The early search for oil and gas reservoirs centered on acquiring an overall view of regional tectonics, followed by a more detailed appraisal of local structure and stratigraphy. These days, however, the quest for reservoir quality calls for a deliberate focus on diagenesis.
In its broadest sense, diagenesis encompasses all natural changes in sediments occurring from the moment of deposition, continuing through compaction, lithification and beyond—stopping short of the onset of metamorphism.1 The limit between diagenesis and metamorphism is not precise in terms of pressure or temperature, nor is there a sharp boundary between diagenesis and weathering.
Thus, the nebulous domain of diagenesis lies somewhere between the ill-defined borders of weathering at its shallow end and low-grade metamorphism at its deep end. These postdepositional alterations take place at the relatively low pressures and temperatures commonly existing under near-surface conditions in the Earth’s lithosphere.
Diagenesis comprises all processes that convert raw sediment to sedimentary rock.It is a continually active process by which sedimentary mineral assemblages react to regain equilibrium with an environment whose pressure, temperature and chemistry are changing. These reactions can enhance, modify or destroy porosity and permeability.
Prior to the onset of diagenesis, porosity and permeability are controlled by sediment composition and conditions that prevailed during deposition. Even before it is laid down, a sedimentary particle may undergo changes between its source whether it was eroded from a massive body of rock or secreted through some biological process—and its point of final deposition.
The water, ice or wind that transports the sediment also selectively sorts and deposits its load according to size, shape and density and carries away soluble components. The sediment may be deposited, resuspended and redeposited numerous times before reaching its final destination.
Diagenesis commences once a sedimentary particle finally comes to rest.The nature and rapidity of postdepositional changes depend on the medium of deposition as well as the type of sediment deposited.As a given lamina of sediment is laid down, it becomes the interface between the transport medium and the previously deposited material, thus separating two distinctly different physicochemical realms.
In its new setting, the sediment contains a variety of minerals that may or may not be in chemical equilibrium with the local environment, and changes in interstitial water composition, temperature or pressure can lead to chemical alteration of its mineral components.
At or below the surface of this new layer, the sediment may be locally reworked by organisms that track, burrow, ingest or otherwise redistribute the sediment, sometimes subjecting it to bacterial alteration.
As deposition continues, the sedimentary lamination is buried beneath the depositional interface, forming successively deeper strata; there, it encounters continually increasing pressures and temperatures accompanied by changing chemical and biological conditions. These new conditions promote further consolidation and cementation of loose sediment and ultimately form lithified rock.
Important factors that influence the course of diagenesis are classified as either sedimentary or environmental. Sedimentary factors include particle size, fluid content, organic content and mineralogical composition. Environmental factors are temperature, pressure and chemical conditions.Particles in a layer of sediment may be subjected to
compaction, in which particles are moved into closer contact with their neighbors by pressure • cementation, in which particles become coated or surrounded by precipitated material
recrystallization, in which particles change size and shape without changing composition
replacement, in which particles change composition without changing size or form
differential solution, in which some particles are wholly or partially dissolved while others remain unchanged
authigenesis, in which chemical alterations cause changes in size, form and composition.
Any one of these transformations can significantly impact porosity and permeability and thus modify reservoir volume and flow rate. These effects are therefore of great interest to petroleum geologists and engineers in their endeavors to optimize production.
Indeed, drilling and production engineers must contend with similar phenomena to counteract the effects of fluid incompatibility, mobilization of clays and reservoir compaction. This article discusses diagenesis as it affects conventional reservoirs, focusing primarily on porosity and permeability changes in siliciclastic and carbonate rocks.
Setting the Stage
Porosity and permeability are initially controlled by sedimentary conditions at the time of deposition but are subsequently altered through diagenesis. The environment of deposition sets the stage for diagenetic processes that follow. Depositional environments for siliciclastic sediments, from which sandstones are formed, differ greatly from those of carbonates, which can form limestones. These rocks also differ in their reactions to changes in their environment.
Siliciclastics are primarily the product of erosion from a parent source. They are transported by some medium—fresh water, seawater, ice or wind—to their depositional site. Sand deposition is controlled by sediment supply, and the supply of coarser grains, in particular, is affected by energy of the transport medium.
For water-driven systems, energy is largely a function of sea level. During periods of relatively low sea level, or lowstand conditions, coarse-grained sediments can be carried beyond the continental shelf to be deposited in basinal marine settings.
Conversely, during rises in sea level, or highstands, most coarse-grained clastics are retained by fluvial systems or deposited at the beach, rather than in deep marine settings. It is the lowstand settings that are responsible for most of the coarse-grained siliciclastics deposited in deep water petroleum basins.
By contrast, the deposition of most carbonates is largely controlled by marine biological activity, which is viable only within a narrow range of light, nutrient, salinity, temperature and turbidity conditions. These requirements tend to restrict most carbonates to relatively shallow, tropical marine depositional settings.
