Abstract Conodonts are apatitic marine microfossils of Cambrian through Triassic age. During incipient metamorphism (50°-300° C) they change color from pale yellow to brown to black due to carbon-fixing within the trace amount of organic matter in their skeletons. As thermal metamorphism continues (300°-550° C), conodonts change from black to gray to white to crystal clear as a result of carbon loss, release of water of crystallization, and recrystallization. The conodont color alteration technique provides a unique link between mineral and organic indexing of thermal metamorphism and is best suited for carbonate rocks. Conodont color alteration index (CAI) isograd maps for three stratigraphic intervals in the Appalachian basin show: (1) Conodont color alteration is directly related to the depth and duration of burial and the geothermal gradient. (2) Tectonics affect color alteration only where folding and faulting act to significantly increase depth of burial. (3) Isograds and overburden isopachs are conformable throughout most of the northern half and in the western part of the southern Appalachian basin; in these areas, isograd values gradually increase eastward except for a major disruption in the area of the Rome trough. (4) South of central Virginia, isograds are disrupted and irregular because late Paleozoic thrusting has severed and telescoped original burial metamorphism isograd patterns. (5) Basin restoration using conodont CAI isograds indicates a maximum shortening in northeast Tennessee of about 115 miles (185 km). (6) The CAI 2 isograd (=brown conodonts) for each stratigraphic interval lies near the eastern limit of oil production for that interval; this limit shifts eastward for each successively younger stratigraphic interval concomitant with decreasing overburden. (7) Gas production is less related to isograds and depends mostly on primary and (or) secondary porosity and permeability. The CAI 4 isograd (= brownish-black conodonts), however, approximates the eastern limit of gas production because the temperature (depth of burial) necessary to produce this high level of organic metamorphism concurrently produces mineral metamorphism that reduces porosity and permeability and the likelihood of commercial reservoirs.

Abstract Dispersed solid organic matter occurs as a minor constituent in most sedimentary rocks. It consists of diverse materials that are like the macerals in coals (though usually in different proportions than in normal coals), and its maturation is chemically and physically like coalification. Reflected-light microscopy enables one to recognize the different organic grains and to select the best type for optical measurement to indicate the indigenous maturation of organic matter in the rock. The organic constituent selected should have the following characters. It (1) is virgin when deposited with the sediment, (2) matures regularly, (3) is not subject to retrograde alteration, (4) resists reaction with adjacent fluids and solids, (5) is not significantly affected by pressure, (6) occurs widely in rocks of diverse lithology and facies, (7) is distinguishable from pre-altered and redeposited material, (8) can be analyzed separately, (9) persists through a broad range of catagenesis and metamorphism, and (10) has properties that can be analyzed throughout the alteration range by a relatively inexpensive technique. Vitrinite grains that are not recycled from previous rocks satisfy the above requirements, and reflectance is normally the property measured. Data from experimental studies in the laboratory and from a number of sedimentary basins show that, for the most part, temperature and duration of heating determine the progress of catagenesis. Regional and vertical studies of organic catagenesis indicate a correlation between rank of solid organic matter and occurrence of oil and gas—even though the fluids migrate extensively. Oil is generated first at about 0.5% vitrinite reflectance (oil immersion) and occurs last, associated with gas condensate, at about 1.3%. Above that rank, abundant methane can be generated from types of kerogen that do not yield oil. Petroleum occurrence is limited in much of the eastern United States because the sedimentary rocks have been too hot in the past (mostly as a result of former deep burial). The regions that are favorable or unfavorable for exploration for petroleum or gas, from the point of view of level of organic maturation, are indicated on a map. On the continental shelf, especially where sediments are less than two kilometers thick, it is most important to determine whether burial temperature has been adequate for petroleum formation. Determination of actual past temperatures is not required for correlation with petroleum occurrence. Measured level of organic catagenesis can be used, however, to estimate actual former temperatures from our knowledge of the general time-temperature-rank function of the reactions and from geologic information on the burial history of the rocks in question.

