. In the Vocontian Basin a rhythmicity with a mean thickness of 10 cm is very widely developed. (4) The absence of critical fossil groups over part of the succession. Aragonitic fossils, including the all-important ammonites, are commonly not
From HOUSE,M. R. & GALE,A. S. (eds), 1995, OrbitalForcing Timescalesand Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 177-197.
177
178
A . S . GALE
~~
*
f
~ Monsterlan6 Basin
..,,JUmbria-Marche [
Fig.1. Map of western Europe with Cretaceous outcrops cross-hatched. Basins mentioned in the text are indicated. Scale bar, 100 km.
preserved over considerable thicknesses in chalks, as in the Upper Cenomanian C. guerangeri Zone (cf. Marcinowksi 1980). The present study is based on field logging and macrofossil dating of typical sections for each of five basins, with additional carbon isotope data for the Anglo-Paris and Umbria-Marche Basins (Jenkyns et al. 1994). The Cenomanian of the Anglo-Paris Basin is completely exposed in five places on the coast of southern England (Kennedy 1969) and in the Boulonnais in NE France (Am6dro 1993). The Cenomanian in this region is between 50 and 90 m thick and comprises a basal condensed glauconitic sand, overlain by rhythmically bedded marls and chalks; the chalk component increases overall throughout the succession. Ammonites and inoceramids are common through most of the Cenomanian and provide the basis for a detailed zonation (see below). The cliff section south of Speeton in east Yorkshire (NE England) is the only locality in the Cleveland Basin studied here. The succession includes 40-50 m of thinly-bedded red and white pelagic limestones and thin flaser marls (Jeans 1973). .ammonites are few but inoceramid bivalves are common. The lowest part of the Cenomanian was not exposed at the time of study. The Cenomanian of the Miinsterland and Lower Saxony Basins in NE Germany (Meyer 1990; Kaplan 1992) is 100-150m in thickness, and comprises chalks and marls. Ammonites are fotmd
only at certain horizons but inoceramids are abundant throughout. Parts of the Cenomanian succession in the three northwestern European basins are remarkably similar, even down to the scale of bed details. For contrast, two deep-water Tethyan successions through the Cenomanian in southern Europe were studied. In the Vocontian Basin, in NE France, a thick (300-1000 m plus) succession of deep-water pelagic and hemipelagic thinly-bedded limestones and marls, including thin sandstone turbidites, was deposited (Porthault 1978; Cotillon 1991; Thomel 1969). The limestones contain an extensive ammonite fauna (Thomel 1969, 1992) which enables accurate correlation with the Anglo-Paris Basin. Two sections, at Vergons and St Lions, in the Alpes-Maritimes were studied. At Gubbio, in the Umbria-Marche Basin (Apennines, central Italy), a thin (50--60 m) succession of well-bedded, deepwater pelagic (nanofossil-planktic foraminiferan) limestones called the Scaglia Bianca is developed in the Cenomanian (Cresta et al. 1989), which has not yielded macrofossils, but can be dated from plankic foraminifera (Premoli-Silva 1977) and the 813C curve (Jenkyns et al. 1994). The limestone-chalk and limestone-marl couplets in the Anglo-Paris, Mtinstedand, Lower Saxony and Vocontian Basins vary in thickness from a decimetre to several metres, and petrographically comprise foraminiferal/calcisphere biomicrite and argillaceous biomicrite (Ditchfield & Marshall 1989). They show pervasive bioturbation, and the genera Thalassinoides, Planolites and Chondrites are conspicuous, especially at bed-boundaries. The chalk-marl couplets in the Anglo-Paris Basin were originally interpreted as dilution cycles (Bottjer et al. 1986), but are now widely considered to be carbonate productivity cycles (Ditchfield & Marshall 1989; Paul 1992), supported by work in progress on nannofossils as productivity indicators by Dawn Windley (University College, London). Primary lamination is found only in the marl units in deep-water settings in latest Cenomanian sediments, where it generally accompanies increased levels of organic matter (NW Germany, Hilbrecht & Hoefs 1986; Vocontian Basin, Crumi6re 1989; Gubbio, Arthur & Premoli-Silva 1982), and a single thin bed in the Middle Cenomanian of the Vocontian Basin. Thus, for much of the Cenomanian it is unlikely that the water column was strongly stratified (cf. Cotillon 1991). Robinson (1986) was the first to suggest that microrhythms in the English Chalk were formed by Milankovitch forcing, and from his calculations showed that the most likely agent was precession (mode at 21 ka). Subsequently, Hart (1987)
CYCLOSTRATIGRAPHY OF THE CENOMANIAN
|
lm
chert parting
f i n e r mad
marl
Fig. 2. Drawing of part of the Cenomanian Scaglia Bianca in the Bottacione Gorge at Gubbio, UmbriaMarche, Italy, at the level of Arthur & Fischer's (1977) 80 m marker 'bullseye'. Although this succession is very well bedded, the stratification cycles identifed by Schwarzacher (1994) as corresponding to the 100 ka eccentricity signal are not obvious, neither are precessional couplets.
attempted to construct a timescale for the Cenomanian Stage at Culver Cliff in the Isle of Wight, by counting what he called 'idealized cycles' which he presumed to be forced by obliquity (cited as 41.5 ka). By adding in an estimation of the number of cycles missing in a hiatus at the base of the Cenomanian, he counted more than 143 cycles, giving the stage a duration of c. 6 Ma. However, a unit including the higher part of the Chalk Marl (top Lower and basal Middle Cenomanian) is highly condensed at Culver, where it is 5.5 m thick and includes 17 couplets. In a thick basinal section, such as Southerham in Sussex, the same unit is 24 m thick and includes 42 couplets. Furthermore, the Albian-Cenomanian boundary in the Anglo-Paris Basin is everywhere a major hiatus representing c. about 2 Ma. This illustrates the need to understand the relative completeness of sections in the context of the entire basin; preserva-
179
tion of an entire succession of couplets cannot be assumed. De Boer (1983, Fig. 6.3) and De Boer & Wonders (1984, Fig. 1) provided a schematic log of the Cenomanian Scaglia Bianca and Livello Bonarelli graduated in precession cycles (Gilbert Units) near Moria in Umbria-Marche, Italy, which gave a total duration of the Cenomanian as 7.59 Ma. Subsequently, Schwarzacher (1994) undertook a detailed study on various Cenomanian sections in the Umbria-Marche area, for which he measured the beds and applied time series analysis to the data. His conclusions were based on identification of the stratification cycle as the 100 ka eccentricity cycle, which gave the duration of the Cenomanian as 5.8-6.2Ma. A problem with the Scaglia Bianca in Umbria-Marche, well appreciated by Schwarzacher, is the difficulty of correctly identifying and counting stratification cycles, which is absolutely critical to this method. In the opinion of this author, who has also logged the Cenomanian in Umbria-Marche, it is not possible to objectively identify a heirarchy of beds and stratification cycles through large parts of the Cenomanian succession here. To illustrate this point a typical section of Scaglia Bianca is given here (Fig. 2); bedding features are separation planes (including stylolytes), which are variably persistent laterally, irregularly-spaced thin marls, and red, pink and grey cherts. Because of difficulties of this type, the author is sceptical about a cyclochronology based on Umbria-Marche Basin sections. Gale (1990) provided a summary cyclostratigraphy for the rhythmically bedded Cenomanian of the Anglo-Paris Basin, and suggested a duration for the Cenomanian of 4.4 Ma, based on precession and eccentricity cycles. The author is now much less convinced of the ubiquity of bundling produced by the 100 ka cycle in this succession. The Cenomanian Chalk in the Anglo-Paris Basin can be divided into five disconformity-bounded units, which were identified by Robaszynski et al. (in press) as sequences. A letter was assigned to each sequence (A-E) (Gale 1990) and each couplet therein sequentially numbered. The same nomenclature has here been applied to other basins.
Stratigraphical framework for the Cenomanian of Western Europe Biostratigraphy
Ammonites provide the basic framework for biostratigraphy of the Cenomanian stage. The zonation used here is based on the work of Hancock (1960), Kennedy (1969) and Wright et al. (1984), with
180
A . S . GALE
additions based on the author's own work. Where faunas are well and continuously preserved, a total of 13 successive assemblages can be recognized through the Cenomanian (Fig.3).
inocerami( zones M. gr. mytiloides
L o w e r Cenomanian (1) Mantelliceras mantelli Zone. This is the interval zone from the appearance of M. mantelli (J. Sowerby) to the appearance of M. dixoni Spath.
planktic forams
b13C
I sequences
P.
helvetica
M. gr. opalensis
W.archaeocretacea
~1 rs SB
Inocerarnus pictus
Rotalipora cushmani
Inoceramus atlanticus
. . . . . . .
I. schond.
R.
reicheli
2
l Ts $8
noceramu: virgatus
R. brotzeni
Inocerarnus anglicus
Fig. 3. An integrated stratigraphy for the Cenomanian, showing, from left to right: 1, radiometric dates from Obradovitch (in press); 2, stages; 3, precession scale in 21 ka units, ammonite zones, inoceramid zones; planktonic foraminiferan zones (after Sliter 1989); 813C stratigraphy (after Jenkyns et al. 1994); sequence stratigraphy (after Robaszynski et al. in press). Peaks 1-3 in the 813C column are from Jenkyns et al. 1994. Although actual values of 813C vary regionally the maximum values of peak 3 (end Cenomanian) are generally slightly in excess of 5.0 ppt 813C. N.c, Neostingoceras carcitanense; S.s, Sharpeiceras schlueteri; M.s, Mantelliceras saxbii; T.c, Turrilites costatus; T.a, Turrilites acutus.
CYCLOSTRATIGRAPHY OF THE CENOMANIAN This zone can be subdivided into three subzones (Gale & Friedrich 1989). (a) Neostlingoceras carcitanense Subzone, characterised by N. carcitanense (Matheron) and Idiohamites (Algerites) ellipticus (Mantell). Commonly found as phosphatized remani6 fauna at base of Cenomanian. (b) Sharpeiceras schlueteri Subzone, characterised by S. schlueteri Hyatt and S. laticlavium (Sharpe), inflated Mantelliceras [M. cantianum Spath, M. lymense (Spath), M. mantelli] with common Hypoturrilites spp. Associated with acme of large Inoceramus crippsi crippsi Mantell. (c) Mantelliceras saxbii Subzone, containing common M. saxbii (Sharpe). (2) Mantelliceras dixoni Zone. This is the interval zone from the appearance of M. dixoni to the appearance of Cunningtoniceras inerme (Pervinqui~re). Three assemblages can be recognized in localities where ammonites are abundant and well preserved. (a) Common Mantelliceras dixoni and M. saxbii; acme of Inoceramus virgatus Schlueter. (b) Fauna containing diverse turrilitids, including Mesoturrilites spp. (c) Diverse Mantelliceras, rarer Acompsoceras, Hyphoplites.
Middle Cenomanian. (1) Cunningtoniceras inerme Zone (see Robaszynski et al. in press). (interval zone from appearance of C. inerme to appearance of Acanthoceras rhotomagense (Brongniart). Common C. inerme, C. cunningtoni (Sharpe). (2) Acanthoceras rhotomagense Zone (interval zone from entry of A. rhotomagense to entry of A. jukesbrownei (Spath). This can be divided into two successive subzones: (a) Turrilites costatus Subzone; (b) Turrilites acutus Subzone. (3) Acanthocerasjukesbrownei Zone (total range Zone of A. jukesbrownei). Upper Cenomanian. (1) Calycoceras guerangeri Zone [total range zone of C. guerangeri (Spath)]. Fauna dominated by Calycoceras spp., Eucalycoceras pentagonum (Jukes-Browne) Thomelites sp. (2) Metoicoceras geslinianum Zone ([otal range zone of M. geslinianum (d'Orbigny)]. See Wright & Kennedy (1981) for full description of ammonite fauna, commonly accompanied by the belemnite Actinocamax plenus in the middle part of the zone. (3) Neocardiocerasjuddii Zone (total range zone of N. juddii (Barrois & Gu6rne). The base of the overlying Turonian can best be defined by the entry of Watinoceras devonense Wright & Kennedy (Kennedy & Cobban 1990).
181
Inoceramid bivalves are of considerable stratigraphical use in the Cenomanian of Europe, and can be used for subdivision into between five and seven zones (Keller 1982; Tr6ger 1989; Wiedmann et aL 1989). In addition to the ranges, acmes of some species are distinctive and provide correlative horizons on a Europe-wide scale. These include the acme of large I. crippsi crippsi in the M. mantelli Zone, and L virgatus abundance in the lower part of the M. dixoni Zone. Inoceramids are extremely useful when ammonites are absent, or too rare to allow precise dating. Other macrofossil groups which are stratigraphically useful on an interbasinal scale include belemnites, calcitic bivalves and brachiopods. The belemnites Actinocamax primus (Archangelsky) and A. plenus (Blainville) are characteristic of the lowest A. rhotomagense and the M. geslinianum Zones, respectively, in all the regions studied except Umbria. The brachiopods Orbirhynchia mantelliana (J. de C. Sowerby) and Modestella geinitzi (Schloenbach) are of use in correlation between the Anglo-Paris, Cleveland, Miinsterland and Lower Saxony Basins. Planktic foraminifera, although widely used in Tethyan pelagic successions, do not offer a very fine subdivision of the Cenomanian Stage (Sliter 1989). The exact correspondence between the base of the Cenomanian as recognized using ammonites and that based on the entry of Rotalipora brotzeni (Sigal) is not established. The zone of R. brotzeni corresponds roughly to the Lower Cenomanian of ammonite workers; the zonation in the Middle and Upper Cenomanian includes successively the zones of R. reicheli Mornod, R. cushmani (Morrow) and W. archaeocretacea Pessagno. The extinction o f R. cushmani has been widely used as a datum, but may well be significantly diachronous from region to region (Gale et al. 1993).
