A Color Guide to the Petrography of Carbonate Rocks:

Grains, textures, porosity, diagenesis

By Peter A. Scholle and Dana S. Ulmer-Scholle


This volume expands and improves the AAPG 1978 classic, A Color Illustrated Guide to Carbonate Rock Constituents, Textures, Cements, and Porosities(AAPG Memoir 27). Carbonate petrography can be quite complicated. Changing assemblages of organisms through time, coupled with the randomness of thin-section cuts through complex shell forms, add to the difficulty of identifying skeletal grains. Furthermore, because many primary carbonate grains are composed of unstable minerals (especially aragonite and high-Mg calcite), diagenetic alteration commonly is quite extensive in carbonate rocks. The variability of inorganic and biogenic carbonate mineralogy through time, however, complicates prediction of patterns of diagenetic alteration. This book is designed to help deal with such challenges. It includes a wide variety of examples of commonly encountered skeletal and nonskeletal grains, cements, fabrics, and porosity types. It includes extensive new tables of age distributions, mineralogy, morphologic characteristics, environmental implications and keys to grain identification. It also encompasses a number of noncarbonate grains, that occur as accessory minerals in carbonate rocks or that may provide important biostratigraphic or paleoenvironmental information in carbonate strata. With this guide, students and other workers with little formal petrographic training should be able to examine thin sections or acetate peels under the microscope and interpret the main rock constituents and their depositional and diagenetic history.

  1. Page viii

    Carbonate petrography — the study of limestones, dolomites and associated deposits under optical or electron microscopes —greatly enhances field studies or core observations and can provide a frame of reference for geochemical studies. Petrography is an especially powerful tool because it enables the identification of constituent grains, the detailed classification of sediments and rocks, the interpretation of environments of deposition, and the determination of the often complex history of post-depositional alteration (diagenesis). The last of these, the ability to determine the timing of diagenetic events such as cementation or secondary porosity development relative to the emplacement of hydrocarbons or metallic ores, makes petrography an important component of geochemical and sedimentologic studies in energy- and mineral-resource exploration applications as well as in academic research.

    The petrographic study of carbonate rocks is particularly useful because carbonate grains, unlike clastic terrigenous ones, normally are produced in close proximity (from less than a meter to hundreds of meters) to the site of their ultimate deposition. In addition, carbonate grains are formed mainly by organisms, and thus the grains convey ecological information about the environment of formation as well as stratigraphical information on the age of the deposit.

    In some ways, carbonate petrography is not a very complex undertaking, especially when compared to the petrography of clastic terrigenous deposits. Most carbonate rocks are dominated by just one or two common carbonate minerals (mainly calcite and dolomite) plus a limited number of accompanying minerals — silica, detrital grains, phosphate, glauconite, and a few evaporite precipitates. The diagram below shows the general compositions of the full spectrum of carbonate minerals found in modern and ancient strata.

    1. Page 1

      Cyanobacterial stromatolites usually are grouped in the Phylum Cyanophyta — Precambrian (Archean)-Recent Classification of other microbes is complex, uncertain, and ever changing (generally placed under the Prokaryotes, but most of these organisms are really best considered as “microproblematica”). Organisms once termed blue-green algae are now generally termed cyanobacteria.

      Ranges of some specific calcimicrobes depicted in this section:

      • Girvanella— Cambrian-mid. Cretaceous (Eocene?)

      • Epiphyton— Cambrian-Devonian

      • Renalcis — Cambrian-Devonian

      • Frutexites — Latest Cambrian-Devonian

      Many are photosynthetic and therefore require light; non-photosynthetic microbes also are important, especially in cryptic settings. Recognition of photosynthetic forms is especially critical in paleoenvironmental studies.

      Wide salinity tolerance from strongly hypersaline to freshwater; rare as dominant sediment formers in modern, normal-salinity marine environments.

      Wide temperature tolerance: sub-glacial to hot springs settings; most common in temperate- to warm-water marine settings.

      Marine stromatolites range from subtidal to intertidal settings — intertidal forms predominate today.

      A progressive shift occurred from normal-salinity environments in the Precambrian to highly stressed environments today, possibly due to the Phanerozoic increase in grazing organisms or interspecific competition. Cenozoic microbial carbonate deposits are predominantly peritidal.

      Marine forms are mainly aragonitic; incorporated detrital components can have any carbonate or terrigenous mineralogy; lacustrine forms are mostly calcitic.

      Most are uncalcified and the remainder have “nonskeletal” or “extraskeletal” calcification largely incidental to their growth. Calcification results from biochemical processes (removal of CO2), but generally is not necessary for, or beneficial to, the organism’s survival.

      Calcimicrobial deposits, thus, have no clearly defined and consistent skeletal morphologies (hence the difficulty of classifying these microproblematica). Calcimicrobial deposits are recognized by overall sediment structure, by externally calcified filaments or spherical bodies, and by trapped sediment. Flat-lying laminated sediment, domal stromatolites, or clotted, finger-like thrombolite structures are characteristic — shapes vary with environmental conditions (water depth, current strength, and others).

      Lamination in stromatolites reflects microbial growth through day-night cycles and tidal cycles; those organic laminae commonly are interspersed with micritic or peloidal carbonate or terrigenous detritus that was deposited during episodic storms.

      Non-stromatolitic calcimicrobes typically form lumpy encrustations or small upright “shrubs”.

    2. Page 33

      Kingdom Protista, Phylum Sarcomastigophora, Subphylum Sarcodina, Superclass Rhizopoda, Class Granuloreticulosea, Order Foraminiferida — Basal Cambrian-Recent

      • Benthic foraminifers: Cambrian-Recent (early forms were exclusively agglutinating)

      • Calcareous benthic foraminifers — Ordovician-Recent; large forms from Late Carboniferous-Recent

      • Planktic foraminifers: Middle Jurassic-Recent

      Despite being single-celled protozoans, this is a very complex group of organisms, with 12 suborders recognized by Loeblich and Tappan (1984) and some 60-80,000 species identified from Phanerozoic strata. So many shape, size, and wall-structure varieties exist, however, that this chapter can provide only the minimal information needed to identify the most important groups.

      Modern foraminifers are fully marine to marginal marine organisms, extending from the intertidal zone to abyssal oceanic depths and from cold-water polar settings to warm tropical environments. Some genera live in marginalmarine hypersaline or subsaline water bodies where they are commonly found in great numbers (but low species diversity).

      Most foraminifers are benthic organisms (of the roughly 4,000 modern species, only about 40 are planktic).

      Some of the largest living benthic species harbor symbiotic algae in their tissues and thus live primarily in the photic zone; the vast majority, however, are not light dependent.

      For reasons related mainly to food supply, most planktic foraminifers live in the upper 300 m of the water column, although after death, their tests fall to the underlying, deeper seafloor.

      Foraminifers can be major rock forming elements in open- or restricted-shelf as well as deeper marine deposits. In some cases, foraminiferal abundances reach tens of thousands of individuals per m3 of sediment.

    3. Page 51

      Often grouped with the tintinnids (pelagic ciliate protozoans of the subclass Spirotheca), although modern tintinnids are organic-walled and calpionellids had calcareous walls. Thus, calpionellids are grouped by other workers as Protozoa incertae sedis.

      Calpionellids — Late Jurassic (Tithonian) to Early Cretaceous (Valanginian; possibly into Albian)

      Tintinnids — Jurassic-Recent (but with possible record extending into the Paleozoic, perhaps even to the Cambrian)

      These open marine organisms are significant contributors to pelagic limestones and chalks in the Late Jurassic. Their distribution is largely restricted to the warm-water Tethyan region, within about 30-35° of the paleoequator.

      All calpionellids apparently were composed of low-magnesium calcite; thus, generally well preserved. The TEM studies conducted by Fischer et al. (1967; cited at end of book’s introduction) showed that some calpionellids built two-layered tests in which the main layer incorporated carbonate detritus (including coccoliths) and was lined by an inner, secreted layer.

      Small size (typically 45 to 150 μm in length and 30 to 90 μm in width), spherical to elongate, U- or V-shaped grains with a large opening rimmed, in some cases, by a narrowed, slightly thickened collar.

    4. Page 75

      Worm remains are known from Precambrian to Recent — most are soft-bodied, but preserved fossil forms include some segmented worms that built solid housing structures. These generally belong to the:

      • Phylum Annelida: Proterozoic-Recent

        • Class Polychaeta: (Proterozoic?) Cambrian-Recent

      The most important sediment-producing or sediment-influencing groups in Phylum Annelida include three groups or families within the order Sabellida:

      • Serpulids and spirorbids (groups that precipitate solid calcareous tubes)

      • Sabellariids (producers of agglutinated tubes)

      • A variety of soft-bodied burrowers and pellet producers

      Most preserved forms lived in fully marine to hypersaline-water settings; rare in freshwater and even rarer in terrestrial settings (although non-calcified forms can produce pellets in those environments).

      Serpulids are most common in shallow to coastal waters (largely as hard-substrate encrusters) but extend into deeper shelf waters as well. Especially common in slightly hypersaline settings (where they may form small reef-like masses) or at hiatus surfaces.

    5. Page 83

      This short-lived but widespread group have been classed by various as sponges, corals, or calcareous algae. Archaeocyaths now are almost unversally considered as subphylum of the Porifera (possibly related to the demosponges); a few workers still group them in a separate phylum (Phylum Archaeocyatha). They range mainly from basal Cambrian to late Early Cambrian (a few forms persisted to Middle and Late Cambrian).

      One of the earliest groups to secrete substantial skeletal calcium carbonate and the first reef-building organism.

      Sessile, benthic, filter feeders. Exclusively marine organisms that lived in tropical, normal salinity (ca. 30-40 ppt) waters at depths from the intertidal zone to a few tens of meters, mainly in areas with relatively low influx of terrigenous sediments (see Debrenne and Reitner, 2001).