Abstract Catagenesis of sedimentary organic matter leads to end products of methane and a carbonaceous, graphite-like residue. Kerogens heated in the laboratory as well as those that have experienced natural catagenesis demonstrate progressively better developed graphite-like atomic arrangement as a function of increased temperature conditions. The height / width ratios of the d<sub>002</sub> X-ray diffraction peak of graphite are useful in expressing various levels of graphitization and thus are potentially useful in paleothermometry.

Abstract Clay mineral assemblages in sedimentary rocks can be indicative of the maximum temperature (below 300° C) and metamorphic grade to which these rocks have been subjected. The chemical composition of the rock and pore fluid, along with the detrital mineralogy are also influential in determining the mineral assemblages that form at higher temperatures. Clay mineral assemblages are also dependent on reaction time at low temperatures. The most useful minerals which can be used as geothermometers are as follows: (1) for shales: illite/smectites, illite, and chlorite; (2) for sandstones: chlorite, chlorite/smectite (corrensite) and dickite; and (3) for volcanics and volcanogenic rocks: chlorite/smectites and zeolites. In pelitic sediments the conversion of smectite to illite and its subsequent recrystallization to 2M muscovite allow a detailed temperature zonation of these rocks into a sequence of temperature ranges. These temperature (grade) sequences are applied to the disturbed belt of Montana. Mesozoic strata involved in Laramide thrust faulting events have mineral assemblages, principally the mixed-layer clays and zeolites, which reflect a low grade metamorphic environment and indicate metamorphic temperatures between 100° and 200° C. These metamorphic mineral assemblages are displayed by strata as lithologically diverse as shales, sandstones, volcanogenic rocks, and bentonites. Field evidence indicates that these strata were never sufficiently buried during sinking of the sedimentary basin to produce the observed mineralogic and chemical changes. Comparison of stratigraphically equivalent sediments on the adjacent undisturbed Sweetgrass Arch show mineral assemblages indicative of low temperature (<<100° C) conditions. Heating to a degree higher than that predicted from normal geothermal gradients is inferred to be caused by burial beneath thrust plates. This is suggested as the mechanism that generated the temperatures indicated by observed mineral assemblages in the disturbed belt.

Abstract The zeolite facies was defined to “bridge the gap” between diagenesis and metamorphism and was largely based upon the studies of Coombs (1954) from Taringatura, New Zealand. The lack of zeolites in many stratigraphic sections which were presumably subjected to similar P s -T conditions as zeolite-bearing rocks has led to the definition of a clay-carbonate facies . On theoretical grounds it can be shown that T and fluid composition exercise a strong control on the mineral assemblages and this has been verified for active geothermal areas such as Wairakei and Broadlands, New Zealand. Estimates of P s , T and fluid composition attending zeolite facies alteration can be made from correlation of mineral assemblages with those predicted from experimental and computed phase equilibria. Such estimates encounter several problems: (1) at low P s and T growth and persistence of metastable phases are more important than at higher P s , T; (2) for P<sub>H<sub>2</sub>O</sub> < P s equilibria involving zeolite dehydration are very strongly affected; (3) experimental studies often yield phases which differ chemically and structurally from the naturally occurring zeolites; (4) variables such as the activity of silica (a<sub>SiO<sub>2</sub></sub>) may be important since volcanic glass is involved in the production of many zeolites; (5) in teconically disturbed areas Zen and Oxburgh and Turcotte have shown that P s -T conditions at the base of thrust sheets may place these rocks outside the stability field of many of the zeolites for significant periods of time; (6) factors such as distribution of porosity and permeability may strongly affect the mineralogy and these are often difficult to evaluate after the fact. For a given area there is reasonable correlation between zeolite assemblages, coal rank and clay mineral assemblages and consideration of all of these types of evidence will lead to the best estimates of P s , T and fluid composition.

