Diagenesis

Synonyms

Diagenesis classified by setting and evolutionary stage ofsedimentary basins: Eogenetic or eodiagenesis (near surface and shallow burial), mesogenetic or mesodiagenesis (deeper burial), and telogenetic or telodiagenesis (uplifted succession) (Choquette and Pray 1970).

Diagenesis classified by process: Syndiagenesis (biogeochemical processes at the sediment-water interface through shallow burial), anadiagenesis (dominantly physicochemical processes under deeper burial or orogenic conditions), and epidiagenesis (biogeochemical processes associated with fluid flow during uplift) (Fairbridge 1967).

Catagenesis (late, deep-burial diagenesis, referred by some as “burial metamorphism” as it incorporates the earliest stage of metamorphism).

Definition

Diagenesis is the sum total of physical, chemical, and biological processes that occur in sediments and sedimentary rocks from immediately after deposition through to the metamorphic realm. No universal definition exists for diagenesis and the term has evolved since it was defined nearly 150 years ago (de Segonzac 1968). It is generally agreed that diagenetic processes occur under Earth surface conditions (~0–30 °C and 1 bar of pressure) to temperatures of ≤250 °C and pressures of up to 2.5 kb (7 km) involving a broad range of fluid compositions from fresh water to concentrated brines (Fig. 1). Diagenesis involves both the conversion of sediments into sedimentary rock and the subsequent changes that occur prior to entering the metamorphic realm. There is no distinct boundary between the culmination of diagenesis and the onset of metamorphism nor between early diagenesis and weathering. The process of diagenetic conversion begins at the sediment-water interface in marine and terrestrial depositional sites, which have sufficient accommodation space to accumulate and preserve sediments, and continues to influence mineral and organic matter throughout its burial in sedimentary basins and associated tectonic history. Given the high reactivity of many minerals and organic matter within sediments and sedimentary rocks and the evolving nature of pore fluid temperature and chemistry, diagenesis is a continually active process with a protracted history of physical and chemical processes that lead to dissolution, chemical transformation of mineral components, precipitation of new mineral cements, and maturation of organic matter. Many resources such as coal, oil, gas, and ore deposits result from diagenetic processes.

Diagenesis, Fig. 1

Matrix illustrating different aspects of diagenesis. The matrix is arranged with increasing depth (temperature) toward the bottom, and different sedimentary components in the columns (Modified from Burley et al. (1985))

The Diagenetic Cycle

The diagenetic cycle begins immediately following deposition on the ocean, lake, or channel/floodplain floor and continues through shallow- to deep-burial, and for many sediments ends with exposure above sea- or base-level where erosion, chemical weathering, and other diagenetic processes take place. Water and organic matter decomposition are the agents of diagenesis providing the fuel for the chemical and microbial processes that act on newly deposited detrital siliciclastic materials and chemically precipitated minerals, such as carbonates, amorphous silica, and evaporites (Ali et al. 2010). The course of diagenesis is dictated by sedimentary factors such as particle size and mineralogic composition, fluid-rock ratio, fluid chemistry and flow rates, organic matter content, the presence of microbial communities, as well as environmental conditions (temperature, fluid chemistry, pressure). In turn, diagenetic processes can change the chemical properties of the pore fluids over time, causing additional changes to mineralogy, fluid composition, and petrologic properties. Figure 1 summarizes many aspects of diagenesis. Figure 2 places these processes in a schematic cross section and provides a framework for the conditions under which diagenesis occurs.

• The early diagenetic realm (eogenesis or eogenetic zone, Fig. 2) occurs during burial to a few hundred meters and involves changes to the sediment, organic matter it hosts, and the interstitial fluids (pore water) by surface-related processes under near surface temperatures (≤30 °C). Diffusion of gases (e.g., oxygen, carbon dioxide, hydrogen sulfide, and methane) can also influence the geochemical environment. Early diagenesis typically involves dewatering of fine-grained sediments by gravitational compaction, bioturbation by burrowing organisms, dissolution of unstable minerals, diffusion of dissolved cations (e.g., Ca²⁺, Mg²⁺, Fe²⁺, Mn²⁺, Sr²⁺), anionic species and gases (e.g., SO₄²⁻, HCO₃⁻, O₂, CO₂, CH₄, H₂S) through the sediment, bacterial decomposition of organic matter, transformation (neomorphism or recrystallization) of unstable minerals, and precipitation of authigenic carbonates, oxides, hydroxy-oxides, aluminosilicates, sulfates, and sulfides. Sedimentation rate controls the degree of physical compaction and dewatering of sediments as well as the depths of bioturbation and cation/molecule diffusion rates. Most early diagenetic chemical reactions are driven by the reactivity of metastable minerals (e.g., amorphous silica, aragonite, high Mg-calcite, dolomite) or phyllosilicates. Many early diagenetic processes involve the kinetic boost by bacteria, referred to as bacterial mediation. Empirical studies of the early diagenetic realm involve outcrop studies and coring of marine and lake sediments and analysis of sediment mineralogy, texture, and grain size distribution, pore-water chemistry, and in recent years, the composition of associated microbial communities.

