Abstract Quartz (SiO 2 ) is the most abundant mineral in terrigenous sedimentary rocks and is exceedingly durable (surviving multiple generations of weathering and deposition). Quartz and silica occur in many varieties—true quartz in the form of megaquartz, chert, microquartz, or chalcedony and various other forms of silica, mainly opal (opal-A and opal-CT [cristobalite]).

Abstract Feldspars (XAl (1-2) Si (2-3) O 8 ) are the most common rock- forming minerals in the Earth’s crust, and they occur in many varieties — ranging from sodium- and calcium-rich (plagioclase) to potassium-rich (K-feldspar or alkali feldspar). K-feldspars may also contain significant amounts of sodium in their crystal lattices. Feldspars are far less resistant than quartz to chemical and physical destruction and thus are altered or removed by weathering, transport and diagenesis, yielding secondary pores or alteration products (illite, white mica/sericite, albite or kaolinite). Even so, they are the second most abundant grains in sandstones, and identifying their mineralogy is crucial for accurate sandstone classification and provenance studies.

Abstract Rock fragments (also called lithic fragments or composite grains) can be derived from a wide variety of lithotypes and commonly have source-specific textures and compositions that can be recognized in thin section. Because of their multicrystalline/granular nature, rock fragments tend to be more common in the coarser grain-size modes of clastic terrigenous rocks (although, under the right circumstances, they can even be seen in mudrocks). Given the composite character of lithic fragments, many petrographers use the Gazzi-Dickinson method of point counting to record the constituent crystals within the fragments, rather than counting the fragments as such (Ingersoll et al., 1984). Rock fragments should be very common in sediments, and they are in many deposits, but because of their multi-crystalline or multi-granular nature, many succumb to the effects of weathering, abrasion or later mechanical or chemical diagenesis. But because the surviving rock fragments yield some of the most direct evidence of contributions from igneous, metamorphic or sedimentary terranes, it is especially important that such grains be accurately identified.

Abstract Accessory minerals include all the many detrital minerals that are found in clastic terrigenous rocks that do not contribute directly to rock classification (thus, primarily minerals other than quartz and feldspar). Although thousands of minerals could potentially fall under that definition; practically, a limited number are found with any great frequency. Accessory minerals as a whole typically make up less than 1% (rarely more than 2%) of most terrigenous sedimentary rocks. Quartz arenites commonly have the fewest accessory minerals (as little as 0.05% in some cases); arkoses are somewhat richer in accessories, and lithic arenites generally have the highest levels. This results from the fact that most accessory minerals, like some feldspars and lithic fragments, lack the abrasion resistance or chemical stability to survive erosion, transport and diagenesis. Accessory minerals can be examined in thin sections; alternatively, they can be concentrated by mechanical (shaker table) or flotation (heavy liquid) methods (see, for example, Munsterman and Kerstholt, 1996; Koroznikova et al., 2008) and can then be viewed with stereoscopic microscopes, SEM or other methods. Thin- section examination shows the grains in the context of rock fabric, but such minerals can be quite scarce in any single section. Disaggregation and concentration is much more effective for evaluating the full assemblage of such minerals in rock or sand samples and also allows identification by x-ray or geochemical methods. Because accessory minerals are so commonly studied as separates, they generally are divided into light and heavy minerals with a boundary drawn by various workers at specific gravities between 2.85 and 3 (they also are commonly divided into opaque and nonopaque minerals). The most commonly encountered detrital light accessory minerals are micas (mainly muscovite and also biotite). Heavy minerals are vastly more numerous, and can be grouped into ultrastable, intermediate stability, and unstable categories. The ultrastable minerals are the ultimate survivors, even more stable than quartz under most conditions — thus, they are found in most clastic terrigenous rocks. The intermediate group has varied levels of survivability, but most such minerals can be degraded or removed under specific conditions; minerals in the unstable group survive only under very favorable conditions (minimal mechanical and chemical stresses). There are so many detrital accessory minerals that occur in clastic terrigenous rocks that it is simply impossible to provide a brief yet usable summary of the mineralogical features and optical characteristics of all these minerals. We provide instead a chart of the relative stabilities of the most common accessory minerals (Table 4.1). In addition, characteristic mineral properties are described in the individual photo captions for each mineral illustrated. For additional information readers are encouraged to consult the references at the end of this chapter or the more general mineralogy texts listed in the bibliography in the introduction to this book.