Abstract Most sedimentary diagenesis involves recrystalJization or overgrowths on original minerals, or the growth of new phases. This new growth may trap fluid as inclusions that provide data not only on the nature, composition, pressure, and density of the fluids present during diagenesis, but particularly on the temperature at which the host crystals grew. As most optical methods of study require inclusions >1-2 μm in diameter, fine-grained products of diagenesis, in the 10-20 μm range, seldom provide useful material. The possibilities of finding inclusions of useful size increase as the size of the host crystal increases. In spite of this limitation, reasonably valid quantitative or qualitative physical and chemical data, both new and from the literature, have been obtained on inclusions from the following specific diagenetic environments: (1) crystal-lined geodes, vugs, and veins in sediments; (2) Mississippi Valley-type ore deposits; (3) carbonate and quartz cements in detrital rocks; (4) saline and sulfur deposits; (5) petroleum reservoir rocks; (6) sphalerite in bituminous coal beds. Most inclusion temperatures in these and other similar environments range from 25 to 150° C, and most of the fluids are moderately to strongly saline brines which commonly contain petroleum and as much as tens of atmospheres of methane-rich gas. Homogenization temperatures of inclusions in some Mississippi Valley-type ore deposits are higher than 150° C but seldom more than 200° C. It is concluded that hot, strongly saline fluids have moved through many, if not most, sediments at some time in their history, and that at least part of the diagenetic changes seen have been caused by such fluids.

Abstract Fission tracks that are present in apatite crystals recovered from sedimentary and crystalline basement rocks can be used to determine the thermal (tectonic) history of a sedimentary basin. Because fission tracks in apatite usually record the time when the rock temperature cooled below 100° C, one can use the apparent apatite fission-track ages as a function of depth, to determine the amount and duration of an uplift event.

Abstract The maximum temperature to which a shale has been heated as a result of burial can, in some instances, be estimated using oxygen isotope geothermometry. The isotopic fractionation, or difference between O<sup>18</sup>/O<sup>16</sup> ratios of two coexisting minerals which have reached isotopic equilibrium with one another, is temperature dependent. Hence, if two coexisting minerals, which have isotopically equilibrated with one another, can be separated from a shale, and if the variation of the equilibrium isotopic fractionation between these minerals is known, the temperature of equilibration can-be estimated. Quartz and coexisting illite or mixed layer illite/smectite is a promising pair for isotope geothermometry of shales. A preliminary equilibrium fractionation curve for this pair is given by: where exp refers to the fraction of layers in the mixed-layer clay which are expandable. The results of three isotope geothermometry studies are summarized. Mineralogic and O 18 /O 16 data for coexisting quartz and illite from the altered volcanic rocks of the active hydrothermal region at Broadlands, New Zealand were used to investigate isotopic equilibration and to serve as a basis for calibration of the quartz-clay isotope geothermometer. Mineralogic and O l8 /O l6 data for coexisting clay-sized quartz and illite /smectite from one of three deep wells in the Gulf Coast indicate that isotopic equilibrium is approached between these two minerals at a well temperature above about 100° C. The illite/smectite apparently exchanges oxygen with pore waters in an approach toward isotopic equilibrium during the reaction: smectite + Al + K 𠆒 illite + Si. The released Si forms quartz which dominates the finest quartz fractions and forms overgrowths on detrital grains. This quartz apparently forms in isotopic equilibrium with pore waters. Isotopic temperatures derived from coexisting quartz and illite from the Precambrian Belt argillites range from 225° C to 310° C and generally increase down-section. This temperature range is compatible with the bulk mineralogy and probable depth to which the rocks have been buried. Oxygen isotope geothermometry can not yet be used routinely. However, it may be more possible to do so after additional information is obtained on factors such as chemical alteration and mineralogic reactions that control isotopic exchange.