Because carbonate deposition is affected by inundation of shallow marine platforms, most carbonate sediment is generated during highstands of sea level and is curtailed during lowstands. These differences in siliciclastic and carbonate deposition can ultimately affect reservoir quality.
Sand deposited during highstands may be eroded and transported downstream during lowstands. In contrast, carbonates deposited during highstands may be uncovered during lowstands, leaving them exposed to meteoric fluids that subject them to chemical changes, reworking and porosity modifications such as karsting.
A variety of outcrops and their unique diagenetic environments have been studied and described extensively, leading geologists to recognize similarities among various settings. Several schemes have been developed for classifying diagenetic regimes.
One method, proposed by Machel, is applicable to all rock types.12 This classification integrates mineralogic, geochemical and hydrogeologic criteria from clastic and carbonate rocks. It is divided into processes that occur in near-surface, shallow and intermediateto deep-burial diagenetic settings, along with fractures and hydrocarbon-contaminated plumes.
A different diagenetic model was outlined by Fairbridge in 1966. It emphasizes the geochemical aspect of diagenesis and recognizes three distinct phases: syndiagenesis, anadiagenesis and epidiagenesis. Each of these phases tends toward equilibrium until upset by subsequent changes in environmental parameters.
Another popular classification scheme relates carbonate diagenetic regimes to the evolution of sedimentary basins. This schema, originally proposed by Choquette and Pray, is now increasingly being applied to clastic processes as well.15
It is divided into three stages, some of which may be bypassed or reactivated repeatedly. Eogenesis is the earliest stage of diagenesis, in which postdepositional processes are significantly affected by their proximity to the surface.
During this stage, the chemistry of the original pore water largely dominates the reactions. The upper limit of the eogenetic zone is normally a depositional interface, but it may be a surface of temporary nondeposition or erosion. The lower limit shares a gradational boundary with the next stage and is not clearly defined because the effectiveness of surface-related processes diminishes gradually with depth, and many such processes are active down to different depths.
However, the maximum limit for eogenesis is estimated at 1 to 2 km, or 20°C to 30°C. The greatest change in the eogenetic zone is probably the reduction of porosity from cementation by carbonate or evaporite minerals.
Mesogenesis is the stage during which sediments or rocks are buried to such depths that they are no longer dominated by processes directly related to the surface. This phase, sometimes referred to as burial diagenesis, spans the time between the early stage of burial and the onset of telogenesis. Cementation is thought to be the major process affecting porosity in the mesogenetic zone, whereas dissolution is probably minor.
Telogenesis refers to changes during the interval in which long-buried rocks are affected by processes associated with uplift and erosion. Telogenetic porosity is strongly associated with unconformities. The upper limit of the telogenetic zone is the erosional interface.
The lower boundary is gradational and is placed at the depth at which erosional processes become insignificant. When a water table is present, the lower limit of the telogenetic zone extends to that point, which commonly serves as an effective lower limit of many weathering processes.
Dissolution by meteoric water is the major porosity-forming process of the telogenetic zone. As with the above schema, most diagenetic classifications are broadly based; some overlap with others and all contain exceptions to the rule.
Agents of Change
Freshly deposited sediments—mixtures of chemically unstable minerals and detrital materials act as building blocks of diagenesis, while water and organic matter fuel the process. Within a depositional system, changes in temperature and pressure can lead to the separation of different chemical compounds in unstable mixtures.
The liberation of unstable materials from one area is accompanied by their introduction elsewhere. Water plays a large role in diagenetic processes, dissolving one grain, hydrating others. The chemical activity may even change the properties of the water medium itself over time.
Water is but one of many agents of diagenesis; organic-rich sediments in various states of decomposition introduce a host of chemical reactions and bacteriological activities that consume all available oxygen.
This, in turn, leads to a chemically reducing environment. Under pressure, the gases of decomposition enrich the water with carbon dioxide and lesser amounts of methane, nitrites and other dissolved organic products.
This fortified water becomes a strong solvent, increasing solubility of carbonates and in some cases acting against silica in sandstones.Clays are also important to the diagenetic equation. They are responsible for forming easily compressible grains, cements and pore-clogging crystals.
Some clays form prior to deposition and become mixed with the sand-sized mineral grains during or immediately following deposition; others develop within the sand following burial. These clays are classified as allogenic and authigenic clays, respectively. Allogenic, or detrital, clays originate as dispersed matrix or sand- to cobble-sized mud or shale clasts.
These particles may be carried by downward or laterally migrating pore waters to infiltrate previously deposited sands. Individual clay particles may be dispersed throughout a sandstone or may accumulate to form thin laminae. Clays can also flocculate into sand-sized aggregates.Another type of aggregate is clay or mud “rip-up” clasts eroded from previously deposited layers.