Abstract Many clastic terrigenous rocks contain variable, but in some cases substantial, amounts of primarily nondetrital constituents. These include biogenic/skeletal grains such as calcareous shells, siliceous tests, phosphatic vertebrate or invertebrate material and organic matter (from plant remains down to plankton and microbial filaments). Nonskeletal, but still biogenic grains, primarily fecal pellets, also can be abundant in some deposits, especially bioturbated ones. Other materials, such as phosphate, gypsum, green marine clays (glauconite, berthierine, chamosite) and ferrous oxides and hydroxides, are found in terrigenous deposits as minerals formed by direct precipitation, through alteration of other minerals or as detrital grains. Most of these grains, in some circumstances, can be sufficiently abundant to be the major constituents of rocks. Even where such grains are not the major rock constituent, however, it is important to recognize them and, if deemed important, one can add a descriptive adjective to any rock name (e.g., glauconitic quartz arenite or radiolarian-bearing arkosic siltstone).

Abstract The term “texture” encompasses a wide range of attributes of sediments/rocks and their constituent grains, including grain size and sorting, particle morphology (form and sphericity, rounding, and surface texture), grain orientation, imbrication and packing. All of those properties have significance in interpreting transport processes and depositional settings of sedimentary rocks, but they also have economic importance in a wide variety of fields ranging from engineering of construction and road materials to understanding and predicting the porosity and permeability relationships of such materials in petroleum exploration/production or hydrologic contexts. Most textural properties mentioned above are best measured in unconsolidated or easily disaggregateable materials where grains can be size-sorted by sieving, settling, laser particle analysis or other techniques or where individual grains can be viewed in three dimensions. These properties generally are far more difficult to measure accurately in thin sections of consolidated rocks. To use just one example, the size of a grain in thin section can never exceed the longest axis of the grain but it can easily be shorter because most cuts through grains are tangential or oblique to that axis. So, in most cases, the basic size-, sorting- and shape-related properties of grains are merely estimated in thin sections through the use of visual comparators. Even there, the most useful comparators are ones that have been specifically corrected for thin-section use. Detailed direct measurements on grain size and shape can, of course, be done using thin-section microscopy, especially through point-counting large numbers of grains, but there too, correction factors must be used to overcome, to the degree possible, the two-dimensional (2-D) view of three-dimensional (3-D) grains afforded in thin sections (see Harrell and Eriksson, 1979; Johnson, 1994). Advances in computerized photomicrographic image analysis can make the process of measurement and measurement correction both more accurate and far less time consuming (see, for example, Schäfer and Teyssen, 1987; Seelos and Sirocko, 2005; Syvitski, 2007). An enormous amount of effort was made in the period from the 1930s to the 1970s to perfect textural measurements and to find reliable statistical measures that could be used to identify specific environments of deposition. That work still finds application in soft and unconsolidated sediments, but it will not be discussed in detail in this book, because it is of lesser applicability to petrographic studies of hard rocks. A number of papers in the bibliography at the close of this section can be used to follow up on unconsolidated sediment studies (especially the excellent overview of statistical measures provided in Folk, 1980) and almost all textbooks on sands and sandstones include discussions of these topics.

Abstract More than 50 different classification schemes for sandstones have been proposed over the past century. Some of those classifications are mainly related to the textural properties of clastic terrigenous deposits (and those are discussed in Chapter 6). Many of the sandstone classifications combine texture and composition, sometimes confusing the important distinction between those two characteristics. Others focus exclusively on composition, although some provide separate terms to describe textures. However, the many classifications typically share a few common characteristics. Most plot compositional data on ternary diagrams, with quartz, feldspar and lithic fragments as the poles—referred to as QFL diagrams. As always, however, the devil is in the details. For example, what qualifies as quartz (monocrystalline quartz, polycrystalline quartz, chert or other quartzose rock fragments) varies from classification to classification, as do the rock names and the percentage boundaries applied in different nomenclatural schemes.