Carbon isotope stratigraphy The major positive fi13C excursion present in the latest Cenomanian and earliest Turonian has been recognized as a useful stratigraphical marker by many workers, and is associated with the widespread deposition of organic-rich shales on the shelves and in the ocean basins at that time (Jenkyns 1980). The similarity of fine detail of this curve between Europe and Pueblo in the Western Interior Basin of the United States (Gale et al. 1993) suggests that this can be a powerful tool in high resolution correlation. Jenkyns et al. (1994) have developed a carbon stratigraphy for the whole Cenomanian, both at Gubbio in Italy and in the Anglo-Paris Basin. This demonstrates an overall increase in 513C throughout the Cenomanian
182
A . S . GALE CLEVELAND BASIN
Southerham, Lewes
Speeton
_
LOWER SAXONY BASIN
ANGLO-PARIS BASIN
Baddeckenstedt
~
Hoppenstedt
~Ii d~conformable M. Cenomanian
~- _-~_ ~
B18 __w•
prominent double |imestone bed group of rnarlier ' beds containing b 0.
mantelliana
group of (6) more carbonate - rich beds containing abundant I. virgatus thin
coarse, condensed diachronous bed transgressive lag
I
metre
~2\ BI.
Fig. 4, Correlation on a bed scale of the lower part of the M. dixoni Zone in the Anglo-Paris, the Cleveland and the Lower Saxony Basins. The same groups of calcareous and marly beds are developed in each region. The triangles represent occurrences of the brachiopod Orbirhynchia mantelliana. The *I at Speeton indicates the level of Jeans' (1973, Fig. 3) I marker. The base of the M. dixoni Zone is taken from the first occurrence ofM. dixoni at Southerham in couplet B 1.
(Fig. 3) and in the position of a minor but widely developed excursion near the base of the Middle Cenomanian, recognized by Paul et al. (in press) in the Anglo-Paris and the Cleveland Basins.
represent transgressive surfaces and sequence boundaries. The sequences display progressive onlap through the Cenomanian as global sea levels rose, with progressive spread of pelagic chalks on to the basin margins and massifs in the Middle and Upper Cenomanian. Sequence stratigraphy has not been investigated for the other basins.
Sequence stratigraphy The sequence stratigraphy of the Cenomanian has recently been described in detail for the AngloParis Basin (Robaszynski et al. in press), who identified six synchronous sedimentary sequences throughout the region. On the basin margins the successions are incomplete, include sands and contain extensively developed hardgrounds which
Integrated stratigraphy Integration of ammonite, inoceramid and planktonic foraminiferan biostratigraphy with the fi]3C curve and sequence stratigraphy (Fig. 3) provides a high-resolution framework into which most marine Cenomanian sections in Europe can
CYCLOSTRATIGRAPHY OF THE CENOMANIAN be fitted. It is now possible to add the cyclostratigraphy to this framework. Correlation of the C e n o m a n i a n
M. mantelli Zone: Couplets A 1 - 5 1 The boundary between the highest Albian and the M. mantelli Zone is marked almost everywhere by a major hiatus, involving as much as 1-1.5 Ma. Only in the Vocontian Basin in the southeast of France was sedimentation across this boundary uninterrupted, and the highest Albian here includes an ammonite zone or subzone missing in the majority of other European successions (Gale et al. unpublished work), with the possible exception of the thicker (currently unexposed) parts of the Mtinsterland and Lower Saxony Basins. At Vergons in the Alpes-Madtimes 51 couplets are present in the M. mantelli Zone (Fig. 3). The Neostlingoceras carcitanense Subzone contains 22 couplets. In the Anglo-Pads Basin only the higher part of the M. mantelli Zone is represented by a rhythmic facies, and a highly condensed glauconitic sandy marl (Basement Bed, Tourtia, Glauconitic Marl) containing phosphatized remani6 fossils of the Neostlingoceras carcitanense Subzone rests unconformably on an eroded surface of M. inflatum or S. dispar Zone sands, marls or clays. By comparison with the Vocontian Basin, the Glauconitic
183
Marl therefore represents the lowest 25-30 couplets of the Cenomanian. Rhythmic sedimentation commenced low in the S. schlueteri Subzone and couplets A31-51 are represented in the thicker successions in the Anglo-Pads Basin, such as at Folkestone. Mantelliceras dixoni Zone; Couplets B 1 - 3 8 The M. dixoni Zone is underlain by an erosional surface everywhere in the Anglo-Paris Basin, and rests with minor disconformity on the M. mantelli Zone. In southern England the base of the M. dixoni Zone progressively onlaps the M. mantelli Zone in a southwesterly direction to rest just above the hardground capping the Late Albian Upper Greensand in the southem Isle of Wight (Kennedy 1969). In the marly facies of the northern Anglo-Paris Basin the basal bed of the M. dixoni Zone is a calcisiltite containing glauconite grains and phosphatized intraclasts (the 'saxbii phosphates' of Kennedy 1969), interpreted by Robaszynski et aL (in press) as a condensed transgressive lag at the base of their third Cenomanian sequence. The subjacent surface is penetrated for up to a metre by dark-filled omission-suite Thalassinoides. In the Lower Saxony Basin, at Hoppenstedt and Baddeckenstedt (Figs 4 & 5) a coarse, hacklyfracturing bed containing inoceramid debris and meandrine sponges is developed at the equivalent
Fig. 5. View of the quarry section in the M. dixoni Zone, here disconformably overlain by Middle Cenomanian, at Baddeckenstedt, Lower Saxony (from a slide by C.J. Wood). See Fig. 4 for measured section.
184
A.S.
level (B1). At Speeton an inoceramid-rich level, just beneath marker bed I of Jeans (1973, Fig. 3) is tentatively correlated with this bed. In the Anglo-Paris Basin the lowest beds of the M. dixoni Zone (BI-10) are of variable thickness, but are everywhere surmounted by a distinctive couplet comprising a dark marl and a thin, prominent limestone (B 11). This is itself overlain by a further dark marl and a thin but less prominent, marly limestone (B12). Couplets B13-18 are relatively carbonate-rich beds characterized by a profusion of whole and fragmentary L virgatus. Couplets B l l - 1 8 are also well developed as conspicuous limestone beds of very even thickness at Speeton in the Cleveland Basin (Fig. 4). In the Vocontian Basin the hardest limestones in the succession (Fig. 6; B l l - 1 6 ) contain common L virgatus and M. dixoni and correlate with this level. B19-22 are widely developed, more clay-rich couplets, characterized by the occurrence of the small rhynchonellid brachiopod Orbirhynchia mantelliana (Band 1 of Gale 1989), and a Mantelliceras dixoni Zone (assemblage B) fauna including Mesoturrilites spp. and Acompsoceras spp. An identical occurrence of O. mantelliana in marlier beds is also found in the Cleveland and Lower Saxony Basins. In the Vocontian Basin the equivalent beds are thick, clay-rich and weakly rhythmic, and contain thin (< 1 cm) sandstone turbidites. In the Anglo-Paris Basin the marly unit with Orbirhynchia is overlain by a massive prominently weathering limestone (B23-24) at localities such as Southerham Grey (Lewes, Sussex; Fig. 4, Fig. 7A,), where it contains abundant Schloenbachia varians and diverse inoceramids. At Hoppenstedt and Baddeckenstedt (Fig. 5) in the Lower Saxony Basin a comparable double limestone bed is developed, which, at the latter locality, is overlain non-sequentially by Middle Cenomanian marly chalks. At Speeton the unit is represented by two prominently weathering limestones; the upper of the two is massive and contains vertical burrows. A group of well-developed limestones is present at B23-24 in the Vocontian Basin (Fig. 5). The successions developed in chalks and marls in the lower M. dixoni Zone of the Anglo-Paris, Lower Saxony, Cleveland and Vocontian Basins are similar in details of the lithological character (percentage of carbonate:clay) and relative thickness of individual beds (Fig. 4). The European basins are also comparable in the number of couplets present, and the occurrences and abundance of certain fossils, although there is minor variation in couplet numbers between localities. The completeness of the Lower Cenomanian
GALE
successions in the Anglo-Paris and Vocontian Basins can be illustrated by comparing couplet numbers between two distinctive biostratigraphical levels, an acme of large Inoceramus crippsi crippsi, associated with common Sharpeiceras schlueteri, up to a strong limestone near the summit of the M dixoni Zone which contains common Turrilites wiestii Sharpe (both easily recognizable in both basins). In both basins 48-49 couplets are present in this interval in the most complete successions, in spite of an order of magnitude difference in thickness. The upper part of the M. dixoni Zone and the lowest part of the Middle Cenomanian is missing from large regions of northwest Europe, on account of non-sequence or erosion over structural highs at the summit of sequence number 3 (Robaszynski et al. in press). The intervening deposits are preserved as an uncondensed succession only locally in the northeast of the Anglo-Paris Basin in southern England and northern France (Southerham Grey, Folkestone, Escalles), in the western Lower Saxony Basin (Wunstorf), and the northern Mtinsterland Basin (Teutoburgerwald).
Cunningtoniceras inerme, Acanthoceras rhotomagense and lower A. jukesbrownei Zones; Couplets B38-45, C 1 - 4 6 In southern England and northern France, where the ammonite succession is best known, there is a significant gap between the highest record of Mantelliceras (B27) and the lowest Middle Cenomanian ammonite Cunningtoniceras inerme (B38), which at Southerham Grey Pit, near Lewes, totals 8-9 m. The most expanded successions exposed in the Anglo-Paris Basin are at Folkestone, Lewes and Escalles in northern France (Gale 1989, 1990; Amrdro 1993; Fig. 5). Here (Fig. 8) a group of thin, relatively clay-rich couplets (B35-38) rests on more thickly bedded marly chalks, and is overlain by a very distinctive suite of limestones and dark marls. The lower part of B41 is a dark, silty marl which contains a distinctive fauna (partly restricted t o this level), including the bivalve Chlamys arlesiensis (Woods) and Oxytoma seminudum (Dames) and the serpulid Glandifera rustica (J. Sowerby) (Gale 1989; Paul et al. in press). The upper part of this marl is full of lightcontrasting burrows (notably Thalassinoides reworked by Chondrites) piped from the chalky limestone above (Fig. 9C). The limestone surmounting couplet B43 yields the lowest specimens of Acanthoceras rhotomagense and Inoceramus tenuis (Mantell); this bed is full of dark Chondrites piped down from the overlying C 1 (Fig. 9A).
CYCLOSTRATIGRAPHY OF THE CENOMANIAN
185
Fig. 6. Section in the Ravin Notre Dame at Vergons, Alpes-Maritimes (SE France) in the upper part of the M. mantelli and M. dixoni Zones. Couplet B 1 is Bed 4B of Thomel (1969). Note the conspicuous group of limestones B11-14 are as well developed here as in northern Europe (compare Fig. 4).
Couplet C1, k n o w n in nineteenth century literature as the Cast Bed (Gale 1989), is a dark, silty marl which at the base contains an abundance of the small scallop Entolium orbiculare (J. Sowerby) and a diverse and characteristic fauna,
including the belemnite Actinocamax primus, the echinoid Hemiaster sp., O. seminudum and diverse serpulids. A remarkably similar succession to that in the Anglo-Paris Basin, in which individual couplets
186
A.S.
GALE
Fig. 7. (A) Section in the Chalk Marl (M. dixoni-C, inerme Zones) at Southerham Grey Pit, near Lewes Sussex, UK, to show expanded succession. The Lower-Middle Cenomanian boundary falls within couplet B38. (B) Cliff section in Chalk Marl and Grey Chalk (Middle and Upper Cenomanian) at Abbot's Cliff, east of Folkestone, Kent, UK, photographed in low winter light by Dr Steve Hesselbo. The pervasive precession-driven cyclicity is well seen.
187
CYCLOSTRATIGRAPHY OF THE CENOMANIAN
COMPTON BAY, ISLE OF WIGHT VENTNOR
SOUTHERHAMBEACHY HEAD, HOLYWELL, CULVER CLIFF, LEWES, EASTBOURNE, EASTBOURNE, ISLE OF WIGHT SUSSEX SUSSEX SUSSEX
ESCALLES, BOULONNAIS
613C + excursion; upper peak pulse fauna - bivalves, belemnites, serpulids abundant
<
dark Chondrites
B42 B41
' :m~-e-
abundant ght C h o n d r i t e s )ulse fauna; Oxytoma I
1 metre
B39 !138 837
ANGLO-PARIS Lower-Middle Boundary
BASIN
Cenomanian
113s
entry ,.m-
ammonite
Cunningroniceras
r - - - -_-_~1
B 3 5 ~ ~ _-_-_-_~ B 3 4 , ~
facies chanlle; marlier above
m
!
1
Fig. 8. Correlation of the basal Middle Cenomanian succession in the northern Anglo-Paris Basin. Note how the base of the Cast Bed (couplet C1) cuts down to rest on lower levels in the B couplets at Beachy Head and in the Isle of Wight (Ventnor and Culver). Distinctive dark Chondrites (Fig. 9A & B) are invariably present in the subjacent bed.
display identical macrofaunal and ichnological characteristics, can be found in both the north Mtinsterland Basin, and in the western part of the Lower Saxony Basin (Wunstorf) in northern Germany (Fig. 10). This part of the succession falls at the level of the 'Primus Event' (after the rare occurrence of Actinocamax primus; see Ernst et al. 1983), and was described in considerable detail by Meyer (1990). Most noticeably, the equivalent of couplet B41 displays the same ichnofabric (light Chondrites and Thalassinoides burrows in a dark marl; Fig. 9D) and contains elements of the fauna characteristic of this bed in southern England. Couplet C1 is readily identifiable because it contains an identical fauna to that of the equivalent Cast Bed in the Anglo-Pads Basin, and also the distinctive darkfilled abundant Chondrites passing into the subjacent bed (Fig. 9B). Couplet C1 cuts down to rest on different levels in the highest B couplets (as in southern England) and, as it does so, the abundance of omission-suite traces in the underlying bed increases. In the two most complete German sections, Rheine and Wunstorf, two additional couplets are present above the highest present in the most complete exposed sections of the Anglo-Pads Basin (B44 and 45). The presence of the lowest of these couplets can be inferred at Folkestone from the presence of rare remani6 chalk-filled ammonites in the base of the Cast Bed (Fig. 9A).