      Constructed small bioherms in association with calcimicrobes. They also are found in lesser abundance, size and diversity in inter-bioherm areas.

      The good preservation of most archaeocyath skeletal material indicates a primary calcitic composition.

      Most archaeocyaths have a solitary cup- or bowl-shaped skeleton that has a pair of porous walls enclosing a large central cavity. The inner and outer walls have a series of spherical perforations and are connected by numerous perforate or imperforate partitions (vertical septa and horizontal tabulae).

      Less commonly, archaeocyaths had branched, massive, or chain-like colonial forms.

      The average size of archaeocyath cups is 1 to 2.5 cm in diameter and 15 cm in height. Cups as small as 2-3 mm or as large as 60 cm in diameter are known, however.

      Closely associated with calcimicrobial encrusters such as Renalcis and/or Epiphyton, and external morphology is commonly outlined by such encrusters.

    6. Page 101

      Phylum Cnidaria, Class Anthozoa

      • Subclass Zoantharia — Ordovician (Cambrian ?)- Recent

        • Order Tabulata — Early Ordovician-Late Permian

      • Possible tabulate corals have been reported from the Early Cambrian (Sorauf and Savarese, 1995). The group was widespread and diverse from Late Ordovician to Middle Devonian, but declined in Late Devonian and into the Carboniferous. The group became extinct during the great end-Permian faunal crisis.

      Tabulate corals were fully marine, sessile organisms and were contributors to stromatoporoid and microbial reefs and bioherms of Ordovician to Carboniferous age. Although substantial contributors to some reefs, tabulates rarely were dominant reef framework formers. Tabulate corals did build smaller, isolated bioherms that are widely distributed in muddy, open shelf carbonate rocks.

      Many tabulate corals were attached to their substrates, others were unattached, rolling free on the sea floor.

      Whether tabulate corals had a symbiotic relationship with zooxanthellate analogs, and therefore were restricted to living in the photic zone, is an open question because such soft tissues are not preserved. Environmental reconstructions, however, indicate that tabulates lived at shallow marine depths within the photic zone, so an analogous symbiotic relationship is possible, but has neither been proven nor refuted.

      Virtually all tabulate corals probably were originally calcitic; very few (probably on the Tetradiidae, a group that is not universally classed with the tabulates) may have been aragonitic. Determination of original mineralogy is based mainly on the quality of structural preservation.

    7. Page 123

      Phylum Bryozoa

      • Subphylum Entoprocta — Middle Cambrian?, Late Jurassic-Recent

      • Subphylum Ectoprocta

        • Class Phylactolaemata — Middle Jurassic-Recent

        • Class Gymnolaemata — Early Ordovician-Recent (dominates Mesozoic-Recent)

          • Order Ctenostomida — Early Ordovician-Recent

          • Order Cheilostomida — Late Jurassic-Recent

        • Class Stenolaemata — Early Ordovician-Recent (dominates Ordovician-Permian)

          • Order Cyclostomida — Early Ordovician-Recent

          • Order Cystoporida — Early Ordovician-Late Triassic

          • Order Trepostomida — Early Ordovician-Late Triassic

          • Order Cryptostomida — Early Ordovician-Late Triassic

      Bryozoans are sessile, filter-feeding organisms with a wide salinity tolerance — most are marine, but a few species (from the Entoprocta, Phylactolaemata, and Ctenostomida) inhabit fresh water and a few others (from the Cheilostomida) are found in brackish-water environments.

      Bryozoans have wide latitudinal (tropical to polar), temperature, and depth ranges (0 to 8.5 km). They can be the main constituents in Mesozoic and Cenozoic temperate- and cold-water shelf carbonates, as well as in deeper shelf and slope settings; in the Paleozoic, they were more conspicuous in tropical to subtropical habitats.

      Many bryozoans require a firm substrate on which to encrust; some are free living, and others have roots extending into sandy substrates. Massive and encrusting varieties are found in high-energy environments; delicate, erect varieties are indicative of low-energy environments.

      Entoproct bryozoans are soft bodied and, therefore, are rarely preserved.

      Most ectoproct bryozoan zooecial walls are composed of calcite (usually low-Mg calcite; a few consist of high-Mg calcite and others are partially aragonitic). Some species have chitinous or gelatinous walls.

    8. Page 141

      Phylum Brachiopoda — earliest Cambrian-Recent

      • Subphylum Linguliformea (Early Cambrian-Recent): shells lack skeletal articulation structures; shells are chitinophosphatic with laminar microstructure; pedicle usually present, emerging between valves or from opening in ventral valve.

        • Class Lingulata (Early Cambrian-Recent): brachiopods with chitinophosphatic shells lacking teeth and sockets; pedicle usually present emerging from shell between valves or from apex of one of the valves.

        • Class Paterinata (Early Cambrian-Late Ordovician): shell rounded to elliptical with straight posterior margin with pseudointerarea; delthyrium often closed by plates; pedicle reduced or absent.

      • Subphylum Craniiformea (Early Cambrian-Recent): calcareous shells; valves lack hinge teeth and sockets; shell usually attached to substrate by cementation of pedicle (ventral) valve.

        • Class Craniata (Mid. Cambrian-Recent): features as above for subphylum.

      • Subphylum Rhynchonelliformea (Early Cambrian-Recent): Brachiopods with calcitic shells that have endopunctate, impunctate, pseudopunctate, or tabular microstructure; crura usually present extended to form a brachidium (spiralia or loops) in some groups; articulated valves with hinge teeth and sockets are the norm, but in some forms, reduced or modified types of articulation structures are present; the vast majority of known rhynchonelliform brachiopods are included in the classes Strophomenata and Rhynchonellata.

        • Class Chileata (Early Cambrian only): short-lived early group, see features in Clarkson (1998, p. 181).

        • Class Obolellata (Early-Mid. Cambrian): short-lived early group, see features in Clarkson (1998).

        • Class Kutorginida (Early-Mid. Cambrian): short-lived early group, see features in Clarkson (1998).

        • Class Strophomenata (Mid. Cambrian-Triassic): Shell usually concavo-convex or planoconvex; shell usually pseudopunctate; straight hinge with simple teeth (often lost); some groups with spines, pedicle opening usually closed by plate(s). Includes the Orders: Strophomenida (six suborders), and Productida (two suborders).

        • Class Rhynchonellata (Early Cambrian-Recent): Biconvex shells with both strophic and nonstrophic hinges; impunctate and punctate shells; crura usually present; brachidium often present. Includes the orders: Orthida (shell usually impunctate); Rhynchonellida; Pentamerida; Athyrida (spiralia present, usually impunctate); Atrypida (spiralia present, impunctate); Spiriferida (spiralia present, punctate and impunctate shells); Spiriferinida (spiralia present, impunctate and punctate shells), and Terebratulida (loop present, punctate shell).

      In general, brachiopods were especially abundant in the Paleozoic where they reached their peak diversity in the Devonian. In many settings, they were among the main rock-forming organisms. Although they are much less abundant in Mesozoic and Cenozoic strata, they retain considerable biostratigraphic value in those deposits.

    9. Page 153

      Phylum Mollusca, Subphylum Cyrtosoma

      • Class Gastropoda — Early Cambrian-Recent

        • Order Thecosomata (pteropods) — Cretaceous-Recent (possible precursors Cambrian?-Permian?)

      Gastropods are the largest class of both living and fossil mollusks (with nearly 8,000 genera), although they are rarely major rock-forming organisms.

      Gastropods (snails) are a remarkably wide-ranging group of organisms. They are found at all latitudes and in normal marine, brackish, hypersaline, and fresh water as well as subaerial environments. They rarely are major sediment formers, however, except in stressed (especially hypersaline or freshwater) settings.

      Warm-water forms generally are thicker shelled than cold-water forms.

      Pteropods are open-marine, predominantly warm-water, nektic organisms that contribute mainly to deep-sea oozes on seafloors shallower than about 3,000 m (because of aragonite dissolution effects).

      Gastropod shells have a thin outer coating of organic material (conchiolin) plus a thick carbonate layer generally consisting of only aragonite. Some families, however, have shells with separate layers of calcite and aragonite. Where present, the calcite layer normally is thicker than the aragonite layer. Gastropod calcite has a low Mg content (typically less than 0.3 mole% Mg; rarely exceeding 1 mole% Mg). Pteropods have aragonite shells.

      Both shell-bearing and non-shell-bearing gastropods exist. The shelled forms are univalves that have an unchambered cone, most commonly coiled about a central axis. Some forms are able to withdraw fully into their shell and have a plate (an operculum) that they can draw behind themselves to close the shell opening; opercula can be composed entirely of conchiolin (proteinaceous organic material that is rarely preserved) or aragonite.

      Diverse coiling patterns exist: high-spired, conical, and planispiral forms are common; some groups (such as the vermetids) have very open spirals and form shells that resemble serpulid worm tubes.

      Adult gastropods typically are about 2-3 cm in length (modern forms of up to 60 cm length are known, however). Fragments typically mm- to cm-sized.

      Pteropods are nektic gastropods and although the majority are shell-less, some have slender, conical, generally uncoiled, thin-walled shells, typically less than 1-2 cm in length.

    10. Page 177

      Phylum Echinodermata

      • Subphylum Echinozoa

        • Class Echinoidea — Late Ordovician-Recent

      Echinoids (sea urchins) live in normal marine environments because they with a very limited range of salinity tolerance (generally only a few ppm).

      They occur mainly as grazers or burrowers in sandy shelf areas or as grazers and bioeroders along rocky shorelines. They occur in deeper waters as well, extending to abyssal depths. Fossil forms are most common in normal marine, open shelf or platform deposits.