Abstract Shale and mudstone are both widely used terms for fine-grained terrigenous clastic rocks (although some use fissility as a requirement for the use of the term “shale”), but there is at present no broadly agreed upon terminology for naming and classifying these rocks (see discussions in Schieber et al., 1998, and Potter et al., 2005). Because in past stratigraphic and sedimentologic studies, the great majority of fine-grained rocks have been designated as shales (such as, for example, the Cretaceous Eagle Ford Shale that is, in many places, a marl or even a limestone) or the Monterey Shale (a diatomite or chert depending on diagenesis), we will use the term shale in this chapter, with the understanding that it includes what some prefer to identify as mudstone. Also, because the most basic definition of shales, that they be dominated by particles smaller than 62.5 μm (e.g., Blatt et al., 1980), implies that shales span the clay-silt boundary, a good many rocks that are identified as siltstones in the literature also qualify as shales (and vice versa). The naming of any rock must on one hand convey a maximum amount of information about the rock, yet at the same time accept incompleteness for the sake of brevity. What would the key properties of a shale be with that purpose in mind? Grain size, for example, while a useful property of sandstones, lacks utility for shale because it is too small to be readily discerned. Mineral composition, though unquestionably useful, is again stymied by small grain size. XRD or whole rock chemistry data would be needed to make it workable. In fact, without advanced instrumentation, the most accessible properties of a shale are probably its color, its relative softness (lithification state), its reaction with hydrochloric acid (is it calcareous?) and textural features such as lamination, bioturbation, etc. Whereas it is not uncommon to see geologists use color charts to describe rock color, simpler qualifiers (gray, red, green, greenish, etc.) are in many instances sufficient. Relative softness can be tested be scratching the sample with a nail, and most geologists also would have some hydrochloric acid handy. So, aside from color, most of what we are able to say about a shale at first encounter is decidedly of qualitative nature, and that circumstance makes textural features very valuable when we set out to describe and categorize shales.

Abstract The term “diagenesis” refers to essentially everything that happens to sediments and rocks after their deposition but prior to metamorphism. There are a variety of diagenetic processes, biological, chemical and physical, that ultimately convert sediments into sedimentary rocks. The earliest of those events are covered in this chapter on near-surface diagenesis; subsequent chapters cover processes and products that occur primarily during later stages of diagenesis (mainly mesogenesis). Those include mechanical and chemical compaction, cementation, dissolution, replacement and structural deformation. All these processes can profoundly affect the porosity, permeability and hydrocarbon reservoir potential of clastic terrigenous deposits, and most of them are a function of initial sediment composition and the changes in pressure, temperature and water chemistry that accompany progressive burial. Less explicitly covered, but potentially no less important, is diagenesis that can occur during one or more episodes of local or regional uplift and consequent exposure (telogenesis). These events also introduce changes in the pressure/ temperature/water chemistry regime of rocks, and thus can cause major diagenetic changes, especially grain dissolution and cementation. In an attempt to address the impacts of the various diagenetic events that rocks may experience, this book includes both a section on the recognition of porosity types as well as one on paragenesis (i.e., the placement of diagenetic events into a temporal sequence related to the burial/uplift history of rocks). Burial diagenesis is critically important in controlling the porosity of clastic terrigenous rocks and is, in the main, porosity destructive—that is, almost all rocks lose porosity with increased burial depth. Nonetheless, several factors can retard or inhibit porosity loss, including early grain-coating cements, that block later overgrowth cementation, regional overpressuring of basins that reduce effective overburden stresses and, under some circumstances, hydrocarbon entry that can reduce rock-water interactions. In addition, the processes of dissolution and fracturing may, under the right circumstance, lead to actual increases in subsurface porosity. So the discussion of porosity destruction, preservation and creation pervades all chapters in the diagenesis section, and emphasis is placed on recognition of key features associated with anomalous porosity retention or creation.

Abstract Compaction is one of the major processes by which sediments lose porosity and begin the transformation to sedimentary rocks. Compaction is driven mainly by overburden loading and involves changes in the packing density of constituent grains. This is accomplished initially through grain reorientation and repacking accompanied by water expulsion from porous sediments. With additional overburden loading, fracturing and cleavage of brittle grains and plastic deformation of ductile grains contribute to increased packing density and concomitant loss of pore space. Further reduction of intergranular pore space, beyond that produced by “mechanical compaction”, results from pressure-solution processes, sometimes termed “chemical compaction”. Chemical compaction includes selective dissolution and interpenetration at grain-to-grain contacts, as well as broader dissolution along solution seams and stylolites. Although more common in carbonate rocks, chemical compaction features are widespread in clastic terrigenous deposits as well (e.g., Heald, 1955; Walderhaug and Bjørkum, 2003).