This lateral persistence of benthic macrofossil and trace fossil occurrences over large areas of northern Europe is remarkable and indicative of the uniformity of bottom conditions in basinal marly chalks on a bed-by-bed scale. Also, these occurrences enable detailed correlations to be made which allow one to test relative completeness on a bed scale between different basins. In the Vocontian Basin (Vergons, St Lion) the C. inerme Zone is developed in irregularly-bedded, finely-rhythmic facies (sub-Milankovitch) for which a cyclostratigraphy cannot be developed. A massive, 1 m thick micritic limestone contains a rich fauna of the earliest A. rhotomagense Zone, including common lnoceramus schoendorfi Heinz is present (top Bed 11 of Thomel 1969). In marlier beds above, the 513C positive excursion of the Anglo-Paris Basin is represented by 10cm of laminated marl (Fig. 11). Twelve limestones (or bundles of thinner limestone-marls of irregular thickness) are separated by dark marls and contain an abundant Turrilites costatus assemblage fauna, dominated by the baculitid ammonite Sciponoceras baculoide (Mantell). These beds correlate directly with 12 couplets in the Anglo-Paris Basin (couplets C2-14) on the evidence of the ammonite fauna (the Vocontian succession includes some Tethyan forms rare or absent in the Anglo-Paris Basin, such as Puzosia and Gaudryceras).
188
A.S.
GALE
CYCLOSTRATIGRAPHY OF THE CENOMANIAN CLEVELAND BASIN
MUNSTERLAND - L. SAXONY BASIN
ANGLO - PARIS BASIN
t
Ir Speeton
189
ii Folkestone
Lengerich
Rheine
Wunstorf C3
C2.
-
_.
c~
[-- --S
61 3C + excursion
-_ _-_-~. pulse fauna B ..... --. ~ , . ~ d a r k Chondrites 4 5 _ S ~ abundant
-_-_-j --_--.
,, ,, .... ,, ,, '-----S---J] ~-_--~--_ -_-
_ 7--/
- "~-5_.~. _light C h o n d r i t e s abundant pulse fauna A
,
~ ~
I
facies
change
metre
Fig. 10. Correlation of the lower beds of the Middle Cenomanian between the Cleveland, Anglo-Paris, Mtinsterland and Lower Saxony Basins. The dark Chondrites which pipe down the base of the Cast Bed (C 1) are correlative across the area. Only in the two thickest successions (Rheine and Wiinstorf) are couplets B44 and B45 are preserved. The base of the C. inerme Zone is based on the lowest record of the zone fossil in couplet B38 at Southerham, Sussex, UK; the lowest occurrence of the overlying A. rhotomagense Zone index is in B43 at the same locality.
Higher A. j u k e s b r o w n e i and C. guerangeri
brownei (Bed VII of Jukes-Browne & Hill (1903)
Zones; Couplets C1-C49
at Dover, the horizon with laminated structures o f Kennedy (1969). This unit (Fig. 12), which includes eight poorly defined couplets (D1-8), was interpreted as the transgressive systems tract at the base o f the fourth Cenomanian sequence by Robaszynski et al. (in press). The basal metre contains numerous oysters of the genus Pycnodonte and several thin dark marls. Overlying this are
In the
Anglo-Paris
Basin,
the
base
of
the
A. jukesbrownei Zone contains abundant I. atlanticus Heinz in rhythmically bedded chalks. These are overlain by a roughly-weathering unit o f calcarenitic chalk containing laminated lenses of calcarenite and large poorly preserved A. jukes-
Fig. 9. Continuity of trace fossil horizons in the Middle Cenomanian between the Anglo-Paris and Miinsterland Basins. Dark-fill Chondrites and Planolites piping down the base of the Cast Bed (C l) at Folkestone (A) into B43, and at Lengerich (B) into B44. Note the pebble-preservation ammonite in the Cast Bed at Folkestone [(A), immediately above the coin), which is a remnant of couplet B44. Light-fill Chondrites, ?Gyrolithes and Planolites pipe down the upper calcareous part of B41 at Folkestone (C) and Lengerich (D). Scale bar, 10 mm.
190
A . S . GALE VOCONTIANBASIN VERGONS
ST. LK)N
ANGLO-PARISBASIN 1
ESCALLES, ROHL(~NNAI~
COMPTON, ISLF OF WIGHT
thin marly limestones; abundant $. bacu/o/de, rarer T. costatus and other ammonites of same fauna (see caption)
entry abundant $. bacu~Ide
marly chalk, weakly rhythmic 813C + excursion
I massive limestone; micrite containing low A. r / m ~ Zone ammonites
I metre
Fig. 11. Correlation of the lower part of the A. rhotomagense Zone between the Vocontian Basin (SE France) and the Anglo-Paris Basin (S England, N. France). In both successions, the base of the zone is represented by a strong limestone, overlain by a marly unit containing a positive excursion in 613C (Anglo-Pads Basin) or a thin laminated dark bed (Vocontian Basin). The overlying limestones or groups of limestones contain a fauna dominated by the straight ammonite Sciponoceras baculoide and commonly including Scaphites spp., Acanthoceras rhotomagense and Turrilites costatus. Individual beds match between the two basins. Bed numbers in the Vocontian Basin are from Thomel (1969). Note the difference in scale between the two basins.
five well-developed couplets (D9-13) of even thickness containing dark marls overlying strongly burrowed omission surfaces. The succession in Mtinsterland and Lower Saxony is very similar (Gale & Kaplan unpublished work) where the Austern (Oyster) or Pycnodonte Event (Ernst et al. 1983; Meyer 1990; Kaplan 1992) lies beneath a grey-contrasting, inconspicuously rhythmic unit. Here, the top two of the overlying five couplets are strongly developed. In northern England the calcarenitic Nettleton Stone overlies a dark marl containing Pycnodonte and is a slightly condensed representative of D2-8. The overlying succession, up to the entry of dark marls at the base of the M. geslinianum Zone, is similar in terms of the number and characteristics of beds in southern England (Anglo-Paris Basin), northern England (Cleveland Basin) and northwest Germany (Mtinsterland and Lower Saxony Basins; Gale & Kaplan unpublished work). A group of six beds, seperated by thin dark marls in exactly
equivalent stratigraphical positions, is found in all three regions. In conclusion, the succession between the base of the A. rhotomagense Zone and the base of the M. geslinianum Zone in the Anglo-Paris, the MiJnsterland and the Lower Saxony Basins can be correlated very accurately by the use of faunalmarker horizons and distinctive beds of limestone or marl, and bed boundaries dominated by trace fossil abundances. This allows a direct comparison of numbers of marl-chalk and marl-limestone couplets between the two areas, and for the same intervals in the Vocontian Basin and the UmbriaMarche Basin (determined from the 613C curve, Fig. 13). All regions (see Table 1) contain between 90 and 95 couplets, even though there is an order of magnitude difference in thickness between these successions. In the Cleveland Basin the highest part of the C. guerangeri Zone is missing (cf. Gale et al. 1993) and only 84 couplets are present.
191
CYCLOSTRATIGRAPHY OF THE CENOMANIAN
)20 )19 :)18
rhythmicity lost in white chalk facies
:)16
D1S D14 DIZ__ S conspicuous rhythms; dark marls overly intensely burrowed surfaces on D8-11
Dll _ _ DIOI__
D8
grey-weathering calcarenitic marls abundant laminated structures and large
D7
Acanthoceras. Pycnodonte common
D9 I _ _
in lowest metre D6
recessing dark marl
D5 DOVER, KENT D4
D2 D1
/ CULVER CLIFF, /ISLE OF WIGHT
EASTBOURNE, SUSSEX 1 metre
SWANAGE, DORSET
COMPTON BAY, ISLE OF WIGHT
ANGLO-PARIS BASIN Fig. 12. Correlation of the succession at the top of the Middle Cenomanian (A. jukesbrownei Zone) in the northern Anglo-Pads Basin. Above a basal marl (D1), eight poorly defined couplets of chalk and marly chalk (D1-8) contain concave-up laminated structures which are probably burrow fills. Overlying this are five well-marked couplets which comprise dark marls and lighter comprise dark marls and lighter chalk (D9-13). The top of the A. jukesbrownei Zone is taken at the highest occurrence of the zonal ammonite in D 13 at Dover.
192
A.S.
LIVELLO BONARELLI
GALE
613C +EXCURSION
UPPER SCAGLIA BIANCA MIDDLE AND UPPER CENOMANIAN, BOTTACIONE GORGE, GUBBtO
o
a13C +EXCURSION (M.CEN)
3C ppt z.-~
i
,
L
i
2.5
PDB
i
Fig. 13. Middle and Upper Cenomanian Scaglia Bianca and Livello Bonarelli at the Bottacione Gorge, Gubbio, Italy, to show the 813C curve (data from Jenkyns et al. 1994). Note the double positive excursion in the Middle Cenomanian, and the end-Cenomanian major positive excursion. The interval between the two excursions appears to contain 90-95 couplets, similar in number to other basins in Western Europe (see Table 1).
CYCLOSTRATIGRAPHY OF THE CENOMANIAN Table 1. Numbers of bedding couplets between higher peak of $13C excursion in Middle Cenomanian and base of ~13C excursions in Upper Cenomanian in five European basins Basin Anglo-Paris Mtinsterland Vocontian Umbria-Marche Cleveland
Thickness
Number of couplets
40-50 40-60 200 17 15, incomplete
94 92-95 92 90-95 84
M. g e s l i n i a n u m - N , juddii Zones; Couplets E l - 1 7 This interval includes the lower part of the Cenomanian-Turonian 513C excursion, which can be widely identified and correlated on a global scale (Gale et al. 1993; Jenkyns et al. 1994). The work of Jefferies (1962, 1963) provides a detailed documentation of the lateral continuity of beds of the Plenus Marls in the Anglo-Paris Basin (Fig. 14). Jefferies was able to recognize eight successive beds and characterize each in terms of its fauna, lithology and erosional/gradational relationship to adjacent beds. The base of Bed 1 rests on the sub-plenus erosion surface (Jefferies 1962) and is piped down some distance (c. 1 m) into the subjacent chalk in Thalassinoides and other burrows. Robaszynski et al. (in press) interpret this surface as a sequence boundary at the base of the highest Cenomanian sequence. Locally, the erosional surface cuts down a considerable distance into the underlying Grey Chalk (C. guerangeri Zone), as seen in the north of the basin by JukesBrowne at Hitchin (Jukes-Browne & Hill 1903), and in the southeast of the basin at Cr6santignes (Jefferies 1962), where a conglomerate of chalk boulders rests on the surface. It is important to point out that the Plenus Marls or Plenus Formation, as used in the North Sea or Cleveland Basins (e.g. Jeans et al. 1991), does not correspond in age or facies to the Plenus Marls in the Anglo-Paris Basin (Gale et al. 1993). In NW Germany, in the Mtinsterland and Lower Saxony Basins, the equivalent of the base of the Plenus Marls is marked by a striking lithological change from underlying thinly-bedded chalks to a dark marl - the Fazieswechsel (facies change) (Meyer 1990; Kaplan 1992). The overlying group of three couplets are directly equivalent to those in the expanded Bed 1 of the Plenus Marls in more complete successions (e.g. Eastbourne). The Chondrites Event, which contains abundant lightcontrasting burrows (mostly Chondrites, rarer Taenidium) is equivalent to the light-Chondrites
193
abundance at the junction of Beds 2 and 3 of the Anglo-Paris Basin (e.g. Gale et al. 1993 Fig. 1). In the Lower Saxony and Mtinsterland Basins the Chondrites Event is overlain by the Plenus Bank, a nodular limestone which contains very rare specimens of Actinocamax plenus in the upper part. The lower part of the Plenus Bank probably correlates with Bed 3 of the Plenus Marls, with Beds 4--5 condensed into the upper part. The top of the Plenus Bank is a major hiatus in many localities, missing out the highest part of the M. geslinianum and the whole of the N. juddii Zones. The beds above the Plenus Marls in the AngloParis Basin are calcisphere- and inoceramid-rich nodular chalks with wispy flaser-marls, called the Melbourne Rock Beds in southern England. The bed stratigraphy of this unit can be followed around the basin, in considerable detail, from the use of marker marls and fossil concentrations (Fig. 14). In the Vocontian Basin a sharp facies change from thinly-bedded pelagic limestones to weakly rhythmic dark marls underlies the organic-rich, laminated Niveau Thomel which represents the maximum of anoxia in the region (Crumi6re 1989). In the absence of ammonites, and lack of carbon isotope data, no precise correlation is currently possible with northern Europe.
Discussion A striking feature of the rhythmic chalk-mad successions in the different European basins is the persistence of beds and groups of beds characterized by particular carbonate:clay ratios. In the M. dixoni Zone a group of six or seven purer limestones (B 11-17) are overlain by a marly succession capped by two limestone beds (B23 and 24). Similarly, the base of the A. rhotomagense Zone is widely characterized by a massive pure limestone (B43-45, depending on locality), succeeded by poorly defined marly couplets (C1-5), overlain by an alternation of thin limestones and marls (C6-17). The formation of these beds was presumably under continent-wide climatic control which affected very precisely both the production of carbonate and the flux of detrital clay over the entire region. Although the differences are obvious to the eye in weathered cliff profiles the actual variation in carbonate:clay ratio is often only a few per cent. Correlation between basins can sometimes be made in extraordinary detail by use of lithology, macro- and trace-fossil events, as in the C. inerme and lower Acanthoceras rhotomagense Zone in the Cleveland, Anglo-Paris, Mtinsterland and Lower Saxony Basins (Fig. 10). More commonly, it is possible only to correlate widely spaced horizons between basins by the use of biostratigraphy or
194
A . S . GALE :
",!
.I
i
t
Compton
Pebble Marl
9L
4 Roveacrinus
Bed; Fagesia catinus
-r~r
la 1
~a Z
Bed B ~ , ; Bed Bed 4 Bed 3 Bed 2
DOVER, KENT
Bed
Bed
1
COMPTON BAY, CULVER CLIFF, ISLE OF WIGHT ISLE OF WIGHT
SUIt-VOilE
SWANAGE, DORSET
light Chondrites
1 me,,e [
E3
\
F.2 sub-I~enus erosion surface
EASTBOURNE, SUSSEX
Fig. 14. Correlation of the Plenus Marls and Melbourn Rock Beds across the Anglo-Paris Basin, from the Dorset coast (Swanage, southern England) to Courcelles-sur-Voire (Aube, SE Paris Basin) in the southeast of the Paris Basin.
~13C excursions and compare the intervening numbers of couplets. Because of the presence of basinwide hiatuses, a complete cyclostratigraphy is necessarily a composite, using parts of different sections. This is comparable to the method called stacking used to construct composite 8180 curves for Neogene ODP cores (Raymo et al. 1990). The Lower Cenomanian contains numerous, and sometimes large, gaps in most of the regions studied, with the exception of the Vocontian Basin.