A similar mechanism is at work in reworked fragments of older shales or mudstone that are deposited as sand-sized or larger aggregates. Allogenic clays can also be introduced into sands as biogenic mud pellets that are produced through ingestion and excretion by organisms. These pellets may be retained in burrows or transported as detrital particles. The biologic activity tends to homogenize the mud and sand
All types of clay can occur as detrital components. Bioturbation, mass flow and soft-sediment deformation are other modes for introducing clays into the fabric of marine sandstones, whereas mechanical infiltration is the mode for continental sandstones.
Detrital clay, of whatever mineral chemistry, occurs as tiny, ragged abraded grains and naturally accumulates in pore spaces, forming tangential grain-coating and porebridging fabrics.
Authigenic clays, unlike allogenic clays, develop within the sand subsequent to burial. Pore-water chemistry and rock composition strongly influence the growth of authigenic clays; connate water chemistry is modified over time by new influxes of water, through dissolution or precipitation of minerals and by cation exchange.
Various components of rock, such as lithic fragments, feldspars, volcanic glass and ferromagnesian minerals—minerals containing iron and magnesium—react with the pore water to produce clay minerals that may in turn undergo subsequent transformation to other, more stable forms of clay.
Authigenic clays can be recognized by their delicate morphology, which precludes sedimentary transport (below left). Authigenic clays in sandstone are typically found in four forms.
Clay coatings can be deposited on the surfaces of framework grains, except at points of grainto-grain contact. In the interstices between grains, the coatings act as pore-lining clays. These clays may be enveloped during subsequent cementation by feldspar and quartz overgrowths. Chlorite, illite, smectite and mixed-layer clays typically occur as pore linings. Pore linings grow outward from the grain surfaces and often merge with the linings on opposing grains in a process known as pore bridging.
Individual clay flakes or aggregates of flakes can plug interstitial pores. These pore-filling flakes exhibit no apparent alignment relative to framework grain surfaces.
Clay minerals can partially or completely replace detrital grains or fill voids left by dissolution of framework grains, sometimes preserving the textures of the host grains they replaced.
Clays can fill vugular pores and fractures.
The interactions among clay, organic matter and water become even more important in the context of sandstone and limestone porosity.
Sandstone Diagenesis
Freshly deposited sand—the precursor of sandstone—contains an assemblage of minerals that vary with local rock source and depositional environment . Sand-sized grains create a self-supporting framework at the time of deposition, finer particles form a detrital matrix and the remaining volume is pore space. Framework grains are detrital particles, chiefly of sand size—between 0.0625 and 2 mm. in diameter—commonly composed of quartz, feldspars and rock fragments.
The detrital matrix consists of mechanically transported fines—particles of less than 0.03 mm —that are predominantly clay minerals. The constituent minerals of this assemblage were formed under a specific range of temperature, pressure, pH and oxidation-state conditions unique to each mineral. These conditions will have a bearing on the physicochemical stability of the mineral assemblage.
Diagenetic processes are initiated at the interface between the depositional medium and the previous layers of sediment. These processes are modified as the layer is buried beneath sedimentary overburden. With time, the sand responds to changing pressure, temperature and pore-fluid chemistry—eventually emerging as a sandstone, minus some of its original porosity but perhaps with gains in secondary porosity.
All sands have intergranular porosity that changes with diagenesis: Macropores become micropores; minerals dissolve and create voids. Other minerals dissolve, then precipitate as cements that can partially or completely occlude pore space. Initial porosity may be as high as 55%.
That pore space is occupied by fluids such as water, mineral solutions or mixtures thereof; some pore fluids are inert, while others react with previously precipitated cements, framework grains or rock matrix. Porosity and permeability are especially important parameters both for diagenetic development and its effects on reservoir rock.
The amount of water or other fluids and their rate of flow through the pore network govern the amounts and types of minerals dissolved and precipitated, which in turn can alter flow paths and rates. Diagenetic processes by which sandstone porosity is lost or modified are outlined below.
Penecontemporaneous porosity loss—Those processes that occur after deposition but before consolidation of the enclosing rock are said to be penecontemporaneous. Certain processes, such as bioturbation, slumping and the formation of soil, fall into this category; although they may not be important on a large scale, they can be responsible for local reductions in sand porosity.
The activities of flora and fauna, such as plant roots, worms or bivalves, can disturb the original fabric of sediment. Root growth and chemical uptake, along with walking, burrowing or feeding activities of fauna, redistribute the sediment.
Slower sedimentation rates allow more time for organisms to rework a sedimentary layer. Bioturbation tends to have more impact in marine environments than in other settings. Slumping, or mass downslope movement, can result in a homogenization of sediments.
This newly formed mixture of sand and clay has substantially less porosity than the original sand layer. Soil creation can be an important diagenetic agent in environments such as alluvial fans, point bars and delta plains. Soil coverings contribute to the acidity of meteoric waters that percolate downward to underlying rock.