      Echinoids are common in both warm- and cold-water settings, although they rarely are major rock-forming organisms (i.e., they rarely exceed 10-15% of the total sediment).

      Modern and ancient echinoids are/were composed of moderate- to high-Mg calcite. Modern forms contain between 2 and 17 mole% Mg; the Mg content varies with generic group and increases with increasing water temperature (see Milliman, 1974, p. 130-134, for details and citations).

      Echinoids, like all echinoderms at some stage in their life cycle, show pentameral (five-fold) symmetry. They have heavily calcified, globular to discoidal, hollow, endoskeletal tests (coronas) that are composed of individual sutured, interlocking or imbricated calcite plates. The calcitic coronal plates are porous and sponge-like; echinoids with rapid growth rates have spongier plates (with more holes and less calcification) than slowgrowing counterparts. Thus, slow growing, cold-water forms can be more heavily calcified than those from warmer waters (Raup. 1958)

      In life, echinoid tests are covered with elongate, moveable spines (in some species extremely short, but in others longer than 8 cm). The spines normally detach after death and can themselves be significant sediment contributors.

      Generally, each plate of an echinoid behaves optically as a single, extensively perforated, calcite crystal (see comments below). Echinoid teeth, however, are polycrystalline.

    11. Page 193

      Phylum Arthropoda

      • Superclass Trilobitomorpha

        • Class Trilobita — Early Cambrian (late Proterozoic?)-Late Permian

      Most trilobites were mobile, benthic, detritus feeding, fully marine organisms with a limited salinity tolerance (they are not found in inferred salinity-stressed settings). A few groups of pelagic trilobites are known.

      Although most common in shallow shelf settings, trilobites, especially eyeless forms, are also found in deeperwater environments. They are major rock-forming elements mainly in shallow shelf deposits of Cambro-Ordovician age.

      Trilobite carapaces were composed of chitin with large amounts of calcium carbonate and variable amounts of calcium phosphate (up to 30% in some species). The carbonate consisted of calcite, probably with moderate to high Mg content.

      Trilobites were characterized by exoskeletal carapaces with three lobes that extended the length of the organism. Carapaces were divided into a head shield (cephalon), an abdominal section (thorax) with 2 to 40 segments (sclerites), and a tail shield (pygidium).

      The shields and segments were sharply recurved inwards along the margins of the organism. Carapaces were shed during growth stages (molting behavior) adding to the large numbers of trilobite grains in many sedimentary deposits.

      Adult trilobites ranged in length from 0.1 to 75 cm; they average about 5 cm in length and 1-3 cm in width.

      1. The segmented nature of the carapaces, coupled with trilobite molting behavior, means that these organisms are normally found as fragmentary remains. Individual segments typically are in the mm to cm length range and are less than a mm in thickness.

      2. The recurved margins of trilobite shields and the multidirectionally curved forms of thoracic segments (sclerites) yielded fragments that commonly have characteristic “hook” or “shepherd’s crook” shapes.

      3. Skeletal fragments have a homogeneous prismatic microstructure, with extremely fine (micrometer-scale) calcite prisms oriented perpendicular to the carapace surface. Typically, the wall appears smooth and uniform with no obvious crystals; trilobite fragments, however, show sweeping (undulose) extinction when rotated under cross-polarized light. Some trilobites may also have finely lamellar layers.

      4. Many specimens show small perforations (canaliculi) that trend perpendicular to the skeletal walls.

      5. Fine growth lines may be visible — they parallel the carapace surface but do not interrupt the continuity of calcite prisms

      6. Trilobite fragments can be visibly multilayered, with thin inner or outer layers over the main carapace wall. Outer layer can be organic rich with a dark coloration in transmitted light.

      7. Homogeneous prismatic wall structure (and consequent extinction behavior) of trilobites is similar to that shown by ostracodes and a few bivalves. Trilobite fragments, however, generally are larger than ostracodes and are more irregular in curvature than either ostracodes or bivalves.

    12. Page 207

      There are thousands of problematic organisms — organisms unassigned to a specific phyletic group, or ones that were assigned to different groups by different workers. We have simply picked a few that are particularly distinctive and/or that are important in rocks of hydrocarbon exploration interest. We list prior phyletic assignments and age ranges below and provide descriptions and keys to recognition in the figure captions.

      Receptaculitids - grouped with sponges, corals, dasycladacean green algae, or problematica — common from Early Ordovician to Late Devonian, with smaller, more globular forms extending into the Permian

      Nuia - grouped with problematic codiacean algae or as an unassigned organism — Late Cambrian-Ordovician

      Palaeoaplysina - grouped with sponges, phylloid algae, or hydrozoans — Mid. Pennsylvanian-Early Permian

      Tubiphytes - variously grouped with cyanobacteria/blue-green algae, red algae, calcareous sponges, foraminifers, hydrozoans — at least Late Carboniferous to Late Jurassic

      Lithocodium - grouped as codiacean algae or loftusiid foraminifers — Late Triassic to Early Cretaceous (Albian)

      Hensonella - grouped as mollusks (scaphopods), coralline red algae, or dasycladacean green algae — Cretaceous (Hauterivian-Albian)

      A cross section of the wall structure of Calathium sp., with its central cavity and moderately well-preserved radiating wall structure. Calathids are the earliest receptaculitids — they had ovoid or tubular skeletons that strongly resemble sponges (one of the groups in which receptaculitids commonly are classed). The sparry calcite-filled areas (and micrite-filled circles) are recrystallized, originally aragonitic, elongate pillars that constituted the skeletal wall (see Nitecki et al., 1999); the rest of the micritic sediment has filled areas of former void spaces or sites of later-decomposed organic tissues.

    13. Page 215

      Phylum Chordata, subphylum Vertebrata

      Vertebrates range from Cambrian to Recent (initially as jawless fish)

      • Jawed vertebrates — Early Silurian-Recent

      • Terrestrial vertebrates — (Late Devonian) Carboniferous-Recent

      • Reptiles — Carboniferous-Recent

      Early vertebrate remains (Cambrian-Silurian) are confined to marine settings; subsequent diversification led to expansion into virtually all environments from polar to tropical and from terrestrial to abyssal marine. Most vertebrate remains in carbonate rocks are fish scales and teeth from lacustrine or marine settings; bone and tooth material from other groups, however, also can be found on occasion.

      Vertebrate remains are rare but can sometimes be found concentrated by wave or current action or by nonsedimentation of other materials at hiatus surfaces.

      Bones and teeth are largely composed of organic proteins (mainly collagen) and calcium phosphate (carbonatehydroxylapatite, sometimes termed collophane when it is microcrystalline). The interior parts of bones (the cancellous portions) have a spongy texture that commonly is filled with precipitated cement (most commonly carbonate, silica or phosphate) during diagenesis.

      Vertebrate organisms have an enormous range of external morphologies, but all have a vertebral column and other hard parts (other bones, teeth, or scales) that typically disarticulate upon death and can become scattered into carbonate and noncarbonate sedimentary deposits.

      Vertebrate debris can range in size from less than a mm to well over 1 m, but is typically in the mm to cm size range.

      Most bones share common features: a dense, smooth, outer or cortical part, and an interior composed of multiple layers of porous or spongy, cancellous material (in life, the porous areas are occupied by marrow).

      Teeth are constructed of three layers: a pulpy cavity, a covering of dentine, and where the tooth is exposed, an additional covering of enamel.

      Dentine is relatively hard and dense and has a mineral content of about 75%; enamel is even denser and is almost 98% hydroxylapatite.

    1. Page 227

      Ooid (oolith) - a spherical to ellipsoidal grain, 0.25 to 2.00 mm in diameter, with a nucleus covered by one or more precipitated concentric coatings (cortical layers) with radial and/or concentric orientation of constituent crystals. Nuclei typically consist of detrital terrigenous grains, skeletal fragments, or pellets and peloids, and coatings can have a variety of compositions. A rock composed dominantly of ooids is termed an “oolite”. That term is commonly misused, however, to describe the constituent ooid grains.

      Spastolith or deformed ooid - An ooid or other coated grain that has been deformed, generally by shearing the concentric laminations away from each other or from the nucleus. In rarer cases, the deformation is tectonic.

      Superficial ooid - An ooid with an incomplete or very thin cortical coating; specifically one in which the thickness of the accretionary coating is less, commonly far less, than the radius of the nucleus.

      Pisoid - A small spheroidal particle with concentrically laminated internal structure, larger than 2 mm and (in some usages) less than 10 mm in diameter. A pisolite is a rock containing abundant pisoids.

      Oncoid - In North American usage, an oncoid is a coated grain of algal (but not red algal) or microbial origin that is coarser than 2 mm in diameter; a spheroidal form of microbial stromatolite showing a series of concentric (often irregular or scalloped) laminations. These unattached stromatolites are produced by mechanical turning or rolling, exposing new surfaces to microbial/algal growth. Common European usage is less genetic, and in that usage a microbial/algal origin is not a prerequisite. An oncolite is a rock composed of oncoids; the term, however, is often used improperly as a synonym for “oncoid”.

      Rhodoid (rhodolith) - An irregularly laminated calcareous nodule composed of encrusting coralline algae arranged in more or less concentric layers about a core; spheroidal but knobby surfaced, and up to several centimeters in diameter; form in warm to cool, clear, shallow sea water down to depths of 150-200 m.

      Calcareous ooids and pisoids are known from the Late Archean to Recent; specific coated grains, such as Girvanella oncoids or red algal nodules (rhodoids) are restricted by the age ranges of the constituent organisms (listed in chapters on organic grains).