In particular, the Albian--Cenomanian boundary is a major level of non-sequence and condensation in many regions, and one ammonite subzone is widely missing from the top of the Albian. Rising sea levels throughout the Cenomanian (Robaszynski et al. in press) progressively increased the available accommodation space on shelves and allowed widespread preservation of a complete cyclostratigraphy in the Middle and Upper Cenomanian.
CYCLOSTRATIGRAPHY OF THE CENOMANIAN The couplets present in the Cenomanian are of rather even thickness and do not show significant bundling. They were probably generated as productivity cycles controlled by the precession signal (mode at 21 ka) in the Milankovitch Band, which is probably the single commonest lithological expression of orbital forcing. For the interval from the base of the A. rhotomagense Zone to the basal N. juddii Zone, 107 such couplets are present in Western Europe, which give a precession timescale of 2.24 Ma. The equivalent interval in the Western Interior of the USA (C. gilberti to N. juddii Zones) has recently been dated at 2.2 Ma (Obradovitch in press) by the Ar-Ar method using sanidines from bentonites. These dates are remarkably close, considering the potential error on the radiometric dates, and support the interpretation that the couplets are precession cycles. In Fig. 3, the cyclochronology is presented as a scale alongside an integrated stratigraphy for the Cenomanian. For this interval of Mesozoic time it is thus possible to investigate rate processes against an orbital timescale. One interesting feature that emerges from this is the remarkably even duration of sedimentary sequences at 800 Ka to 1 Ma (Gale 1991), each of which contain 45-50 precession couplets. The timescale also demonstrates that the latest Cenomanian transgression, in the high M. geslinianum and N. juddii Zones, took place very rapidly. In conclusion, it is possible to construct Milankovitch timescales for particular intervals based on study of bedding rhythms, with certain provisos:
(1) That high-resolution cycle correlation between basins can be obtained with the aid of biostratigraphy, augmented wherever possible,
195
by other independent means, enabling accurate comparison of couplet numbers. (2) Hiatuses can certainly be very hard to identify in rhythmic hemipelagic and pelagic sediments, and completeness on a couplet scale cannot be assumed (cf. Weedon 1991) and must be demonstrated both within basins and from the comparison of several different basins. This work shows that it is possible to preserve complete successions of couplets over long time intervals in relatively shallow watershelf sediments, which was probably enabled in the later Cenomanian by unusually large accommodation space during a time of overall sea-level rise. (3) It becomes increasingly probable that decimetre to metre scale bedding couplets in Cretaceous pelagic and hemipelagic sediments are dominantly a product of the 21 ka precession cycle (cf. Fischer 1991), commonly (but not invariably) modified into bundles or stratification cycles within the short eccentricity cycle (El, 100 Ka). In their study of a long time-series in the Lower Cretaceous (Berriasian-Barremian) of the Vocontian Basin, Rio et al. (1989) found an average duration of 21 ka for bedding cycles using various independent radiometric timescales. From atmospheric GCM simulations, Park & Oglesby (1990, 1991) concluded that the midCretaceous climate system was more sensitive to precession than obliquity in most regions. During this research I have benefited from greatly from discussions with many Cretaceous and other colleagues, including C. J. Wood, U. Kaplan, J.-E Crumi~re, J. M. Hancock, A. G. Fischer, M. R. House and D. G. Smith. G. Ernst kindly showed me the sections in Lower Saxony, and U. Kaplan those in Miinsterland.
References
AMI~DRO,E 1993. La lithostratigraphie et les biofacies: des outils de correlation dans les craies Cenomani6nnes du detroit du Pas de Calais. Annales de la Soci(t( G(ologique du Nord, 2 (2nd series), 73-0. ARTHUR,M. A. & DEAN,W. E. 1991. A holistic approach to cyclomania: examples from Cretaceous pelagic limestone sequences. In: EINSELE,G., RICKEN,W. & SEILACHER, A. (eds) Cycles and events in stratigraphy. Springer-Verlag, Berlin. 126-166. & FISCHER, A.G. 1977. Upper CretaceousPaleocene magnetic stratigraphy at Gubbio, Italy. 1. Lithostratigraphy and sedimentology. Bulletin of the Geological Society of America, 88, 367-371. - & PREMOLI-SILVA,I. 1982. Development of widespread organic-rich strata in the mediterranean Tethys. In: SCrILANGER,S. O. & CITA, M. B. (eds) Nature and Origin of Cretaceous Carbon-rich Facies. Academic Press, London, 7-54.
BOTTJER, D. J., ARTHUR,M. A., DEAN, W. E., HATTIN, D. E. & SAVRDA,C. E. 1986. Rhythmic bedding in Cretaceous pelagic carbonate environments sensitive recorders of climatic cycles. Paleoceanography, 1, 467-481. COTn.LON, E 1991. Varves, beds, and bundles in pelagic sequences and their correlation. In: EINSELE, G., RICKEN, W. • SEILACHER,A. (eds) Cycles and events in Stratigraphy. Springer-Verlag, Berlin, 820-839. CRESTA, S., MONECHI,S. & PARISI,G. 1989. MesozoicCenozoic stratigraphy in the Umbria-Marche area. Memoire descrittive della Carta geologica d'Italia, 39. CRUMI~RE, J.-E 1989. La seri6 c6nomano-turonienne en fosse vocontienne orientale et signification du black shale "Thomel". Geotrope, 1, 84-94. DE BOER,P. L. 1983. Aspects of Middle Cretaceouspelagic sedimentation in southern Europe: production
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and storage of organic matter, stable isotopes and astronomical influences. Geologica Ultraiectina, 31. - & WONDERS, A. A. 1984. Astronomically induced rhythmic bedding in Cretaceous pelagic sediments near Moria, Italy. In: Berger, A. L. (ed.) Milankovitch and climate. Reidel Publishing Company, Dordrecht. - & MARSHALL, J. D. 1989. Isotopic variation in rhythmically bedded chalks: paleotemperature variations in the Upper Cretaceous. Geology, 17, 841-845. EINSELE, G. & RICKEN, W. 1991. Limestone-marl alternation - an overview. In: EINSELE,G., RICKEN, W. & SEILACHER, A. (eds) Cycles and events in stratigraphy. Springer-Verlag, Berlin, 23-47. ERNST, G., SCHMID, E & SEIBERTZ, E. 1983. Eventstratigraphie im Cenoman und Turon yon N-W Deutschland. Zitteliana, 10, 11-46. FISCHER, A. G. 1991. Orbital cyclicity in Mesozoic strata. In: EINSELE,G., RICKEN,W. & SEILACHER,A. (eds) Cycles and events in stratigraphy. Springer-Verlag, Berlin, 48--62. , HERBERT, Z . & PREMOLI-SILVA,I. 1985. Carbonate bedding cycles in Cretaceous pelagic and hemipelagic sequences. In: PRATT, L. M., KAUFMANN, E. G. & ZELT, E B. (eds) Fine Grained Deposits of the Cretaceous Western Interior Seaway: Evidence of Sedimentary Processes. SEPM Second Annual Midyear Meeting, Golden, Colorado. Field trip no. 9, August 15th 1985. GALE, A. S. 1989. Field meeting at Folkestone Warren, 29th November, 1987. Proceedings of the Geologists'Association, 100, 73-80. 1990, A Milankovitch scale for Cenomanian time. Terra Nova, 1, 420-425. 1991. Inter-basinal correlation of cyclothems and chalk-marl couplets in the Cenomanian of western Europe: rates of eustatic change determined from precession cycles. Terra Abstracts, 3, 278. & FRIEDRICH,S. 1989. Occurrence of the ammonite Sharpeiceras in the Lower Cenomanian Chalk Marl of Folkestone. Proceedings of the Geologists' Association, 100, 80-82. , JENKYNS,H. C., KENNEDY,W. J. & CORFIELD,R. M. 1993.Chemostratigraphy versus biostratigraphy: data from around the Cenomanian-Turonian boundary. Journal of the Geological Society, London, 150, 29-32. HANCOCK, J. M. 1960. Les ammonites du C6nomanien de la Sarthe. Comptes rendu congres Societds Savantes - Dijon 1959: Colloque sur le crdtacd supdrieur frangais, 249-52. HART, i . B. 1987. Orbitally induced cycles in the Chalk facies in the United Kingdom. Cretaceous Research, 8, 335-348. HILBRECHT, H. & HOEFS, J. 1986. Geochemical and palaeontological studies of the ~13C anomaly in boreal and north Tethyan Cenomanian-Turonian sediments in Germany and adjacent areas. Palaeogeography, Palaeoclimatology, Palaeoecology, 53, 169-189. JEANS, C. V. 1973. The Market Weighton Structure: tectonics, sedimentation and diagenesis during the Cretaceous. Proceedings of the Yorkshire Geological Society, 39, 409--444.
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LONG, D., HALL, M. A., BLAND, D. J. & CORNFORD, C. 1991 The geochemistry of the Plenus Marls at Dover, England: evidence of fluctuating oceanographic conditions and of glacial control during the development of the CenomanianTuronian ~13Canomaly. Geological Magazine, 128, 603-632. JEFFERIES, R. P. S. 1962. The palaeoecology of the Actinocamax plenus subzone (lowest Turonian) in the Anglo-Paris Basin. Palaeontology, 4, 609-647. 1963. The stratigraphy of the Actinocamax plenus subzone (Turonian) in the Anglo-Paris Basin. Proceedings of the Geologists' Association, 74, 1-34.
JENKYNS, H. C. 1980. Cretaceous anoxic events: from continents to oceans. Journal of the Geological Society of London 137, 171-188. , GALE, A. S. & CORFIELDR. M. 1994. Carbon- and oxygen-isotope stratigraphy of the English Chalk and Italian Scaglia and its palaeoclimatic significance. Geological Magazine, 131, 1-34. JUKES-BROWNE, A. J. & HILL, W. 1903. The Cretaceous rocks of Britain. II. The Lower and Middle Chalk of England. Memoirs of the Geological Survey of the United Kingdom, 568 pp. KAPLAN,U. 1992. Die Oberkreide Aufschlusse im Raum Lengerich/Westfalen. Geologie und Paliiontologie in WesOealen,21, 7-37. KELLER, S. 1982. Die Oberkreide der Sack-Mulde bei Alfeld (Cenoman-Unterconiac): Lithologie, Biostratigraphie und Inoceramen. Geologisches Jarbuch (A), 64, 3-171. KENNEDY,W. J. 1969. The correlation of the Lower Chalk of south-east England. Proceedings of the Geologists' Association, 80, 459-560. - • COBBAN, W. A. 1991. Stratigraphy and interregional correlation of the Cenomanian-Turonian transition in the Western Interior of the United States near Pueblo, Colorado, a potential boundary stratotype for the base of the Turonian stage. Newsletters on Stratigraphy, 24, 1-33. MARCINOWSKI, R. 1980. Cenomanian ammonites from the German Demoratic Republic, Poland and the Soviet Union. Acta Geologica Polonica, 30, 215-325. MEYER, T. 1990. Biostratigraphische und sedimentologische Untersuchungen in der Pliinerfazies des Cenoman yon Nordwestdeutschland. Mitteilungen aus dem Geologischen Institut der Universit~it Hannover, 30. OBRADOVITCH, J. In press. A Cretaceous time scale. in: CALDWELL,W. G. E & KAUFMANN,E. G. (eds) Evolution of the Western Interior Foreland Basin. Special Paper of the Geological Association of Canada. PARK, J. & OGLESBY, R. J. 1990. A comparison of precession and obliquity effects in a Cretaceous paleDclimate simulation. Geophysical Research Letters, 17, 1929-1932. & 1991. Milankovitch rhythms in the Cretaceous: a GCM modelling study. Global and Planetary Change, 4, 329-355. PAUL, C. R. C. 1992. Milankovitch cycles and microfossils: principles and practice of palaeoecological analysis illustrated by Cenomanian chalk-marl
CYCLOSTRATIGRAPHY OF THE CENOMANIAN rhythms. Journal of Micropalaeontology, 11, 95-105. , MITCHELL,S. E, MARSHAL,J. D., LEAREY, P. N.,. GALE, A. S., DUANE, A. M. & DITCHFIELD,P. D. In press. Palaeoceanographic events in the Middle Cenomanian of Northwest Europe. Cretaceous Research. PORTHAULT, B. 1978. Foraminiferes caracteristiques du Cenomanien a facies pelagique dans le Sud-Est de la France. G~ologie Mediterran~enne, 5, 183-194, PREMOLI-SILVA, I. 1977. Upper Cretaceous-Paleocene magnetic stratigraphy at Gubbio, Italy, II. Biostratigraphy. Bulletin of the Geological Society of America, 88, 371-374. RAYMO, M. E., RUDDIMAN, W. E, SHACKLETON,N. J. & OPvo, D.W. 1990. Evolution of Atlantic-Pacific 81aC gradients over the last 2.5 m.y. Earth and Planetary Science Letters, 97, 353-368. RaCKEN, W. 1991. Time-span assessment-an overview. In" EINSELE,G., RICKEN, W. & SEILACHER,A. (eds) Cycles and events in stratigraphy. Springer-Verlag, Berlin, 773-794. Rio, M. FERRY, S. & COTILLON, E 1989. P6dodiciti6s dans les s6ries p61agiques alternantes et variations de l'orbite terrestre. Example du Cr6tac6 inf6rieur dans le Sud-Est de la France. Comptes Rendus de l'Academie des Sciences, Paris, 309 ser. 2, 73-79. ROBASZYNSKI, E, JUIGNET, P., GALE, A.S., AMEDRO,E & HARDENBOL, J. In press. Sequence stratigraphy in the Cretaceous of the Anglo-Paris Basin, exemplified by the Cenomanian Stage. In: JAQUIN, T., DE GRACIANSKY, P. & HARDENBOL, J. (eds) Sequence Stratigraphy. Society of Economic Paleontologists and Mineralogists, Special Publication. ROBINSON, N. D. 1986. Fining-upward microrhythms with basal scours in the Chalk of Kent and Surrey, England and their stratigraphic importance. Newsletters on Stratigraphy, 17, 21-28.