Abstract Four truths about holes in sandstone underlie current research and unsolved problems of sandstone diagenesis. First, primary intergranular porosity and permeability of sand are greatly reduced and subject to total destruction in the early stages of burial diagenesis. Compaction, cementation, recrystallization, and replacement are widely recognized porosity-reducing mechanisms. Rate of porosity loss with depth is related mainly to original sand composition. Second, at later stages of diagenesis, secondary porosity can be produced by dissolution, of detrital and authigenic minerals. Porosity can be restored and enhanced at depth. Porosity of many major hydrocarbon reservoirs is mainly secondary, a fact just recently documented in the literature. Secondary porosity must be formed before hydrocarbon migration if it is to serve as reservoir porosity. Secondary porosity can be destroyed diagenetically, but it will persist to greater depths than will primary porosity. Third, chemical diagenesis of sandstones is a kinetic process; mineral matter is dissolved, transferred, and precipitated by aqueous solutions moving through sandstones. The main source of water is from dewatering of shales interbedded with sandstones. Reconstruction of the chemical evolution of moving water, of its flow paths through a basin, and of the time of migration are the keys to predicting subsurface distribution of sandstone porosity. Mathematical modeling of hydrodynamics and mineral reactions by means of computer simulation is a promising approach to porosity prediction. Fourth, the course of sandstone diagenesis in a given basin is programmed by the preburial, prediagenetic factors of provenance, depositional environments, and tectonic setting. These interrelated factors influence sand composition and texture, which in turn govern mineral reactions and fluid-flow rates. Diagenetic processes determine porosity in terms of origin, amount, subsurface distribution, pore-size distribution, pore shape, surface area, and attendant permeability. Thus, diagenetic history must be taken into account by geologists and engineers in the petroleum industry. Rewards from studies of diagenesis will be sharper porosity prediction in exploration and more efficient management of rock-fluid interactions in producing reservoirs.

Abstract The range of physical and chemical conditions included in diagenesis is 0-200° C, 1-2000 kg/cm 2 , and water salinities from fresh to brines twice as concentrated as the Dead Sea. Numerical values of these parameters vary not only with depth, but areally at any single depth. In addition, gross differences in heat flow in different tectonic settings can cause the average temperature at a given depth in a eugeocline to be double that in a miogeocline. Despite the large variation in physical and chemical conditions during diagenesis, it is striking that calcite and quartz are the dominant chemical precipitates in pore spaces of sandstones. To a first approximation, this reflects two facts. (1) Most preserved sands were deposited in shallow marine environments and sea water is at least saturated with respect to calcium carbonate. (2) The average sand contains about 65% quartz and rates of flow of subsurface waters are so slow that quartz is the main buffer for the silica content of these waters. The slow rate of movement of subsurface waters places important restrictions on when cementation of a sand can occur. To lithify a sand, the circulating water must be supersaturated with respect to the solid to be precipitated in the pore space and the number of pore volumes of water that must flow through the sand must be very large. If the sand is areally widespread, calculations for well-sorted pure quartz sand indicate that cementation of the unit by quartz by horizontal flow of subsurface water is impossible within geologically reasonable periods of time. Plugging of pore spaces by quartz must result from vertical circulation, probably when the depth of sand burial does not exceed a few hundred meters. This cementation can occur either soon after deposition of the sediment or at any subsequent time when tectonic forces elevate the sediment to a shallow depth.