      Modern calcareous ooids are known with aragonite or Mg-calcite compositions (or combinations of both), and there is evidence that these same compositions existed throughout Phanerozoic time, perhaps with specific temporal preferences (e.g., Sandberg, 1983; Wilkinson and Given, 1986). Laminae of organic material are found interlayered in most ooid cortices and help preserve structure during diagenesis.

      Calcareous cave and soil pisoids typically have low-Mg calcite compositions. Other ooids/pisoids (covered in later chapters) can have ferruginous (especially hematite or chamosite), siliceous, bauxitic, phosphatic, evaporitic (gypsum, halite) or other coatings.

      Ooids and other coated grains require conditions suitable for inorganic or microbial precipitation and for biological encrustation of grains. They also require repeated rotation of grains to allow the formation of concentric coatings. Thus, the best environments for ooid formation are tidal deltas and bars, or beaches (marine or lacustrine) where surficial grains are kept in daily motion. Because reefs or bioherms “compete” with ooids in high-energy settings, biologically stressed areas (with abnormal salinities or temperatures) can favor ooid formation by inhibiting organism growth and enhancing rates of carbonate precipitation. Because of their rounded shape ooids are easily reworked into adjacent environments (especially eolianites).

      Other coated grains (superficial ooids, pisoids and oncoids) can be formed in soils and caves (vadoids; cave pearls), in relatively deep-water, current-scoured platform areas (rhodoids), in shelf areas prone to periodic storm action, in partially protected lagoons, and in a wide variety of other settings.

    2. Page 245

      Intraclast - A fragment of penecontemporaneous, commonly weakly consolidated, carbonate sediment that has been eroded and redeposited, generally nearby, within the same depositional sequence in which it formed (Folk, 1959 and 1962).

      Lumps - In modern sediments, irregular composite aggregates of silt- or sand-sized carbonate particles that are cemented together at points of contact: in ancient carbonates, similar-appearing lobate grains that are composed of carbonate mud (micrite). After Illing (1954); no longer widely used.

      Grapestone - Sometimes used to describe aggregates of silt-sized carbonate crystals (or grains), but more properly applied to grape-like clusters of such aggregates bound together by cements or organic encrustations.

      Extraclast - A detrital grain of lithified carbonate sediment (lithoclast) derived from outside the depositional area of current sedimentation (Folk, 1959).

      Calclithite - A rock formed chiefly of carbonate clasts (extraclasts) derived from older, lithified limestone, generally external to the contemporaneous depositional system. Commonly located in arid settings, along downthrown sides of fault scarps. Term coined by Folk (1959).

      Intraclasts and extraclasts are found in deposits of any age from Archean to Recent. Intraclasts are especially common in Precambrian to Mid. Ordovician strata, where they form widespread flat-pebble conglomerates. Such deposits probably reflect the abundance of microbial deposits and the scarcity or absence of macrofaunal grazers and burrowers during that time period (e.g., Garrett, 1970).

    3. Page 253

      Pellets - Small (typically 0.03 to 0.3 mm long), spherical to ovoid or rod- shaped grains composed of carbonate mud (micrite). Most pellets lack internal structure and are uniform in size and shape in any single sample; in the strict sense, pellets are the fecal products of invertebrate organisms (see Folk, 1959).

      Peloids - Allochems formed of cryptocrystalline or microcrystalline calcium carbonate with no restrictions on the size or origin of the grains (McKee and Gutschick, 1969). This term allows reference to grains composed of micritic material without the need to imply any particular mode of origin — it is therefore a useful “term of ignorance” covering possible pellets, indistinct intraclasts, micritized ooids or fossil fragments and even some microbial or inorganic precipitates that are not necessarily even “grains” in the sense of primary constituents as opposed to interstitial early diagenetic “cements”.

      Pellets and peloids occur in Precambrian through Phanerozoic strata; pellets are important sediment constituents mainly in Phanerozoic strata. Structured crustacean pellets are especially prominent in Jurassic and Cretaceous rocks (although they are known from middle Paleozoic to Recent strata).

      Pellets and peloids are composed of aggregated carbonate mud and/or precipitated calcium carbonate. Thus, their original composition is (or was) aragonite or calcite (of any Mg level) or a mixture of both. Pelletal glauconites and phosphorites also are common.

      Fecal pellets are produced wherever worms, crustaceans, holothurians and other grazing, burrowing, or swimming invertebrates (or vertebrates) exist, but most pellets are destroyed prior to burial. Rapid cementation, usually bacterially mediated, aids preservation, as does rapid sedimentation in low-energy settings. Thus, lagoons (especially hypersaline ones), low-energy tidal flats, and sheltered or relatively deep-water platforms are common sites of pellet preservation. Fecal pellets of pelagic zooplankton, especially copepods, are common in Cretaceous to Recent deep-sea deposits.

      Fecal pellets must be distinguished from microbial peloids or inorganic, peloidal marine cements, especially those composed of high-Mg calcite. Such precipitates are especially common in reef cavities, subtidal to intertidal stromatolites, hot springs or other travertine deposits, and submarine vent areas.

      Peloids have varied origins and environmental associations. Algal or fungal boring and micritization of grains are common in a variety of open marine to restricted or coastal settings with relatively slow or intermittent sedimentation rates. In particular, areas subject to occasional storms that move grains from active areas of formation to quiet sites of destruction are especially prone to peloid formation. Such sites include backbarrier or back-bar grass flats, lagoons, and protected deeper shelf settings.

    4. Page 259

      A wide variety of non-carbonate grains can be found as constituents of carbonate rocks. In some cases, these grains are isolated and very subordinate particles; in other examples, they can be important rock-forming elements. Most of these minerals can also be found as diagenetic precipitates in carbonate rocks, but in this section only examples of true grains are illustrated (although some were synsedimentary diagenetic materials that effectively acted as sediment grains).

      It is beyond the scope of this book to examine these non-carbonate grains in detail, but a few of the more important types — clastic terrigenous grains, glauconite “pellets”, and ferruginous and phosphatic ooids —will be discussed briefly. The criteria for their recognition and the compositional characteristics of these grains are given in the individual figure captions.

      The recognition of non-carbonate grains in carbonate rocks is important for the interpretation of depositional environments and for the proper classification of mixed-composition rocks. Most specialized classifications of carbonate rocks simply use adjectives such as “quartzose”, “very quartzose”, “feldspathic” or “glauconitic” to note the presence and relative abundance of non-carbonate grains (the term “silicified” is appropriately used where a significant amount of silica replacement has affected the deposit).

      Well-rounded detrital quartz sand grains are scattered throughout this dolomitized carbonate mudstone. The quartz grains are at various stages of extinction, but none show birefringence colors higher than first order. Such excellent rounding typically indicates a precursor sedimentary source or long-term abrasion in a high-energy setting.

      Abundant angular to subangular, detrital quartz grains (and subordinate feldspar) in a sideritic carbonate. Such angular grains are more common as terrigenous contributions to carbonate sediments than the very well rounded grains of the previous example. The siderite crystals are clear to brownish and have high relief.

      Large, ovoid, pelletal glauconite grains in a glauconitic marl. The light green color in plane-polarized light and speckled, dark green appearance under cross-polarized illumination are characteristic for glauconite. Although “glauconite” basically is an iron- and magnesium-rich illite-type mineral, the term really refers to a family of related minerals (“glaucony” is sometimes used as a generic term for such materials when detailed mineralogical data is lacking). Glauconite grains can have varying degrees of mineral ordering, as well as a complex range of interlayered clay minerals (especially smectite).

    5. Page 265

      Carbonate mud is the equivalent of clay in terrigenous rocks and can form pure deposits (variously termed micrites, carbonate mudstones, lime mudstones, or calcimudstones on the carbonate side, and claystones or shales on the clastic terrigenous side). Clay-sized particles also act as matrix material that supports larger grains or are lodged interstitially between a self-supporting framework of larger grains. Decades ago, when both life and muds seemed pure and simple, both clays and carbonate muds were viewed as miniature versions of larger grains, acting primarily as reworked particles. It is now known that some clays are detrital, some are altered or neoformed on the seafloor, and some are precipitated during the long diagenetic history that accompanies burial, adding considerable complexity to the interpretation of terrigenous shaly deposits. The same is true on the carbonate side. Carbonate muds can be part of the spectrum of disintegration products of carbonate organisms, some can be formed by direct inorganic precipitation, and some may be formed in association with microbial metabolism. Furthermore, some may be primary sediment that responded to hydrodynamic forces during sediment formation and some may be precipitated interstitially, at or near the seafloor (through organic or inorganic processes), or during later diagenesis. It is even possible for grains to break down into smaller carbonate particles during diagenesis, or for diagenetic conversion of former carbonate mud to a mosaic of coarser calcite crystals (microspar). Although we have learned much over the past decades about mud-sized materials, we are far from having a full understanding of them. We also have not yet developed reliable criteria for the consistent distinction between organically produced and inorganically precipitated materials, or even between detrital particles and authigenic precipitates.

      Micrite - An abbreviation of “microcrystalline calcite”. The term is used both as a synonym for carbonate mud (or “ooze”) and for a rock composed of carbonate mud (calcilutite). Micrite consists of 1 to 4 μm-diameter crystals and forms as an inorganic precipitate or through breakdown of coarser carbonate grains. Micrite is produced within the basin of deposition and shows little or no evidence of significant transport (Folk, 1959).

      Microspar - Generally 5- to 20-μm-sized calcite produced by recrystallization (neomorphism) of micrite; can be as coarse as 30 μm (Folk, 1965). Restricted to recrystallization products, not primary precipitates.

      Pseudospar - A neomorphic (recrystallization) calcite fabric with average crystal size larger than 30-50 μm (Folk, 1965).

