Glauconitic peloids and chamositic ooids - favorable factors, constraints, and problems

Earth-Science Reoiews, 20 (1984) 211-243

211

Elsevier Science Publishers B.V., Amsterdam--Printed in The Netherlands

Glauconitic Peloids and Chamositic Ooids Favorable Factors, Constraints, and Problems F.B. Van Houten and M.E. Purucker

ABSTRACT Van Houten, F.B. and Purucker, M.E., 1984. Glauconitic peloids and chamositic ooids - favorable factors, constraints, and problems. Earth-Sci. Rev., 20: 211-243. Glauconitic minerals constitute a family ranging from green smectite to a 10 A dioctahedral mica (glauconite). Chamositic minerals include a 7A trioctahedral serpentine (berthierine) and a 14 A trioctahedral chlorite (chamosite). These green iron-rich, neoformed or transformed clay minerals are most commonly concentrated in sand-size granules. Recent berthierine and Recent and ancient glauconitic minerals occur mainly in structureless peloids, most of which are believed to have been fecal pellets. In contrast, most of the ancient chamositic minerals are in multi-coated ooids generally assumed to have been made by gentle rolling on the sea floor. Glauconitic and chamositic granules accumulated most commonly in marine shelf environments during episodes of reduced influx of sediment. In modern deposits chamositic peloids predominate on the inner shelf, whereas glauconitic peloids are most abundant on the middle and outer shelf. In general, ancient glauconitic and chamositic deposits had a rather similar environmental distribution; in detail, however, they reflect more varied and overlapping marine habitats. Glauconitic greensands and chamositic ironstones commonly occur above a coarsening- or shoaling-upward facies sequence. Many of them are cross-bedded and burrowed, and some are interbedded with a ferruginized or phosphatized hardground. Although differing in detail, their temporal distributions throughout Phanerozoic time were rather similar. Both attained a maximum when cratonic blocks were widely dispersed and sea level was high in Early Paleozoic and Late Mesozoic time. In addition, recurring development of chamositic ooids commonly coincided with repeated regional transgressions. This review of current information and differing interpretations leads to significant questions that are essential subjects for future research. Moreover, some of these relate to unsolved problems of phosphorite genesis. INTRODUCTION

Each advance in knowledge about the habitats of modern sediments revives a basic theme in sedimentology - - to what extent are they keys to the past? An instructive test of the adequacy of a Holocene model to explain 0012-8252/84/09.90

© 1984 Elsevier Science Publishers B.V.

212 ancient analogues is afforded by the geologic record of glauconitic peloids (greensands or glauconies) and chamositic ooids (ironstones) in sequences of detrital and carbonate sedimentary rocks. In this essay we focus on the association of sedimentary facies that favored accumulation of chamositic ooids and glauconitic peloids, and on regional and global factors favoring the development of Phanerozoic ironstones and greensands. We have reviewed pertinent information and ideas about the mineralogy of chamositic and glauconitic clay minerals and the morphology of granules composed of them, but have not pursued the detailed geochemical relations involved. The presence of glauconitic and chamositic granules is evidence of ancient

Van Houten received his doctorate from Princeton University in 1941, and returned to that faculty in 1946. As a sedimentologist concerned mainly with the relation between sedimentation and tectonics he has studied Late Paleozoic and Early Mesozoic deposits in rift basins in eastern North America and northwestern Africa, Cretaceous rocks in intracratonic basins in northeastern Africa, and Cenozoic deposits in orogenic basins in the Rocky Mountains, the Colombian Andes and the Tunisian Atlas. Analytical studies of these sequences have focused on the origin of red beds, zeolites, and oolitic ironstones. Present address: Department of Geological and Geophysical Sciences, Princeton University, Princeton, New Jersey 08544, U.S.A.

Purucker was educated at the California Institute of Technology and at Princeton University where he received his Ph.D. in geology in 1983. One of his research interests is the origin of iron-rich sedimentary rocks in Proterozoic and Phanerozoic deposits. Another interest focuses on the history of the Earth's magnetic field and the intricacies of the magnetic recording mechanism. Present address: Department of Geological and Geophysical Sciences, Princeton University, Princeton, New Jersey, 08544, U.S.A.

213 shelf deposits (Heckel, 1972), yet this leaves much to be reconstructed because among sedimentary facies those of ancient shelves remain difficult to interpret. Two related questions reflect the trend of our concern. If glauconitic and chamositic minerals developed from rather similar clay mineral precursors in generally similar shelf environments why were they so consistently incompatible throughout Phanerozoic time? Conversely, if glauconitic peloids and chamositic ooids were produced by somewhat different specific conditions, how did they come to be intimately associated in some deposits? Many investigators have recorded their diverse opinions about these clay minerals. To some, the usefulness of glauconitic minerals as indicators of physical conditions is limited (Wermund, 1961). To others, rather different specific marine environments can be reconstructed for the origin of glauconitic and chamositic deposits (Odin and Matter, 1981; Berg-Madsen, 1983). Despite this disagreement a consensus confirms that each of these minerals has a meaning of its own. The stratigraphic record provides basic data for our attempt to resolve some of the persisting problems. Review of the geologic history of chamositic oolitic ironstones is facilitated by comprehensive compilations of them (Kimberley, 1978, 1981; Zitzmann, 1977, 1978), as well as by numerous studies of the facies in which they developed. In contrast, a similar detailed tabulation of ancient glauconitic deposits cannot be derived from the available literature. With but few exceptions most of them have been described vaguely or not at all, partly because they contain only scattered granules and are of no potential economic value, and older reports overlooked the possibility that greensands might contain subordinate chamositic granules. Nevertheless, Nikolaeva (1977) has compiled a useful general roster of glauconitic deposits. MINERALOGY Nomenclature

Names applied to green iron-rich clay minerals have been defined and used differently. The short glossary that follows adheres closely to definitions of Millot (1970), Bayliss (1975), Bailey (1980), Odin and Matter (1981) and Brindley (1982). We apply the generic terms, glauconitic and chamositic, to groups or "families" of minerals and to granules containing them where use of a generic name is appropriate. Glauconitic minerals, generally with more than 15% total Fe203, include green smectite, mixed-layer glauconite and smectite, and glauconite (Fig. 1). This is the glaucony facies of Millot (1970), Odin and Matter (1981), and Berg-Madsen (1983). Green ferric illite

214 iron-rich clay minerals

granular M

Glauconite

"~ ~

G/auconitic Smectite

"~ Green (3~ Smecfite 3 ~o .p. ~

Berthierine Chamosite

granular

non!

x

peloids 2

A

M

A

X

?