• Sediments continue to undergo geochemical modification in the burial diagenetic realm (mesogenesis or mesogenetic zone, Fig. 2) involving processes that are no longer directly related to the surface including chemical compaction (pressure solution), dissolution, precipitation of authigenic minerals (i.e., cementation), and organic matter maturation. This diagenetic realm is commonly partitioned into the intermediate burial (between 500 and 2000 m depth) and deep burial (≥ 2 km) zones. Elevated temperature and pressure play a major role in driving the physicochemical reactions in the burial diagenetic realm. The chemical composition of pore waters evolves with depth, typically becoming more saline, due to fluid-rock interactions such as dewatering of clays and dissolution of buried evaporites. Maturation of organic matter releases organic acids and contributes toward lowering pH and changing redox conditions of interstitial fluids (Surdam and Crossey 1987; Harrison and Thyne 1992). Consequently, minerals, which were initially stable in the shallow diagenetic realm, become metastable with burial driving dissolution and precipitation of authigenic minerals such as calcite, dolomite, anhydrite, silica, zeolites, and clay mineral transformations. Overall, cementation is thought to dominate over dissolution in the burial diagenetic realm (Ali et al. 2010).

• Exposure of sedimentary rocks above base level by tectonic uplift (telogenesis) or sea-level fall exposes them to atmospheric processes and physical erosion and chemical weathering. Meteoric (fresh, dilute) waters can penetrate deeply into exposed or tectonically uplifted sedimentary rocks through permeable units or fracture porosity and develop extensive karst and dissolution features in carbonate-dominated successions (telogenetic zone, Fig. 2).

• Tectonic forces and sea level fluctuations lead to cycling of sediments and sedimentary rocks between the diagenetic realms resulting in multiple cycles of dissolution, cementation, mineral transformation, and recystallization until a steady state or equilibrium is reached. The extent and nature of diagenesis will reflect the reactivity of the initial sediment, its fluid-rock interaction history including fluid-rock ratios and exposure time to each diagenetic zone.

• Where multiple diagenetic events have influenced a sedimentary rock, the paragenetic sequence of these events (Fig. 3) can be reconstructed by integrating microscope (thin sections to scanning electron microscopy) and geochemical studies of the sedimentary rocks and their fluid inclusion waters as well as associated organic matter. Diagenetic transformation of sedimentary rocks can substantially impact their porosity and permeability and in turn their fluid conduit and aquifer or reservoir potential (Montañez 1997).

  Diagenesis, Fig. 2

  

  Summary of diagenetic realms. Sediments deposited in Earth surface environments pass through the diagenetic cycle post deposition, which can include the eogenetic zone (early diagenetic realm), the mesogenetic zone (burial diagenetic realm), and the telogenetic zone (exposure) (Modified from Ali et al. (2010))

  Diagenesis, Fig. 3

  

  Photomicrographs of carbonate. (a) Stained (Alizarin Red and potassium ferricyanide) thin section showing calcite cement sequence (paragenesis) in a calcite spar cement that occludes a primary pore developed between grains (pink); thin marine cement and freshwater cement rims (pink staining are developed on grains. White triangular cement lining pores are dolomite. (b) Cathdolouminescent zoning in calcite cement occluding porosity in a fracture (or between grains). The alternating bright luminescent and nonluminescent banding of the cement records the migration of multiple generations of calcite-precipitating fluids

