The composition of the major rock-forming constituents of siliciclastic sediments (including mudrocks) is not directly diagnostic of environment. Krynine (1942) suggested that many sandstone types were characteristic of particular tectonic-sedimentary environments, for example arkoses
Fig. 3.11 The textural spectrum in limestones and the two basic classifications used for limestones (Dunham1962; Folk1962)
(feldspathic sandstones) supposedly represent nonmarine sediments derived from granitic orogenic complexes and greywackes (lithic sandstones) represent early “geosyncli- nal” sedimentation. These interpretations have been dis- carded, although Dickinson and Suczek (1979) and Schwab (1981) have shown that sandstone composition may closely reflect the plate-tectonic setting of the basin. However, this does not necessarily translate into depositional environment. For example, volcaniclastic forearc sediments may be deposited in fluvial, lacustrine, marginal-marine, shelf, or submarine-fan settings, depending on the continental-margin configuration. Dalrymple (2010b) provided an updated summary of these ideas.
Certain minor components of sandstones may be strongly suggestive of the depositional environment. For example, glauconite pellets form only in shallow marine environments (Odin and Matter 1981) and are rare as detrital or resedi- mented grains. Carbonaceous debris from plants is typical of nonmarine environments. Paleosol development is indicated by lenses of calcium carbonate (caliche, calcrete) and other minerals. Abundant red iron staining indicates oxygenated environments, typically either the preservation of oxidized states in detrital particles (van Houten 1973) or the pro- duction of oxidized colors during early diagenesis (Walker
1967). Red beds are therefore mostly indicative of non- marine or high intertidal environments (Turner 1980), although there are exceptions (e.g., Franke and Paul1980). Within a given basin, detrital composition may reflect variations in source area or depositional environment. These data may therefore be used as a paleocurrent indicator (Sect.6.5.1) and may assist in the stratigraphic correlation of units formed under the same hydraulic conditions. Davies and Ethridge (1975) showed that detrital composition varied between fluvial, deltaic, beach, and shallow marine envi- ronments on the Gulf Coast and in various ancient rock units as a result of hydraulic sorting and winnowing processes and chemical destruction. Thus, although composition is not environmentally diagnostic, it may be useful in extending interpretations from areas of good outcrop or core control into areas where only well cuttings are available.
For carbonate sediments, in contrast to siliciclastics, petrographic composition is one of the most powerful environmental indicators (James et al. 2010). Carbonate facies studies, therefore, require routine thin-section petro- graphic analysis, nowadays supplemented by the use of the scanning electron microscope and cathodoluminescence (Wilson 1975; Scholle 1978; Flügel 1982; Machel 1985). Most carbonate grains are organic in origin, including micrite mud derived from organic decay or mechanical attrition, sand-sized and larger particles consisting of organic fragments, fecal pellets, grapestone, and ooliths, all of which are produced in part by organic cementation processes, and boundstones or biolithites, formed by framework-building
organisms. Most carbonate particles are autochthonous; therefore, an examination of the composition of a carbonate sediment is of crucial importance.
These and other differences between carbonate and sili- ciclastic sediments are summarized in Fig. 3.12. Textural classifications of carbonate sediments were discussed in the previous section (Fig. 3.11) and have been expanded to include compositional details of framework-building organ- isms by Embry and Klovan (1971) and Cuffey (1985). Interpretation of ancient carbonate sediments is complicated by the fact that the organisms that generate carbonate par- ticles have changed with time (James et al.2010, Fig. 12). However, there are many similarities in form and behavior between modern and extinct groups, so that actualistic modeling can usually be carried out with caution (see also Sect. 4.3). Another problem is that diagenetic change is almost ubiquitous in carbonate rocks and may obscure pri- mary petrographic features (this is briefly discussed in Sect. 4.3.1). Ginsburg and Schroeder (1973) showed that some carbonates are converted contemporaneously from reef boundstones or grainstones to wackestones by continual boring, followed by infill of fine-grained sediment and cement. Mountjoy (1980) suggested that many carbonate mud mounds may owe their texture to this process.
Most environmental interpretation of carbonates is based on thin-section examination using a microfacies description system such as that erected by Wilson (1975; Fig.3.4of this book). Wilson was able to define a set of standard facies belts for the subenvironments of a carbonate platform and slope (Fig. 3.13), each characterized by a limited suite of
Fig. 3.12 The differences between carbonate and clastic sediments. Adapted from James (1984a)
microfacies reflecting the variations in water depth, water movement, oxygenation, and light penetration. A classic example of a carbonate petrology study, the analysis of modern sediments of the Bahama Platform, is described briefly in Sect. 4.3 (see Figs. 3.7 and 3.8). Each facies assemblage can readily be related to environmental vari- ables, such as the quiet-water pelletoidal facies and the high-energy skeletal sand and oolite lithofacies.
Many carbonates consist of dolomite rather than lime- stone. In most cases, the dolomite is clearly a replacement, as indicated by the presence of dolomite rhombs penetrating allochemical particles such as shell fragments or ooliths. In other cases, there may be evidence for a primary or penecontemporaneous origin, which may be environmentally useful information. Primary dolomite crusts form in associ- ation with evaporites in areas where seawater is evaporated at high rates, such as on supratidal flats and in shallow tidal lagoons. Modern examples include parts of Bonaire, sabkhas on the Persian Gulf, Deep Springs Lake, California, and Coorong Lagoon, Australia (Blatt et al.1980, pp. 512–522; Friedman1980). The dolomite-evaporite association, toge- ther with the evidence of certain evaporite or desiccation textures (Sect.3.5.7) is strongly environmentally diagnostic. The dolomitization processes depends on “seepage reflux- ion” or “evaporative pumping” of sea water to the surface under hot, arid conditions (Adams and Rhodes 1960). However, this is only one mechanism for the production of dolomite. Most dolomite probably forms at depth by a dia- genetic mixing of sea water and fresh water, a process known as the Dorag model (Badiozamani1973; Land1973). In rare cases dolomite can also occur as detrital grains.