XX

X

XX

X

X

XX

x5

ooids M

A

X

x4

x

X

x4

XX

X

XX

Fig. 1. Morphology and mineralogy of iron-rich clay minerals. M = modern, A = ancient; XX = abundant, X = common, x = rare. 1 Film facies, including cement, thin crusts, and grain coating. 2 Some have single rims of radially arranged flakes. 3 Includes nontronite; specific mineral seldom identified. 4 Protoooids with only a few laminae. 5 Sample from Barataria Bay, Louisiana (Wermund, 1961, p. 1681).

with less than 10% Fe203 (Kossovskaya and Drits, 1970) is not a glauconitic mineral. Nevertheless, there apparently is a transition between illitic and glauconitic minerals recorded in intermediate stages of substitution of Fe for A1 (Berg-Madsen, 1983). C h a m o s i t i c m i n e r a l s are principally berthierine and chamosite in ooids and non-glauconitic peloids (Fig. 1). G l a u c o n i t e or g l a u c o n i t i c m i c a - - an iron- and potassium-rich 10 A mica (illite)-type clay mineral. Also called "mineral glauconite". Present in some Mesozoic and Cenozoic peloids, but predominates among Paleozoic ones. Glauconitic smectite - - mixed-layer glauconite and smectite with varying proportions of interlayering, and with some degraded illite among the expandable layers (McConchie and Lewis, 1980). Most common in Mesozoic to Recent peloids. Green smectite -iron-bearing dioctahedral smectite common in Late Cenozoic to Recent granules. Compositionally intermediate between montmorillonite and nontronite (Odom, 1976, p. 237). The latter mineral has been identified in a few samples. Berthierine -iron-aluminum 7 A trioctahedral serpentine. Has been called septachlorite and septachamosite (Deer et al., 1966, p. 241). It is the

215 chamosite of many older publications. Predominates in Mesozoic to Recent chamositic peloids and ooids. C h a m o s i t e - - iron-rich 14 .~ trioctahedral chlorite. Has commonly been applied to both 7 ,~ and 14 ,~ minerals, especially before X-ray analysis was available. Some authors who have not distinguished between berthierine and chamosite call both minerals chlorite or leptochlorite (Bubenicek, 1961, 1971). Others assign only the 14 A one to chlorite (Millot, 1970; Dreesen, 1982) Most common in Paleozoic oolitic ironstones. The name thuringite has been discarded because the mineral is a ferric chamosite (Bayliss, 1975). This variety is more common in older Paleozoic ironstones that have undergone advanced diagenesis or low-grade metamorphism. M i n e r a l growth and m i c r o - e n v i r o n m e n t s

Transport of iron to sites of development of chamositic and glauconitic minerals has been attributed to several different mechanisms (see Kimberley, 1981, p. 44). It may have been transported from source areas by rivers, in solution, colloidal suspension, or as coatings adsorbed on clay particles. It may, instead, have been transported by groundwater from swamps and marshes to nearshore environments. In contrast, it may have been dissolved from ferriferous clay minerals carried to deeper anoxic marine realms, and then spread shoreward by transgression. Carroll (1958) concluded that clay from the source area can carry enough adsorbed iron oxide to provide the iron needed for chamositic minerals. Detrital kaolinite present in many ironstones suggests that it may have been the carrier that escaped transformation. Clearly, the type of iron transport assumed bears on the mode of mineral growth proposed. Some glauconitic minerals may have been produced by transformation of degraded detrital mixed-layer clay minerals by adsorption of potassium and iron (Burst, 1958; Hower, 1961; Millot, 1970). More commonly, however, they apparently result from de novo growth, or neoformation, of dissolved material precipitated in pores in a substrate that is progressively altered and replaced (Odin and Matter, 1981). The precipitate forms as crystallites of green and glauconitic smectite, or as a silicic ferruginous gel which transforms diagenetically to a glauconitic smectite (Lisitsyna and Butuzova, 1982) in a slightly reducing micro-environment. Once formed the mineral presumably was susceptible to aging to a better ordered micaceous clay mineral as more potassium was added from sea water. Odin and L6tolle (1980) suggest that in this process the smectite stage may develop in a few thousand years, whereas development of glauconitic mica may require about 105 years. In contrast, Sorokin et al. (1980) and Valeton et al. (1982) believe that the initial gel-like precipitates may have ranged from a high K and Fe mica type

216 to a low K smectite type which reflect specific geochemical conditions at the place of origin. Laboratory synthesis of glauconitic minerals under essentially "natural" conditions (Harder, 1978, 1980) demonstrates "that glauconite may form by the precipitation of Fe-hydroxides and adsorption of silica and K from dilute solutions...". In this very early diagenetic process Fe and A1 in detrital minerals are dissolved in reducing micro-environments and precipitated in more oxidizing ones in a variety of substrates in the sediment. The origin of chamositic.minerals has long been the subject of speculation (see Kimberley, 1981, p. 58, Tables IV, V); there is general agreement only that they are unstable in the presence of free (molecular) oxygen (Spear, 1968). An essentially primary precipitation from seawater under mildly reducing conditions has been proposed (Taylor, 1951; Hallimond, 1951; Dunham, 1960; Schoen, 1964). Direct precipitation of both glauconitic and chamositic minerals probably did fill voids, such as those in microfossils. Nevertheless, the consensus apparently favors an early diagenetic anoxic modification of a precursor that had accumulated in'granules on the oxic sea floor. This primary substrate has been described as a precipitated Fe-A1-Si gel (Harder, 1964; Borchert, 1965; Spear, 1968; Talbot, 1974), a detrital poorly ordered kaolinite-Fe203 hydroxide colloid (Carroll, 1958; Karpov et al., 1967; Schellmann, 1969; Bhattacharyya, 1983), a precipitated ferric oxide or alumino-goethite (Bubenicek, 1971; Millot, 1970; Velde et al., 1974; Gygi, 1981), or detrital lateritic particles (Siehl and Thein, 1978; Harder, 1980). In addition, some chamositic minerals may have replaced aragonitic ones in carbonate sequences that were mantled by volcanic ash (Dreesen, 1982), but this was a subordinate process. Berthierine (7 A) probably is the first-formed iron-rich clay mineral (Fig. 1)~ and it apparently ages diagenetically to 14 chamosite (Odin and Matter, 1981; Velde et al., 1974). Nevertheless, the persistence of berthierine in Precambrian deposits (French, 1973; Klein and Fink, 1976) and the presence of both berthierine and chamosite in a Silurian ironstone (Schoen, 1964) are evidence against a pervasive process of this sort. Both glauconitic and chamositic minerals have been subjected to later diagenetic alteration. The former may be phosphatized, or locally replaced by pyrite (McConchie and Lewis, 1980). Chamositic ooids may be replaced by a phosphatic mineral, siderite, pyrite, or magnetite, which progressively destroys their concentric fabric (Dunham, 1960, p. 249). In at least one ironstone, however, completely pyritized chamositic ooids still have their sheaths (Maynard, 1983, p. 51). As a result of extensive diagenetic alteration the detailed mineralogy of chamositic ironstones is normally more varied (Borchert, 1965, figs. 13, 14) than that of most greensands. Glauconitic and chamositic minerals have commonly been characterized in terms of oxidation-reduction potential (Eh) and hydrogen ion activity

217

(pH) of the local environment (for example, Borchert, 1965), as summarized in a widely cited diagram by Krumbein and Garrels (1952). In this scheme the very general geochemical parameters range from mildly reducing to mildly oxidizing within an ordinary range of pH. Consequently, the broad limits fail to define specific conditions of deposition. Berner (1981) recognized these deficiencies for analysing modern sediments, and proposed a classification based on the presence or absence of oxygen and sulfide dissolved in the sediment. Maynard (1982) then pointed out that the deficiencies applied to the analysis of ancient sediments as well, and he extended Berner's classification of geochemical environments to them. Within it glauconite and chamositic minerals, commonly associated with pyrite, prevail in the carbon-poor anoxic sulfidic zone, developing under slightly more oxic conditions than siderite. Future investigations of the geochemical controls of both glauconitic and chamositic minerals should attempt to relate them to the categories of this new classification.