      Modern marine shelfal carbonate mud is mainly aragonite (with some high-Mg calcite); deep-sea chalk oozes are low-Mg calcite as are most lacustrine calcareous muds. The composition of carbonate muds produced from breakdown of skeletal material are clearly controlled by the mineralogy of those organisms. Paleozoic shells were generally more calcitic than the aragonite-dominated shelled fauna of the modern world. Furthermore, even the mineralogy of “inorganic” marine precipitates (muds as well as cements) is now known to have varied throughout geologic time (Lasemi and Sandberg, 1984, 1993).

      1. Modern carbonate mud consists largely of the breakdown products of organisms (due to decomposition of organic binding materials and abrasion or maceration of shells). Macroscopic algae (especially green algae) are major contributors of needle-shaped, mud-sized, aragonitic particles in tropical platform and platform margin settings. Modern inorganic aragonitic precipitates, in the water column or on the seafloor, also are needle-like (with individual crystals typically 3-5 μm in length) and may contribute to carbonate muds.

      2. The calcitic micrite of older carbonate rocks was neomorphically formed from mixed mineralogy precursors to form an equant mosaic of 1- to 4-μm crystals. The precursor material acted as detrital particles and so may show geopetal fabrics, scattered coarser particles, and other indications of mechanical sedimentation. Inclusions or molds of precursor minerals may be seen within micritic calcites (especially using SEM).

      3. Neoformed microcrystalline cement and microbial precipitates may show clotted or peloidal fabrics and can grow in any position within interparticle pores or larger cavities (non-geopetal fabrics).

      4. Microspar and pseudospar typically have patchy distributions grading into normal micrite; crystal outlines tend to be elongate (loaf-shaped) or have irregular, sutured boundaries).

    6. Page 273

      Primary sedimentary structures are physical and/or biological features formed during the process of sediment deposition. Generally such structures are best seen in outcrop, core, or polished hand sample, but smaller features such as borings or fenestral fabrics are both readily apparent in thin section and important to recognize. Their identification can improve interpretations of depositional environments and can also help to decipher patterns and timing of diagenesis. The characteristic features for the recognition and interpretation of primary sedimentary structures are provided in the figure captions. Diagenetic sedimentary structures, such as hardgrounds, soil crusts, or stylolites, are covered in the appropriate diagenetic chapters.

      Borings - Openings created in relatively rigid rock, shell, or other material by boring organisms. The rigid host substrate is the feature that distinguishes borings from soft-sediment burrows.

      Burrow porosity - Feature created by organic burrowing in relatively unconsolidated sediment, in contrast to borings. Most burrows collapse, become filled with sediment, or are back-filled by the burrow-forming organism itself.

      Fenestrae (fenestral fabric) - Primary or penecontemporaneous gaps in rock framework larger than grainsupported interstices. Such features may be open pores or may have been partially or completely filled with internal sediment and/or sparry cement. Fenestrae occur as somewhat rounded features of spherical, lenticular, or more irregular shapes; their large size in comparison to normal interparticle openings and their multigranular roofs, floors, and other margins are key characteristics. Fenestrae are commonly somewhat flattened parallel with the laminae. They may, however, be round or very irregular, and some are elongate in a vertical dimension. Although isolated fenestrae occur in sedimentary carbonates, it is more common to find many in close association. Fenestrae are generally associated with microbial mats and result from shrinkage, gas formation, organic decay, trapping of air through swash-zone wave action, or other synsedimentary processes (Choquette & Pray, 1970).

    1. Page 283

      Consistent classification and concise naming of rocks and sediments are essential for effective communication throughout the international scientific community. An ideal classification scheme combines objective, quantifiable description of readily observable features that are grouped into named categories. At the same time, it is desirable to have groupings that incorporate a maximum level of genetic or interpretive significance (groupings that reflect mechanisms of formation, environments of deposition, and the like). Although many classifications have been proposed for carbonate rocks and sediments, only two — the Folk (1959/62) and Dunham (1962) classifications — have successfully met the test of time (along with two others that are variants of the Dunham scheme). All four schemes are based on the distinction of three fundamental components: grains (skeletal fragments, ooids, pellets/peloids, intraclasts, and non-carbonate detritus), matrix or carbonate mud, and open pores or sparry-calcite-filled primary interparticle porosity (see diagram on previous page). The differences between the classifications are mainly that Folk uses the relative percentages of grains and matrix, Dunham as well as Embry and Klovan use mud- versus grain-supported fabrics, and Wright uses a more genetic division into biological, diagenetic, and depositional fabrics. This chapter summarizes the features of each classification and provides petrographic examples of carbonate rocks with their Folk and Dunham names; dolostone classifications and examples are covered in the chapter on dolomites.

      The Folk classification uses multiple descriptive terms. The fundamental name is based on the four grain types and the relative abundances of grains (allochems), matrix, and cement or pore space. Eleven basic terms are generated (top diagram on facing page), including ones for pure mud rocks (micrites), muddy rocks with spar patches (dismicrites) and organically-bound rocks (biolithites). Because of their special environmental significance, intraclasts and ooids are favored in the naming process (see top diagram caption).

      To describe the features of carbonate rocks that reflect the degree of sorting and rounding, Folkʼs terminology includes textural modifiers (middle diagram on facing page). In general, deposits classed on the left side of the diagram were formed in “low-energy” settings; rock types farther to the right represent deposition in increasingly high-energy depositional settings.

      A third component of a full Folk name relates to the average grain or crystal size of the rock. That terminology is summarized in the bottom diagram on the facing page.

      A carbonate rock named under the Folk classification can include any or all of the terms generated in these three categories, plus any additional descriptive terms the user desires. Thus, for example, the same rock could be termed a “biosparite” or a “rounded biosparite” or a “coarse calcarenite: rounded biosparite” or a “coarse calcarenite: rounded rudist-coral biosparite” or a “slightly dolomitized coarse calcarenite: rounded rudistcoral biosparite”, depending on the level of detail desired.

      1. Quantifiable, descriptive (objective) terminology.

      2. Although primarily descriptive, rock terms convey considerable genetic (environmental) information.

      3. Multiple optional terms — for grain size, faunal composition, alteration, non-carbonate constituents, and other features — allow informative names at any desired level of detail.

      4. Used worldwide, especially by petrographers in academic settings.

    2. Page 293

      Although most of this book focuses on the identification of grains and cements, what is often of prime interest to hydrocarbon explorationists is understanding the absence of those materials — in other words, the origin and history of open primary or secondary pore space. This chapter, therefore, will deal with recognition of different types of porosity; the chapters on diagenesis will cover the mechanisms and relative timing of porosity creation, retention, reduction, or destruction.

      A number of classifications of porosity in carbonate rocks have been proposed (see citations at end of section), but only the Choquette and Pray (1970) scheme has met with widespread acceptance. Thus, it will be the only one described and applied in this book. This classification combines terms that encompass four separate categories of observations. The main term (called the “basic porosity type”) codifies the location and type of pore space. That term is prefaced with a genetic modifier or modifiers that relate to the process, direction or stage (enlarged, reduced or filled) of porosity evolution, and the time of pore formation; an additional term describing pore sizes can also be added. Finally, an abundance term can be appended at the end of the name to describe the percentage of pore space. In practice, most geologists simply specify the basic porosity type along with the one or two modifiers that are best suited to their needs.

      The basic porosity types are illustrated in two diagrams (below and at the top of the next page). The basic porosity types are organized according to whether they are fabric selective, not fabric selective, or either fabric selective or not. The modifying terms are shown in the middle diagram (next page). Examples of the major porosity types (and some more minor ones) are given in subsequent illustrations.

      A final note: the proper classification of porosity requires accurate observation of the amount and nature of pore spaces. Some porosity is either too large or too small to be recognizable in thin section (see upper photograph on the title page of this chapter), but most is visible at thin-section scales. To recognize and measure porosity properly under the microscope, one MUST use thin sections prepared from rock chips that were pressureimpregnated with color-dyed epoxy. Grains or crystals commonly are plucked out of sections during cutting and grinding; only with colored impregnation media can one distinguish pre-sectioning “real” pores from ones created during section preparation. Intensely dyed epoxy also lends emphasis to porosity and helps to reveal micropores that could otherwise be overlooked. To do quantitative or semiquantitative measurements of porosity using microscopy, one must mathematically correct the observations made in two-dimensional space (see, for example, Halley, 1978); modern digital image analysis methodologies can also be applied to this process (e.g., Anselmetti et al., 1998).

      Choquette & Pray (1970) basic fabric-selective porosity types

      A diagrammatic representation of the basic fabric-selective porosity types used in the Choquette and Pray (1970) carbonate porosity classification. What is meant by fabric selectivity is that the porosity is controlled by the grains, crystals, or other physical structures in the rock and the pores themselves do not cross those primary boundaries.

      Choquette & Pray (1970) basic non-fabric-selective or variable porosity types

      A diagrammatic representation of the basic non-fabric-selective or variably fabric-selective porosity types used in the Choquette and Pray (1970) carbonate porosity classification. These are all porosity patterns that actually or potentially can cross-cut primary grains and depositional fabrics. They also include porosity types that potentially can be much larger than any single primary framework element.

  2. Page 303
    1. Page 303

      Diagenesis encompasses any physical or chemical changes in sediments or sedimentary rocks that occur after deposition (excluding processes involving high enough temperatures and pressures to be called metamorphism). Diagenesis, thus, can begin at the sea floor (syngenetic or eogenetic alteration), continue through deep burial (mesogenetic alteration), and extend to subsequent uplift (telogenetic alteration). Diagenesis can obscure information about primary features, but diagenesis also can leave behind substantial information about the history of post-depositional settings, pore water compositions, and temperatures.