Clay particles generated through the formation of soil may be carried in suspension by meteoric water to infiltrate previously deposited sand layers. There, individual clay particles may disperse throughout a sandstone, accumulate to form thin laminae or attach as clay coatings on framework sand grains.
Porosity loss during burial—Deeper burial is accompanied by the primary causes of porosity loss: compaction and cementation.25 Compaction reduces pore space and sand thickness. Cementation can reduce pore space or can hinder sand compaction and dissolution at grain contacts.
During compaction, sand grains move closer together under the load of overburden or tectonic stress, destroying existing voids and expelling pore fluids in the process. Chemically and mechanically unstable grains, such as clays and volcanic rock fragments, tend to compact faster than more stable grains, such as quartz. Compaction mechanisms include grain rotation and slippage, deformation and pressure dissolution.
rain slippage and rotation are typical responses to loading in which a slight rotation or translation of grains permits edges of nondeformable grains to slip past adjacent grain edges, creating a tighter packing arrangement. The amount of porosity that can be lost depends, in part, on grain sorting, roundness and overburden pressure. Porosity loss from compaction has been estimated to range from 12% to 17% in various outcrop studies.
Ductile grain deformation—As ductile grains deform under load, they change shape or volume. Originally spherical or ovoid at the time of deposition, ductile grains are squeezed between moreresistant framework grains and deform into adjacent pore spaces. This reduces porosity while decreasing stratal thickness.The extent of compaction and porosity loss depends on the abundance of ductile grains and the load applied.
Compaction-induced deformation is also affected by cementation, timing and overpressure. Sandstones containing ductile grains undergo relatively little compaction if they are cemented before burial of more than a few meters or are strongly supported by pore fluid pressure in an overpressured subsurface setting.
Whereas the load from increased overburden pressure is typically carried by grain-to-grain contact, in an overpressured condition some of the stress is transferred to fluids within the pore system. Fluids normally expelled with increased pressure become trapped and carry some of the load.
Brittle fossilized sediments also deform under a load. Thin skeletal grains from fauna such as trilobites, brachiopods and pelecypods are subjected to bending stress because of their length. When these grains break, they allow overlying grains to sag into tighter packing arrangements.
Pressure dissolution—Points of contact between mineral grains are susceptible to dissolution, typically in response to the weight of overburden. Mineral solubility increases locally under the higher pressures present at grain contacts. Stylolites are the most common result of this process.
Pressure dissolution can reduce bulk volume and hence porosity. Dissolved material may be removed from the formation by migrating interstitial waters; alternatively, it may be precipitated as cement within the same formation. Grains composed of calcite, quartz, dolomite, chert and feldspar are commonly subjected to pressure dissolution.
Replacement—This process involves the simultaneous dissolution of one mineral and the precipitation of another. In this reaction to interstitial physicochemical conditions, the dissolved mineral is no longer in equilibrium with pore fluids, while the precipitated mineral is.
This substitution changes the mineral composition of the original sediment by removing unstable minerals and replacing them with more-stable ones. This process of equilibration can occur over the course of succeeding generations, whereby one mineral begets another as environmental conditions change.
Replacement opens the way to an assortment of porosity and permeability modifications. For example, replacement of silicate framework grains by carbonate minerals can be followed by dissolution of carbonate minerals, eventually resulting in porosity exceeding that of the original sediment.
On the other hand, porosity and permeability can be reduced by replacement of rigid feldspar minerals with ductile clay minerals, which are easily compacted and squeezed into pore throats between grains.
Some minerals are particularly susceptible to replacement. Others, such as pyrite, siderite and ankerite, are on the other end of the spectrum: They replace other cements or framework grains.
The degree of susceptibility to replacement normally follows an ordered mineral stability series in which minerals removed from their zone of stability are readily replaced (above). However, even the most stable minerals such as muscovite or quartz are not immune to replacement. Cementation—Cements consist of mineral materials precipitated chemically from pore fluids. Cementation affects nearly all sandstones and is the chief—but not the only—method by which sands lithify into sandstone.
Cementation can bolster porosity if it supports the framework before the sandstone is subjected to further compaction. In this case, remaining porosity is not lost to compaction, and excellent reservoir properties can be preserved to considerable depths. However, because cementation reaction rates generally increase with temperature, subsequent increases in depth can promote cementation and corresponding decreases in porosity with depth.
On the other hand, cementation can lock fine-grained particles in place, preventing their migration during flow that might otherwise block pore throats and reduce permeability. The amount and type of cement in a sandstone depend largely on the composition of the pore fluids and their rate of flow through the pores, as well as the time available for cementation and the kinetics of cement precipitating reactions.
It is common for certain minerals to form cements in sandstones. Over 40 minerals have been identified as cementing agents, but the most common are calcite, quartz, anhydrite, dolomite, hematite, feldspar, siderite, gypsum, clay minerals, zeolites and barite (right).