Abstract Four basic types of porosity occur in sandstones: intergranular, dissolution, micro and fracture. The first three types are related to rock texture and can be considered end members of a ternary classification diagram. Fracture porosity may be associated with any other porosity type. All sandstones initially have intergranular porosity, which, if not destroyed, often is associated with good permeability, large pore apertures, and prolific hydrocarbon production. Dissolution porosity results from leaching of carbonate, felaspar, sulfate, or other soluble material. Sandstone reservoirs with dissolution porosity range from excellent to poor, depending on amount of porosity and interconnection of pores. Isolated dissolution pores result in low permeability. Sandstones with significant amounts of clay minerals have abundant microporosity, high surface area, small pore apertures, low permeability, high irreducible water saturation, and an increased sensitivity to fresh water. Fracture porosity, which contributes no more than a few percent voids to storage space, will enhance the deliverability of any reservoir. Open fractures, either natural or induced, are essential for economic deliverability rates from reservoirs with essentially only micropores or isolated dissolution pores. Porosity type and/or pore geometry change with diagenesis: macropores become micropores, minerals dissolve to create voids, and pores are partly to completely occluded by precipitation of minerals. It is important to have an understanding of pore geometry, that is the size, shape, and distribution of pores in a reservoir. Pore geometry influences the type, amount, and rate of fluid produced. Porosity type seldom is homogeneous in rocks. As a result, log interpretation problems may occur in sandstones containing significant micropores and interconnected macropores. Micropores may hold irreducible water while macropores may hold producible oil, gas, or water, depending on height above the oil- or gas-water contact. Log calculations may indicate high water saturation and a nonproductive interval, although the reservoir may be capable of water-free hydrocarbon production because the water is not producible.

Abstract Secondary porosity plays an important role in the diagenesis of some sandstones. The volume of secondary porosity equals or exceeds that of primary porosity in the sandstones of many sedimentary basins worldwide, and a significant percentage of the world's reserves of natural gas and crude oil are contained in secondary sandstone porosity. Prudhoe Bay Field and the Jurassic fields of the North Sea are examples of the many giant hydrocarbon accumulations in secondary sandstone porosity. Chemical, physical, physicochemical, biochemical and biophysical processes result in secondary sandstone porosity through leaching and shrinkage of rock constituents, or through the opening of fractures and porous burrows and borings. Secondary sandstone porosity can originate anywhere in the sedimentary crust: (1) before effective burial in the environment of deposition (eogenetic) 2 ; (2) at any depth of burial above the zone of metamorphism (mesogenetic) 2 ; and (3) during exposure following a period of burial (telogenetic) 2 . Secondary porosity may occur in sandstones of any mineralogical or textural composition and of any Phanerozoic age. It is most common in sandstones that have undergone relatively long lasting, deep burial and have lost their primary porosity. Most of the secondary porosity in ancient sandstones originated as a result of mesogenetic leaching of the carbonate minerals calcite, dolomite and siderite. This decarbonatization removes depositional carbonate constituents and diagenetic carbonate such as cements or replacements. Most of the mesogenetic decarbonatization results from the decarboxylation of organic matter in strata adjacent to the sandstone during the course of organic maturation. The process of decarboxylation leads to the generation of carbon dioxide which, in the presence of water, produces carbonic acid. This acid reacts with the carbonate minerals. In most instances it is possible to differentiate microscopically between primary and secondary sandstone porosity. Thus it is possible to trace the loss of primary porosity during burial. In the presence of water and hydrostatic pressure, primary sandstone porosity cannot exist beyond specific limits of temperature-time exposure except for a small volume of irreducible lamellar porosity between grains. The limiting temperature-time exposure increases with increasing mineralogical stability of the sandstones and, subordinately, with increasing grain size. The mesodiagenesis of sandstones can be divided into four stages: (1) immature—mechanical compaction; (2) semi-mature—chemical compaction of primary porosity; (3) mature—only secondary porosity present; and, (4) supermature—no effective primary or secondary porosity. Decarbonatization may create considerable quantities of secondary porosity during semi-mature mesodiagenesis. However, the average addition of carbonate to the sandstone in this diagenetic stage usually exceeds the average carbonate removal. Decarbonatization culminates during mature mesodiagenesis at which stage it greatly outweighs carbonatization. Much secondary sandstone porosity, therefore, originates after effective primary porosity has been lost. Fractures and irreducible lamellar porosity apparently provide sufficient access for decarbonatizing fluids to start the leaching process even in sandstones of low permeability. Enormous volumes of carbonate move upwards in solution from diagenetically mature sandstones and are, at least in part, reprecipitated in immature and semi-mature sandstones. Within a subsiding prism of clastic sediments much of the carbonate content is being recycled upwards in this fashion and sandstones at shallower depths are being enriched in carbonate. Primary migration of hydrocarbons commonly follows closely after the secondary porosity has been formed, because in the maturation of organic matter, the main phase of hydrocarbon generation follows after the culmination of decarboxylation. This close association of source and reservoir in time and space favours the accumulation of hydrocarbons in secondary porosity. In the presence of pore water, secondary porosity is gradually reduced during deep burial although at a much slower rate compared to primary porosity. The main geological and economic significance of secondary sandstone porosity is that it extends the depth range for effective sandstone porosity far below the generally accepted depth limit for effective primary porosity. Generation and primary migration of hydrocarbons occurs mainly below the range of effective primary porosity.