Morphology and origin of granules Glauconitic and chamositic clay minerals occur in a film facies of cement (intergranular filling), grain coating, and thin crusts, as well as in a granular or peloidal facies of authigenic sand-size particles (Fig. 1). Multi-coated granules are ooids. Those without concentric sheats are peloids if they are of inorganic origin or their mode of origin has not been established, and pellets if of organic fecal origin (Bathurst, 1975). Modern and ancient glauconitic peloids (glaucoliths of Berg-Madsen, 1983) are spheroidal to ellipsoidal or ovoid, lobate or botryoidal (colloform), as well as tabular, vermiform or accordion-shaped (Pryor, 1975; Triplehorn, 1966; Boyer et al., 1977). Additional morphological types include internal moulds of microfossils and spongy and fragmentary grains (McConchie and Lewis, 1980), and moulds of inorganic pore space marked by imprints of crystal faces of calcite cement (Berg-Madsen, 1983). Many of the glauconitic peloids bear internal and surface cracks. These have been attributed to shrinkage due to loss of interlayered water (Odom, 1976), but most of them probably are the result of displacive growth of glauconitic minerals in initially smaller substrates (Boyer et al., 1977; Odin and Matter, 1981). The precursory peloidal substrate may be biogenic carbonate debris, detrital mineral or lithic grains, including volcanic glass, or more commonly fecal pellets (coprolites) of detrital clay made by ingesters of organic-rich bottom mud or by filter feeders of suspended clay (McRae, 1972; Whitehead, 1979). Kaolinite apparently was a common initial clay mineral in Recent glauconitized fecal pellets (Pryor, 1975). Change in the composition of the clay minerals by chemical reactions in the gut of mud feeders may have prepared

218 unstable precursors of glauconitization (Pryor, 1975; Syvitski and Lewis, 1980). Moreover, organic matter in the pellets provided a favorable mildly reducing environment on a moderately oxidative sea floor. Some glauconitic peloids may also be replacements of internal moulds of minute tests, but probably the glauconitic mineral more commonly filled the void by direct precipitation from solution, without a substrate and growing by addition of successive layers (Seed, 1968; Odom, 1976). According to Kohler and K6ster (1976) and Valeton et al. (1982), the widely-held peloidal s~bstrate origin of glauconitic peloids is inadequate to account for those in Cretaceous and Cenozoic deposits in Germany (see also Nikolaeva and Matyunina, 1981). As an alternative the shape, size, and amount of cracking, as well as variations in chemical and mineral composition of the peloids are attributed to specific environments that prevailed during periodic precipitation of coagulated gel-like lenses (also Sorokin et al., 1980). Presumably, the dehydration cracking that broke up the gelatinous precipitate produced soft peloids which were reworked and redeposited locally without significant distortion or compaction. This sort of physical process has also been proposed for the origin of the granular facies of Proterozoic iron formations (Dimroth and Chauvel, 1973; Hall and Goode, 1978). Glauconitic peloids are generally considered to have an essentially uniform internal cryptocrystalline fabric produced by randomly oriented overlapping clay platelets (McRae, 1972; Pryor, 1975; Triplehorn, 1966; Whitehead, 1979). This fabric may be marked by microscopic veinlets, by scattered patches of micaceous texture or of pyrite, and by clots of different shades of green (Muravev and Voronin, 1976). The scanning electron microscope reveals several kinds of nanostructure of the authigenic glauconitic minerals (McConchie and Lewis, 1980; Odin and Matter, 1981; Berg-Madsen, 1983): globular or caterpillar structures, oriented lamellar to micaceous structures (Van Wie, 1971), and boxwork and rosette structures in which the platelets may be stacked edge to edge or face to edge in a house of cards fabric (Hein et al., 1974; Odom, 1976; Valeton et al., 1982). As glauconitization proceeds toward completion, the fabric of the initial substrate is obliterated. A thin rim (corona) of radially oriented flakes around some glauconitic peloids probably was precipitated in situ (Zumpe, 1971; Odom, 1976; McConchie and Lewis, 1980), whereas a single tangential rim on a few others may have accreted by rolling (Triplehorn, 1966). Muravev and Voronin (1976) have described both types of "crustified borders" on glauconitic peloids. Peloids with a few distinct internal rims apparently record several stages of accretion (Zumpe, 1971) during the filling of minute tests (Odom, 1976).

219

Essentially all modern berthierine granules are peloids (Fig. 1), and these are presumed to be largely of fecal origin. Only rarely are they proto-ooids with but a few sheaths, such as those granules forming off the Mahakam Delta in Kalimantan (Allen et al., 1979) and in Loch Etive, Scotland (Rohrlich et al., 1969). Along the Niger, Ogooue, and Orinoco deltas proto-ooids of goethite occur near the coast, whereas berthierine peloids are forming farther offshore (Porrenga, 1967; Giresse,1969). In marked contrast to the preponderance of peloids among Recent chamositic granules, ancient ones are predominantly ooids and only much less commonly non-laminar peloids. Ooids in ancient ironstones were built either by simple or by aberrant accretion (Hemingway, 1974). The predominant simple ones consist of concentric sheats around an inconspicuous nucleus of clay or ferric oxide, and less commonly around an ooid or fossil fragment, or a quartz grain. A few have internal shrinkage cracks which rarely reach the surface. Many chamositic ooids and ferric-oxide (commonly hematitic) replacements of them have a primary discoidal (flaxseed or bean) shape (Chowns, 1968, 1970) with only slight or no equatorial thickening of the laminae. Closely packed ooids in some ironstones are elongate parallel to bedding; in other beds the elongation is randomly oriented, suggesting jostling of the ooids during post-burial bioturbation. Simple concentric ooids commonly are rather well sorted for size (Dunham, 1960) as well as for shape (Chowns, 1968; pers. comm., 1981). Some still-plastic ooids were partly unwrapped, compressed, or torn during reworking and burial to form spastoliths. Locally several ooids were bound together by a common multi-layer coat to form a composite ooid. The rarer aberrant ellipsoidal ooids have a pronounced equatorial bulge of thickened laminae and a polar discontinuity of some sheaths. This distinctive shape may have resulted from preferential equatorial laminar accretion (Knox, 1970; Wright, 1977). Very rarely ooids with a nucleus of quartz are uniquely asymmetrical or eccentric, with discontinuous laminae coating only part of the grain. A few eccentric ooids became concentric by the addition of entire laminae. Chamositic and ferric-oxide discoidal ooids (Chowns, 1970) are commonly associated with phosphatic kaolinitic peloids or rarely ooids, as well as with distinctly spherical concentric ferric-oxide ooids, many of which have a nucleus of quartz sand, shell fragment, or broken ooid. In a few deposits discoidal and spherical ooids developed in somewhat different local environments (Schellmann, 1969; James and Van Houten, 1979), and are essentially unmixed. Other ironstones contain chamositic and ferric-oxide ooids apparently made in the same area under temporarily changing physico-chemical conditions, so that locally layers of ferric-oxide ooids and chamositic