Methods Applied to Study of Siliciclastic and Carbonate Rocks

Petrographic study of thin sections employing the optical microscope forms the basis for understanding the composition (mineralogy), spatial and textural relationships between detrital grains, authigenic minerals, and pore geometries (Fig. 3). Optical examination is aided by specialized preparation (impregnation of the rock with colored epoxy to highlight porosity) and staining techniques for recognition of specific carbonate minerals (Fig. 3a) and feldspars as well as calcite cements of differing Fe²⁺ concentration (Nelson and Read 1990; also see Humphries (1992) for a guide to thin section preparation methods). Cathodoluminescence study of thin sections (Fig. 3b) is commonly used to further identify textures and to define a paragenetic sequence for a given sample or suite of samples. A paragenesis is the sequence in which different minerals or generations of a given mineral (e.g., calcite) were precipitated. Carbonate cements exhibit varying shades of brown, red, yellow, and orange under cathodoluminescence due to incorporation of varying amounts of Fe²⁺ and Mn²⁺ into the crystal lattice. Cathodoluminescent zoning in carbonate cements reflects the Mn²⁺ to Fe²⁺ ratio of a given generation of carbonate cement, and in turn that of the fluid from which it precipitated. Authigenic feldspar overgrowths and cements are nonluminescent, whereas detrital cores and grains exhibit blue luminescence. Discrete generations of carbonate cements typically are identified by their coupled staining and cathodoluminescence (Fig. 3).

X-ray diffraction (XRD) is the preferred method for structural identification of fine-grained minerals, particularly in mudrocks. Moore and Reynolds (1989) provide a summary of preparation and analysis: Brindley and Brown (1984) and Newman (1987) provide detailed descriptions of clay minerals.

Geochemical analysis may include elemental and isotopic (both stable and radiogenic) approaches. While a comprehensive survey of all methods is beyond the scope of this overview, several common techniques will be mentioned. Electron microbeam techniques are powerful tools for in situ chemical and mineralogical examination of sedimentary rocks of all types. Scanning electron microscopy (SEM) is widely used to understand three-dimensional morphology, pore geometries, sequence of mineral growth, and more recently, microbe/mineral interactions (Banfield and Nealson 1997). Electron microprobe analysis (EMPA) provides quantitative chemical analyses of phases typically of 10 micron and greater size though specialized instruments may target smaller regions. This tool is generally restricted to major and minor element analysis. The secondary ion microprobe (SIMS) has gained wide use in geoscience applications (see review in Stern 2009) and provides quantitative trace element and stable isotopic analyses in spatial context. Methods involving laser ablation in tandem with inductively coupled plasma mass spectrometry are providing new approaches to both elemental and isotopic analysis of sedimentary materials with excellent opportunities for application to diagenesis research (see examples in Rasbury et al. 2012). Integration of petrographic features with elemental and isotope analysis (carbon, oxygen, sulfur, and nitrogen) of authigenic phases, at times coupled with fluid inclusion study, has been effective in understanding conditions of formation and paleofluid composition in some cases (see summaries in Sharp 2007). Banner (2004) provides a comprehensive overview of applications of radiogenic isotopes to sedimentary systems.

Carbonate Diagenesis

Carbonates are minerals with a structural CO₃²⁻ group and Ca²⁺, Mg²⁺, or a combination of these two elements. Weight percent Fe²⁺, Mn²⁺, and Sr²⁺ occur in less common carbonates. Most carbonate sediments form in marine settings commonly with coexisting polymorphs of aragonite and low- (≤5% MgCO₃) to high-Mg calcite (5–20% MgCO₃), as well as rarely dolomite (Moore 1989). Terrestrial carbonates form as low-Mg calcite and siderite in lacustrine, wetland, and spring deposits, low-Mg calcite and sphaerosiderites in soils, and low- to high-Mg calcite and aragonite in travertines and cave speleothems. This entry focuses on marine carbonates given their prevalence in the modern and geologic record; see reviews by Ford and Pedley (1996) and Capezzuoli et al. (2014) for freshwater carbonates.

Carbonate mineralogy in the marine realm is determined thermodynamically by mineral stability (solubility) and kinetically by substrate mineralogy, biological effects, differences in precipitation rates, and the presence/absence of organic compounds and inorganic inhibitors (e.g., SO₄²⁻, PO₄³⁻, Mg²⁺) (Burton 1993). Biological effects include microbial mediation and vital effects. With the exception of low-Mg calcite, all of these carbonate mineralogies are thermodynamically metastable (high Mg-calcite, aragonite, nonstoichiometric dolomite) at Earth surface conditions and will convert to stable low-Mg calcite or stoichiometric dolomite during subsequent exposure to evolved marine pore waters, introduction of meteoric fluids, and/or interaction with saline brines during burial (Walter 1985).

