Along with calcite and aragonite, dolomite makes up approximately 2 percent of the Earths crust. The bulk of the dolomite constitutes dolostone formations that occur as thick units of great areal extent in many sequences of chiefly marine strata. (The rock dolostone is referred to by only the mineral namei.e., dolomiteby many geologists.) The Dolomite Alps of northern Italy are a well-known example. Other relatively common occurrences of the mineral dolomite are in dolomite marble and dolomite-rich veins. It also occurs in the rare igneous rock known as dolomite carbonatite.
From the standpoint of its origin, the dolomite of dolostones is one of the most interesting of all the major rock-forming minerals. As discussed below, a large percentage of the dolomite in thick marine dolostone units is thought by many geologists and geochemists to have been formed by replacement of CaCO3 sediment rather than by direct precipitation.
Ferrous iron commonly substitutes for some of the magnesium in dolomite, and a complete series very likely extends between dolomite and ankerite [CaFe(CO3)2]. Manganese also substitutes for magnesium, but typically only to the extent of a few percent and in most cases only along with iron. Other cations known to substitutealbeit in only relatively minor amountswithin the dolomite structure are barium and lead for calcium and zinc and cobalt for magnesium.
Nearly all the natural elements have been recorded as present in at least trace quantities in dolostones. It is, however, unclear which ones actually occur in the dolomite; some of them may occur within other mineral constituents of the analyzed rocks. Indeed, only a few of these elementse.g., strontium, rubidium, boron, and uranium (U)are known definitely to occur within the dolomite structure.
Dolomite effervesces with dilute hydrochloric acid, but slowly rather than vigorously as calcite does; in general, it appears to smolder slowly, and in some cases it does so only after the rock has been powdered or the acid warmed, or both. This difference in the character of the effervescence serves as the test usually used to distinguish dolomite from calcite in the field. In the laboratory, staining techniques, also based on chemical properties or typical compositions, may be used to distinguish between these minerals. The stains generally employed are especially valuable for investigating rocks made up of alternate lamellae of dolostone and limestone composition.
In a somewhat simplified way, the dolomite structure can be described as resembling the calcite structure but with magnesium ions substituted for calcium ions in every other cation layer. Thus, the dolomite structure can be viewed as ideally comprising a calcium layer, a CO3 layer, a magnesium layer, another CO3 layer, and so forth. However, as described for the potassium feldspars, dolomitesunlike calcitesmay also exhibit order-disorder relationships. This results because the purity of some of the cation layers may be less than ideali.e., some of the calcium layers may contain magnesium, and some of the magnesium layers may contain some calcium. The term protodolomite is frequently applied to Holocene dolomites (those formed during approximately the last 11,700 years) that have less than ideal dolomite structures. Most dolomites of ancient dolostones, however, appear to be well ordered. Modifications that may reflect diverse calcium-versus-magnesium layering aberrations are treated extensively in professional literature.
Dolomite crystals are colourless, white, buff-coloured, pinkish, or bluish. Granular dolomite in rocks tends to be light to dark gray, tan, or white. Dolomite crystals range from transparent to translucent, but dolomite grains in rocks are typically translucent or nearly opaque. The lustre ranges from subvitreous to dull. Dolomite, like calcite, cleaves into six-sided polyhedrons with diamond-shaped faces. Relations between lamellar twinning and cleavage planes of dolomite, however, differ from those of calcite (see figure), and this difference may be used to distinguish the two minerals in coarse-grained rocks such as marbles. Dolomite has a Mohs hardness of 31/2 to 4 and a specific gravity of 2.85 0.01. Some dolomites are triboluminescent.
The dolomite of most dolostones is granular, with the individual grains ranging in size from microscopic up to a few millimetres across. Most dolomite marbles are coarsely granular with individual grains ranging between 2 and 6 millimetres (0.079 and 0.24 inch) in greatest dimension. Vein dolomite grains may be up to several centimetres across. Saddle-shaped groups of dolomite crystals, most of which occur on fracture surfaces, measure from 0.5 to 2 centimetres (0.20 to 0.79 inch) across.
Dolomite occurs widely as the major constituent of dolostones and dolomite marbles. As mentioned above, the origin of dolomite-rich rocks in marine sequences remains an unresolved problem of petrogenesis.