      Diagenesis can reduce porosity and permeability, or it can increase them. In general, though, the trend is toward progressive loss of porosity and permeability with increased time and depth of burial, and that shift is commonly quite substantial. The top diagram (opposite page) shows the highly generalized average porosities of modern carbonate sediments, typical ancient carbonates, and the “exceptionally porous” rocks that constitute hydrocarbon reservoirs. Modern sediment porosities range from 35-45% for grainstones to 70% or more for mudstones or chalks. Typical ancient carbonates have less than 5% porosity, and even reservoir rocks average far less than half the porosity of their modern carbonate equivalents. Thus, understanding diagenetic processes, the factors that inhibit porosity loss, and the relative timing of oil migration versus porosity evolution are critical to exploration for hydrocarbons and carbonate-hosted mineral deposits.

      Diagenesis typically involves a variety of physical and chemical processes — the most common of these are:

      1. Cementation (the filling of open pore space, of primary or secondary origin, with newly precipitated materials)

      2. Dissolution (the leaching of unstable minerals forming secondary pores, vugs, or caverns)

      3. Replacement of one mineral by another (or “inversion”, the replacement of one polymorph of a mineral by another)

      4. Recrystallization or strain recrystallization (changes in crystal size, strain state, or geometry without change in mineralogy)

      5. Physical or mechanical compaction (including dewatering and deformation or reorientation of grains)

      6. Chemical compaction (dissolution mainly along surfaces such as stylolites or solution seams)

      7. Fracturing

      The terminology applied to such a complex range of carbonate diagenetic processes and products is understandably also complex and is generally applied with disconcerting inconsistency. Folk (1965) provided what is still the most concise, yet inclusive, terminology for diagenetic fabrics. Pore-filling cements are described based on their mode of formation (passive or displacive precipitation), crystal morphology (based on length-width ratios as shown in the middle diagram, facing page), crystal size (see table in limestone classification chapter), and relationship to foundation (overgrowth, crust, or spherulitic growth without obvious nucleus).

    2. Page 313

      Synsedimentary diagenesis in the marine realm is relatively uncomplicated (by comparison with meteoric and burial diagenesis) because it generally operates over short time spans (only years to thousands of years, in most cases) and involves a restricted range of pore fluid chemistries. Nevertheless, through a combination of physical, chemical and biological processes, coupled with access to a nearly unlimited supply of dissolved materials in seawater, marine diagenesis can often bring about remarkable change in carbonate sediments and produce some very complex fabrics. Furthermore, the subsequent overlay of meteoric or burial diagenetic alterations can greatly complicate the recognition of marine diagenetic fabrics in ancient carbonate rocks. That is especially true because the aragonitic or Mg-calcitic cements that result from marine diagenesis are essentially just as unstable in meteoric or burial-stage pore fluids as primary grains of those compositions.

      The intensity or extent of marine cementation is a function of the supply of solutes from seawater. Solute supply, in turn, depends on sedimentation rates and the effectiveness of water transport from the surface into the interior of a sediment pile. Mechanisms of water movement include, among others, wave forcing, tidal pumping, thermal convection, and diffusive transport. Areas of very slow sedimentation (e.g., hiatus surfaces, low-sedimentation-rate platform interiors, or low-productivity deep sea settings) can have substantial marine cementation (including hardgrounds) because they all have long times of contact between seawater and a thin package of sediment, even with no special mechanism for water pumping. In high-sedimentation rate areas, on the other hand, substantial marine cementation occurs mainly in reef front or coastal settings where wave or tidal action can force seawater through the sediments to a considerable depth. Likewise, atoll margins and steep carbonate platform flanks are sites of extensive marine cementation because of convective water input coupled, in some cases, with low sediment accumulation rates. Hot or cold seeps on the sea floor also represent sites of exceptional water throughput and extensive cementation.

      Grain and matrix dissolution are widespread in certain marine environments, particularly in cold- and deep-water areas. Modern oceanic waters have an aragonite compensation depth or ACD at roughly 1,500 m (the ACD is the depth below which aragonite does not accumulate because the rate of dissolution exceeds the rate of aragonite supply). Aragonite also is extensively dissolved in cool and cold-water shelf areas. The modern calcite compensation depth (CCD) lies at roughly 4,500 m (but that depth, as well as that of the ACD, varies with latitude, productivity, and other factors, and undoubtedly has varied significantly with geologic time).

      Bored (biodegraded) grains with cement infill of borings and generation of micrite envelopes (also discussed in the sections on pellets/peloids and sedimentary structures-borings).

      Isopachous crusts of fibrous to bladed, peloidal, or aphanocrystalline high-Mg calcite cement. The aphanocrystalline crusts consist of equant, less than 4 μm-sized rhombs that look much like micrite.

      Isopachous crusts of fibrous aragonite cement within grain cavities and as intergranular cements (predominantly found in warm-water, slightly hypersaline settings and tropical beachrock deposits).

      Marine-cemented hardground formation in selected areas (see above) — associated, in many cases, with phosphate and glauconite cementation, boring and faunal encrustation, and intraclast formation.

      Large botryoids of cavity-filling aragonite and high-Mg calcite cement.

      Internal sediment fills of primary cavities or neptunian dikes in framework-supported sediments.

      Coastal beachrock and spray-zone cements.

      Microbe/cement associations in marine methane and thermal seeps.

      Modern marine cements in warm-water settings consist mainly of high-Mg calcite (~12-18 mol% Mg), but with extensive aragonite as well. In colder-water areas (temperate, polar and deep marine), high-Mg calcite cements predominate, but become scarcer and less Mg-rich at higher latitudes. Many ancient carbonate deposits certainly had aragonite and high-Mg calcite cements, perhaps with secular variations in their abundance (e.g., Wilkinson and Given, 1986), but low-Mg calcite marine cements may also have formed at some times. In older limestones, original aragonite and high-Mg calcite cements generally have been converted diagenetically to low-Mg calcite and must be recognized by micro-inclusions, geochemical analysis (especially Mg and Sr contents), relict morphologies or crystal outlines, or, as a last resort, characteristic patterns of preservation or alteration (former aragonitic cements, for example, typically have poor primary fabric preservation.)

      Characteristic morphologies of marine cements

      A diagrammatic depiction of some common types of modern marine high-Mg calcite and aragonite cements. Most of these morphologies will be illustrated in this section. Adapted from James and Choquette (1983).

    3. Page 331

      Meteoric diagenesis represents alteration that occurs at or near the earth’s surface in strata influenced or pervaded by waters of recent atmospheric origin. The meteoric environment is typically divided into unsaturated (vadose) and saturated (phreatic) zones divided by a water table (see top diagram, facing page). The interfaces between surficial meteoric fluids and strata filled with other pore fluids (seawater or basinal waters) are “mixing zones” that can have special diagenetic characteristics.

      Many, perhaps most, shallow marine carbonate deposits undergo meteoric diagenesis, either as a consequence of buildup of sediments above sea level, or through drops in sea level that expose platform carbonates. In addition, meteoric water can circulate well below the land surface to alter carbonate deposits far older than the exposure interval. Meteoric processes commonly act over time periods of hundreds to millions of years.

      Meteoric diagenetic patterns typically are complex and variable for the following reasons: 1. regional and temporal variations in starting material; 2. variations in rainfall and water throughput rates (in part, related to permeability variations); 3. variations in water chemistry (from locality to locality or vertically through the water column at any one site, especially at mixing interfaces); 4. variations in the duration of exposure or alteration during multiple episodes of exposure; and 5. the effects of plants and plant-derived acids that vary regionally and also changed through geologic time as a consequence of evolution of different plant groups.

      The vadose zone is characterized by extensive dissolution of unstable carbonate minerals (aragonite and high-Mg calcite), often with reprecipitation of more stable carbonate (low-Mg calcite). As a consequence, primary porosity commonly is filled during meteoric diagenesis, and secondary porosity is created.

      Unless there is a thinning or collapse of the rock section, meteoric diagenesis is relatively porosity neutral, at least at the scale of grains, with dissolution at one site supplying solutes for reprecipitation elsewhere. Meteoric diagenesis does, however, have a strong effect on permeability (e.g., permeability reductions through cementation of interconnected primary pores or permeability increases through solution enlargement of fractures).

      Many vadose cements have fabrics reflecting the selective distribution of water in that environment — pendant (microstalactitic or gravitational) cements hanging from undersides of grains and meniscus cements concentrated at grain contacts. Whisker crystals (also termed needle-fiber cements), calcified filaments, blackened pebbles, root structures (rhizoliths), microspar, and Microcodium also are common features.

      Phreatic zone cements are typically isopachous rims or complete pore fillings of equant calcite.

      Freshwater meteoric calcites are depleted in Sr2+, Mg2+, δ18O, and δ13C, relative to their marine precursors. Most, but not all, meteoric settings are oxidizing, resulting in typically low Fe2+ and Mn2+ contents in meteoric cements (reflected in non-ferroan staining and no cathodoluminescence response).

    4. Page 351

      This chapter deals only with the diagenesis of calcitic components of limestones — the formation of dolomite, silica and other minerals is covered in subsequent chapters.

      Burial diagenesis represents alteration that occurs below the zone of near-surface water circulation (i.e., below the meteoric phreatic mixing zone or below the zone of active seawater circulation). Burial diagenesis plays a major, often THE major role, in the diagenesis of sediments from the point of view of length of time spent in that environment (commonly millions to hundreds of millions of years) and in terms of porosity changes.

      Burial diagenetic features are among the most difficult to identify with assurance for a variety of reasons: 1. the transition between surficial (meteoric or marine) pore fluids and burial realm fluids is ill-defined, variable, indistinct, and rarely well understood (so often it is not clear where surficial diagenesis ends and burial diagenesis begins); 2. the burial realm is “out of sight and out of mind”, which means that the processes and products formed there can only be remotely and incompletely observed; 3. deposits found in the burial diagenetic zone must have passed through marine or meteoric diagenesis zones (or both), making it difficult to determine precisely whether a particular fabric is exclusively a product of burial diagenesis.