Abstract Secondary porosity in sandstones can be classified according to origin and pore texture. Five significant genetic classes of secondary porosity are defined by the following processes of origin: (1) fracturing; (2) shrinkage; (3) dissolution of sedimentary grains and matrix; (4) dissolution of authigenic pore-filling cement; and, (5) dissolution of authigenic replacive minerals. Hybrid pores are characterized either by the coexistence of several genetic classes of secondary porosity or by the coexistence of primary and secondary porosity. Secondary porosity appears in five major groups of pore textures: (1) intergranular pores; (2) oversized pores; (3) moldic pores; (4) intra-constituent pores; and, (5) open fractures. Some secondary porosity mimics the entire range of pore sizes and pore textures of primary sandstone porosity. Other secondary porosity bears a general resemblance to the textures of primary porosity but differs in detail. Secondary porosity may also appear in textures that are entirely different from those of primary porosity. In most instances it is possible to identify the occurrence of secondary porosity in thin section using a set of simple petrographic criteria that include: (1) partial dissolution; (2) molds; (3) inhomogeneity of packing; (4) oversized pores; (5) elongate pores; (6) corroded grains; (7) intra-constituent pores; and, (8) fractured grains. In medium- and coarse-grained sandstones secondary porosity can, in some instances, be observed by the naked eye. The detailed analysis of the petrological attributes of secondary porosity may require the use of advanced analytical techniques, such as cathode luminescence petrography, scanning electron microscopy, pore-cast examination, microprobe analysis and stable-isotope analysis. In most sandstones, however, a surprising amount of important information can be obtained simply by the careful and methodical use of conventional petrographic microscopy.

Abstract It is increasingly apparent that temperature effects have been overemphasized in evaluating many diagenetic reactions associated with burial. The classic concept of burial metamorphism is far too simplistic to explain the wide variation in reactions and reaction sequences observed in many diagenetic terranes. This is particularly true when considering the complex problem of diagenetic alteration of volcanogenic sandstones. For example, changing the geothermal regime from area to area cannot explain the common observation that individual mineral ranges broadly overlap and are not related to stratigraphic position. The diagenetic reactions of interest in volcanogenic sandstones such as: glass→clay, glass→zeolite, zeolite→zeolite, plagioclase→clay, or zeolite, among others, involve a fluid phase and commonly ionic species in the fluid phase. Consider for example the typical diagenetic reaction of heulandite→laumontite (Ca ++ + Ca 3 K 2 Al<sub>8</sub>Si<sub>28</sub> O 72 -24H 2 0→Ca 4 Al<sub>8</sub>Si 16 0 48 - 16H 2 0 + 2K + + 12SiO 2 + 8H 2 0). Obviously the problem is not one of just thermal stability, but is one of chemical or ionic stability as well. Such factors as fluid flow and composition are as significant as depth of burial in controlling the distribution of diagenetic mineral phases in volcanogenic sandstones. Variations in fluid flow, and more importantly fluid composition, can explain many of the perplexing questions that previously were inadequately explained by thermal variations alone. Fluid effects are most pronounced in the early stages of diagenesis when the fluid/solid ratio is high, and in the later stages during fracturing and/or dewatering.