Calcite is a common carbonate cement, as are dolomite and siderite. Framework grains of carbonate rock fragments typically act as seed crystals that initiate calcite cementation. Quartz typically forms cement overgrowths on framework quartz grains and tends to develop during burial diagenesis at temperatures above 70°C.
Given sufficient space for enlargement, the overgrowth crystal will continue to grow until it completely masks the host grain surface. Adjacent grains compete for diminishing pore space, interfere with each other and generally produce uneven mutual borders forming an interlocking mosaic of framework grains and their overgrowths.
Authigenic feldspar occurs in all types of sandstones, mainly as overgrowths around detrital feldspar host grains but occasionally as cement or newly formed crystal without a feldspar host grain. Though common, feldspar cements are less abundant than carbonate, quartz and clay cements.
Authigenic clay cements are common in reservoir rocks of all depositional environments. The most common clay mineral cements are derived from kaolinite, illite and chlorite.
Enhanced Porosity in Sandstones
All sands initially have intergranular pores. Primary porosity, present when the sediment is deposited, is frequently destroyed or substantially reduced during burial. However, other diagenetic processes may also be at work, some of which may enhance porosity.
Porosity that develops after deposition is known as secondary porosity. It is typically generated through the formation of fractures, removal of cements or leaching of framework grains and may develop even in the presence of primary porosity. Secondary pores can be interconnected or isolated; those pores that are interconnected constitute effective porosity, which contributes to permeability. In some reservoirs, secondary pores may be the predominant form of effective porosity.
Secondary porosity can be important from a petroleum system perspective. Most hydrocarbon generation and primary migration take place below the depth range of effective primary porosity. The primary migration path and the accumulation of hydrocarbons are commonly controlled by the distribution of secondary porosity.30
Secondary porosity may develop during any of the three stages of diagenesis—before burial, during burial above the zone of active metamorphism or following uplift. However, burial diagenesis is responsible for most secondary porosity.
In sandstones, such porosity generally results from replacement of carbonate cements and grains or, more commonly, from dissolution followed by flushing of pore fluids to remove the dissolution products. Lesser amounts of porosity also result through leaching of sulfate minerals, such as anhydrite, gypsum and celestite. In general, secondary porosity is attributed to five processes.
Porosity produced through fracturingwhether it is caused by tectonic forces or by shrinkage of rock constituents: Should these fractures subsequently fill with cement, that cement may be replaced or dissolved, giving rise to second-cycle fracture porosity.
Voids formed as a result of shrinkage caused by dehydration of mud and recrystallization of minerals such as glauconite or hematite: Shrinking affects grains, matrix, authigenic cement and authigenic replacement minerals. Pores generated through shrinkage vary in size from a few microns across to the size of adjacent sand grains.
Porosity created by dissolution of sedimentary grains and matrix: Frequently, the soluble constituents are composed of carbonate minerals. Dissolution produces a variety of pore textures, and pore size may vary from submicroscopic voids to vugs larger than adjacent grains.
Dissolution of authigenic minerals that previously replaced sedimentary constituents or authigenic cements: This process may be responsible for a significant percentage of secondary porosity. Replacive minerals are typically calcite, dolomite, siderite, zeolites and mixed-layer clays.
Dissolution of authigenic cement: As with dissolved grains, most dissolved cements are composed of carbonate minerals: calcite, dolomite and siderite, though others may also be locally important. These cements may have occupied primary or secondary porosity. This is perhaps the most common cause of secondary porosity.
The size, shape and distribution of pores in a sandstone reservoir affect the type, volume and rate of fluid production. Three porosity types distinctly influence sandstone reservoir production: Intergranular pores are found between detrital sand grains.
Some of the most productive sandstone reservoirs have predominantly intergranular porosity. Dissolution pores result from removal of carbonates, feldspars, sulfates or other soluble materials such as detrital grains, authigenic mineral cements or replacement minerals (above right).
When dissolution pore space is interconnected with intergranular pores, the effectiveness of the pore system is improved. Many excellent reservoirs are a product of carbonates that have dissolved to form secondary intergranular porosity. However, if there is no interconnection, there is no effective porosity, leaving the pores isolated, with no measurable matrix permeability.
Microporosity comprises pores and pore apertures, or throats, with radii less than 0.5 μm. In sandstones, very small pore throats are associated with microporosity, although relatively large pores with very small pore throats are not uncommon. Micropores are found in various clays as well, and argillaceous sandstones commonly have significant microporosity, regardless of whether the clay is authigenic or detrital in origin.
Unless the sandstones have measurable matrix permeability, small pore apertures and high surface area result in high irreducible water saturation, as is often seen in tight gas sandstones. Porosity is seldom homogeneous within a given reservoir. It is often possible to find variations in porosity type across the vertical extent of a reservoir.