      Several factors mitigate for and against extensive burial diagenesis. Burial diagenesis is hindered by water circulation rates that typically are lower in subsurface settings than in near-surface environments (because of slower circulation mechanisms as well as reduced permeabilites). Higher temperatures and increased pressures at depth, however, tend to accelerate many diagenetic processes. Elevated pore-fluid pressures (reducing grain-to-grain stress) and early hydrocarbon input retard mechanical and chemical burial diagenesis.

      Statistical evidence (top diagram, facing page) indicates that burial diagenesis is very important in porosity reduction. Most rocks, especially limestones, show a consistent loss of porosity with progressive burial.

      The burial-diagenetic zone is characterized by a mix of physical and chemical diagenetic processes, most leading to porosity destruction, but in some cases yielding net porosity increases.

      Burial-related mechanical compaction features include dewatering structures, compactional drape around shells and nodules, plastic or brittle grain deformation, and fractures.

      Embayed grain contacts, fitted fabrics, solution seams, and stylolites are common chemical compaction features that form mainly in burial settings.

      Burial-stage calcite cements are low-Mg calcite. Most crystals grew slowly, and thus are relatively imperfection-free, clear (limpid) crystals as compared with marine and even meteoric precipitates. Morphologies include bladed, prismatic overgrowths of earlier cement crusts; equant calcite mosaics; drusy calcite mosaics with crystal sizes increasing toward pore centers; very coarse to poikilotopic blocky calcite spar; and outer, inclusionpoor stages of syntaxial overgrowths. Although these fabrics are common in mesogenetic precipitates, none is unequivocally or exclusively formed during burial diagenesis. Bathurst (1971 and 1975) and Dickson (1983) provide more detailed discussion of geometric criteria for recognition of burial cements.

      Many burial-stage cements are formed from relatively reducing pore fluids and, thus, may have elevated Mn2+ and Fe2+ contents. The iron is easily detected with staining techniques; the manganese/iron ratio is qualitatively identifiable with cathodoluminescence (CL). The typical CL pattern found in burial stage calcite cements is a transition from nonluminescent to brightly luminescent to dully luminescent response. This is generally interpreted as a transition from oxidizing (pre-burial or early burial) conditions with little or no Mn2+ or Fe2+incorporation into the calcite lattice, to reducing conditions with Mn2+ and Fe2+ incorporation, and finally to reducing conditions in which Fe2+ availability and incorporation exceed Mn2+ availability and incorporation. More complex CL stratigraphies, however, are common.

    5. Page 371

      Dolomite is a rhombohedral mineral, CaMg(CO3)2; dolostone is the appropriate term for a rock composed of that mineral. Dolomite is best identified through staining, and by its rhombic, often zoned, untwinned habit.

      Dolomite is a complex and relatively poorly understood mineral. Thermodynamically, dolomite should be a stable, widespread precipitate from seawater, but kinetic factors (hydration of Mg2+ ions in seawater, the high ionic strength of seawater, the relative efficiency of aragonite and high-Mg calcite precipitation, inhibition effects of SO42- ions) mitigate against its formation. Modern dolomite therefore is relatively scarce. In addition, ordered dolomite is slow-growing, and thus is difficult to synthesize in the laboratory under earth-surface conditions.

      True dolomite (stoichiometric, ordered dolomite; top diagram, facing page) is well ordered, with one cation layer entirely composed of Mg2+ and the next entirely composed of Ca2+. If perfectly formed, that also ensures a 50:50 (stoichiometric) balance between Ca2+ and Mg2+ in the dolomite structure. Most modern dolomites, however, are poorly ordered and Ca-rich (termed “protodolomite” by some workers). Those crystals are relatively unstable and “ripen” or eventually neomorphose to more stable, ordered dolomite crystals.

      Many models have been proposed for dolomitization (see excellent summaries in Morrow, 1982b; Land, 1985; Tucker, 1990; and Purser et al., 1994). All center around three basic factors: a source of Mg (generally seawater), a way to move large volumes of that water through the sediment package, and a way to reduce the kinetic inhibitions to dolomite precipitation. Sabkha and brine reflux models call upon evaporative concentration of seawater (with removal of sulfate through bacterial reduction or inorganic sulfate precipitation); marine-fresh water mixing zone and Coorong models rely on dilution of seawater; the burial model uses elevated temperatures, modified pore water compositions, and, in some cases, thermochemical sulfate reduction to reduce inhibitions on dolomite precipitation. Organogenic dolomitization relies on intense bacterial sulfate reduction and methanogenesis in organic-rich sediments in a wide range of settings (Mazzullo, 2000). Modern dolomite has been found in small volumes in many settings, ranging from hypersaline sabkhas to normal salinity tidal flats, and subsaline lagoonal environments. Modern dolomite is predominantly a replacement product; subsurface dolomites are found as either replacements or as primary pore-filling precipitates. Some authors have speculated that dolomites of other ages (especially the Precambrian) were primary precipitates, but that hypothesis has not been confirmed.

    6. Page 393

      Sulfate and chloride minerals occur as cements, displacive and replacive nodules, and interbedded strata in carbonate rocks. They precipitate from evaporatively concentrated waters in arid-region lakes, ponds and lagoons along marine shorelines and, more rarely, in deeper shelf and basinal settings with restricted marine inflow. Evaporite deposits are products of arid environments; however, evaporitic solutions are highly mobile due to their high density. Evaporative brines thus may migrate into adjacent or underlying strata and precipitate diagenetic sulfate or chloride minerals (generally as displacive crystals and nodules, or as carbonate replacements) in units that may otherwise be unrelated to arid settings. Even after deposition and substantial burial, evaporite minerals can be remobilized and reprecipitated in distant, stratigraphically unrelated units. Therefore, careful petrographic analysis is needed to determine both the conditions of primary deposition and the timing of diagenetic events in evaporite-bearing limestones and dolomites.

      Barite, celestite and anhydrite also can occur as hydrothermal precipitates in carbonate rocks.

      Calcite solution-fill replacement (calcitization) of gypsum and anhydrite results from the dissolution of evaporites by sulfate-poor pore fluids. These pore fluids become saturated to supersaturated with respect to Ca2+; if there is enough bicarbonate in the pore fluids, calcite may precipitate.

      Anhydrite crystals have high birefringence (up to third order); in thin section, the other common sulfate and halide minerals have much lower birefringence. Anhydrite’s birefringence also can appear to “twinkle” like that of calcite, but the effect is less strongly developed than in calcite. Anhydrite crystals normally are colorless, but may contain inclusions of precursor phases. Anhydrite may form large tabular crystals or felted, fibrous crystal masses (generally as nodules). The larger crystals may exhibit pseudo-cubic cleavages.

      Gypsum, celestite and barite can be extremely difficult to differentiate from each other in thin section. They all have low relief and birefringence (gray to white). Gypsum tends to form colorless, elongate, tabular to lenticular crystals or fibrous masses or aggregates of crystals. Gypsum also tends to form poikilotopic cements that encase numerous grains – siliciclastic or carbonate. Gypsumʼs cleavage is lozenge-shaped; therefore, if cleavage planes are visible, they are diagnostic for gypsum. Gypsum crystals may form rosettes and twins that are called swallow- or fish-tailed selenite. These larger crystals form displacively below the sediment/water interface in unconsolidated sediments; such crystals contain abundant inclusions of the sediment. Selenite crystals also can grow upward from the sediment-water interface into standing saline water bodies.

      Celestite ranges from colorless to blue in thin section. Blue crystals of celestite can be pleochroic, which helps to differentiate it from gypsum and barite. Celestite forms fibrous to rounded aggregates of crystals. When it is found in fibrous masses, the crystals are normally more elongate than similar crystals of gypsum. Cleavage, when visible, is pseudo-cubic.

      Barite normally is colorless and forms globular concretions, granular to earthy masses, fibrous or bladed crystals. Cleavage, when visible, is pseudo-cubic. Because barite and celestite form a solid solution series, they are extremely difficult to tell apart in thin section. Generally, other chemical techniques must be used to be confirm identifications. Like gypsum, barite also forms crystal rosettes.

      Halite is difficult to see in thin section, because it is isotropic and highly soluble. Because halite is isotropic, it can easily be overlooked if the cleavages are not prominent or if it doesnʼt contain inclusions (i.e., it may be indistinguishable from the glass on which the section is mounted). Impregnating the sample with blue epoxy makes the halite stand out from the porosity. If the thin section is not properly prepared (cut and ground in oil, not water), however, halite is unlikely to be preserved. Halite crystals are normally colorless and exhibit low relief, but they may appear dusty due to the great abundance of solid and liquid inclusions. Halite can occur as a poikilotopic cement in either carbonate or siliciclastic strata.

      Anhydrite – CaSO4, orthorhombic

      Gypsum – CaSO4 • 2H2O, monoclinic

      Celestite – SrSO4, orthorhombic, complete solid solution series exists with barite

      Barite – BaSO4, orthorhombic, commonly contains up to 3% lead

    7. Page 407

      Silica, a general term used for a variety of crystal forms or morphologies of SiO2, is a widespread diagenetic mineral in carbonate rocks. Silica may occur as cement or it may be found as a replacement of original or diagenetically altered sediment. Silica typically replaces or infills carbonate minerals, evaporites and organic material (e.g., petrified wood).

      The major source of silica for diagenesis is biogenic opal; therefore, silica is especially prevalent in deep-marine sediments from active upwelling zones and shallower-water carbonates from nutrient-rich carbonate shelves. Sponge spicules, diatoms and radiolarians are the most common biogenic contributors and are diagenetically unstable when compared to siliciclastic grains. Other, generally less significant sources of silica in carbonate rocks include volcanic ash, by-products of chemical weathering in soil zones (silcretes), and hydrothermal fluids. Some bedded cherts from saline lakes may be related to hydrous sodium silicate precursors (e.g. Eugster, 1967).