220 ooids are interbedded. In some of the thick ironstones chamositic ooids predominate in the lowermost and uppermost parts, which commonly are sideritic, whereas ferric-oxide ooids predominate in the middle. Moreover. within some ooids sheaths of ferric oxide or kaolinite (and rarely a phosphatic mineral) alternate, coat by coat (see Taylor, 1951, fig. 4), implying an essentially primary origin of the oxide and silicate sheaths. Many of the iron-bearing ooids, appropriately analysed, contain amorphous silica. Sheaths of simple, concentric chamositic ooids are constructed of tangentially arranged flakes of clay minerals (Rohrlich, 1974; Wright, 1977). Largely by analogy with the origin of tangential carbonate ooids by rolling in shallow agitated water, the gelatinous coats of most of the chamositic ooids are believed to have been accreted mechanically during gentle rolling on the soft sea floor. It is not clear, however, whether this accretion was a direct addition of tangentially arranged platelets, or resulted from modification of radially accreted coats by repeated rolling and minor abrasion. Moreover, the role of micro-organisms in the growth of ooids is not known, but it may be significant (Hallimond, 1951). Evidence of blue-green algae involved in producing chamositic ooids in lakes (Shterenberg et al., 1968) is suggestive. In some ancient ooids zones of tangentially constructed laminae alternate with zones of randomly oriented crystals (Chowns, 1968), indicating interruption of the normal process of accretion. Hemingway (1974) has concluded that composite ooids with a succession of common sheaths around several granules could not have been built by precipitation during agitation on the sea-floor. Accordingly, he supports an alternative hypothesis of origin that chamositic ooids developed within the sediment by early diagenetic growth (see also Carozzi, 1960; Gygi, 1981). Ooids produced this way presumably were freed by intraformational scour and concentrated in beds by bottom currents. This hypothesis overlooks the fact that individual granules in composite carbonate ooids probably were made by gentle rolling on the sea floor, and that laminated spherules that grew in situ in bauxite and laterite have radially arranged platelets in their sheaths. The latter observation not only implies that tangentially coated ooids did not grow Within mud, but it also eliminates particles developed in laterites and bauxites as direct precursors of tangentially coated ooids in ironstones (Kimberley, 1980; Bhattacharyya and Kakimoto, 1982). In contrast to the prevailing view that ooids in oolitic ironstones developed in shallow marine environments, Nahon et al. (1980) propose that they are in situ products of lateritic weathering. According to this unlikely interpretation, ironstones "represent paleoexposure crusts in a tropical climate ... related to worldwide eustatic changes of sea level" (p. 1297).

221 FACIES MODELS

Introduction

Comparison of ancient glauconitic and chamositic deposits is facilitated by construction of facies models, or stratigraphic sequences, which portray successions of associated sedimentary rocks as a record of vertical facies changes (Fig. 2). These patterns, in turn, imply lateral facies relations according to Walther's principle of superposition of strata (Middleton, 1973). The importance of vertical shoaling-upward sequences and cyclic sedimentation associated with chamositic ironstones was recognized by Hemingway (1951; see also Taylor, 1951; Dunham, 1960; Hallam, 1966). Maynard (1983) has reviewed examples of vertical sequences containing oolitic ironstones. Muravev (1976) summarized the predictable associations of glauconitic strata in repeated vertical successions. Odin (1975; also Odin and L6tolle, 1980) has portrayed a more general, regional model of the development of glauconitic and chamositic minerals in a sequence of geochemical facies. In this model the ferriferous clay minerals are most common in an oxidate facies, less common in a carbonate and hydrolyzate facies, and rare in a facies of insoluble residues. Most of the ancient concentrations of iron-rich clay minerals accumulated in marine shelf environments during relatively long periods of reduced influx of sediment and of incipient transgression, after briefer intervals of pro-

~hardground transgression

greensand

B

stillst and

transgression

Fig. 2. Facies model of marine deposits with chamositic oolitic ironstone (A) and glauconitic peloidal greensand (B). Shoaling- and coarsening-upward sequence is produced by progradation or regression; coarsening-upward sequence is also produced by migration of sandwave across associated muddy facies. Transgression spreads offshore facies shoreward.

222 gradation or shoaling (Fig. 2). Within some sequences maximum transgression introduced anoxic conditions that commonly produced pyritic black shale (Borchert, 1965; Hemingway, 1974). In these asymmetrical sequences (megacycles of Bubenicek, 1961) the granules developed most abundantly in condensed deposits (Hallam and Bradshaw, 1979) that commonly are burrowed, contain fossil debris, and are interbedded with ferruginized and phosphatized hardgrounds, or associated with unconformities (Jenkyns, 1971). The role of sediment-starved intervals in glauconitic sequences was recognized long ago by Goldman (1922), Heim (1924, 1934) and Hadding (1932). Development of glauconitic (Lindstr6m, 1963, 1979; Fiarsich, 1971; Singh and Kumar, 1978; Whitehead, 1979; Gebhard, 1982) and chamositic (Urban, 1966; Geyer and Hinkelbein, 1971; Bitterli, 1979; Fi~rsich, 1979; Bhattacharyya, 1980; Gygi, 1981) minerals in condensed sequences and associated hardgrounds has been emphasized.

Examples The chamositic facies model presented here (Fig. 2; Table I) is welldisplayed in m u d d y and sandy coarsening-upward sequences, some of which

TABLE I Coarsening- or shoaling-upward facies sequences with chamositic oolitic ironstones Age Late Miocene Late Eocene Late Eocene Middle Eocene Late Cretaceous Late Cretaceous Early Cretaceous Late Jurassic Late Jurassic Middle Jurassic Early Mid-Jurassic Late Early Jurassic ,1 Late Early Jurassic ,1 Late Early Jurassic ,1 Early Jurassic Early Jurassic Late Devonian Middle Devonian Early Silurian Early Ordovician

Locality northeastern Columbia northeastern Columbia southwestern Tunisia north-central Louisiana southeastern Egypt western Siberia Atlantic shelf, New Jersey southern England northwestern Switzerland northwestern Switzerland northeastern England southwestern Germany Luxembourg northeastern France northeastern England southeastern Spain central Libya central Pennsylvania Alabama and Georgia eastern Newfoundland

Reference James and Van Houten, 1979 Kimberley, 1980 Nicolini, 1967 Jones, 1969 Bhattacharyya, 1980 Nagorskiy, 1981 Cunliffe, 1982 Talbot, 1973, 1974 Bitterli, 1979; Gygi, 1981 Lusser, 1980 Knox, 1970 Urban, 1966 Thein, 1975; Teyssen, 1983a, b Bubenicek, 1971 Hemingway, 1974 Geyer and Hinkelbein, 1971 Van Houten and Karasek, 1981 Kaiser, 1972 Chowns and McKinney, 1980 Maynard, 1983

,1 Successive coarsening-upward oolitic ironstones within the major ore deposit.