Early Diagenetic Realm

For shallow-water carbonates, extensive diagenesis occurs proximal to the sediment-fluid or atmosphere interface due to repeated flushing by seawater or freshwater (eogenetic zone, Fig. 2). Their metastable mineralogy makes them highly susceptible to post-depositional mineralogical and geochemical alteration as the pore fluid chemistry evolves with fluid mixing and chemical reactions (Moore 1989; James and Choquette 1990). Diagenesis at the sediment-seawater interface begins with boring by endolithic algae, sponges, and fungi, which replace primary skeletal carbonate with clay-size carbonate precipitates (micrite), a process called micritization. Within the first few centimeters to meters of burial, dissolution of high-Mg calcite and aragonite releases bicarbonate (HCO₃⁻) that along with fluid CO₂ degassing promotes neomorphism and recrystallization of grains, carbonate cementation, and calcite overgrowths on biotic templates. Neomorphism involves the molecular-scale replacement of an unstable phase of carbonate by thermodynamically more stable low-Mg calcite. Recrystallization of grains and early cements involves the wholesale dissolution of the original carbonate and simultaneous precipitation of a more stable phase resulting in a possible volume change and loss of original textures. Recrystallization occurs in both carbonates and siliciclastic deposits. These early diagenetic processes lead to preservation of much of the original fabric and retard compaction in shallow-water marine carbonates. Dissolution of aragonitic fossils and matrix creates secondary moldic or vuggy porosity.

In contrast to shallow-water carbonates, the predominance in deep-sea sediments of stable low-Mg calcite causes delay in dissolution and cementation. The presence of small amounts of calcite overgrowths and cements on shells and shell recrystallization indicates localized regions of carbonate dissolution and diffusion/advection to areas of calcite cementation (Moore 1989). Oxidation of organic matter buried with carbonate sediments provides HCO₃⁻ to pore fluids for carbonate precipitation. Pore water profiles of the oxygen (δ¹⁸O) and calcium (δ⁴⁴Ca) isotopic composition and trace element concentrations are used to model the rates of carbonate recrystallization in deep-sea sediments (Fantle et al. 2010). Gravitational compaction of deep-sea sediments in the shallow burial realm (within the first 200 m of burial) is a major influence on lithification and leads to substantial loss of porosity (from ~80% to ≤50–60%).

The diagenetic evolution of carbonate sediments, which were deposited on the slopes surrounding shallow-water carbonate platforms, falls between that of deep-sea sediments and shallow-water marine carbonates. This reflects the mixing of a pelagic low-Mg calcite source with a platform-shed source of reactive aragonite and high-Mg calcite as well as overall higher concentrations of organic matter (Swart 2015). Thus, periplatform carbonates may undergo extensive dissolution and cementation during initial burial making them more resistant to compaction than deep-sea sediments.

Meteoric diagenesis: The vast majority of carbonates are deposited during sea level highstands when shallow water platforms are inundated (Ali et al. 2010). These carbonates are subsequently exposed to freshwater, which is undersaturated with respect to carbonates, during subsequent drop in base level whether driven by eustasy or tectonic uplift (telo- or epidiagenesis). Meteoric diagenesis occurs over a series of near-surface (eogenetic) diagenetic zones (Fig. 2) including soil formation and karstification and vadose, phreatic, and mixing zones (see reviews by Moore 1989; James and Choquette 1990). Overall, freshwater reactions with metastable carbonates result in extensive stabilization (through neomorphism or recrystallization) of carbonate grains, superimposed generations of dissolution and low-Mg calcite cementation, as well as substantial redistribution of pore space. The degree of diagenetic alteration is dictated by the reactivity of the initial mineralogy, the pore fluid chemistry, fluid-rock ratios, and flow rates.

Diagenesis in the freshwater undersaturated vadose zone is characterized by seepage flow through intergranular porosity and conduit flow through solution-enlarged fissures (Fig. 2). Vadose waters are typically CO₂-charged and chemically aggressive given exchange with the atmosphere and with soil-formed CO₂. Rapid conduit flow leads to extensive, non-mineral selective dissolution. Rates of calcite precipitation are far slower than fluid flow rates, thus as dissolution proceeds fluids become oversaturated with respect to low-Mg calcite and precipitate cements down-flow (Moore 1989). Freshwater that percolates slowly through the vadose zone by diffuse flow promotes mineral-controlled diagenesis, which is controlled by the differing solubilities of the carbonate polymorphs. Aragonite is 1.5 times more soluble than low-Mg calcite at a given temperature, whereas high-Mg calcite, whose solubility increases with Mg content, can be up to an order of magnitude more soluble than low-Mg calcite (Walter 1985). Dissolution of more soluble carbonate under the slow diffuse flow rates leads to oversaturation with respect to low-Mg calcite proximal to the site of dissolution driving widespread fabric-retentive neomorphism, recrystallization, and precipitation of vadose calcite cements (e.g., pendant and meniscus cements, crystal silt, fine crystalline equant cements unevenly lining primary porosity). The end-product can be wholesale transformation of metastable carbonates to thermodynamically stable low-Mg calcite while retaining much primary texture and fabric. This conversion process, however, is slow (on the order of 10⁶ year) reflecting the overall low water-rock ratios of the vadose zone.