      Except for silcrete formation or hydrothermal alteration, silica diagenesis is rarely a very early- or a very late-stage diagenetic event in carbonate rocks. Rather, it is most typically a product of burial diagenesis. This is due to the timing of the conversion of biogenic opal-A, first to opal-CT lepispheres, and then to stable microquartz or megaquartz. These silica reactions are dependent on temperature (and/or burial depth) and time. In pelagic deposits (away from hydrothermal input), opal-A to opal-CT conversion begins at 20-30°C and may take 10 million years to go to completion; opal-A is rarely found in sediments older than 20 Ma (Hesse, 1990). The conversion of opal-CT to quartz most likely starts at temperatures of ~ 50°C and depths of 500 m, but continues to higher temperatures. Opal-CT is not found in sediments older 144 Ma and chert is relatively scarce in young Cenozoic deposits.

      Amorphous silica— also known as opal; isotropic; high negative relief; colorless to gray or brown; normally contains irregular cracks or fractures; occurs as cements, nodules or replacements (especially wood).

      Equigranular quartz — equant crystals; in polarized light, the maximum birefringence should be first-order white to pale straw-yellow (unless the thin section is thicker than normal); larger individual crystals are normally hexagonal and may be doubly terminated; no cleavage; normally colorless, but may contain inclusions. Fabric is termed cryptocrystalline (chert) when crystals are <5 μm, microcrystalline for crystals 5-20 μm, and megaquartz for crystals >20 μm. Quartz may occur as individual crystals or in large nodular masses replacing or displacing sediment.

      Fibrous quartz — elongate fibers of quartz; same birefringence as equigranular quartz (but birefringence decreases with increasing water content); colorless to brown; common banding or zoning (bands may consist of alternating forms of chalcedony); commonly forms cements, small to very large nodules, and may pseudomorphs other grains nodules or minerals.

    8. Page 417

      Pyrite (FeS2) is the most abundant iron sulfide mineral found in carbonate sediments. Pyrite is an isometric mineral that commonly forms crystals that are cubic, pyritohedral or octahedral, but it may also form anhedral replacement masses. In sediments, pyrite also occurs as framboids or spheres composed of aggregates of minute crystals. Pyrite is opaque in thin section and is readily identified by reflected light microscopy due to its brassy to golden yellow color (simply holding a strong light source above the thin section as it sits on the microscope stage and blocking transmitted light illumination will generally suffice for identification).

      Hematite (Fe2O3) is normally an opaque mineral. In reflected light, hematite is deep red to rusty red. It rarely forms crystals and occurs typically as amorphous masses. Hematite commonly forms through weathering and oxidation of pyrite or other iron sulfides, and it is not unusual to find pyrite and hematite together.

      Goethite (FeO(OH)) is an opaque orthorhombic mineral, whereas limonite (FeO(OH)•nH2O) is a cryptocrystalline or amorphous, hydrated form of this compound. Both minerals are reddish brown to yellowish brown in reflected light, and they can be difficult to tell apart from hematite. They are weathering products of either iron sulfides or hematite.

      Sphalerite (ZnS) is an isometric mineral that is isotropic in cross-polarized illumination, has a high positive relief, and ranges from colorless to pale yellow or light brown. A slight birefringence may be present when the crystals have been strained. Crystals are usually not well formed, but where present, crystal faces may be curved. Well-developed lamellar twinning is common in sphalerite. Sphalerite is found in Mississippi Valley-type mineralized carbonate rocks and other hydrothermal deposits.

      Fluorite (CaF2) is an isometric mineral that forms cubic crystals, although anhedral masses are common in carbonate rocks. Fluorite normally is colorless in thin section, but strongly colored samples may be pale purple to green. Halite and fluorite are easily confused since they are both isotropic, form euhedral cubic crystals and have negative relief. Fluorite can be distinguished from halite based on its well-developed octahedral cleavage, lower negative relief, and color spots that are produced by inclusions within the crystals. Most fluorite was precipitated from hydrothermal fluids and may be associated with Mississippi Valley-type mineralization.

      The two most common phosphatic minerals in carbonate rocks are fluorapatite (Ca5(PO4)3F) and hydroxylapatite (Ca5(PO4,CO3,OH)3(F,OH)x). When intergrown, the minerals formed are francolite (crystalline form) and collophane (cryptocrystalline form). Collophane is the more common mineral — it is isotropic to very weakly birefringent with colors that range from yellowish to brownish. Most early diagenetic phosphate is made of collophane. Francolite has a higher relief and low birefringence (gray to low white); it is colorless to pale brown, and may be slightly pleochroic. Diagenetic phosphatic minerals can form amorphous nodular masses, cements or replacements. Diagenetic phosphates form mainly in areas with substantial primary sedimentary phosphate accumulation — areas with low sediment accumulation rates and high nutrient inputs.

      Glauconite (K,Ca,Na)1-0.56(Fe3+, Mg, Fe2+,Al)2(Si, Al)4O10(OH)2) is a clay mineral found only in marine deposits. It forms pellets or granules in areas of slow sedimentation. It also precipitates as an early diagenetic mineral replacing clasts or filling porosity in shallow to deep marine settings that have high nutrient levels and low sediment accumulation rates. Glauconite is green to olive green in color and has a greenish birefringence; it can look similar to chlorite, but chlorite is usually more platy and has anomalously low birefringence.

      Hydrocarbons can be found as interstitial material in carbonate rocks or as fluid inclusions within carbonate cements. In some cases, hydrocarbons effectively terminate cementation by blocking the entry of aqueous fluids responsible for diagenesis. Bitumen, asphalt and hydrocarbon-filled inclusions all are products of this complex interplay of hydrocarbon-bearing and aqueous fluids. Evidence of hydrocarbon entry includes residues and inclusions, as well as curved meniscus cements and the preservation of unstable carbonate phases, such as aragonite, in very old rocks.

      An SEM image of a pyrite framboid. Framboids are almost perfectly spherical bodies of small, interlocking pyrite crystals. These spherical aggregates typically form discrete bodies, but they are also found as clusters or multiple spheroids. They are authigenic in origin and form in reducing environments or in reducing microenvironments associated with decomposing organic matter.

    9. Page 429

      Although light-microscope petrography is an extremely valuable tool for the identification of minerals and their textural interrelationships, it is best used, in many cases, in conjunction with other techniques.

      Precise mineral determinations are greatly aided by staining of thin sections or rock slabs, by x-ray analysis, or by microprobe examination. Where noncarbonate constituents are present in carbonate rocks, they often are better analyzed in acid-insoluble residues than in thin section. Where detailed understanding of the trace element chemistry of the sediments is essential, x-ray fluorescence, inductively coupled plasma mass spectrometry, ion microprobe, electron microprobe, atomic absorption or cathodoluminescence techniques may be applicable; and where it is desirable to know the temperatures, water sources, and/or pore fluid compositions involved in cementation, fluid inclusion geothermometry, stable isotope geochemistry, strontium isotope geochemistry and a number of other analytical techniques may provide useful information.

      In addition, many sediments may be too fine-grained for adequate examination with the light microscope. The practical limit of resolution of the best light microscopes is in the 1-2 μm range. Many carbonate and noncarbonate matrix constituents fall within or below this size range. Furthermore, because most standard thin sections are about 30 μm thick, a researcher typically sees 10 or 20 such small grains stacked on top of one another in a micritic limestone, with obvious loss of resolution. Smear mounts or strew mounts (slides with individual, disaggregated grains smeared or settled out onto the slide surface) are an aid in examining small grains where the material can be disaggregated into individual components. In most cases, however, scanning and transmission electron microscopy have proved to be the most effective techniques for the detailed examination of fine-grained sediments.

      The bibliography for this chapter (and those in many previous chapters as well) provides references to techniques useful in supplementing standard petrographic analysis. Although many of the techniques require sophisticated and expensive equipment, others, such as thin-section staining, production of acetate peels, or concentration of insoluble residues, can be done in any laboratory and at very little cost.

      Because of the potential desirability of supplemental techniques, it is often useful to prepare epoxy-cemented thin sections without coverslips. These sections can be examined under a light microscope, either by placing a drop of water and a coverslip on the sample during viewing, or by using mineral oil or index of refraction oils with or without coverslips. Such examination involves some loss of resolution, but does allow the cleaning and drying of the surface of the section and subsequent staining, cathodoluminescence, or microprobe examination. One can even partially or completely immerse the thin section in acetic or hydrochloric acid and decalcify the section, thereby sometimes enhancing organic structures or insoluble-mineral fabrics. Finally, uncovered thin sections can be ground thinner in cases where examination of very fine-grained sediments is needed.

      Clearly, one can spend years analyzing a single sample using all possible techniques. Efficient study requires a thorough understanding of all available tools and proper application of the most useful and productive of these.

      Staining techniques are among the fastest, simplest, and cheapest methods for getting reliable mineralogical, and some qualitative elemental, data on carbonate phases. The following list of minerals and their diagnostic stains is derived from the work of Friedman (1959), Dickson (1965 and 1966), Milliman (1974), and others. The original papers, listed in the bibliography, will provide details about the exact application and methods.

      Aragonite - can be distinguished from calcite by the use of Feigelʼs Solution. Aragonite turns black whereas calcite remains colorless for some time. Mixing Feigelʼs Solution requires 7.1 g of MnSO4•H2O; 2 to 3 g of Ag2SO4; 100 cc of distilled water and a 1% NaOH solution. Difficult to prepare and store.

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