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measurement of time. In: EINSELE, G., RICKEN, W SEILACHER, A. (eds) Cycles and events in stratigraphy. Springer-Verlag, Berlin, 855-863. 1994. Cyclostratigraphy of the Cenomanian in the Gubbio district. In: DE BOER, P. L.& SMITH, D. G. (eds) Orbital forcing and cyclic sequences. Special Publication of the International Association of Sedimentologists, 19, 87-97. SLITER, W. V. 1989. Biostratigraphic zonation for Cretaceous planktonic foraminifers examined in thin section. Journal of Foraminiferal Research, 19, 1-19. THOMEL, G. 1969. Etudes stratigraphiques et paldontologiques du Cdnomanien subalpin entre Digne et Menton. Thesis, University of Nice. -1992. Ammonites du Cdnomanien et du Turonien du sud-est de la France. Serre Editeur, Nice. TROGER, K.-A. 1989. Problems of Upper Cretaceous inoceramid biostratigraphy and paleobiogeography in Europe and Western Asia. In: WIEDMANN,J. (ed.) Cretaceous of theWestern Tethys. Proceedings of the 3rd International Cretaceous Symposium, Ttibingen 1987, Schweizerbart'sche, Stuttgart, 911-930. WEEDON, G. 1991. The spectral analysis of stratigraphic time-series. In: EINSELE, G., RICKEN, W. & SEmACHER, A. (eds) Cycles and Events in Stratigraphy. Springer-Verlag, Berlin, 840-854. WIEDMANN,J. KAPLAN,U., LEHMANN,J. & MARCINOWSKI, R. 1989. Biostratigraphy of the Cenomanian of NW Germany. In:WIEDMANN,J. (ed.) Cretaceous of the Western Tethys. Proceedings of the 3rd International Cretaceous Symposium, Ttibingen 1987, Schweizerbart'sche, Stuttgart, 931-948. WRIGHT,C. W. & KENNEDY,W. J. 1981. TheAmmonoidea of the Plenus Marls and the Middle Chalk. Palaeontographical Society (Monograph), London. - - & HANCOCK,J. M. 1984. The Ammonoidea of the Lower Chalk. Introduction. Palaeontographical Society (Monograph). -
Cyclostratigraphy, Quo Vadis? ALFRED
G. F I S C H E R
Department o f Geological Sciences, University of Southern California, Los Angeles, CA 90089-0740, USA
Abstract: Climatic oscillations forced by orbital variations left their imprint not only on Pleistocene ice regimes, as postulated by Croll (1875, Climate and Time in their Geological Relations, Appleton, New York) and by Milankovitch (1941, Belgrade, Serbian Academy of Science, 133), but also on the sedimentary record as postulated by Gilbert (1895, Journal of Geology, 93, 515-533), and not only in glacial times but also, and particularly sharply, in times of greenhouse climates. The amenable facies are those characterized by relatively continuous sedimentation, though in carbonate platforms interruptions constitute the Milankovitch signal. The recording mechanisms include variations in biogenic components and biotic activity, and direct expressions of physical and chemical sedimentation. So far, cyclostratigraphy has largely remained a demonstration of principle. Applications to geochronology are in their beginnings. The chronological refinement promised by cyclostratigraphy demands quantities of stratigraphic data orders of magnitude beyond those of conventional stratigraphy. Progress in application will therefore depend on combining automated scanning devices (well logs, core and sample logging techniques of various sorts) capable of providing simple variation graphs of long sedimentary sequences, with detailed paleontological and magnetostratigraphic studies. The existence of the cycle hierarchy will provide some protection against chronological errors induced by cryptic gaps in the record, as will the reproduceability of results in distant sequences and different facies, and the possible fingerprinting of individual cycles by magnetic character.
G. K. Gilbert suggested in 1895 that the Earth's sediments contain within themselves a record of orbital variations: a built-in chronometer. Geologists were slow to respond, but gradually the existence of such pulse beats has become apparent. Already they have become the standard for Quaternary history and their application to the refinement of older histories is merely a matter of time.
Pelagic rhythmicities - Quaternary and older Quaternary ice sheets oscillated in complex rhythms that followed the cycles of the orbital variations, as suggested by Croll (1875) and further developed by Milankovitch (1941). The key was found in the isotopic composition of foraminiferal tests in deep-sea cores (Imbrie 1985), and the SPECMAP isotope curve, constructed from these, now serves as a standard for Quaternary stratigraphy and chronology. Marine Pleistocene events recorded in unaltered carbonate can probably be dated to within 20 000 a. It is to be recalled, however, that the effort required to bring us to this level
has not been small: it has involved many people for half a century, has involved much marine coring and required the installation, operation and maintenance of many mass spectrometers. I should think that the effort, viewed realistically, has probably cost on the order of $50 million, and the slice of geological time that has come into sharp historical focus is c. 0.3% of Phanerozoic time. We are now concerned with the other 99.7%! The applicability of oxygen isotopes to geochronology is tied to large fluctuations in ice volume, and is therefore (and for other reasons) not likely to be very useful for carrying the chronology far into earlier times. Not only isotopes, but carbonate content and various other parameters of Quaternary pelagic cores, show similar rhythmicities. Some of these are related to variations in influx of muddy meltwater from ice sheets (Roof et al. 1991), others to aeolian-dust transport linked to glacial times (Tineke et al. 1991; Weedon & Shimmield 1991), and yet others to carbonate productivity (Arrhenius 1952; Herbert & Mayer 1991). Many of these sequences show stronger precessional signals than do the isotope curves, and the calcium carbonate curve changes phase from the Pacific to the Atlantic Oceans. Those differences imply a more direct response to insolation
From HOUSE,M. R. & GALE,A. S. (eds), 1995, Orbital Forcing Timescales and Cyclostratigraphy, Geological Society Special Publication No. 85, pp. 199-204.
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regimes and to regional changes in atmospheric and oceanic behaviour, as contrasted to the global (ice volume) signal of the isotope curves. In pelagic carbonates these same patterns extend back through time, into the Cretaceous and presumably into the late Jurassic - the very beginnings of pelagic chalk sedimentation. Retrojections of astronomically calculated eccentricity cycles serve to provide a chronological frame back into the Pliocene (Hilgen 1991) and potentially through much of Neogene time, but for earlier times the record will have to stand on its own merits. A particularly interesting aspect of many pelagic sequences is a coherent double cyclicity: as carbonate productivity changed in the upper waters (mixed layer), oxygen availability changed in the bottom waters, recorded in preservation of organic carbon as well as in the intensity of bioturbation and in the compositionof the trace-fossil fauna: oxygen deficit in bottom waters coincided with minimal carbonate production above. This redox cyclicity is strongly developed in Cretaceous sediments (Savrda & Bottjer 1989; Fischer et al. 1991), but may also be observed in the British Jurassic and in the Eocene of Angola (Fischer, pers. obs.). Similar redox cyclicity has been reported from the Mio-Pliocene of Japan (Sakamoto 1992). It seems to have occurred mainly in the oxygen minimum layer in the upper thermocline, but reached down to abyssal bottoms in the great cul-de-sac of the Cretaceous North Atlantic (Arthur & Dean 1986). The cyclicity, normally apparent at the precessional level, may be arrested for longer episodes during 'anoxic events', such as those of the MidCretaceous (Jenkyns 1980), and for much longer periods in restricted parts of epicontinental seas such as America's Western Interior Seaway of Cretaceous time (Pratt et al. 1985). The response of carbonate producers (coccolithophoracean algae, calpionellid infusorians, globigerinacean foraminifera) to orbital cyclicity is not wholly clarified. The carbonate production patterns of Pacific plankton during Pleistocene time suggest to the author that maximal blooms of carbonate producers occur at times when the eccentricity is low, and probably especially during the precessional phase when aphelion occurs in summer: those are the times of lowest seasonal contrast, when zonal circulation and wind velocities are at their maximum, when hemispheric contrasts are least, gyres are most active, and when the mixed layer is likely to be thickest and well stirred. If this be correct, then maximal burial of organic matter occurred in the opposed phases, reflecting the high-eccentricity phase of the eccentricity cycle and the summer-perihelion phase of the precession. These are the times of maximal disruption of zonal winds by monsoonal disturbances.
Eustasy and platform emergence Other facies serve to trace the pattern into earlier times: the large sea-level oscillations of the Pleistocene have their counterpart in the Late Carboniferous and Early Permian, as recorded in the classical cyclothem transgressions of the North American Interior. But there is a difference: Pleistocene sea-level oscillations, resulting from ice flux, were driven mainly by the c. 40 000 a obliquity cycle and the c. 100 000 a eccentricity cycle, and marine deposits show good evidence of the c. 20 000 a precession. In the Carboniferous, on the other hand, longer cycles, possibly that of the 400 ka eccentricity, seem to have played a larger role, while high-frequency cyclicity seems poorly developed (see curves of Boardman &Heckel 1989). Did the Carboniferous climates have greater inertia than Pleistocene ones? Was the ice under orbital control? Small sea-level oscillations are strikingly displayed in the emergence patterns of Triassic carbonate platforms and show the 5 : 1 precession : eccentricity ratio (Schwarzacher & Haas 1986; Goldhammer et al. 1987). Do they reflect fluctuations in mountain glaciers and perhaps small polar ice caps? Or was sea level driven orbitally by some unresolved factor? Shoaling and emergence cycles, less strikingly rhythmic, are known from carbonate platforms extending back through the early Paleozoic into the Proterozoic. Not only platforms proper, but also their carbonate aprons, show cyclicity. Tuning suggests orbital timing for Cambrian sequences. Shoaling or emergence cycles have been observed in Cretaceous carbonate platforms, but the patterns and the case for orbital timing are less compelling than are those of the Trias.
Deep-water fans For some time there have been suggestions that turbidite sediments of deep-water fans may contain Milankovitch signals (Foucault et al. 1987; Van Tassel 1987). Whereas, on the one hand, the sands of these sequences were viewed as products of stochastic slump-failure on delta fronts, the rise of sequence stratigraphy suggested the possibility of control by sea-level changes with apparent ties to orbital cyclicity. A recent paper (Weltje & De Boer 1993) on a Pliocene pro-delta fan on Corfu uncovers yet another aspect: bundles of sands, inerpreted as individual fan lobes, have Milankovitch timing. Also they show correlated changes in sand maturity and lignite content. It would seem, then, that pro-delta sedimentation is also sensitive to orbitally driven climatic change in the drainage basins feeding
CYCLOSTRATIGRAPHY, QUO VADIS.9
deltas, and that the pro-delta facies may represent another good target for cyclostratigraphic research.
Oscillations in precipitationevaporation balance The alternation of wet and dry times of Lake Playa complexes in the Eocene (Fischer et al. 1991) and in the Triassic-Jurassic (Olsen 1986) show striking Milankovitch cyclicity with strong precessional signals. Deep evaporite embayments, such as that of the Permian Castile formation (Delaware Basin; Anderson 1982), also furnish an excellent record in which the periodicities are datable by varves. These contain the earliest rigorously established precessional signals recognized to date.
Uncertainties and the status quo The existence of recognizable, orbitally-induced rhythms in stratigraphic sequences has now been established, in principle, not only for times of glaciation but also for such greenhouse episodes such as the Middle Cretaceous: the Milankovitch hierarchy of precession-obliquity-eccentricity appear to be particularly well developed in the greenhouse state. In the marine record it has now been traced back into the pelagic Cretaceous (Fischer et al. 1991) and into Triassic platforms (Schwarzacher &Haas 1986), in the lacustrine record into the Triassic (Olsen 1986). The precessional record (and a hint of the 100 ka eccentricity) have been recognized in the Permian (Anderson 1982). Cyclic patterns exist in Palaeozoic sediments, but their identification becoms more difficult, for at least two reasons: (1) the uncertainties of geochronology increase the farther we reach back into history, making it more difficult to assign periods to such rhythms; (2) whereas, it appears most unlikely that the eccentricity cycles have undergone change during Earth history, the ratios of precession, and especially obliquity cycles, to the eccentricity cycles must have changed along with the deceleration of Earth spin. Berger et al. (1992) have estimated the shift, but the actual rate of deceleration remains poorly constrained. On the one hand, this currently represents an obstacle to the recognition of orbital cycles in the Palaeozoic and Proterozoic; on the other hand, the possibility of using the change in ratios within the cycle hierarchy as a means of tracking the deceleration of spin represents a challenge. Evidence of orbital cyclicities in the Palaeozoic include rhythms in the Carboniferous (Heckel 1986), Devonian (Goodwin & Anderson 1985) and Cambrian (Bond & Kominz 1991).
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It thus appears that a good many sedimentary facies are likely to contain a record of orbital cyclicity. It also appears that many of the orbital signals can be obtained by means simpler than mass spectrometry. At that, we have, so far, only a beginning: the time series of cycles established so far have generally been short, and the long ones being established (Olsen 1986) are not tied to the standard relative (biostratigraphic) chronology, though circuitous ties by way of magnetostratigraphy may become possible. Thus, we have, as yet, orbital or Milankovitch timing for only a few biostratigraphic and magnetostratigraphic zones (Pliocene, Hilgen 1991; Eocene, Schwarzacher 1987b), and none as yet for stages, except possibly for the Cenomanian (Gale 1989). Furthermore, such data as we have refer only to some one section and have not yet been put to the vital test of reproduceability in other parts of the world. The task of cyclostratigraphy is a huge one, for it calls for obtaining quantities of information orders of magnitude above the level of conventional stratigraphy. This calls for new techniques: we are only now learning how to extract the periodic signals efficiently and how to work with them.