The water table (Fig. 2) delineates the boundary between the vadose and underlying saturated phreatic zone. High flow rates and turbulent exchange at the water table promotes CO₂ dissolution into water or high CO₂ degassing rates making this region a dynamic zone of dissolution and low-Mg cementation (Moore 1989; James and Choquette 1990). The phreatic zone hosts the bulk of dissolution and cementation by freshwater. Exposure of marine carbonates due to sea level fall leads to deep penetration of meteoric water given that the freshwater phreatic lens will develop to depths of ~40 times the height of the water table above sea level (Ghyben–Herzberg principle). Diagenetic processes in the shallow phreatic zone are orders of magnitude faster than in the vadose zone reflecting overall higher flow rates and fluid-rock ratios. In this zone, low-Mg calcite cements form isopachous rims of equant to bladed crystals on grains and within moldic porosity and microspar cements partially to fully occlude porosity. Relatively rapid flow rates lead to large-scale transport of dissolved carbonate and Ca²⁺ down-flow resulting in extensive dissolution updip and cementation downdip.

Mineralogic stabilization within the phreatic zone (Fig. 2) occurs rapidly (10⁴–10⁵ year) relative to the vadose zone. Once stabilized, subsequent diagenetic modification is limited to minor cementation within conduits that focus fluid flow. The mixing zone (Fig. 2), at the interface between the freshwater phreatic zone and the underlying deep phreatic zone with saline continental or marine interstitial fluids, is a chemically active region given that physical and diffusive mixing of waters leads to nonlinear changes in carbonate mineral saturation. The mixing zone is a region of abundant secondary porosity and subsurface karst development (e.g., Boulder zone in the Tertiary Florida aquifer or the ceynotes of the Yucatan Peninsula). The different zones of meteoric diagenesis have characteristic stable (O and C) isotopic and trace element signatures that when coupled with stratigraphic and petrographic studies have been used to reconstruct the diagenetic history of carbonate successions. The reader is directed to Moore (1989) and Swart (2015) for a comprehensive review of the geochemistry of meteoric diagenesis. The deep phreatic zone (Fig. 2) is a shallow burial locale of substantially slower diagenesis given the slow fluid flow rates and low fluid-rock ratios. Low-Mg calcite and dolomite cements that form in this zone are typically large and blocky reflecting their slow growth rates under very low saturation states.

The meteoric diagenetic realm is dynamic with fluctuations in base level, whether driven by eustasy or tectonics, that induce repeated large-scale vertical migration of the marine, vadose, phreatic, and mixing zones through a sedimentary succession over time (Read and Horbury 1993). Throughout Earth history, sea level has fluctuated 10 s to 100 s of m over a spectrum of timescales (10⁴–10⁷ -year) driven by orbitally forced waxing and waning of ice sheets, atmospheric pCO₂ variability, and tectonic and mantle processes. These shifts in the meteoric diagenetic zones impart a mineralogic and geochemical signature that is characteristic of the magnitude of relative sea-level fluctuations and the climate regime under which they occurred (Read and Horbury 1993). Zonation in low-Mg calcite cements, discernible petrographically under transmitted light or by cathodoluminescence (Fig. 3) as well as by their geochemical compositions (Fe, Mn, Mg, concentrations and δ¹⁸O and δ¹³C isotopic compositions), record changes in fluid conditions (redox, temperature, salinity) through time from which the spatial migration of the near-surface diagenetic zones can be reconstructed (Niemann and Read 1988; Nelson and Read 1990). In many cases, zoned phreatic cements capture multiple cycles of fluid change within a pore providing a strip chart of the eustatic fluctuations that affected a given carbonate succession (Read and Horbury 1993) even when deeply eroded by post-depositional processes (Bishop et al. 2009).