Dolomiteactually protodolomiteis known to have formed fairly recently in restricted environments such as on supratidal flats that occur in The Bahamas and Florida Keys. Also, no dolomite has been synthesized in an environment comparable to natural conditions. Thus, the explanation for the formation of dolomite in these marine units remains in question. It is now thought that dolostones may be of various origins. Indeed, several different models have been suggested for dolomite formation, each based on diverse considerations, combined with empirical and/or experimental data.
Except for models invoking formation of dolomite by direct precipitation, a process thought by most geologists to apply to only a small percentage of all dolostones, each model is based on the assumption that the dolomite of dolostones has been formed by conversion of CaCO3 sediment or sedimentary rocks to dolostone. Thus, the models have been formulated to account for this conversion, which is known as dolomitization.
The most widely discussed models for dolomitization, either partial or complete, involve four chief variables: time, location with respect to the sediment-seawater interface, composition and derivation of the solutions involved, and fluxing mechanisms. The time ranges from dolomitization that occurs penecontemporaneously with deposition to that which takes place subsequent to relatively deep burial of the precursor sediments. The location ranges from at or very near the sediment-seawater interface to well beneath some overlying sediments that were deposited at a later time. The solutions supply the magnesium needed and must have the appropriate pH and concentrations of other necessary ions; these solutions are generally considered to be seawater (either normal seawater or brines concentrated by evaporation), connate water, meteoric water, or some combination of these waters. (Connate refers to water that becomes enclosed within sediments upon their deposition; meteoric water is derived from the atmosphere as rain or snow, which often occurs in pore spaces within rocks.) Another important variable is the presence of dissolved sulfate (SO4 2) ions, as this retards the dolomitization process. The fluxing mechanisms are generally attributed to density differences of the solutions involved and the permeability characteristics available for percolation through the precursor sediment. In addition, the presence of a geothermal heat source in a basin may enhance both fluid flux and the rate of dolomitization. There also are additional direct and indirect controlse.g., climate, biochemical processes, and HDO:H2O and/or D2O:H2O ratios in the water. (The symbol D represents deuterium, the hydrogen isotope with a nucleus containing one neutron in addition to the single proton of the ordinary hydrogen nucleus.) Bacteria may also play a role in the formation of dolomite. In any case, it has been shown that some dolostones have gained their current characteristics as a consequence of certain combinations of these conditions and processes.
Criteria involving factors such as the identity of associated rocks and the coarseness of the grains of dolostones have been suggested for use in attributing one versus the other hypothesized models to certain occurrences of dolostone. None, however, has been accepted as an absolute criterion by many carbonate petrologists.
The desire for an understanding of dolomitization of sedimentary strata has been based on economic as well as scientific interests. In many places, dolomitization has led to increases in permeability and porosity and thus increased the potential of such rock strata as good oil, gas, and groundwater reservoirs and, in some cases, even as hosts of certain kinds of ore deposits.
The other fairly common dolomite occurrences include the following: Dolostones have been metamorphosed to both dolomite and calcite marbles; dedolomitization processes account for the latter. Some dolomite marbles are nearly pure dolomite. Dolomite carbonatites are of the same general origin as calcite carbonatites. The dolomite present in dolomite veins has also been ascribed diverse origins; some appears to have been deposited by percolating connate or meteoric groundwater, and some seems more likely to have been deposited by hydrothermal solutions charged with magmatic volatiles.
Dolomite is a common rock-forming mineral. It is a calcium magnesium carbonate with a chemical composition of CaMg(CO3)2. It is the primary component of the sedimentary rock known as dolostone and the metamorphic rock known as dolomitic marble. Limestone that contains some dolomite is known as dolomitic limestone.
Dolomite is rarely found in modern sedimentary environments, but dolostones are very common in the rock record. They can be geographically extensive and hundreds to thousands of feet thick. Most rocks that are rich in dolomite were originally deposited as calcium carbonate muds that were postdepositionally altered by magnesium-rich pore water to form dolomite.
Dolomite is also a common mineral in hydrothermal veins. There it is often associated with barite, fluorite, pyrite, chalcopyrite, galena, or sphalerite. In these veins it often occurs as rhombohedral crystals which sometimes have curved faces.
Dolostone: Dolostone from Lee, Massachusetts. The "sugary" sparkle displayed by this rock is caused by light reflecting from tiny dolomite cleavage faces. This specimen is approximately 4 inches (10 centimeters) across.