Extraction of signals Gilbert's 'pulse-beat' has been verified, but it is more complex than he imagined. It may be subtle, and it can be extracted only from rather continuously deposited sequences. Initially, the approach to rhythmicity was via spacing and thickness of beds (Gilbert 1895; Schwarzacher & Fischer 1982), but these are largely dependent on outcrop conditions and are influenced by weathering and tectonism. Study of fresh rock in cores (Fischer et al. 1991; Larson et al. 1993) hold advantages. Chemical profiles by wet chemical analysis (Herbert & Fischer 1986) or quantitative profiles of faunal or floral change (Tornaghi et al. 1989) are an excellent approach, but are time consuming. More rapid are various scanning methods, such as darkness variations plotted by image analysis (Herbert & Fischer 1986). However, the most rapid approach is that utilizing profiles already collected for other purposes, such as well logs (Fischer & Roberts 1991), and the GRAPE records (Herbert & Mayer 1991) and magnetic susceptibility logs (Tineke et al. 1991), now routinely run on deep-sea cores. Drilling bore holes and logging them by new techniques, such as elemental profiles by neutron activation (Jarrard & Arthur 1989; Larson et al. 1993) would be an efficient way of gathering data on rhythmicity. However, proper utifization of these data demands an understanding of how these
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signals were generated; a reading of how and why the depositional system changed. This calls for combining geophysical, geochemical and geobiological data. The processing of data is another field which is still undergoing development (Schwarzacher 1987a,b; Weedon 1993). In sediments deposited at comparatively uniform rates the space dimension is close to the time dimension and rhythmicities are easily recognized, in the field and in simple curves. Variable sedimentation rates commonly distort periodicities to degrees that still leave them recognizable to the eye but unresolved by spectral analysis. In visual recognition of frequency hierarchies the author has found mirror-plots of curves to be particularly useful. Once likely frequencies are recognized in this way they may furnish the basis for 'tuning' the stratigraphic dimension into something closer to time, and more amenable to spectral analysis (Ten Kate & Sprenger 1993). In sum, it is now plain that rhythmic strands related to long-period variations of the orbit runthrough Earth history. At this stage they are more plainly resolved in the Mesozoic and Cenozoic than in earlier times, but that is probably only a matter of limited understanding. We are only now beginning to recognize these rhythmicities and to learn how to extract them.
Potential The recognition that orbital cycles can be followed into the Past by the record they have left in sediments is in itself an important contribution to geological understanding, and one that should be pursued in its own right. But is also holds important implications for such matters as our understanding of evolutionary forces, and potential as a tool for other lines of biospheric inquiry. Cyclochronology
The most obvious use for these rhythmicities is in geochronology. Existing time scales (Harland et al. 1989) remain crude: the frame of timing has generally been extrapolated from the relatively few soundly dated rocks to stage boundaries, and the errors introduced in that operation are compounded by corelation of stage boundaries and further by extrapolation to zonal or other boundaries within stages. Stage boundaries in the Cenozoic are now believable to within 1 ma, but reaching farther into the past brings rapid increase in uncertainty, to 5 ma or more in the mid-Mesozoic and twice as much or more in the Palaeozoic (Harland et al. 1989). The orbital signals hold promise of bringing
chronology down to the 100 000 a level even where precessional signals are not preserved, and to the 10 000 a level where they are. The prospect now faced is how to bring such cyclochronologies into reality. This will proceed in two ways: (1) by anchored timescales; and (2) by floating timescales. In Quaternary sequences the identification of specific orbital cycles is possible via the SPECMAP isotope curve which is anchored to the Present, or directly by matching them against the astronomically-calculated insolation curve. The eccentricity component of that curve represents a nine-body problem, only approachable by approximations progressively rectified by a theoretically unlimited number of corrective steps. The details of that curve have a high confidence level for several million years, and maybe aids in correlating stratigraphic sequences of Pliocene and late Miocene age, even when these are not anchored to the Present. Hilgen (1991) has refined Neogene geochronology in this manner (Table 1). However, when projected for tens of millions of years that confidence disappears: even minor, debatable corrections alter the details of the curve, so that
Table 1. Correction of Early Pliocene chronology by use of cyclostratigraphy tied to projected astronomic curve (after Hilgen 1991) Corrected date Brunhes/Matuyama 0.78 Top Jaramillo event 0.99 Base Jaramillo event 1.07 Cobb Mountain event 1.10 Top Olduvai event 1.79 Base Olduvai event 1.95 Top Reunion event 2.14 Base Reunion event 2.15 Gauss/Matuyama 2.62-2.59 First occ. N. atlantica 2.72 Top Kaena event 3.02 Base Kaena event 3.11 Last occ. G. altispira 3.17 Top Mammoth event 3.22 Base Mammoth event 3.33 Gilbert/Gauss 3.58 First occ. G. crassaformis 3.58 Last occ. G. margaritae 3.79 Top Cochiti event 4.18 Base Cochiti event 4.29 Top Nunivak event 4.48 Base Nunivak event 4.62 Top Sidufjall event 4.80 Base Sidufjall event 4.89 Top Thvera event 4.98 Base Thvera event 5.23 First occ. G. puncticulata 4.50
Former dates 0.73-0.72 0.90-0.89 0.97-0.94 1.19 1.76-1.67 1.91-1.87 2.07-2.01 2.07-2.04 2.48-2.47 2.55 2.92-2.91 3.01-3.00 3.05-3.04 3.08-3.07 3.17-3.15 4.41-3.40 3.40 3.59-3.56 3.82-3.80 3.92-3.90 4.07--4.05 4.25-4.20 4.44-4.32 4.57-4.47 4.89--4.72 5.00--4.94 4.13-4.10
CYCLOSTRATIGRAPHY, QUO VADIS? shape-correlation of unanchored stratigraphic data to the curve is no longer feasible (Berger & Loutre, pers. comm.). It will, therefore, be necessary to work with floating chronologies, based on the recognition of the several cycles of the orbital hierarchy but not referred to a curve anchored to the Present. Such floating chronologies will serve to fix the duration of biostratigraphic or magnetostratigraphic zones within the less precise frame of radiometric time. Critics of cyclostratigraphy harp on about the incompleteness of the geological record (Algeo & Wilkinson 1988; Anders et al. 1987), but surely the average continental shelf facies is not suitable for the construction of cyclochronologies. Pelagic sequences on marine rises and ridges tend to contain major stratigraphic gaps. But the match of the Pleistocene record of pelagic and hemipelagic sediments suggests that the estimates of stratigraphic gaps in this facies have been exaggerated.
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Cyclostratigraphy will have to reckon with the existence of cryptic stratigraphic gaps, but has two levels of insurance. One lies in the cycle hierarchy: let the odd 20 ka precessional cycle be missing: the 100 and 4 0 0 k a cycles should help to keep cyclostratigraphy on track. The other lies in reproduceability: some level of reservation should be retained about cyclostratigraphies based on only one sequence, but the level of confidence should rise if distant sequences and different facies yield comparable results. The ideal case will be that in which individual cycles can be identified and correlated by some manner of fingerprints. Distinctive magnetic signatures (Napoleone & Ripepe 1989) deserve careful investigation as possible fingerprints of this type. Eventually a patchwork of such floating chronologies will close into a coherent geochronology with a resolution one or two orders of magnitude better than the existing one.
References ALGEO, Z. J. & WILKINSON,B. H. 1988. Periodicity of (Piobbico core). Journal of Sedimentary Petrology, mesoscale Phanerozoic sedimentary cycles and the 61, 1164-1172. role of Milankovitch orbital modulations. Journal of & ROBERTS,L. T. 1991. Cyclicity in the Green River Geology, 96, 313-322. Formation (Lacustrine Eocene) of Wyoming. ANDERS, M. N., KRUEGER,S. W. & SADLER, P. M. 1987. Journal of Sedimentary Petrology, 61, 1146-1154. A new look at sedimentation rates and the FOUCAULT, A., POWlCHROWSKI,L. & PRUD'HOMME, A. completeness of the stratigraphic record. Journal 1987. Le control astronomique de la sedimentation of Geology, 95, 1-14. turbiditique: exemple du Flysch a Helminthoides ANDERSON, R. Y. 1982. A long geoclimatic record from des Alpes Ligures. Academie des Sciences Comptes the Permian. Journal of Geophysical Research, Rendus, ser II, 305, 1007-1011. 87, 7285-7294. GALE, A. S. 1989. A Milankovitch scale for Cenomanian ARRHENIUS, G. 1952. Sediments from the East Pacific. time. Terra Nova, 1,420-425. Swedish Deep-sea Expedition, Reports, 5, 1-227. GILBERT, G. K. 1895. Sedimentary measurement of ARTHUR, M. A. &DEAN, W. E. 1986. Cretaceous geological time. Journal of Geology, 3, 121-125. paleoceanography. In: TUCHOLKE, B. E. & VOGT, GOLDHAMMER,R. K., DUNN,P. A. & HARDIn,L. A. 1987. P. R. (eds) Decade of North American Geology, High-frequency glacio-eustatic oscillations with Western North Atlantic Basin Synthesis Volume. Milankovitch characteristics recorded in northern Geological Society of America, 617--630. Italy. American Journal of Science, 287, 853-892. BERGER,A., LOUTRE,M. E & LASKAR,J. 1992. Influence GOODWIN, P. W. & ANDERSON, E. J. 1985. Punctuated of the changing lunar orbit on the astronomical aggradational cycles: a general hypothesis of frequencies of pre-Quatemary insolation patterns. episodic stratigraphic accumulation. Journal of Paleoceanography, 4, 555-564. Geology, 93, 515-533. BOARDMAN, D. R. II & HECKEL, P. H. 1989. GlacialHARLAND, W. B., ARMSTRONG,R. L., Cox, A. V., CRAIG, eustatic sea-level curve for early Late L. E., SMITH, A. G. & SMITH, D. G. 1989. A Pennsylvanian sequence in north-central Texas Geological Timescale 1989. Cambridge University and biostratigraphic correlation with curve for Press, Cambridge. mid-continent North America. Geology, 17, HECKEL, P. H. 1986 Sea-level curve for Pennsylvanian 802-805. eustatic marine transgressive-regressive depoBOND, G. C. & KOMINZ,M. A. 1991. Some comments on sitional cycles along midcontinent outcrop belt, the problem of using vertical facies changes to infer North America. Geology, 14, 330-334. accomodation and eustatic sea-level histories with HERBERT, T. D. & FISCHER, A. G. 1986. Milankovitch examples from Utah and the southern Canadian climatic origin of Mid-Cretaceous black shale Rockies. Kansas Geological Survey Bulletin, 233, rhythms in central Italy. Nature, 321,739-743. 273-289. --& MAYER, L. A. 1991. Long climatic time series CROLL, J. 1875. Climate and Time in their Geological from sediment physical property measurements. Relations. Appleton, New York. Journal of Sedimentary Petrology, 61, 1089-1108. FISCHER,m. G., HERBERT,T. D., NAPOLEONE,G., PREMOLI HILGEN, F. J. 1991. Extension of the astronomically SILVA,I. & RIPEPE,M. 1991. Albian pelagic rhythms calculated (polarity) time scale to the
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Miocene/Pliocene boundary. Earth and Planetary Science Letters, 107, 349-368. IMBRIE, J. 1985. A theoretical framework for the Pleistocene ice ages. Geological Society of London Journal, 142, 417-432. JARRARD, R. D. & ARTHUR, M. A. 1989. Milankovitch paleoceanographic cycles in geophysical logs for ODP Leg 105, Labrador Sea and Baffin Bay. In: SRIVASTAVA, S. P., ARTHUR, M. A., CLEMENT, B. et al. (eds) Proceedings of the Ocean Drilling Project, Scientific Results, 105, 757-772. JENKYNS, a. C. 1980. Cretaceous anoxic events: from continents to oceans. Geological Society of London Journal, 137, 171-188. LARSON, R. L., FISCHER, A. G., ERBA, E. & PREMOLI SILVA, I. (eds) 1993. APTICORE-ALBICORE, A Workshop on Global Events and Rhythms of the Mid-Cretaceous. Joint Oceanographic Institutions, Inc., 1755 Massachusetts Ave. NW, Suite 800, Washington DC 20036, USA. MILANKOVITCH,M. 1941. Kanon der Erdbestrahlung und seine Anwendung auf das Eiszeitenproblem. Serbian Academy of Science, Belgrade, 133. NAPOLEONE, G. & RIPEPE, M. 1989. Cyclic geomagnetic changes in mid-Cretaceous rhythmites, Italy. Terra Nova, 1, 437-442. OLSEN, P. 1986. A 40-million-year lake record of early mesozoic orbital forcing. Science, 234, 842-848. PRATT, L. M., KAUFFMAN,E. G. & ZELT, E (eds) 1985. Fine-grained deposits and biofacies of the Cretaceous interior seaway: evidence of cyclic sedimentation processes. Society of Economic Paleontologists and Mineralogists, Field Trip Guide Book, 4. ROOF, S. R., MULLINS,H. Z., GARTNER,S., HUANG,T. C., JOYCE, E., PRUTZMAN, J. &TJALSMA, L. 1991. Climatic forcing of cyclic carbonate sedimentation during the last 5.4 million years along the West Florida continental margin. Journal of Sedimentary Petrology, 61, 1070-1088. SAKAMOTO, T. 1992. Sedimentary rhythm of Nakayama Formation (middle Miocene to early Pliocene) in the Sado Island, Japan Sea. Abstracts, 19th International Geological Congress, 2, 288. SAVRDA,C. E. & BoTrJER, D. J. 1989. Trace-fossil model for reconstructing oxygenation histories of ancient marine bottom waters: application to Upper Cretaceous Niobrara Formation, Colorado. Palaeogeography, Palaeoecology, Palaeoclimatology, 74, 49-74. SCHWARZACHER,W. 1987a. The analysis and interpretation of stratification cycles. Paleoceanography, 2, 79-95.
FISCHER 1987b. Astronomeivcally controlled cycles in the lower Tertiary of Gubbio (Italy). Earth and Planetary Science Letters, 84, 322-26. --& FISCHER, A. G. 1982. Limestone-shale bedding and perturbations in the Earth's orbit. In: EINSELE, G. & SEILACHER, A. (eds) Cyclic and Event Stratification. Springer Verlag, Berlin, 72-95. & HAAS, J. 1986. Comparative statistical analysis of some Hungarian and Austrian Upper Triassic peritidal carbonate sequences. Acta Geologica Hungarica, 29, 175-196. TEN KATE, W. D. & SPREN~ER, A. 1993. Orbital cyclicities above and below the CretaceousTertiary boundary at Zumaya (N Spain), Agost and Relleu (SE Spain). Sedimentary Geology, 87, 69-101. TORNAGHI, M.-E., PREMOLISILVA,I. & RIPEPE, M. 1989. Lithostratigraphy and planktonic foraminiferal biostratigraphy of the Aptian-Albian "Scisti a Fucoidi", Piobbico core, Marche, Italy: background for cyclostratigraphy. Revista Italiana de Paleontologia et Stratigrafia, 95, 223-264. TINEKE, N. E, KROON, D., TEN KATE, W. D. & SPRENGER, A. 1991. Late Pleistocene periodicities of oxygen isotope ratios, calcium carbonate contents, and magnetic susceptibilities of western Arabian Sea margin Hole 728A(ODPLeg 117). In: PRELL,W. L., NIITSUMA,N. et al. Ocean Drilling Program Science Research Proceedings, 117, 309-320. VAN TASSELL, J. 1987. Upper Devonian Catskill delta marine cyclic sedimentation: Brallier, Scherr and Foreknobs Formations of Virginia and West Virginia. Geological Society of America Bulletin, 99, 414-126. WEEDON, G. P. 1993. The recognition and stratigraphic implications of orbital-forcing of climate and sedimentary cycles. Sedimentology Review, 1, 31-48. --&SHIMMIELD, G. B. 1991. Late Pleistocene upwelling and productivity variations in the northwest Indian Ocean deduced from spectral analysis of geochemical data from Sites 722 NS 744. Proceedings of the Ocean Drilling Program, Scientific Results, 117, 431-443. WELTJE, G. &DE BOER, E L. 1993. Astronomically induced paleoclimatic oscillations reflected in Pliocene turbidite deposits on Corfu (Greece): implications for the interpretation of higher-order cyclicity in ancient turbidite system. Geology, 21,307-310.