The composition of evaporite minerals is not a good guide to the depositional environment of evaporites. A sample of normal seawater if evaporated to dryness yields a sequence of precipitates in the following order: calcite,
gypsum, halite, epsomite, sylvite, and bischofite. However, the composition of the final deposit in nature may vary considerably because of the effects of temperature, the availability of earlier formed components for later reaction, and the rate and nature of replenishment of the water supply. Evaporites may form in a variety of marine and nonmarine environments (Schreiber et al. 1976; Schreiber 1981; Ken- dall2010), and it is not their composition as much as their internal structure and lithologic associations that are the best clues to the depositional environment (e.g. the association with penecontemporaneous dolomite mentioned previously). The formation of many chemical sediments, such as chalk, chert, phosphates, and glauconites depends on factors such as organic activity and ocean water oxygenation and temperature (Gorsline1984). These factors are controlled in part by large-scale oceanic circulation patterns, which are being studied in order to develop predictive models of facies development (Parrish1983).
Certain other chemical deposits contain useful environ- mental information. Chert is common as a replacement mineral in carbonate sediments, where it forms nodules and bedded layers commonly containing replacement casts of fossils, ooliths, etc. Knauth (1979) suggested that such chert was formed by the mixing of fresh and marine waters in a shallow subsurface, marginal-marine setting. Chert also occurs in abyssal oceanic sediments in association with mafic and ultramafic igneous rocks. Radiolarians, sponge spicules, and diatoms are common. These are some of the typical components of ophiolites, which are remnants of oceanic crust and indicate a former deep water environment (Grunau1965; Barrett1982).
Iron-rich rocks occur in a variety of settings. Their chemistry is controlled partly by Eh and pH conditions. Pyrite and siderite are common as early-diagenetic crystals and nodules in the reduced environment of organic-rich Fig. 3.13 The standard facies belts of Wilson (1975), as modified by Schlager (2005, Fig. 4.3)
muds, particularly in fluvial or coastal (deltaic, lagoonal) swamps. Occasionally, such deposits may be present in rock-forming abundance, as in the pisolitic bog iron ores. Plant remains, impressions, and replaced (petrified) wood are commonly associated with these forms of iron. Pyrite is also associated with unoxidized, disseminated organic par- ticles in the black muds of anoxic lake and ocean basins.
Organic-carbon-rich black shales are of considerable importance as petroleum source beds. Arthur et al. (1984) reviewed their origins and significance in terms of oceanic organic productivity, sedimentation rate, and oxygenation of ocean bottom waters (Fig.3.14). Ultimate controls on sedi- mentation are plate-tectonic configurations, which determine
global sea level and climate, and which control the nature of oceanic circulation patterns. Such shales cannot, therefore, be interpreted strictly on a local basis, but must be consid- ered in a regional or even a global context.
The increasing importance of shale-gas as an energy source has focused considerable attention on the facies analysis and petrology of mudrocks. Contrary to common assumptions, most mudrocks contain significant proportions of non-mud components, particularly quartz and calcium carbonate, with clay minerals comprising less than 50 % by volume, based on thin-section point counts (Shaw and Weaver 1965; McQuaker and Adams 2003; Fig. 3.15). When examined in detail, shales are not homogeneous, but
Fig. 3.14 Oxygen content at the sediment-water interface and generalized relationships between types of benthic organisms, sedimentary structures, water chemistry, mineralogy, organic content, and sediment color. Notes: 1*, certain polychaete organisms and foraminifera may inhabit surficial environments where benthic metazoans are excluded; 2*, degree of lamination depends in part on seasonal variations in clastic input and productivity. (Arthur et al.
1984, reproduced by permission of the Geological Society, London)
vary significantly in petrographic composition and internal structures, which, in turn, significantly affects their reservoir properties. It is increasingly being recognized that mud deposits are in many cases not simply passive pelagic accumulations in quiet, deep-water environments, but are the deposits of mudfloccules transported by traction currents or sediment gravityflows, and commonly exhibiting bedforms comparable to those formed in sand-bed deposits (Plint
2010; Shieber et al. 2013).
Hematite and chamosite iron ores are locally important in the Phanerozoic. Most are oolitic and display typical shallow-water sedimentary structures. As van Houten (1985) noted, these ironstones are particularly abundant in the Ordovician and Jurassic record, but display widely varying paleogeographic and paleoclimatic settings, and so their use as facies indicators is uncertain at present. The iron is probably an early diagenetic replacement. Precambrian iron formations are widespread in rocks between about 1.8 and 2.6 Ga. Their mineralogy is varied and unusual (Eichler
1976); it may reflect formation in an anoxic environment (Cloud 1973; Pufhal2010).
Manganese- and phosphate-rich rocks are locally impor- tant. Both commonly occur as crusts and replacements on disconformity and hardground surfaces, where they are taken as indicators of nondeposition or very slow sedimen- tation. Phosphates are particularly common on continental margins in regions of upwelling oceanic currents. More detailed discussions of these sediments are given by Blatt et al. (1980) and Pufhal (2010).
Coal always indicates subaerial swamp conditions, usu- ally on a delta plain, river floodplain or raised swamp. However, lacustrine coals and coal formed in barrier-lagoon settings have also been described. Calcrete (or caliche) occurs in alluvial and coastal environments and is an
excellent indicator of subaerial exposure (Bown and Kraus
1981; James1972; Wright1986a).