Intermediate-to-Deep Burial Diagenetic Realm

Ultimately, most sedimentary deposits are buried sufficiently (>1 km) by overlying sediments and rocks that they are introduced to the intermediate-to-deep-burial diagenetic realm (herein referred to as “burial diagenetic” or mesogenetic zone). The burial diagenetic zone (Fig. 2) begins proximal to the liquid oil window in hydrocarbon source rocks (Machel 2005). Formation temperatures are elevated above ~50 °C reaching conditions of over 100–200 °C, which enhances diagenetic reactions. The burial diagenetic realm is characterized by further chemical compaction and dissolution, mineral stabilization, and precipitation of cements. Depending on the clay mineral content and previous cementation history, chemical compaction can become a predominant process with the degree of pressure solution influenced by pore-water composition, mineralogy, degree of previous cementation, and the presence of organic matter. Grains or primary marine cements or earlier diagenetic precipitates, which were not previously stabilized, become thermodynamically metastable in the deep subsurface due to changing fluid conditions driving further recystallization and mineral replacement, including widespread dolomite replacement of limestone host rocks (Machel 2005).

Interstitial (formation) fluids in the deep-burial diagenetic environment range from brackish (< 35 weight percent) to hypersaline reflecting their interaction history with sedimentary rocks, including evaporites and siliciclastics, along their flow path and mixing with organic-rich fluids, which formed during hydrocarbon maturation. Fluid flow rates vary from near stagnant to moderately fast if an external hydraulic gradient is established by regional groundwater flow systems or dewatering of shale basins. At temperatures proximal to and within the oil window (~80–120 °C), thermal maturation of organic matter (decarboxylation) and other mineral reactions occurring in situ or in shale basins introduce organic-acid, CO₂-rich brines that can lead to extensive non-fabric-specific dissolution (Montañez 1994, 1997; Machel 2005). As these fluids evolve down-flow through wholesale dissolution and pressure solution, they become saturated and precipitate Fe- and Mg-rich calcite cements and Ca-rich dolomites (Montañez 1994). Dolomite, anhydrite, and silica (chert) replacement of limestones is common in the deep-burial diagenetic realm. Exposure of evaporite-rich deposits to hydrocarbon-bearing fluids at elevated temperatures promotes thermochemical sulfate reduction, which produces significant amounts of H₂S and water and precipitation of low-Mg calcite and pyrite, and possibly saddle dolomite (Swart 2015). At temperatures <80–100 °C, bacterial sulfate reduction dominates when dissolved sulfate is exposed to hydrocarbons.

Ultimately, tectonic processes may uplift deeply buried carbonate successions and expose them to meteoric conditions. For those carbonates, which were previously stabilized and lithified in near-surface to deep-burial diagenetic zones, the extent of dissolution and cementation in this telogenetic zone (Fig. 2) will be greatly reduced. Rather, diagenesis will occur as karstification and development of fluid conduits associated with fractures and joints.

Dolomitization

The origin of dolomite, particularly early diagenetic dolomite, remains elusive despite over a century of empirical and experimental studies. This reflects that dolomite (CaMg(CO₃)₂) is rare in modern environments but very common in ancient carbonate successions. Furthermore, constraints on the chemical and hydrologic conditions of dolomite formation are limited and the kinetics of dolomite precipitation are poorly understood (Machel 2004). Although seawater is ~1000 times oversaturated with respect to dolomite, it rarely forms in the marine realm given the need for a source and mechanism of steady Mg²⁺ supply (Swart 2015) and mechanism to overcome kinetic inhibitors (e.g., SO₄²⁻). Thus, dolomite precipitation (i.e., primary dolomite) and replacement of CaCO₃ (i.e., dolomitization) at earth surface conditions require a mechanism to raise the saturation state of dolomite and provide the necessary high fluid-rock ratios for steady Mg²⁺ supply. Highly evaporative environments (arid climate tidal flats) and refluxing of evaporatively sourced brines through shallow-water carbonates (Montañez and Read 1992a), possibly in mafic parent-material or seasonally very dry soils (Capo et al. 2000) and fluids in which the alkalinity is increased by oxidation of organic matter provide such a mechanism for “early” dolomitization (Fig. 1; Machel 2004; Swart 2015). The fluid mixing zone at the interface between the freshwater phreatic lens and underlying deep phreatic zone has been invoked as a locale of dolomite formation (Fig. 2) given that fluid mixing leads to undersaturation with respect to calcite but high supersaturation with respect to dolomite. Dolomite, however, is absent from most modern-day mixing zones and ancient examples have been contested (Machel 2004; Swart 2015). Bacterial mediation is likely involved in the precipitation of many near surface dolomites (Vasconcelos et al. 1995), in particular in highly evaporative environments and/or organic-rich deposits. Early formed dolomites are typically poorly ordered and Ca-rich (non-stoichiometric) and volumetrically minimal. These disordered and Ca-rich dolomites undergo stepwise recrystallization throughout burial, becoming more ordered and stoichiometric with time and with exposure to dolomitizing fluids (Montañez and Read 1992b). Progressive recrystallization overprints their original geochemical signature although a memory of the primary composition can be preserved.