Index
Aalensis Subzone, Yorkshire, 68
absolute particle abundances, calculation of, 78-83 Acanthoceras jukesbrownei Zone, Europe, 181, 186-90, 191 Acanthoceras rhotomagense Zone, Europe, 181, 186-9, 193 Achanarras Limestone, 7 acritarchs, Kimmeridge Clay, 89, 92 Albian deposits, 134, 137, 183 Albian Stage, Europe, 183 Albian-Cenomanian boundary, Europe, 179, 193 allocyclicity, 133 Alpes-Maritimes, 183 American Upper Cretaceous, 133 ammonite zonal geochronology, Yorkshire, 69 Andean foreland basin, fluvial sediments, 20 Angles, succession, 135, 136-7, 138, 144-50, 154-6, 166 Anglo-Paris Basin, 177-91,193-4 annual band, 2 annual growth bands, 2, 7, 8 annual and lesser orbital cycles, 2, 7 Appeninnes, 137 Aptian Urgonian cycles, 172 Archiane valley, 166, 170-2 Ardrche platform, 165, 166 Ardrchoise platform, 172 Arnsbergian marine transgressions, Scotland, 65 Asymmetricus Zone, S. France, 43, 44, 47 Atlantic Ocean sediments, 88, 134, 136, 174, 175, 199 Atlas Mountains, 12 atmospheric gases, cyclical movements, 4 Autissiodorensis Zone, Kimmeridge Bay, 76, 99 autumnal equinox, 7 Bahamas, 174 bankfull flood discharge, 27-8, 29 Barremian carbonates sedimentation rates, 135, 136, 137, 138 Vocontian Trough, 136, 165, 172-5 basin-slope correlations, 165-6 high frequency cycles, 174-5 medium frequency cycles, 172-4 slope-outer platform correlations, 166-72 Barremian-Bedoulian boundary, Angles, 136 basin filling, model of, 29-31 basin parameters, estimates of, 26-31 Bell Shale corals, lunar month periodicity, 7 Berriasian Carbonate platforms, Vocontian Trough, 144 sedimentation rates, Vocontian Trough, 136, 137, 138 Berriasian-Barremian bedding cycles, 195 Berriasian-Barremian interval, Angles, 138 biological variation, 97 bioturbation, 96, 140 bisaccate pollen, Kimmeridge clay, 78, 81, 86, 87, 90, 92, 93, 102, 105
bivalves, 181 banding in shells, 2, 5, 7 black debris, Kimmeridge Clay, 78, 86, 88, 90, 92, 93, 95, 98, 103, 104, 106 Black Sea sediments, 88 Boreal Basin/Tethys sea-way, 159, 162 Bottacione Gorge, 179, 192 Bou Tchrafine, 43, 45, 46-7, 48 brown wood, Kimmeridge Clay, 86, 88, 90, 94, 98 bundles of cycles, 11, 12, 136 Caha Mountain Formation sediments, 22, 26, 30, 31 Calcicalathina oblongata Biozone, France, 146 calcium carbonate carbonate cycles, Vocontian Trough, see under Valanginian content, 143-4, 146-57 Kimmeridge Clay, 86, 88, 95, 96, 98, 102, 103, 104, 110 see also Barremian carbonates calcium carbonate curve, 199 calcrete formation, climatic effects on, 25 calendar band, 2 Calycoceras querangeri Zone, Europe, 178, 181, 189-91 Cambrian days in lunar month, 7 sediments, 200, 201 carbon dioxide feedback, 116, 118, 130 carbon isotope stratigraphy, 181-2 carbonate cycles, Vocontian Trough, see under Valanginian carbonate percentages and temperature control, 12 carbonates, see Barremian carbonates; calcium carbonate Carboniferous, 5, 7, 51,201 climate modelling, 117 climates, 200 continental ice sheets, 54 Munster Basin, 21 periodicities, 65 Scotland, 54, 55, 65 sea level changes, 200 Cast Bed, 186, 187, 188 Castile Formation, 201 Castlehaven Formation, 22 celestial equator, 10 Cenomanian, 201 correlation in western Europe, 177-9, 182-95 stratigraphical framework, 179-82 Cenomanian Chalk, 179 Cenomanian-Turonian ~13C excursion, 191 Cenomanian/Turonian transition, 134 Cenozoic stage boundaries, 202 Central Atlantic, 134 Central North Sea Dome, 72, 73 Central Pacific, 143 Cerastoderma edule, 5, 6, 8 Chalk Marl, 179, 185 Chandler Wobble, 8 205
206
INDEX
channel depth estimation, 28, 29 charcoal, Kimmeridge Clay, 96 chemical variation as cause of cyclicity, 96, 97 Chloritic Sandstone Formation (CSF), 22, 26 Chondrites Event, 191 chorate dinocysts, Kimmeridge Clay, 78, 86, 87, 90, 92, 94, 98 circadian biorhythmicity, 6 circadian change of sea level, 4 clay/quartz, Kimmeridge Clay, 86, 88, 95, 98, 100, 102, 103, 104, 105 Cleveland Basin, 177, 178, 181, 182, 184, 188, 191, 193 cliff of Chames, 166, 168, 169 climate system models, 116-19, 120-3 climatic belt shifts, 109, 111, 158 climatic changes causes, 9, 10, 25, 26, 54, 100, 101,116 effects of, 1, 62 sedimentation and, 20, 24, 25, 27, 32, 97, 10l, 138 climatic modelling, Late Jurassic, 116-25 Clinocardium nuttalli, tidal patterns, 5, 6 Col du Puech 'de la Suque, GSSP for Givetian/Frasnian boundary, 43 compaction correction factors, 57, 58 condensation of succession, 177 conodont zonal durations within the Givetian, 47, 48 constraints in use of sedimentary cycles, 133-40 corals, lunar month periodicity, 7 Corfu, pro-delta fan, 200 cosmic year, 14 Cretaceous Atlantic anoxic events, 88 black shales, 88 carbonate platforms, 200 climate system, 195 greenhouse episodes, 201 orbital frequencies, 148, 149 sea-surface temperatures, 130 sedimentary cycles, 117, 133-4, 135, 137, 138 analytical methods, 68 Vocontian, 143, 146, 157 Cretaceous Greenhorn Formation, Colorado, 88-9 Cretaceous/Tertiary boundary, 14 Cretarhabdus crenulatus Biozone, France, 146 Culver Cliff, 179 Cunningtoniceras inerme Zone, Europe, 181,185, 186-9, 193 cuticle, Kimmeridge Clay, 78, 86, 87, 89, 90, 93, 94, 98, 106, 108 daily band, 2 daily cycle, 2 decadal band, 2 deceleration of Earth spin, 201 declination and diurnal inequality of tides, 3 decompaction factors, 27 deep water fans, 200 deglaciation, 116 degraded palynomorphs, Kimmeridge Clay, 86, 88, 94, 98 Delaware Basin, 201 Delaware river estuary, 93 dendrochronology, 2, 7 depositional basins, estimation of discharge, 29
depositional cycles, Early Toarcian, see Early Toarcian depositional cycles, Yorkshire Devonian, 14 duration, 37, 38 sediments, 7, 9, 39, 42-3, 44, 45, 201 see also Munster Basin Devonian sequences Devonian year, days in, 2 diffusivity assessment, Munster Basin, 25-6, 32 models for basin-fills, 20, 25-32 dinocysts, Kimrneridge Clay, 86, 87, 89, 90, 92, 93, 98, 102, 106, 107, 111 Dispansum Subzone, Yorkshire, 68 Disparilis Zone, Cornwall, 43, 47 diurnal inequality of tides, 3, 4 diurnal tidal system, 4 Djebel Oust Basin, 134 Dolomites, 137, 169 drainage basin parameters, estimation of, 27-9 dune height, 28 Early Toarcian depositional cycles, Yorkshire, 67-9 Milankovitch band cyclicities, 67, 72, 73 tectonic origin, 72-3 time series analysis, 67, 69-72 Earth spin, deceleration, 201 Earth year, 5-6 Earth-Moon system, 8, 12 tidal cycle and, 3 Earth-Moon-Sun system insolation and, 9, 10 tidal effects, 6 variations in oscillation, 11 Earth-Sun distance, solar energy related to, 1 East African Rift, 72 East Greenland Basin, sedimentation rates, 23 eccentricity, 1, 9, 10, 11, 12 periodicities, Late Carboniferous, 65 Eifelian, Mech Irdane, 46 Eifelian/Givetian boundary, GSSP for, 43 El Nifio (ENSO) effect, 8 Elatina Formation, 8 Elatina laminates, 7 Empire, Oregon, sediments, 6 energy balance models (EBM), 116, 117 English chalk microrhythms, 178 Eocene, 201 precipitation-evaporation balance, 201 sediments, 7, 8, 200 equator to pole temperature gradient, 125 equilibrium tide, 3 equilibrium time constant (Teq), 20, 30--2 equinoxes, 2, 7, 10, 116 equinoxial tides, 7 erosion patterns, 1, 9 Exodus Zone, Yorkshire, 106 Falciferum maximum flooding event, 68 Falsiovalis Zone, Cornwall, 44 Famennian sediments, France, 43 feedback processes, 116, 117, 130 floating chronologies, 203 Fontcalent, sediments, 137, 138 foraminiferal, Kimmeridge Clay, 89, 96
INDEX Forest Marble, England, 4 fossil group extinctions, 13-14 fossils, sea level changes recorded by, 4 Francis Creek sediments, 5 Frasnian duration, Munster Basin, 23 sediments, France, 43, 47 galactic band, 2 general circulation model (GCM), 115, 116-19, 120-3, 130 'Gilbert', the, 139 Givetian duration, Munster Basin, 23 timescale, establishment of, 37-8 Bou Tchrafine, 43, 46-7, 48 by radiometric data, 37 estimations of conodont zonal durations, 47 Marble Cliffs, Cornwall, 43-5, 48 Pic de Bissous, 38-43, 46, 47, 48 using couplets, 45, 46, 47 Givetian/Frasnian boundary, 43 glacial-interglacial cycle, Mesozoic, 174 glacial/interglacial climate change model, 116 glaciation periods, 14, 116, 117, 133 glacio-eustatic cycles, 54, 174 Limestone Coal Formation, Scotland, 51, 54, 64, 65 Vocontian, 175 Glauconitic Marl, 183 Global Stratotype Section and Point (GSSP), Givetian boundaries, 37 Goddard Institute for Space Studies (GISS) model, 117 Great Bahama Bank sediments, 93 Green River Formation, 7, 8 greenhouse episodes, 201 Grey Chalk, Anglo-Paris Basin, 185, 191 Grey Pit, 186 growth banding, 2, 7, 8 growth rate formula for nodule formation, 68 Gubbio, 136, 178, 179, 181, 192 Gulf of California, polynomorph distribution, 93 Gulf of Mexico, sediments, 134 Gun Point formation (GPF), 22, 23, 24, 25, 26, 30, 31 tectonic cyclicity, 32 Hale cycle, 8 Hartley coal, 53, 57, 59-65 Hauterivian deposits, 134, 137, 138, 165, 166 Hemiansatus Zone, France, 47 Hemipelagic Zone, France, 165 Hermanni Zone (Hermanni-Cristatus Zone), Cornwall, 43 high frequency cycles, constraints in use of, 133-40 Holocene fan deposits, 27 model for interglacial period, 174 Hornelen Basin, cyclic sediments, 20 Hugii Tethyan Zone, France, 165 Hugii Zone, France, 166, 167 hydrological balance, 117 hydrological cycle of the monsoon, 117 hydrological cycles, model for, 119, 122, 127 ice ages, 1, 9, 174
207
ice albedo feedback, 116 ice flux and sea level, 200 ice sheets, 116, 199 Jurassic model, 118, 125, 130 ice-melting, 13 Index Limestone, 53, 54, 55, 58, 59, 60, 61, 62, 65 inoceramid bivalves, 181 insolation changes, 1, 9, 10 climatic effects, 25, 54, 101,109, 158 model predictions, 122, 124 insolation forcing and carbonate productivity, 158 interannual band, 2 Ireton Shale, 7 Iridium Clay, 14 Jet Rock Member, 68 Jura, 172 Jurassic, 8, 12 climate variations, 115, 116, 125-30 modelling of Milankovitch cycles, 116-25 sediments, 4, 7, 67, 72, 135, 137-8, 172, 200 kerogen sedmentation, Kimmeridge Bay, 89, 90, 91 Kilsyth Trough, 54 see also Limestone Coal Formation, Scotland Kimmeridge Clay Formation, Dorset, 75-6, 77, 106 Kimmeridge Clay, Yorkshire, 89, 91, 96, 106 Kimmeridge Oil Shale, 89 Kimmeridgian, 67 climate modelling, 119-25 palynofacies cycles, see palynofacies of Kimmeridgian cycles Kincardine Basin, 54 see also Limestone Coal Formation, Scotland Knightswood Gas Coal, 55, 58, 59, 60, 61, 62, 64 La Charce (Dr6me), 158 Lake Playa, 201 Late Albian Upper Greensand, 183 Late Jurassic climate variations, 115, 116, 125-30 modelling of Milankovitch cycles, 116-25 Lias, 67, 134 mudstones, Yorkshire, 106, 110 Limestone Coal Formation,Scotland, 51, 62-5 data set and sedimentology, 52-4 glacio-eustatic cycles, 51, 54, 64, 65 spectral analysis, 54-62 Livello Bonarelli, 179, 192 Lockatong Formation, 7 Lombardian Basin, 134 Lower Saxony Basin, 177, 178, 181,183, 184, 188, 191 lunar banding, 7 lunar day, 3, 5, 12 lunar month, 6-7, 12 lunar nodal cycle, 8, 9 lunar perigee, 8 lunar tides, 3, 9 lycopodium spiking, 79-83 Lyme Regis, 67 magnitude-frequency of aggradational events, 23 Majolica Formation, 136 Mantelliceras dixoni Zone, Europe, 181, 182, 183-5, 193
208
INDEX
Mantelliceras inflatum Zone, 183 Mantelliceras mantelli Zone, Europe, 180-1,183, 184 Mantelliceras saxbii Subzone, Europe, 181