Abstract Cementation, the authigenic precipitation of minerals in pore spaces within rocks, is one of the most important processes in the lithification of clastic terrigenous deposits (we also include displacive, authigenic mineral precipitates within this general term). Cements can have a wide range of crystal sizes (terminology shown in Table 11.1) and fabrics. They can form throughout the history of sedimentary deposits, starting with surficial (eogenetic) processes in marine and nonmarine settings and continuing through all stages of burial (mesogenetic) diagenesis as well as uplift-associated (telogenetic) diagenesis. Hundreds of different minerals are found as cements in the panoply of different sandstones and mudrocks. However, most typical sandstones and mudrocks contain perhaps one to five cementing minerals, making identification far less complex than it might appear. All cements form by precipitation of materials from aqueous solution, and variations in subsurface fluid temperatures, pressures and chemistries (pH, salinity, specific ionic abundances, etc.) are the major controls on which minerals are precipitated or dissolved. Solutes can be derived from many sources. Some may come directly from seawater or via influx of meteoric waters; others may come from reflux of evaporitic brines. Additional solutes come from circulating basinal fluids, through chemical dissolution of soluble minerals, pressure solution along stylolites and solution seams, maturation of organic matter or dehydration of gypsum beds. These subsurface fluids can be moved through the sedimentary section at basinal scales via compactional dewatering or thermal convection, commonly aided by permeability “highways” created by fractures and faults. In all settings, cements may form through local dissolution or alteration of unstable minerals, relatively small-scale diffusive transport and nearby reprecipitation of more stable minerals. In very low-permeability settings, that may be the only viable mechanism for cementation. On the other hand, the near-complete exclusion of water (or water contact with grains), as in some hydrocarbon reservoirs, may inhibit both water movement and cement formation.

Abstract Detrital feldspars most commonly undergo dissolution or alteration and replacement in subsurface settings, yet in many sandstones one also can find examples of authigenic feldspar cements. Such cements, in almost all cases, are either albite (the sodic end-member feldspar) or orthoclase (the K-spar end member)—see Chapter 2 for more details. Such feldspar cements commonly occur as thin or irregular overgrowths (not always in optical continuity with their detrital host grain) or as fills of microfractures within feldspar grains; however, they can form major pore-occluding cements in some clastic terrigenous rocks. The overgrowths can be monocrystalline or polycrystalline (sometimes even consisting of a mosaic of micron-sized rhombs—Worden and Rushton, 1992; De Ros et al., 1994). Despite their generally minor role in cementation of sandstones, the recognition of authigenic feldspars is important in understanding burial-related pore-fluid chemistry variations through time, especially because such overgrowths (K-feldspar overgrowths in particular) can be radiometrically dated using K–Ar and 40 Ar/ 39 Ar methods (e.g., Hagen et al., 2001; Mark et al., 2005 and 2008). In this regard especially, it should be noted that feldspar compositional zonation and overgrowth can be inherited from igneous, metamorphic or sedimentary source rocks, and such “overgrowths” must be carefully distinguished from authigenic cement precipitation during burial (just as recycled quartz overgrowths must be distinguished from ones produced in situ during burial).

Abstract Clay minerals are a complex family of aluminosilicates, and generalized chemical formulas for them can be found in Table 11.2 (page 247). They can be either platy or fibrous with a high degree of chemical substitution. The word “clay” also has grain size connotations, so when referring to clay cements, it is best to call them “clay minerals”. Clay minerals occur in siliciclastic rocks as detrital and/or diagenetic components, which are commonly difficult to differentiate. Diagenetic clay minerals, the focus of this chapter, form in several ways: 1. Alteration of unstable silicate minerals, such as feldspars; 2. Pseudomorphic or neomorphic transformation of detrital or precursor diagenetic clays; or 3. Direct precipitates. They are of great interest to the oil and gas industry, because they can have a significant impact on sandstone reservoirs, commonly lowering porosity and permeability and increasing the possibility of formation damage. However, in many cases, clays (especially chlorite) form eogenetic to early mesogenetic grain-coating cements that can impede later cementation and preserve exceptional porosity at depth (e.g., Pittman et al., 1992; Ehrenberg, 1993; Anjos et al., 2003; Berger et al., 2009; Gould et al., 2010; Ajdukiewicz and Larese, 2012).

Abstract Zeolites are a large and complex group of hydrated aluminosilicate minerals that are among the most abundant authigenic silicates in sedimentary deposits. They are so complex that only a superficial summary can be given in this short introduction, and readers are urged to consult the vastly more detailed references provided at the end of the chapter. Both water molecules and cations can be substituted or replaced in most zeolites without disrupting the crystal structure, accounting for their great compositional diversity as well as their widespread use in ion-exchange applications. Their classification and nomenclature has undergone, and continues to undergo, extensive review and revision (e.g., Coombs et al. 1998; McCusker et al., 2001). Although more than 30 different zeolites have been noted from sediments and sedimentary rock, only five are common (analcime, clinoptilolite, heulandite, laumontite and phillipsite) and three others (chabazite, erionite and mordenite) are relatively common (Gottardi and Galli, 1985; Hay and Sheppard, 2001).