223 have offshore (transgressive) carbonate deposits in the lower part (Taylor, 1951; Talbot, 1973). Chamositic ironstones also developed over a prograding, shoaling-upward carbonate cycle in a condensed sequence commonly surmounted by a hardground. This sort of asymmetrical sequence occurs in Early Jurassic deposits in southeastern Spain (Geyer and Hinkelbeim 1971), in Middle (Lusser, 1980) and Late (Bitterli, 1979; Gygi, 1981) Jurassic strata in northwestern Switzerland, and in Late Eocene deposits of southeastern Tunisia (Nicolini, 1967). A few examples of the glauconitic model in detrital successions (Fig. 2), some of which contain subordinate limestone, include Middle Cambrian deposits of southeastern Sweden (Berg-Madsen, 1983), Cretaceous sequences in western interior Canada (Stott, 1982), Late Cretaceous and Early Cenozoic deposits of the northern Atlantic coastal plain (Owens and Sohl, 1969) and southwestern Russia (Muravev, 1976), and the Middle Eocene formations of the northern Gulf Coast (Fisher, 1964). Some greensands are associated with transgressions in shoaling-upward carbonate sequences, as in the Middle Cambrian and Early Ordovician limestones of northern Denmark and southern Sweden (Lindstr6m, 1979; Berg-Madsen, 1983) and in Early Carboniferous successions in east-central United States (Whitehead, 1979). In this carbonate facies glauconitic peloids are commonly mixed with phosphatic granules that accumulated under generally similar, but specifically different, conditions (Odin and L6tolle, 1980). Discussion

Most glauconitic and chamositic deposits are relatively local concentrations compared with the more widespread distribution of facies with which they are associated. Nevertheless, several of the greensands are extensive (for example, Ghosh, 1984). Granules and crusts of these clay minerals developed most readily in condensed sequences during the initial stage of shoreline retreat, as a transgression advanced across a seafloor strewn with fecal pellets, tests, fossil fragments, and lithic and mineral grains, as well as precursor ooids. Alternating sheaths of clay mineral and ferric oxide, and rarely of opal, have been attributed to shuffling between different local environments (Dunham, 1960), but variation in sheaths probably was more readily produced by temporarily changing local conditions. As the ooids and glauconitic peloids developed they were swept from their place of origin by currents and concentrated in accreting sandwaves, bars and shoals, or in channels, where repeated interruptions permitted increased burrowing and development of ferruginized or phosphatized crusts. Within this general framework hardgrounds, or omission surfaces, which developed most readily in carbonate sequences, formed when sedimentation

224

ceased altogether. The burrowed and ferruginized tops of a few chamositic ironstones are overlain by a thin lag deposit, or bonebed, at the base of the succeeding sequence (Fig. 2). In contrast, some greensands and ironstones overlie a hardground. In them the abundance of ooids and peloids decreases upward in the overlying transgressive mudstones. Moreover, in some sequences thin lenses of oolitic ironstone occur in a subtidal muddy facies (Chowns, 1968; Chowns and McKinney, 1980; Van Houten and Karasek, 1981; Foos, 1983) of the transgressive phase. In many of the glauconitic deposits peloids are scattered throughout a bed of sandstone or limestone, or they are concentrated in a less regular pattern in the vertical succession of transgressive deposits below the sandy facies of a shoaling upward sequence. Teyssen (1983a,b) has proposed a tidal current model to account for the concentration of ooids in the Jurassic minette ironstones. This interpretation implies a rather high rate of transport and deposition in a meso- to macro-tidal regime. In this reconstruction some of the ironstones were built by time-velocity symmetrical currents, with a strong periodic component and relatively weak steady one, that deposited cross-bedded sandwaves whose oolitic foreset facies migrated over an associated nearshore bottomset facies. This mode of origin involved no shift of submarine environment. According to Teyssen, other ironstones were build by time-velocity asymmetrical currents with a relatively strong steady component that transported oolitic sandwaves seaward in intershoal channels, depositing them on an offshore muddy facies. R E G I O N A L E N V I R O N M E N T OF A C C U M U L A T I O N

Modern deposits Most investigators believe that berthierine and glauconitic peloids on the sea floor today are largely of authigenic origin (Lisitsyna and Butuzova, 1979; Odin and Matter, 1981). Logvinenko (1982; also Logvinenko et al., 1980), in contrast, maintains that most of the glauconitic peloids have been reworked from older formations, citing anomalously old radiometric ages of analysed glauconitic grains. There undoubtedly are local concentrations of detrital (inherited) glauconitic peloids, but the older than expected age for Recent glauconitic grains does not necessarily contradict their authigenic origin. In these modern peloids, as in ancient ones, the apparent age is a compromise between the age of the inherited clay (substrate) and the time of glauconitization (Odin, 1982). A reworked (clastic, detrital) origin of modern berthierine peloids is improbable because older relevant formations contain ooids. In the modern oceans almost all of the berthierine and glauconitic peloids

225 REGIONAL ENVIRONMENT

Fluvial to Deltaic Lacustrine saline Intertidal ~ d~L. - o~St

"

i~0o° ; ~o~

c

Shallow Offshore

Pn ~ '

L,-tt~-~i ' !~i!i!i~'l

c-:~ /

!~

io-5om ,G,; I i I

Deeper Offshore 50-200m

C

0

'~ ~

54o,.

OOIDS modern ancient

PELOIDS m o d e r n I ancient

1 :

,

i

i

i i~ i!

l

L/

'

Fig. 3. Relative a b u n d a n c e o f glauconitic a n d c h a m o s i t i c granules (see Fig. 1) in regional facies. C h a m o s i t i c minerals h o r i z o n t a l p a t t e r n ; C = a b u n d a n t , c = rare. G l a u c o n i t i c minerals vertical p a t t e r n ; G = a b u n d a n t , g = rare. N o n t r o n i t e - n(rare); P r o t o - o o i d s - p (rare).

are accumulating in regions of retarded detrital sedimentation on shelves along cratonic margins. A few glauconitic deposits also occur on carbonate banks on shallow shelves, as off southeastern United States and north-central and east-central Australia (Odin and Matter, 1981, fig. 10). Berthierine peloids are limited to tropical and subtropical realms, whereas glauconitic ones range from latitude 50°S to 65°N. Presumably development of these granules has been favored by the Holocene global rise in sea level. Peloidal berthierine forms in low-energy environments (Fig. 3) on the inner embayed shelf (estuarine environment under fluvial influence, according to Odin and L6tolle, 1980). In addition, chamositic proto-ooids are forming very locally in coastal embayments and in a saline lake (Rohrlich et al., 1969; Lemoalle and Dupont, 1973; Allen et al., 1979). Glauconitic peloids, on the other hand, are most abundant on the middle and outer shelf (Fig. 3), especially where upwelling is intense (Bezrukov and Senin, 1970), and they may also be developing at oceanic depths of a few kilometers (Odin and Stephan, 1981). In a few localities, however, the abundance of glauconitic peloids does not increase steadily seaward (Hein et al., 1974).