Marble Cliffs, Cornwall, 43-5, 48 Marbri6re Nord, 38--42, 43, 45, 46, 47, 48 marine bands, occurrence of, 55 marine current intensity, 101,106, 111 marine palynomorphs, Kimmeridge Clay, 78, 86-7, 90, 92, 98, 102, 103, 104, 105-6, 110 marl-limestone couplets, 133, 134-40 'maximum seasonal forcing' model, 125, 128, 129 mean annual basin discharge, estimation of, 27-8, 29 Mech Irdane Givetian base, 46 GSSP for Eifelian/Givetian boundary, 43 Mediterranean Pliocene, 143 Melbourne Rock Beds, 193, 194 Mesozoic, 1, 7, 116 cycle recognition, 68 depositional sequences, Europe, 165, 174, 193 stage boundaries, 202 Metoicoceras geslinianum Zone, Europe, 181,191-3 Midland Valley, see Limestone Coal Formation, Scotland Milankovitch band, 1, 2, 7, 9, 12, 13 Milankovitch climate variation modelling, see Late Jurassic climate variations Milky Way, 14 periodic extinctions related to, 13, 14 millenial band, 2 'minimum seasonal forcing' model, 119, 124, 126, 127, 130 Miocene, 202 Sicilian anhydrites, 9 monsoon circulation, 109 Jurassic model, 122, 125 monsoons, 117, 122, 125 Mont Ventoux Chain, 135 Montagne Noire, 38, 39 see also Pic de Bissous Montagnette, the, 166, 169-70 monthly band, 2 monthly cycle, 2 Mouniers Marls, 166 Munster Basin Devonian sequences, 19-20, 32 basin equilibrium and response, 29-32 cyclicity, 23-5 location and geological setting, 20-2 palaeohydrology, 25-9 subsidence within, 21, 22 time-sediment accumulation rates, 23 MUnsterland basin, 177, 178, 181,184, 186-8, 191,193 Namurian deposits, Scotland, see Limestone Coal Formation, Scotland nautiloids, lunar banding, 7 NCAR model, 117 neap tides, 6 Neocardiocerasjuddii Zone, Europe, 181,191-3 Neogene, 200 ODP cores, 193 SPECMAP isotope curve, 202 Neostlingoceras carcitanense Subzone, Europe, 181, 183 Nettleton Stone, 191
Newark Basin, 2 Niveau Thomel, 193 nodule formation, growth rate formula, 68 North American cyclothem transgressions, 200 North Atlantic Water Passage model, 89 North Sea sediments, 5 thermal doming, 72, 73 obliquity, 1, 9, 10, 11,100 periodicities, 7 changes through time, 12, 13, 76 values, 46, 65, 76, 139 obliquity signal, recognition of, 134 ocean circulation, 10, 111 ocean core evidence for temperature changes, 2 ocean temperatures, Jurassic model, 119 Old Red Sandstone Basin, see Munster Basin Devonian sequences Old Red Sandstone (ORS), 19, 22 Ordovician, 14 organic matter, burial of, 200 organic sediment deposition, Kimmeridge Clay, 88 Orinoco Delta sediments, 89, 93 oxygen isotope records, 12, 199 Pacific Ocean, 96, 143, 199 Palaeozoic, 201 carbonate platforms, 200 sediments, 201 stage boundaries, 202 Paltum to Fallacioscum Zones, Yorkshire, 72 Paltum to Tenuicostatum Zones, Yorkshire, 69, 70-1 palynofacies of Kimmeridgian cycles analytical methods, 78-86 causes of cyclicity, 96--7 classification, 76-8 cycle durations, 76, 97, 99-100 depositional models, 88-91 orbital interpretation, 98, 100-11 origin or particles, 94-5 palaeoenvironmental interpretations, 91-6 particle abundances, 86-8 Pendleian deposits, Scotland, 52 marine transgressions, Scotland, 65 Pennsylvanian sedimentary cycles, 13 perihelion, 2, 10, 117 periodic extinctions, 13-14 Permian, 14, 201 Castile Formation, 201 sea level oscillations, 200 Phanerozoic climatic cycles, 54 periodicities, l, 139 phytoplankton blooms, 88 Pic de Bissous, 39, 40, 45 Givetian timescale, 37 microrhythmicity in the Givetian, 41, 42, 48 sediments, 45 Piobbico, 137 plankton carbonate production, 200 productivity, 88, 106, 108, 109, 110, 111
INDEX Pleistocene, 8, 14, 68, 116, 122, 143 carbonate production, 200 climatic cycles, 54 glacial fluctuations, 133 marine events, 199 sea level oscillations, 200 sediments, 203 Plenus Bank, 191 Plenus Marls, 177, 191,193, 194 Pliensbachian, Yorkshire, 72 Plio-Pleistocene, 134, 143 Pliocene pro-delta fan, 200 sediments, 54, 116, 127, 143, 200, 202 pole-equator insolation gradient, 10 pollen grain sedimentation, 92-3 pollen, Kimeridge Clay, 78, 81, 86, 87, 89, 90, 92, 93, 102, 105 Polygnathus hemiansatus, 46, 47 Pont-de-Laval, 166 Porcupine Basin volcanic centres, 72 power spectral analysis, 54, 56--62, 86, 91 prasinophytes, Kimeridge Clay, 86, 92, 94, 98, 102 pre-Cambrian, 6, 7, 8 pre-Cretaceous, 1 pre-Pleistocene, 8 pre-Pliocene, 54, 116 pre-Quaternary sediments, 143 precession, 1, 2, 9, 10, 11, 12 periodicities, 7 changes through time, 12, 13, 76 values, 45, 46, 65, 76, 139 precession frequency calculations, 46 precession signal, recognition, 134 precipitation, 25 Jurassic model, 121,125, 126, 127 precipitation minus evaporation budget, 117 precipitation-evaporation balance, 201 preservational conditions, 106 'Primus Event', 188 Proterozoic, 7, 8 carbonate platforms, 200 cycle recognition, 201 proximate dinocysts, Kimeridge Clay, 78, 86, 87, 92 Pueblo, 181 Purple Sandstone Formation, 22, 30 Quaternary fluvial systems, Munster Basin, 25 ice sheets, 199 sediments, 88, 116, 199, 202 Rattray volcanic centre, 72, 73 Recent, 6, 8 recurrence interval (r) of aggradational events, 23-4 redox cyclicity in sdiments, 200 Reynella Siltstone, 8 rhythmicity irregular or fine scale, 177 loss of, 177 rhythmites, 2, 4 Rochers de Combeau, 166, 169, 170, 171 rock relief, tidal effects on, 3 Rotalipora brotzeni Zone, Europe, 181
209
Rotalipora cushmani Zone, Europe, 181 Rotalipora reicheli Zone, Europe, 181 Roundhole Point, Cornwall, 43 Rouvillei Zone, Morocco/France correlation, 46, .47 rngose corals, banding in, 2
salinity indicators, 92 Saynoceras verrucosum Zone, Vocontian Trough, 146 Scaglia Bianca, 136, 178, 179, 192 sea level changes, 9, 62, 175, 193, 200 causes, 200 effects of, 1, 138-9, 194 Jurassic model, 125, 130 record of, 4-5 sediment accumulation rates in non-marine basins, 23 sediment yield in drainage basins, 27, 28 sedimentary cycles, constraints, of, 13340 sedimentary responses, weak, 130 sedimentary rhythmicity, 1, 11-13 sedimentation, Milankovitch band cycles and, 9-10 semidiurnal inequality, 5 semidiurnal tidal cycles, 3, 4-5, 6 Sharpeiceras schlueteri Subzone, Europe, 181,183 shells, growth patterns, 2, 5, 6, 7, 181 Sherkin Sandstone Formation, 22 shoreline proximity indicators, 92 Sicilian anhydrites, 9 Siwaliks molasse, fluvial sediments, 20 Slehany Formation, 22, 26 snow cover, Jurassic model, 125, 127 solar cycles, 2 solar day, 3, 5-6, 7 solar energy, 1 solar frequency band, 2, 8 solar radiation changes and climate, 116 Jurassic model, 118-19 solar tides, 3 solar year, 8 Southeastern Basin, France, 134, 135 Southerham, Chalk Marl, 179 Southern England, position during Kimmeridgian, 106 Southwest Ireland, 21 see also Munster Basin Devonian sequences SPECMAP isotope curve, 199, 202 spectral analysis of time series, 54-62 spore sedimentation, 92-3 spores, Kimmeridge Clay, 78, 81, 86, 89, 90, 92, 93, 98, 102, 105 spring tides, 6 Stoliczkaia dispar Zone, 183 Subalpine Basin, 144 sulphate reduction index (SRI), 91,107 Sun-Earth-Moon system, see Earth-Moon-Sun system sunspot cycle, 8 surface soil wetness, Jurassic model, 123, 125, 129 surface temperatures, Jurassic model, 120, 124 surface water productivity changes, 127 T. callidiscus ammonite Zone, Vocontian Trough, 152 Tafilalt condensed shelf developments, 47 Tafilalt Platform, 43 tasmanitids, Kimmeridge Clay, 94 tectonic band, 2
210
INDEX
tectonic cyclicity Gun Point Formation, 32 Limestone Coal Formation, 62 Marble Cliffs Cornwall, 45 Toarcian of Yorkshire, 72, 73 tectonism, sedimentation and, 20 temperature changes Jurassic model, 125, 128 ocean core evidence for, 2 terebratum Zone, Morocco/France correlation, 47 terrestrial debris, Kimmeridge Clay, 78, 86, 93, 94, 95, 102, 107, 108 terrestrial palynomorphs, Kimmeridge Clay, 78, 86, 87, 89, 90, 93, 94, 96, 102, 103, 104, 105, 106 Tertiary, 1, 7 Tethyan ammonite Zone, France, 166 Tethyan successions, 134, 178 Tethys, 159, 162 thermal doming, North Sea, 72, 73 thickness-time conversion, 68-9 tidal cycles, 2, 3, 8 days in lunar month and Earth year and, 5 effects of, 3-5, 7, 9 tidal nodes, 3 time off perihelion, the, 10 time series analysis, methods of, 55-8 Toarcian depositional cycles, see Early Toarcian depositional cycles, Yorkshire Todilto Formation, 7, 8 Total Oxygen Content, Kimmeridge Clay, 86, 88, 91, 95, 96, 103, 104, 105, 107, 108, 110 Triassic, 7, 117, 169 carbonate platforms, 200, 201 temperature oscillations, 117 Triassic-Jurassic, precipitation-evaporation balance, 201 Turonian, 181 Turrilites acutus Subzone, Europe, 181 Turrilites costatus Subzone, Europe, 177, 181 UKUniversities Global Atmospheric Modelling Programme (UGAMP)model, 115, 117, 118-19 Umbria-Marche Basin, 178, 179, 191,193 uncompacted volume of sediment, 27
Upper Jurassic-Lower Cretaceous succession, Fontcalent, 135, 137, 138 Urgonian formations, France, 172, 173, 174 Valanginian carbonate cycles, Vocontian Trough, 143-4, 158-62 analysis, 146, 148-58, 162 carbonate productivity, 158 durations, 156, 157 lithostratigraphy and biostratigraphy, 144-6, 147 origin of variations, 146 sediments, Angles, 135, 136, 137, 138 stage duration, 148 Valanginian-Hauterivian interval, France, 143 Valentia Slate Formation, 22 Varcus zone, France, 43 Varcus/Hermanni Zone, France, 47 Variabilis to Striatulum Zones, Yorkshire, 69-70, 71 Variabilis Zone, duration, Yorkshire, 68 vegetation patterns, 1,109, 111 Vendian, 14 Vergons, 183 vernal equinox, 7 Vocontian Basin, 177, 178, 183, 184, 189, 191,193, 195 Vocontian sediments, 135 Vocontian Trough carbonate cycles, see under Valanginian platform and pelagic carbonates relationship, 165-75 volume of compacted sediment, 27 Washing Ledge Shales, 76 Washing Ledge Stone Band, 76, 99 water discharge, power of, 28, 29 water temperature indicators, 92 water turbulence indicators, 92 water vapour in atmosphere, Jurassic model, 119 Watnoceras archaeocretacea, 181 weathering processes, alteration of, 9 Wessex Basin, 72 Western Interior Basin, USA, 133, 134, 181, 193 Western Interior Seaway, 200 Wood Bay Group, Spitsbergen, 20 Wurmian ice age, 174
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Tbe the -mm ,rid ef the Earth-Sun ~ b i t a l patterns produce gravitational and temperature effects that alter the physlaml environment on the Earth~ surfa(e. These give an interpenetrating effect of forcing cycles ranging from twice-dally tides, day-night alternations, various tidal patterns and the annual solar pattern. All of these have been used palaeontologically to give precision to short-term age determination in the past. It is cycles of the Milankovitch band that are showing promise of enabling new practical times(ales to be established for parts of the geological record. These depend on changes in the Ea,i;~ Sun distance and changes in the tilt of the Earth's axis with respect to the Earth's orbit round the Sun. There is increasing evidence that small-scale sedimentary rhythmic couplets may represent the effect of such changes. The disentangling of the interpenetrating cycles to produce an orbital forcing timescale is an exciting problem and challenge for palaeobiology and sedimentology. These should enable numerical dates to be given to bio- and chronostratigraphical timescales and eventually enable many Earth pro(essesto be analysed in real time. The papers in this volume represent major new developments in sedimentological, palaeontological, geochemical and stratigraphiail research being undertaken in this field. • review of Phanerozok: examples of ~lostratigraphy • up-to-date reference source • 204 pages • 117 illustrations • 12 chapters plus introductory review • index
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