Abstract Diagenesis of volcanogenic sandstones is characterized by the alteration of glass and feldspar (plagioclase) to carbonates, clays and zeolites. These alteration products are commonly distributed throughout the body in well defined zones, but the chemical and physical controls on the spatial location of these zones is poorly understood. A quantitative model describing these processes has been constructed by combining the equations describing chemical reaction with those governing mass transfer. This approach leads to a set of differential equations with associated boundary conditions constrained by a set of mass action equations. In this framework the identity and spatial distribution of the reaction products can be shown to originate as a consequence of counter-current mass flows in which different aqueous species diffuse in opposite directions due to imposed boundary conditions. In the case of opposing diffusion currents containing components which are capable of reacting to form a solid precipitate, a well-defined precipitation zone will form. The positions of these zones are a function of the boundary conditions and the strength of the sources and sinks. The calculation emphasizes the importance of knowing whether or not the system was open or closed with respect to fluid flow and/or gas exchange and shows that the Peclet number is a key parameter as it not only affects the spatial distribution of the alteration products but also effects the magnitude of the precipitation flow.

Abstract The potential of a sandstone to serve as a reservoir for producible hydrocarbons is closely related to its diagenetic history, which, in turn, is dependent upon its composition. Convergent plate-margin basins, which formed on the flanks of contemporaneous or precursor andesitic volcanic arcs, received large volumes of detritus reworked from either the arc of its tectonically uplifted plutonic roots. Arc-derived sands of the Bristol, Gulf of Alaska, Queen Charlotte, and Grays Harbor-Chehalis basins (northeast Pacific) are mineralogicaUy similar suites dominated by plagioclase, and by volcanic rock fragments. Diagenesis of these mineralogically immature sands at shallow to moderately deep burial (15,000 feet or 5,000 m) produced a systematic sequence of authigenic cements that include (1) localized, early calcite pore fill; (2) clay rims; (3) laumontite or phyllosilicate pore fill; and (4) late calcite pore fill/replacement and siliceous overgrowths. Each diagenetic stage reduced reservoir porosity and permeability and, correspondingly, increased bulk density and interval velocity of sand sections. Well-developed clay rims in fine-grained sand reduced maximum permeability to a few tens of millidarcys; samples containing stage 3 laumontite or phyllosilicate typically have less than 10 millidarcys permeability. Authigenesis of stage 2 and 3 cements is, in part, temperature controlled. Reservoir properties thus degenerate systematically with increasing depth of burial and increasing degree of heating, and they may produce a diagenetically controlled economic basement well within the range of normal drilling. Generation of oil is also temperature controlled, and the optimum thermal window lies mainly within sediments characterized by stage 3 diagenesis and consequent poor reservoir quality. Prediction of a diagenetic economic basement is necessary for realistic early assessment and systematic exploration of convergent arc-related plate-margin basins.

Abstract The Stanley and Jackfork Groups of the Ouachita Mountains consist of 18,000 feet of interbedded sandstones and shales deposited during the Late Mississippian and Early Pennsylvanian. Interest in their hydrocarbon potential has led to study of textures, compositions, and diagenetic alterations of these sandstones. The data and conclusions presented in this study are based on petrographic examination and porosity and permeability measurements of 187 samples collected from outcrop. The Stanley sandstones are generally poorly sorted, very fine-grained feldspathio- and quartz wackes. They average 8% feldspar, 14% matrix, and 5% silica cement. Porosities range from 0.5-26% and permeabilities from 0.05-23 md. Jackfork sandstones are predominantly moderately to poorly sorted, fine- to very fine-grained quartz arenites. They contain an average of 2% feldspar, 5% matrix, and 9% quartz cement. Porosities range from 0.5-14% and permeabilities from 0.05-9 md. Pressure solution, silica cementation, and replacement of plagioclase by calcite have acted to reduce reservoir potential in both units, whereas corrosion and dissolution of framework grains have added secondary porosity. The presence of halloysite and kaolinite characterizes sandstones affected by surface leaching. Well sorted quartz arenites have poor reservoir quality as a result of extensive silica cementation. Characteristics associated with the retention or secondary development of reservoir potential include poor sorting, small mean grain sizes, and high matrix content.