Extensive dolomitization, however, likely occurs in the intermediate- (100 s m) to deep-burial realm (up to 1000s of m) (Fig. 1). This reflects the elevated temperatures needed to overcome kinetic inhibitors, the very high fluid-rock ratios, the hydrologic plumbing to provide the Mg²⁺ and CO₃²⁻ supply, and the time required to produce volumetrically large (massive) dolomites (Machel 2004). Normal seawater may be capable of extensive dolomitization at slightly elevated temperatures (50–80 °C) and moderate depths (few 100 meters) in regions of large-scale thermal circulation through isolated carbonate platforms (Kohout convection). Regional or basinwide dolomitization, however, likely develops through multiple generations of dolomite replacement of limestones and precipitation of dolomite cements (Montañez 1992, 1994). Hydrologic models invoked for regional dolomitization, in addition to thermal convection, include tectonically induced flow of basinal brines, funneled compaction flow, and rapid upward advection of hydrothermal fluids along faults (summarized in Machel 2004), although the latter process is likely limited in spatial scope. Integration of stratigraphy, petrographic (textures and cathodoluminescent zoning) features, and geochemical compositions of dolomite and its fluid inclusion waters has been effectively used to unravel the paragenesis of multigeneration massive dolomites (e.g., Montañez 1994) and to reconstruct the hydrology of dolomitizing environments (studies summarized in Machel 2004; Swart 2015; and Montañez 1992, 1994). Massive dolomites (dolostones) in the subsurface tend to have higher porosities than limestones making them excellent reservoirs (Montañez 1997). Enhanced porosity is attributed to (1) a 12% porosity increase during replacement given the smaller molar volume of dolomite relative to calcite, (2) dissolution of original calcite during burial dolomitization or during subsequent migration of low pH brines, and/or (3) thermochemical sulfate reduction.

Siliciclastic (Sandstones and Mudrock) Diagenesis

The diagenesis of siliciclastic grains encompasses the full range of grain sizes found in sedimentary systems. Siliciclastic systems usually are composed of a primary population of detrital grains and primary porosity. A summary of the controls on initial composition of clastic grains may be found in Johnsson (1993). Common physical diagenetic changes in siliciclastic sediments include compaction, bioturbation, and tectonic deformation. All three of these processes may result in the rearrangement of primary grains, hence influence the porosity and permeability of the unconsolidated sediment. Compaction acts to reduce primary porosity, and may reduce both inter- and intragranular volume. Deformation may result in crushing of primary grains as well as cross-cutting fractures in lithified rocks. Chemical diagenetic changes in siliciclastic systems involve fluid-rock interaction and result in mineral precipitation, dissolution, or recrystallization. Examples include recrystallization of metastable or unstable phases, precipitation of new, authigenic phases as grain replacement or intergranular cement, and dissolution of unstable phases resulting in secondary porosity. While many cements and mineral reactions appear to result from predictable, equilibrium processes, it is common for siliciclastic sediments and rocks to be in an overall state of chemical disequilibrium due to the presence of a wide range of primary constituents and the relatively low temperatures (hence slow reaction rates) particularly during early diagenesis. Figure 1 highlights some important depth-related changes in clastic rocks. Original grain composition, increasing temperature and pressure, the hydrodynamic conditions, and the presence of reactive organic matter all play an important role in determining the extent of diagenetic modification.

Early Diagenesis

During early diagenesis, the precipitation of carbonate minerals (especially, calcite, ankerite, and siderite), quartz, and aluminosilicates such as kaolinite, smectite and illite predominate. If oxidizing conditions are encountered (as through active meteoric recharge in the near-surface (eogenetic) environment), oxides and hydroxides of iron and manganese may form as well (these include ferrihydrite, goethite, limonite, todorokite, and pyrolusite). High water throughput or excess acidity during early diagenesis may also result in selective dissolution of primary grains (especially carbonates) or alteration of lithic and feldspar grains to produce secondary porosity in addition to authigenic clays. Some primary clastic assemblages, for example, volcanic materials including glass, result in high pH conditions that foster the formation of zeolites in addition to clay minerals. An extensive theoretical treatment of early diagenesis is provided in Berner (1980).