Abstract Carbonate minerals, particularly calcite, dolomite, ankerite and siderite, are important as cements and replacements in sandstones and mudstones. Perhaps surprisingly, carbonate cements can be difficult to work with in terms of quantifying conditions of formation (including time, temperature and water chemistry). There are two reasons for this. First, some carbonate minerals (especially dolomite in the context of cements in clastic terrigenous deposits) can recrystallize at one or more stages, resetting their diagnostic fabrics and elemental and isotopic composition. Second, all carbonate cements are relatively soft minerals with multiple cleavages. That means that fluid inclusions, when subjected to temperatures higher than those at which they formed, can stretch or leak, resetting apparent temperatures of formation (Goldstein, 1986 and 2001). That said, there is abundant evidence that carbonate cements form at virtually any stage of diagenesis, from synsedimentary to deep burial (see overview in Morad, 1998).

Abstract Sulfates and halides can form significant cements in sandstones but typically only in very specific settings. Early formed gypsum, anhydrite and halite are common in arid, highly evaporitic settings and form not only cements but also bedded deposits and displacive or replacive precipitates. Additional sulfates and halides, such as glauberite (a sodium calcium sulfate), barite (BaSO 4 ), celestine (SrSO 4 ), sylvite (KCl) and many others, can form in surficial settings, but most such minerals are rare and may not survive subsequent diagenetic alterations because of their generally high solubility. Eogenetic sulfates and halides are especially common precipitates in arid-region playa lakes, dunes and interdune flats, as well as continental and coastal sabkhas (e.g., Kinsman, 1969; Benison and Goldstein, 2000; Warren, 2006). Gypsum is the most common of these minerals, and extensive gypsum cementation, mostly within 10 m (33 ft) of the ground surface, can produce gypsiferous soils, gypsum crusts and fully cemented gypcretes (Watson, 1985; Nettleton, 1991; Hartley and May, 1998). Gypsum cements in terrigenous strata can consist of coarsely poikilotopic crystals (encompassing numerous detrital sand grains), isolated lenticular crystals or crystal clusters (desert roses), gypsarenites with graded cement sizes and alabastrine gypsum with individual crystallites that average less than 50 μm in diameter (Watson, 1985). Although gypsum is the major precipitate in such surficial deposits, in settings with especially hot, arid, evaporitic and low humidity conditions, anhydrite (anhydrous calcium sulfate)andevenhalitecanbefoundcementingsurficial and near-surficial sands and muds, especially in sabkhas and saline pans (Kinsman, 1969; Casas and Lowenstein, 1989; Warren, 2006).

Abstract The type and abundance of minerals that we observe in the earth’s crust and its sedimentary cover is governed by elemental abundances and thermodynamic mineral equilibria. Iron is fourth in abundance (e.g., Mason, 1966) by weight (∽ 5%) after oxygen, silica, and aluminum (∽82.5% cumulatively), and whereas that relationship readily explains the preponderance of silica and clay minerals as cements in sedimentary rocks, the story for iron is a bit more complicated. Due to its multiple redox states, iron can form (or be part of) minerals in oxidizing as well as reducing environments, and the main “sinks” in the sedimentary rock record are Precambrian banded iron formations, Phanerozoic ironstones and continental red beds. Iron is generally supplied to sedimentary basins in the form of iron hydroxide coatings on fine particles (Carroll, 1958) and as iron silicates in the sand fraction (Walker, 1967). In fluvial sediments, the commonly oxidizing pore water conditions result in intrastratal alteration of detrital iron silicates (pyroxene, hornblende, biotite) and in the precipitation of iron hydroxides (limonite, goethite; yellow-brownish color) within pore spaces. Over time, iron hydroxides in both fine and coarse sediments are converted to hematite that gives the rocks their characteristic red color (Walker, 1967). The timing of hematite formation can, under some circumstance, be dated using paleomagnetic information (e.g., Lu et al., 1994) or geochronologic data (e.g., Reiners et al., 2014).

Abstract This last section on cementation provides images that feature some of the less common mineral cements that were not covered in previous chapters. It is by no means an all-encompassing chapter because there are simply too many minor detrital mineral overgrowths or volumetrically insignificant, fully authigenic precipitates to cover in detail.