Abstract Tertiary-Holocene continental volcanic sediments in southern Guatemala were deposited in three basins each approximately 150 km long and 50 to 100 km wide. Boundaries between these basins mark the positions of transverse breaks in the underlying lithospheric slab. The rates of sediment accumulation and subsidence, and the composition, texture, and thickness of sediments can be expected to vary greatly from basin to basin. Exploration for hydrocarbon reservoirs in similar fore-arc areas should take into account that large scale correlations along depositional strike may be virtually impossible over distances greater than 50 or 100 km. The nature, degree, and timing of the diagenesis of Guatemalan volcaniclastics poses problems as to the hydrocarbon production potential of similar areas in other parts of the world. Guatemalan continental volcaniclastics are feldspathic litharenites. Three successive episodes of diagenesis have resulted in the precipitation of the following sequence of minerals in rock pores:—hematite-goethite;—montmorillonite plus hematite;—montmoril- lonite plus heulandite. These minerals occur as secondary pore-linings and pore-fills. Diagenesis has resulted in significant reduction of original permeability within 2000 years of deposition. Data indicate that the diagenesis of these rocks may not be dependent on depth of burial or temperature. The major diagenetic control was the chemistry of the groundwater. Introduction of groundwater resulted in the solution of unstable components (glass, pyroxene) and penecontemporaneous precipitation from pore-fluids of dissolved ionic species. Specific cements formed in the sands at any time are considered to be a function of the chemistry of the pore-fluids. Boundaries between different diagenetic assemblages in these rocks are probably determined more by ground water chemistry than by temperature or pressure.

Abstract The Eocene and Paleocene Umpqua Formation in the southern part of the Oregon Coast Range comprises a thick sequence of lithic arenites, siltstones, mudstones, conglomerates, coals, and, in the basal part, basaltic volcanic rocks intercalated with detrital sedimentary rocks. The sediments were deposited in a basin that developed during the evolution of the Mesozoic-Cenozoic arc-trench system of western North America. Environments ranging from deltaic to moderately deep marine are reflected in the rocks. Lithic arenites, the dominant sandstone type, have framework constituents of quartz, feldspars, micas, microfossils, plant fragments, heavy minerals, and volcanic, metamorphic and sedimentary lithic fragments. Digenetic minerals include phyllosilicates (chlorite and clay minerals), calcite, iron oxides, quartz, and zeolites. PhyQosilicate cements occur in three varieties: clay coats on framework grains, pore-filling chlorite with a radiating habit, and chlorite as unoriented microcrystalline aggregates. The radiating chlorite is found only in the upper and middle members of the formation, and the zeolite is found in only the lower member. Sandstone porosity has been reduced by the cements and by compaction and mechanical deformation of soft grains. A progressive sequence of diagenetic features from youngest to oldest evident in the upper and middle members is: (1) calcite pore-fill cement and the development of clay coats around framework grains, (2) precipitation of radiating pore-fill chlorite, or alteration of volcanic fragments to form unoriented microcrystalline aggregates of phyllosilicates, and (3) precipitation of silica cement in the center of pores not already completely filled. In the lower member, pore-space was not present for precipitation of the radiating chlorite. Zeolites occur in the lower member, indicating that low-grade metamorphic conditions were attained.