Carbonate Diagenesis
Most carbonate sediments are produced in shallow, warm oceans by marine organisms whose skeletons or shells are built from the calcium carbonate they extract from seawater. Unlike detrital sand deposits, carbonate sediments are usually not transported far from their source, so their size, shape and sorting have little to do with transport system energy. The size and shape of pores in carbonate sediments are more influenced by skeletal materials, which can be as varied as the assemblages of organisms that created them.
Carbonate sediments—composed chiefly of calcite, aragonite (a less stable crystal variation, or polymorph, of calcite), magnesian calcite or dolomite—are made from minerals that are highly susceptible to chemical alteration.
The impact of biological and physical depositional processes, in combination with the diagenetic overprint of metastable chemical deposits, can make the distribution of porosity and permeability in carbonates much more heterogeneous than in sandstones.
In fact, calcium carbonate dissolves hundreds of times faster than quartz in fresh water under normal surface conditions. The dissolution and precipitation of calcium carbonate are influenced by a variety of factors, including fluid chemistry, rate of fluid movement, crystal size, mineralogy and partial pressure of CO2.
The effects of mineral instability on porosity may be intensified by the shallow-water depositional setting, particularly when highstand carbonate systems are uncovered during fluctuations in sea level.
Most diagenesis takes place near the interface between the sediment and the air, fresh water or seawater. The repeated flushing by seawater and meteoric water is a recipe for diagenetic change in almost every rock, particularly as solutions of different temperature, salinity or CO2 content mix within its pores.
Porosity in near-surface marine diagenetic regimes is largely controlled by the flow of water through the sediment. Shallow-burial diagenesis is dominated by compaction and cementation with losses of porosity and permeability. The intermediate- to deep-burial regime is characterized by further compaction and other processes, such as dissolution, recrystallization and cementation.
Near-surface regime—Most carbonate rocks have primary porosities of as much as 40% to 45%, and seawater is the first fluid to fill those pore spaces. Filling of primary pores by internal sediments and marine carbonate cements is the first form of diagenesis to take place in this setting, and it leads to significant reductions in porosity.
Updip from the marine setting, coastal areas provide an environment in which seawater and fresh water can mix. In these groundwater mixing and dispersion zones, carbonate dissolution creates voids that enhance porosity and permeability, sometimes to the extent that caves are formed. Other processes are also active to a much lesser degree, such as dolomitization and the formation of aragonite, calcite or dolomite cements.
Further inland, near-surface diagenesis is fueled by meteoric waters, which are usually undersaturated with respect to carbonates. Rain water is slightly acidic because of dissolved atmospheric CO2.
Where the ground has a significant soil cover, plant and microbial activity can increase the partial pressure of CO2 in downward-percolating rainwater. This increases dissolution in the upper few meters of burial, thus boosting porosity and permeability through rocks of the vadose zone.
In evaporitic settings, hypersaline diagenesis is driven by fresh groundwater or storm-driven seawater that has been stranded upon the land’s surface. These waters seep into the ground and are subjected to evaporation as they flow seaward through near-surface layers of carbonate sediment.
As they evaporate beyond the gypsumsaturation point, they form finely crystalline dolomite cements or replacive minerals. In some petroleum systems, these reflux dolomites form thin layers that act as barriers to migration and seals to trap hydrocarbons.
Shallow-burial regime—Near-surface processes can extend into the shallow-burial setting, but the dominant process is compaction. Burial leads to compaction, which in turn squeezes out water and decreases porosity. Compaction forces sediment grains to rearrange into a self-supporting framework. Further burial causes grain deformation, followed by incipient chemical compaction in which mineral solubility increases with pressure.
In this way, loading applied to grain contacts causes pressure dissolution. Expelled fluids will react with surrounding rock. Intermediate- to deep-burial regime—With depth, several diagenetic processes become active. Chemical compaction becomes more prevalent with additional loading.
Depending on composition, clay minerals in the carbonate matrix may either enhance or reduce carbonate solubility. Pressure dissolution is further influenced by pore-water composition, mineralogy and the presence of organic matter. If the material dissolved at the contacts between grains is not removed from the system by flushing of pore fluids, it will precipitate as cement in adjacent areas of lower stress.
Dissolution is not just a pressure-driven process; it can also result from mineral reactions that create acidic conditions. In burial settings near the oil window, dissolution is active where decarboxylation leads to the generation of carbon dioxide, which produces carbonic acid in the presence of water.
Acidic waters then react with the carbonates. If the dissolution products are flushed from the system, this process can create additional voids and secondary porosity.
With burial comes increasing temperature and pressure, and changes in groundwater composition. Cementation is a response to elevated temperatures, fluid mixing and chemical compaction; it is a precipitation product of dissolution common to this setting.
Burial cements in carbonates consist mainly of calcite, dolomite and anhydrite. The matrix, grains and cements formed at shallow depths become thermodynamically metastable under these changing conditions, leading to recrystallization or replacement of unstable minerals. In carbonates, common replacement minerals are dolomite, anhydrite and chert.