Burial Diagenesis

During burial, shales undergo dewatering including compaction-driven expulsion and loss of water from dehydration of silicates (clay minerals). In sandstones, authigenic growth of quartz (commonly as syntaxial overgrowths on existing clean quartz grain surfaces), kaolinite (often accompanying detrital feldspar alteration), smectite, illite, and calcite dominates. At greater depths, mixed-layer clays become more illitic and chlorite, dolomite, and ankerite are more common. Detrital feldspars alter toward more sodic compositions (albitization). The geochemistry of authigenic phases, as well as of fluid inclusions within these minerals, provides information on the burial conditions.

Novel Approaches to Studying Diagenesis

The study of carbonate diagenesis is experiencing a rejuvenation with the introduction of new visualization, geochemical proxy and modeling approaches. Although petrographic and SEM study coupled with geochemical analysis of carbonate and siliciclastic rocks have become fundamental tools for reconstructing diagenetic pathways and environments, a new suite of stable (boron, sulfur, magnesium, calcium, lithium) and radiogenic (Sr, Nd, U) isotopes are being applied to carbonates to reconstruct paleo-seawater pH, oxygenation, and temperature as well as to place constraints on past carbon, sulfur, and magnesium cycling, weathering rates, and crystallization rates. Such studies have implications for the evolution of pCO₂ and pO₂ and life-environment interactions. With the advent of better methods for determination of trace element concentrations and their isotopic compositions through applications of ICP-MS, the use of trace metals as paleo-redox and paleoproductivity proxies has led to a rich literature (Lyons et al. 2014; Tribovillard et al. 2006; Gill et al. 2011; Jones and Manning 1994). Of particular interest for diagenetic studies is clumped isotope paleothermometry (Δ₄₇), which uniquely permits independent measurements of carbonate cement precipitation temperatures in carbonate and siliciclastic sedimentary rocks and the δ¹⁸O_fluid in which they formed. Although the application of clumped isotope thermometry to diagenetic studies is only beginning, initial results show promise for constraining diagenetic conditions and the evolution of fluid-flow conduits (Huntington et al. 2011; Budd et al. 2013). These applications of sedimentary geochemistry require assessment of separate diagenetic effects from primary signals in rocks across the broadest range of time scales. A surge in computational power over the past two decades is permitting the integration of sedimentology, petrography, geochemistry, geodynamics, and structural geology into forward modeling of the diagenetic evolution of sedimentary successions and the impact of diagenesis on their petrophysical properties and hydrologic potential (e.g., Whitaker et al. 2014; Agar and Geiger 2015).

Summary

Diagenetic modification of freshwater and marine carbonates and siliciclastic deposits by physical and chemical processes is a dynamic process that begins immediately following deposition (0–30 °C) through to the metamorphic realm (~250 °C). Sediments are converted to sedimentary rocks, and permeability and porosity develop by the integrated effects of dissolution, cementation, mineral overgrowth, neomophism, and recrystallization. The extent of diagenetic modification is governed by the original grain composition and reactivity, temperature and pressure, hydrodynamic conditions, and the presence of reactive organic matter. Bacterial mediation is an important component of carbonate and siliciclastic diagenesis.

The diagenetic cycle includes the early (eogenetic) through burial (mesogenetic) diagenetic realms as well as the telogenetic zone associated with tectonic uplift or sea-level exposure of sedimentary rocks. Although dolomite has been shown to form at Earth’s surface conditions with microbial mediation, extensive dolomitization occurs in the burial realm where kinetic inhibitors are overcome and elevated temperatures and high fluid-rock ratios promote the process. Maturation of sedimentary hosted organic matter through the diagenetic cycle has produced the world’s coal, oil, and gas deposits. Modern diagenetic studies are coupling new visualization, geochemical proxy, and modeling approaches with fundamental petrographic and geochemical tools in order to (1) quantitatively constrain the diagenetic history of carbonate and siliciclastic successions, (2) assess the potential of sedimentary rocks as archives of past surface conditions, and (3) reconstruct the evolution of petrophysical properties and hydrologic conduits in sedimentary units, as well as their resource potential.

Cross-References

• Fluid–Rock Interaction