Ancient deposits The general environment of ironstone accumulation has consistently been interpreted as warm and humid with mature soil developed in vegetated uplands (for example, Taylor, 1951; Bubenicek, 1961). Most of the chamositic deposits and much less commonly glauconitic ones, developed in nearshore facies either seaward of low-energy deltas or along interdeitaic

226 coasts in broad embayments of cratonic margins, or in inland seas largely surrounded by land (Wopfner and Schwarzbach, 1976; Kimberley, 1978), as in foredeeps and intracratonic (epicratonic) basins. Many of the glauconitic greensands accumulated in open seas along cratonic margins of major oceans or on structurally high platforms in them, ranging rather widely from shallow to deeper offshore facies (Velde et al., 1974). This regional facies distribution conforms to that of their modern counterparts in a very general way (Fig. 3), as in a few examples of nearshore chamositic ironstones grading seaward into glauconitic deposits, and of ironstones that either succeed or are overlain by offshore glauconitic deposits in a stratigraphic section. In contrast to a very general agreement with modern facies distribution some of the ancient glauconitic and chamositic minerals accumulated in more varied marine habitats (Fig. 3). Interpretation of the ancient environments ranges from intertidal and subtidal inner shelf (Bubenicek, 1971; Thein, 1975; Chafetz, 1978; Singh and Kumar, 1978; Chowns and McKinney, 1980; Cunliffe, 1982) to deep offshore (Gygi, 1981) for both kinds of granules, and includes sequences on carbonate shelves. Moreover, Sorokin et al. (1980) interpret the favorable setting for glauconitic peloids as a delta front environment where river and sea water mix, as well as a quiet water marginal marine realm adjacent to coastal lowlands. These authors also report that the mineral varies from glauconitic mica (low-A1) in an open sea facies to nearly pure smectite (high-A1) in a coastal facies in a very consistent way. This kind of local control of glauconitic and chamositic minerals is corroborated by Early Paleozoic greensands of Baltoscania (Berg-Madsen, 1983) and Late Cretaceous greensands of the Atlantic Coastal Plain (Owens and Sohl, 1973). In both sequences low-Al glauconitic minerals apparently developed in cool, deeper (750 m) water, whereas high-A1 glauconitic minerals accumulated in cool, shallow water. Berthierine is also favored by shallow-water conditions, but in warm seas. In this reconstruction the high-A1 mineral of Berg-Madsen (1983) is a low-Fe glauconitic mica, whereas the high-Al mineral of Owens and Sohl (1973) and Sorokin et al. (1980) is a glauconitic smectite. A subordinate mode of locally transported (parautochthonous) chamositic and glauconitic granules in offshore deposits involves the role of storms. In this setting granules were swept from intertidal and nearshore subtidal areas by storm surges and concentrated in thin offshore deposits (Dreesen, 1982; Gebhard, 1982; Gygi, 1981). In addition, a minor development of chamositic ooids associated with abundant ferric-oxide ooids in a chamositic matrix has been reported in Middle (Bronevoi et al., 1967) and Late (Shnyukov and Fesyunov, 1968) Cenozoic alluvial deposits of southwestern Russia. Another distinctive mode is an intimate association of glauconitic and

227 TABLE II Deposits containing both chamositic and glauconitic granules Age

Locality

Reference

Middle Pliocene Middle Eocene

southwestern Russia north-central Gulf Coast Louisiana, Texas southern Germany central Russia northern Atlantic coastal plain, New Jersey northern Algeria east-central England northwestern Venezuela southwestern Russia

Zitzmann, 1978, pp. 236-238 Jones, 1969; Kaiser, 1974

Early Eocene Late Cretaceous Late Cretaceous Mid-Cretaceous Early Cretaceous Early Cretaceous Middle Devonian

Ziegler, 1975 Chilingar, 1956 Owens and Sohl, 1973 Velde et al., 1974 Hallimond, 1925 Bartok et al., 1981 Karpov et al., 1967

chamositic granules in several deposits (Table II). The mixing of these granules in some of the examples records the derivation of one or the other from an older formation or from a different environment. The evidence suggests, however, that at least locally development of each of the iron-rich clay minerals was not as sharply segregated in the marine realm as the modern model suggests (Odin and Matter, 1981; figs. 18, 19). GEOLOGIC RECORD

Distribution

Strakhov (1969, fig. 87) and Cooke and McElhinney (1979, fig. 5) have traced the general Phanerozoic distribution of chamositic ironstones and recognized that the pattern reflects major episodes in earth history. The geographic and latitudinal distribution of oolitic ironstones on cratonic blocks of Laurasia and Gondwana has been charted by Van Houten and Bhattacharyya (1982, figs. 2, 3). Chamositic (with associated greenalite) peloids began to develop in Early Proterozoic time (French, 1973; Hall and Goode, 1978) and ooids followed in Late Proterozoic time (Canavan, 1965; Harms, 1965) as minor constituents in iron formations. Chamositic oolitic ironstones (Fig. 4) began to accumulate in Middle (?) to Late Cambrian time and rapidly became widespread, with major development from earliest Ordovician (Petranek, 1964) to latest Devonian time. During this episode the ironstones were deposited mainly along margins of scattered cratonic blocks. Subsequent

228 PHANEROZOIC Glauconitic Peloids

MA 0

RECORD Chamositic Ooids

I

Sea Level

Lo

¢O l

,0o I

Hi

/

I

2OO

:

300 •

Ib

/

/

400 •

500

c 600

Pr~c~mbrian

Oangeea assembled

/

Fig. 4. Comparative Phanerozoic record of glauconitic peloids and chamositic ooids. Tally of oolitic ironstones is based on actual number reported (mainly from Kimberley, 1978, and Zitzmann, 1977, 1978). Plot of glauconitic deposits is an estimate of relative abundance based on incomplete data. Sea level curve and time scale mainly after Vail et al. (1977). Odin (1982) has summarized more recently calculated Phanerozoic ages.

consolidation of Pangaea in Late Paleozoic and Early Mesozoic time was followed by renewed block dispersal and a second remarkable ironstone episode. It began in Late Triassic time, expanded during the Jurassic Period with an extensive transgression of Tethys, and continued through Cretaceous to Middle Cenozoic time, accompanied by a major global rise and subsequent waning of sea level. Apparently no ironstone accumulated in Paleocene time, however. During the long congenial interval most of the ironstones developed in intracratonic basins on Laurasia which has been assembled during successive Paleozoic orogenies. In addition, a few ironstones accumulated on large drifting blocks of Gondwana. Small-scale Late Cenozoic renewal of chamositic oolites associated with minor transgressions during generally falling sea level (Fig. 4) ended in Middle Pliocene time with one of the largest Phanerozoic oolitic ironstones accumulating in the Caucasian foredeep in southwestern Russia (Zitzmann, 1978, pp. 236-238). The oldest record of glauconitic peloids is in Early Proterozoic deposits of western Australia (Hall and Goode, 1978), and they are present in Late

229

Proterozoic deposits of India (Singh, 1980). Glauconitic peloids became increasingly common as cratonic blocks were rifted and dispersed from the Late Proterozoic supercontinent (Morel and Irving, 1978), sea level began to rise (Fig. 4), and fecal pellet makers presumably became abundant during the Cambrian Period (Brasier, 1980). The incomplete record, based largely on North American and Eurasian data, suggests that glauconitic peloids continued to accumulate at least locally through most of Phanerozoic time. They were especially abundant in Middle Cambrian to Early Ordovician time during an episode of very high stand of the sea (Brasier, 1980), accumulating mostly on large cratonic blocks and their passive margins. Then development of greensands waned during later Paleozoic and Early to Middle Mesozoic time. A second major greensand episode associated with a global rise in sea level in Cretaceous time (Jenkyns, 1980) persisted into the Early Cenozoic Era, especially along the margins of the opening Atlantic, Tethys, and Gulf of Mexico basins. Minor renewal in Late Cenozoic time continued, with Pleistocene interruptions, to the Holocene.