Dolomite replacement has a marked effect on reservoir quality, though in some reservoirs it can be detrimental to production. While some geologists maintain that dolostone porosity is inherited from limestone precursors, others reason that the chemical conversion of limestone to dolostone results in a 12% porosity increase because the molar volume of dolomite is smaller than that of calcite.
The permeability, solubility and original depositional fabric of a carbonate rock or sediment, as well as the chemistry, temperature and volume of dolomitizing fluids, all influence dolomite reservoir quality.
In chemically reducing conditions, burial diagenesis can generate dolomite by precipitating it as cement or by replacing previously formed metastable minerals in permeable intervals flushed by warm to hot magnesium-enriched basinal and hydrothermal waters.
Temperatures of 60°C to 70°C are sufficient for generating burial dolomites, and these conditions can usually be met within just a few kilometers of the surface. In the deep subsurface, dolomitization is not thought to be extensive because pore fluids and ions are progressively lost with continued compaction.
Few, if any, carbonate rocks currently exist as they were originally deposited. Most are the result of one or more episodes of diagenesis.
Secondary Porosity in Carbonates
As it does in sandstones, diagenesis in carbonates can enhance reservoir properties through development of secondary porosity. Porosity in limestones and dolomites may be gained through postdepositional dissolution. In eogenetic or telogenetic settings, dissolution is initiated by fresh water. In mesogenetic settings, dissolution is caused by subsurface fluids generated through maturation of organic matter in the deep burial environment.
During eogenesis, development of secondary porosity is aided by a number of processes. Dissolution is dominated by meteoric fresh waters, which are undersaturated with respect to calcium carbonate.
However, the extent of dissolution is determined by other factors, such as the mineralogy of sediments or rocks, the extent of preexisting carbonate porosity and fracturing, the acidity of the water and its rate of movement in the diagenetic system.
During telogenesis, uplift exposes older, formerly deep-buried carbonate rocks to meteoric waters, but with less effect than during the eogenetic phase. By this time, what were once carbonate sediments have matured, consolidated and lithified to become limestones or dolostones.
These older rocks have, for the most part, become mineralogically stabilized. Soluble components of the eogenetic sediment—such as ooids or coral and shell fragments composed of aragonite—have probably dissolved during earlier phases.
Having mineralogically evolved toward a less soluble low-magnesium calcite, these rocks are more resistant to dissolution than their eogenetic precursors. Further dissolution requires exposure to fluids that are undersaturated with respect to calcite. When this happens, vugs and caverns may form.
Similar processes can, in part, explain how secondary porosity forms in some dolomites. When partially dolomitized limestones undergo telogenesis, meteoric fluids dissolve any remaining calcite in the rock matrix or its component particles. The calcitic components, being more susceptible to freshwater dissolution than dolomite is, leave new pores in their place.
However, many carbonate reservoirs that were formed in deep waters have yet to be uplifted or exposed to meteoric waters; even without the benefit of eogenesis or telogenesis, they are porous and permeable. Dissolution nonetheless requires that carbonates be exposed to fluids that are undersaturated in calcium carbonate.
Most connate fluids in deeper environments, however, are saturated with respect to calcium carbonate, leaving them incapable of dissolving carbonate rocks and creating secondary porosity. In fact, just the opposite is the case:
These fluids tend to precipitate calcite or dolomite cement and may even be capable of dolomitization in some cases.Creating secondary porosity in the mesogenetic region requires a different mode of dissolution.
With burial, the organic matter in source rocks matures and is eventually converted to hydrocarbons. During this process, organic acids, carbon dioxide and hydrogen sulfide are expelled from the source rock; the gases combine with subsurface waters, producing carbonic acid and sulfuric acid, respectively.
These acids migrate laterally and vertically, dissolving carbonates and creating porosity along their paths. Once the acids are spent, the fluids precipitate carbonate cements.
Diagenesis at Different Scales
Failure to recognize diagenetic effects on porosity can lead to serious errors in reservoir volume calculations and impact flow rate predictions. Different aspects of diagenesis may be recognized at the well site and investigated in the laboratory. Developments in logging technology are continually improving the resolution of wireline and LWD tool measurements, and these data are valuable for high-grading formation intervals that merit further attention in the laboratory.
In this capacity, well logs are quite good at providing a fairly detailed overview of the lithology along the length of the borehole. The laboratory is reserved for studying extremely small intervals—usually samples of core or formation cuttings—in ultramicroscopic detail.
At the wellsite, the intricacies of various diagenetic processes culminate in one central issue: porosity. In particular, the ability to recognize secondary porosity is critical to the formation evaluation process. Depending on the extent of connectivity, secondary porosity can increase or decrease reservoir producibility.
Although conventional porosity logs do not directly measure secondary porosity, log analysts can use an alternative method to obtain this information. If neutron, density and compressional sonic data are obtained, then nonconnected secondary porosity can be detected and quantified.