Discussion The two long congenial Phanerozoic episodes with their general high stand of sea level were effective in fostering the development of glauconitic greensands and chamositic ironstones. Nevertheless, the several exceptions, especially well-documented by the record of ironstones, emphasize that no one constraint or favorable condition was a necessary factor in producing the iron-rich clay minerals. For example, some oolitic ironstones accumulated when cratonic blocks were consolidated and global sea level was relatively low, as in mid-Permian, Late Triassic, earliest Jurassic, and Late Cenozoic time. Moreover, a deposit of chamositic ooids interpreted to be replacements of aragonitic ones (Dreesen, 1982) accumulated in northwestern Europe when the climate there was semi-arid and extension and volcanism were active during Late Devonian time. The less well established pattern of glauconitic deposits confirms a general association with high stand of sea level, as in Late Cambrian-Early Ordovician and Late Cretaceous time. More specifically, however, there is little evident correlation between repeated transgressions and widespread development of greensands (Fig. 4). SUMMARY AND PROBLEMS

Mineralogy and morphology The many published discussions of the origin of glauconitic and chamositic minerals and granules introduce, and commonly confuse, various

230

processes ranging from (1) direct precipitation on an anoxic seafloor, (2) replacement of an organic-rich substrate on an oxic sea floor, (3) precipitation in anoxic sediments a few centimeters below the bottom, followed by exhumation of the granules, to (4) growth of a precursor on an oxic sea floor followed by burial and replacement. The minerals produced are heterogeneous in detail, reflecting a complex interaction of both local chemical conditions at the time of deposition and effects of organic matter on diagenetic processes. Formation of glauconitic and chamositic minerals required mildly reducing conditions, and both are commonly associated with and locally replaced by minor amounts of pyrite. Harder (1980, p. 220) suggests that production of one or the other of these was controlled more by the Si content of sea water, or its derived pore water, than by temperature and pressure at the sea floor. Glauconitic minerals, favored by high Si and K content, probably were largely products of neoformation by replacement of several kinds of substrates or by filling voids, as in tests of microfauna. Growth in agglutinated fecal pellets and minute tests was aided by the presence of organic matter. The common fecal pellet origin required abundant animals that produced coprolites of the appropriate size, and which were replaced, reworked and buried without substantial modification of shape. Identification of a particular kind of substrate, ranging from nearshore shell debris to offshore tests of planktonic animals, provides information not only about the habitat of the substrate, but also about events leading to the offshore environment of glauconitization (Giresse et al., 1980; Jeans et al., 1982). During the process original fabric was progressively destroyed, making identification of the original substrate increasingly difficult. Some glauconitic peloids may have been made by dehydration cracking of a primary gelatinous precipitate which consistently broke into sand-size fragments. This process has not been confirmed, however, and current concensus favors replacement of substrates. Nevertheless, the relevance of precipitation and cracking to production of Precambrian granules before abundant coprolites and minute tests were available suggests that this possible mode of origin deserves consideration. Modern berthierine peloids probably are products of neoformation by replacement of granular substrates in a mode similar to glauconitization (Odin and Matter, 1981), and perhaps requiring shallow burial under reducing conditions. In contrast, the origin of ancient chamositic ooids, lacking a direct modern analogue, is not clearly understood and the subject of conflicting interpretations. Chamositic ooids probably were not made on an anoxic sea floor by direct precipitation in concentric sheaths. More reasonably, an initial Al-rich clay mineral, either a flocculated degraded detrital one or a neoformed one, formed an ooidal substrate; but no ooids of this sort are

231

known among modern granules. Berthierine probably developed in the precursor clay-mineral ooids by transformation involving shallow burial and addition of Fe. Neoformation was a less likely process because it commonly obliterates an inherited fabric (Dunham, 1960, p. 255; Giresse et al., 1980). Evidence from laboratory experiments and electron microprobe analysis of ancient ooids (Bhattacharyya, 1983) supports the hypothesis that berthierine in ooids developed by progressive early diagenetic transformation of detrital kaolinite. Assuming that the precursor ooids were made of detrital, or possibly precipitated, kaolinite on an oxic sea floor, the necessary ferrous iron may have been added (1) at the place of ooid origin during shallow burial, thus requiring exhumation and transport to the place of concentration, (2) after transport to and burial in accreting deposits, or (3) at either site when anoxic iron-bearing water of an oxygen minimum zone transgressed shoreward, as suggested by Borchert (1965). The latter idea implies that a deeper-water black shale commonly lay offshore, beyond a chamositic ironstone. If chamositic ooids and some glauconitic peloids developed in a kaolinitic substrate, construction of the ooids cannot be attributed entirely to the electrostatic behavior of kaolinite. It may have favored development of ooids, but some other factor, such as agitated shallow water low in silica, probably participated. Most of the chamositic ooids apparently required small fecal pellets for cores. If so, a few micropellets, detectable by detailed size analysis, should be present in oolitic ironstones as well as in modern peloidal deposits. Gentle agitation of free-rolling cores on a soft, muddy sea floor presumably built the ooidal precursors by mechanical accretion. The spheroidal to ellipsoidal shape of most of the ooids, especially those with a nucleus of ferric oxide or clay mineral, records the manner of addition of laminae. The shape of ooids accreted around a quartz grain commonly reflects the shape of that nucleus. Moreover, chamositic ooid to spastolithic ooids probably accumulated in more muddy environments with weak currents, whereas spherical ferric oxide-rich ooids developed in aerated, agitated water where clay was winnowed (Schellmann, 1969). Significant questions concerning the origin of glauconitic and chamositic granules persist: (1) Assuming that some clay mineral peloids have been produced by shrinkage cracking of a colloidal precipitate, as commonly proposed for Precambrian granules, can these be distinguished from fecal pellets by their texture, by their O18/O 16 content, or some other consistent feature? (2) If modern berthierine and glauconitic minerals replace similar substrates, what specific constraints determine the pathway of mineral development? (3) Why are there no thick concentrations of non-green illite or smectite fecal pellets in ancient deposits? Was glauconitization required to

232 produce abundant durable pellets? What prevented glauconitic minerals or their precursors from forming ooids? (4) Why does Berthierine form only peloids and rare proto-ooids today; why are there no ancient deposits of abundant chamositic peloids? Does this discrepancy mean that the formation of ancient chamositic ooids is without a modern analogue? (5) Does 7 berthierine normally alter ("age") diagenetically to 14 ,~ chamosite? If so, what conditions favor the process? (6) What was the origin of the prevalent clay mineral and ferric oxide cores of chamositic ooids? Were they micro-fecal pellets, or possibly the product of flocculation, as in the Eocene "sawdust" sandstone in western Kentucky (Pryor and Van Wie, 1971)? (7) Were the commonly uncrushed bean-shaped chamositic ooids and the crushed spastotiths made in different ways, or were both kinds derived from a single initial type? (8) Were all the ooids produced by mechanical accretion, or was in situ growth an important process? What was the rate of accumulation of each of the sheaths? (9) How does the relative role of fecal pellets and ooids, and of primary precipitation and diagenetic replacement in these minerals compare with that in phosphorites? Facies models

The coarsening- and shoaling-upward models synthesize data recording development of glauconitic and chamositic granules in marine environments during long intervals of minimal aggradation. Commonly, they accumulated in burrowed condensed sequences associated with hardgrounds. This reconstruction, as all facies models, is a generalization from many examples, each with its own combination of specific features. Improvement of these models can best be achieved through detailed analysis and reconstruction of many additional examples. This will reveal differences in local environments that favored development of one or the other of these clay minerals and their precursors. It will also identify differences in the position of chamositic ironstones and glauconitic greensands relative to the prograding phase and the incipient transgressive phase of vertical sequences, and to associated phosphatic deposits. Rigorously constrained reconstructions will aid in establishing whether a particular vertical sequence resulted from local autocyclic (intrabasinal) control, such as the growth and abandonment of lobes of a low-energy delta or the migration of tidal sandwaves across an associated muddy facies, or from regional to global allocyclic (extrabasinal) control of sea level. Moreover, this may involve deposition by an autocyclic mechanism within a pattern of regional transgression and regression.