Both the neutron and density porosity tools respond to primary, interparticle matrix porosity, as well as secondary, vuggy and fracture porosity. However, these measurements do not distinguish between primary and secondary porosities. Alternatively, the compressional sonic slowness measurement responds to primary porosity only.
Porosity data can be compared after processing. The neutron and density porosity logs are corrected for environmental effects and matrix conditions, then crossplot porosity is calculated. Sonic porosity from an acoustic tool is calculated utilizing the same rock and fluid properties used in evaluating the neutron-density data.
Any difference between sonic and crossplot porosities represents the fraction of noneffective secondary porosity (left). However, if the secondary porosity is connected, the sonic porosity and neutrondensity porosity values will generally match and this method will not distinguish between the two.
On the other hand, full suites of modern logs are fairly adept at detecting and evaluating secondary porosity. Borehole imaging logs can indicate the type and occurrence of secondary porosity. Nuclear magnetic resonance tools may be used to determine the pore throat size and connectivity and infer secondary porosity from those relationships.
Total porosity, computed from NMR data, is generally matrix independent and is an indicator of incompatible matrix inputs for neutron-density and sonic porosity logs. Multidimensional resistivity logging tools can measure vertical and horizontal anisotropy for inference of secondary porosity. Dipole and multiarray sonic data also help estimate elastic anisotropy and reservoir connectivity. Spectroscopy logs can constrain lithologies to delineate zones of secondary porosity, thus providing an effective matrix density, or grain density, for computing porosity.
The details of cementation, dissolution and other diagenetic minutia are, for the most part, revealed in the laboratory by various forms of microscopy or chemical analysis. Often thin section study and scanning electron microscopy are combined to evaluate different porosity types, determine rock texture and anticipate potential reservoir problems.
Thin-section petrography is a basic technique for examining the textural characteristics of mineral grains in a rock. Rock samples are ground to extremely thin slices, which are polished and impregnated with dyed epoxy resin to enhance porosity identification.
These thin-section slides of rock samples are studied under filtered polarizing light using a petrographic microscope. Geoscientists use polarized light microscopy to observe optical properties caused by anisotropic materials that reveal details about the structure and composition of the rock. In some cases stains are applied to aid in identifying mineralogy such as feldspar grains and carbonate cements.
Thin-section examinations are routinely supplemented with other sophisticated technology, including SEM, X-ray diffraction and cathodoluminescence. The SEM analysis couples extreme depth of focus with a wide range of magnification for identifying minerals or investigating pore morphology and pore throat geometry.
The SEM technique enables geoscientists to photograph the distribution of detrital and authigenic minerals and study the effects of cement and grain coatings on porosity.
X-ray diffraction can reveal much about the crystallographic structure and chemical composition of mudrocks and sandstones and their clay fractions. This technology works on the principle that each crystalline substance produces its own unique diffraction pattern.
When a rock is ground to a powder, each component will produce a unique pattern independent of the others, providing a fingerprint of the individual components to enable identification of the mineralogical makeup of the sample.
In addition, quartz, feldspar and carbonate minerals in sedimentary rocks emit visible, ultraviolet and infrared light when bombarded by high-energy electrons. These emissions can be captured and displayed as color images in a cathodoluminescence microscope or petrographic scope. The CL imaging results can help geoscientists evaluate the provenance of detrital mineral grains.
Other chemical investigative methods, such as stable isotope analysis, are facilitating investigations of pore waters and their effect on cementing minerals in sedimentary rocks. For example, geoscientists can determine the marine or nonmarine origin of pore waters by analyzing the concentrations of carbon, oxygen and sulfur isotopes.
Exploration and production companies have used core analysis and petrophysical examination routines for years. The industry has made many gains from these studies, including improved drilling fluids, compatible completion fluids and a range of acid stimulation treatments—all designed specifically to overcome the effects of diagenetic cements.
Other advanced formation evaluation technologies, though less common, are making inroads into the oil patch. One such advance is X-ray computed tomography .
By aiming a focused X-ray beam at a rock sample, geoscientists can obtain “virtual slices” that can be resolved to a scale of microns.
As a research tool, micro-CT scans are used for pore space characterization. An increasingly important tool in nondestructive testing, its application can be extended to laboratory testing of unconsolidated or friable formation samples. Micro-CT imaging may eventually lead to more-accurate predictions of porosity-permeability trends and calculations of capillary pressure, relative permeability and residual saturation.
Diagenesis has been a subject of research and discussion since the 1860s, and the industry has taken an increasing interest in the topic since the 1940s. In the field of diagenetic research, E&P companies have made great strides by adopting evaluation techniques based on SEM, XRD and CL analysis. With the alternative to forgoing comprehensive diagenetic evaluations becoming less attractive, operators are embracing these and other technologies in their attempts to glean more information and gain greater understanding of their reservoirs.