233

Regional environments of accumulation Berthierine and glauconitic peloids accumulating on the sea floor today quite consistently occur on the inner shelf and the middle to outer shelf, respectively. As a result, depth of water has commonly been cited as a major controlling factor. Ancient glauconitic peloids and chamositic ooids, on the contrary, occur in a broader range of shallow marine environments, and locally are found together in a single ancient deposit. In general, chamositic granules were limited to warm, shallow marine environments. Glauconitic peloids ranged more widely, from nearshore to deeper offshore realms, normally favored by cooler climate or upwelling water (Odin and Matter, 1981, fig. 1D), and containing a decreasing amount of aluminum seaward. Clearly, the role of distance from shore and general depth of water needs further evaluation. The apparently dissimilar modern and ancient records emphasize the importance of accurate environmental reconstruction of each deposit. They also raise the general question - - why is the present distribution of glauconitic and chamositic peloids limited so consistently to particular bathymetric realms? To what extent is it the results of incomplete information, or the effect of unique Holocene marine conditions? The anomalous presence of chamositic ooids in alluvial Cenozoic deposits of southwestern Russia raises additional questions about the role of environmental control, and requires a detailed analysis of this mode of origin and its associated microfacies. Glauconitic greensands commonly developed in middle to outer shelf environments. Nevertheless, some formed nearshore and others possibly in the deep sea (Odin and L6tolle, 1980; Odin and Stephan, 1981). Variation in the environmental distribution as well as the composition of glauconitic minerals is in part the result of the duration and rate of glauconitization. The range of congenial marine environments is illustrated by Cretaceous greensands in Great Britain (Jeans et al., 1982) and in Russia (Muravev, 1976; Sorokin et al., 1980). These examples have been interpreted as indicating that glauconitic minerals occur in relatively local concentrations with little relation to a consistent position in the regional lithofacies pattern. The authors conclude that the favorable condition, such as volcanism, prevailed periodically, influencing the entire basin. In Great Britain the glauconitic minerals developed in particles of lava and ash from submarine basaltic volcanoes, as well as in more acidic ash from subaerial volcanoes. A volcanic substrate undoubtedly has fostered some glauconitic deposits, but it does not account for the widespread authigenic peloids today or for many ancient greensands with no associated volcanism. The fact that glauconitic and chamositic granules occur together at least locally in ancient deposits and that those associated are mainly peloids,

234 suggests that development of chamositic ooids required a special condition within the general environment in which both glauconitic and chamositic minerals accumulated. Presumably this included gentle wave and current action and abundant micro-granules for cores in a shallow marine realm devoid of destructive benthos.

Geologic record Two major episodes of accumulation of glauconitic and chamositic granules during Phanerozoic time were characterized by mild maritime climate in many places, dispersed cratonic blocks and open ocean gateways, as well as widespread transgression of cratons, accompanied by locally decreased influx of sediment into shallow seas. During these favorable episodes development of glauconitic peloids was widespread but most of the numerous deposits comprise only scattered granules; relatively few are thick greensands. Most of the less common chamositic deposits, in contrast, consist of one or several successive thick ironstones, suggesting that once the favorable conditions producing chamositic ooids were established they normally were very effective and persisted for a long time. Breakup of the Proterozoic supercontinent about 600 Ma. B.P. was limited to dispersal of Laurasian cratonic blocks. The earliest transgression of the several larger ones during the Cambrian Period produced widespread glauconitic deposits. About 50 Ma. after initial dispersal and early transgression, chamositic deposits began to develop, and became widespread on smaller Laurasian blocks during the Late Cambrian-Early Ordovician major high stand of sea level. Breakup of Pangea began in Late Triassic time, about 200 Ma. B.P., with the separation of Laurasia and Gondwana. Dispersal of Gondwana blocks followed in later Mesozoic time. Chamositic deposits developed widely during the early stages of this episode of dispersal, mainly on Laurasia and Gondwana flanking the opening Tethys. About 50 Ma. later, glauconitic deposits reached their maximum development, with abundant greensands accumulating along continental margins of the opening Atlantic basins. The geologic record reveals two additional motifs, well-documented by chamositic ironstones, which merit detailed investigation. Ironstones developed repeatedly in a few regions, as in northern Europe and northern Africa during Early and Middle Paleozoic time, and in northwestern Europe during the Jurassic Period. Detailed analysis of these successive ironstones should reveal useful information about persisting favorable environments and the role of their tectonic framework. Borchert (1965) recognized a second motif in the recurring association of ironstones and black shales with episodes of warm climate and high sea level.

235 Throughout Phanerozoic time repeated 30-40 Ma. intervals of transgression and postulated oceanic anoxia (Fischer and Arthur, 1977) were conducive to major development of glauconitic and chamositic minerals only during the two 170 Ma. congenial periods. Available data suggest, however, that the spread of anoxic water across shallow-shelves was much more effective in inducing production of chamositic ooids than of glauconitic peloids. Development of ooids was not limited to these anoxic events, however; some combination of local conditions produced similar results during several of the alternate oxic episodes as well. Throughout this review reference has been made to the association of chamositic, glauconitic, and phosphatic minerals and granules, both in environments of accumulation and in Phanerozoic tectonic frameworks. Many of the conditions of origin and concentration, as in hardgrounds (for example, Jarvis, 1980), as well as facies associations, are shared (Cooke and McElhinney, 1979; Notholt, 1980; Odin and L6tolle, 1980). In detrital sequences chamositic and glauconitic deposits generally have only a minor associated phosphatic component. In carbonate-rich sequences, in contrast, the chamositic minerals wane, glauconitic ones persist, and phosphorite increases. During Phanerozoic time there was a very generally similar pattern of bloom, but in detail the patterns differ in significant ways (Cook and McElhinney, 1979, fig. 1,5; Sheldon, 1980). The geologic record reviewed here raises intriguing questions: (1) Why did many of the Paleozoic chamositic oolites accumulate on small dispersed cratonic blocks, whereas most of the Paleozoic glauconitic deposits accumulated on large cratons? (2) Why was glauconitic development tied so closely to the major early Paleozoic and Late Cretaceous transgressions whereas production of chamositic ooids flourished much more often during Phanerozoic time? (3) Why did chamositic deposits develop so abundantly during the Early Jurassic (Liassic) transgression of Tethys without associated abundant development of glauconitic deposits? (4) Why did glauconitic greensands flourish in the Paleocene Epoch while chamositic ooids failed to develop? Has this any implication for Cretaceous-Cenozoic boundary events? (5) What combination of factors favored accumulation of the vast Middle Pliocene Kerch chamositic ironstone in the Caucasian foredeep during increasingly inhospitable Late Cenozoic climate? (6) Why were glauconitic peloids and chamositic ooids so consistently incompatible throughout Phanerozoic time, if they apparently were capable of developing together in at least a few deposits? (7) What was the genetic relation between chamositic and glauconitic deposits and phosphorites? What was the role of cyclic and episodic control in their development?

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