The physical properties of dolomite that are useful for identification are presented in the table on this page. Dolomite has three directions of perfect cleavage. This may not be evident when the dolomite is fine-grained. However, when it is coarsely crystalline the cleavage angles can easily be observed with a hand lens. Dolomite has a Mohs hardness of 3 1/2 to 4 and is sometimes found in rhombohedral crystals with curved faces. Dolomite produces a very weak reaction to cold, dilute hydrochloric acid; however, if the acid is warm or if the dolomite is powdered, a much stronger acid reaction will be observed. (Powdered dolomite can easily be produced by scratching it on a streak plate.)
Dolomite is very similar to the mineral calcite. Calcite is composed of calcium carbonate (CaCO3), while dolomite is a calcium magnesium carbonate (CaMg(CO3)2). These two minerals are one of the most common pairs to present a mineral identification challenge in the field or classroom.
The best way to tell these minerals apart is to consider their hardness and acid reaction. Calcite has a hardness of 3, while dolomite is slightly harder at 3 1/2 to 4. Calcite is also strongly reactive with cold hydrochloric acid, while dolomite will effervesce weakly with cold hydrochloric acid.
Dolomite occurs in a solid solution series with ankerite (CaFe(CO3)2). When small amounts of iron are present, the dolomite has a yellowish to brownish color. Dolomite and ankerite are isostructural.
Kutnahorite (CaMn(CO3)2) also occurs in solid solution with dolomite. When small amounts of manganese are present, the dolomite will be colored in shades of pink. Kutnahorite and dolomite are isostructural.
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The most common use for dolostone is in the construction industry. It is crushed and sized for use as a road base material, an aggregate in concrete and asphalt, railroad ballast, rip-rap, or fill. It is also calcined in the production of cement and cut into blocks of specific size known as "dimension stone."
Dolomite serves as the host rock for many lead, zinc, and copper deposits. These deposits form when hot, acidic hydrothermal solutions move upward from depth through a fracture system that encounters a dolomitic rock unit. These solutions react with the dolomite, which causes a drop in pH that triggers the precipitation of metals from solution.
Dolomite also serves as an oil and gas reservoir rock. During the conversion of calcite to dolomite, a volume reduction occurs. This can produce pore spaces in the rock that can be filled with oil or natural gas that migrate in as they are released from other rock units. This makes the dolomite a reservoir rock and a target of oil and gas drilling.
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Dolomite is an earthy mineral. It is a Carbonate mineral, also known as Calcium Magnesium Carbonate. Generally it is white to off white in color. Its chemical formulae is CaMg(CO). This mineral crystallizes in the Trigonal system. Gujarat in INDIA is rich in Dolomite deposits. Dolomite is a mineral that has a special saddle-shaped crystal. Dolomite is a common rock-forming mineral.
Dolomite originates in the same sedimentary environments as limestone warm, sand place, shallow, saltwater environments. In agriculture, dolomite is used as a soil conditioner and as a feed additive for livestock. Our Company is a manufacturer & supplier of dolomite in India. We maintain high-quality standards throughout the process of dolomite. We are a leading company in the field of mineral suppliers. We are a manufacturer, supplier, and exporter of Dolomite in India. We offer Dolomite in the best quality range.
In many ancient limestones, the minerals calcite and aragonite have been partly or completely replaced by the mineral dolomite, producing replaced by the mineral dolomite, producing the rock dolostone (or traditionally, the rock dolomite). Such recrystallization commonly obliterates the original limestone texture and constituents. Where relict original limestone texture and components survive, a dolomite can be classified in exactly the same manner as a limestone and interpretation of environmental setting is possible. However, complete dolomitization renders proper naming and genetic interpretation tenuous, if not impossible.
The exact mechanisms and timing by which limestone alters to dolomite are not thoroughly understood. The proportion of dolomite relative to limestone increases markedly backward in geological time. A limestone to dolomite ratio of 10:1 characterizes Mesozoic carbonates; in the Precambrian this ratio is 1:3. Modern carbonates, on the other hand, are almost exclusively limestone. This is one of the most obvious trends visible in the earth's sedimentary rock record. Its cause has been hotly debated! Some sedimentary petrologists suggest that the progressive increase in the amount of dolomite relative to limestone backward in time implies an atmosphere and/or ocean chemistry earlier in earth history more compatible with dolomite formation (either as primary dolomite, or secondary dolomite altered from limestone). An alternative view is that the bulk of all dolomites are secondary; that is, that any limestone, given enough time, will eventually react with dolomitizing fluids and become dolomite. In this view, the backward increase in amount of dolomite at the expense of limestone simply directly reflects increases in the amount of time available for dolomitization to occur. The longer the time, the more likely the process is to happen.
The exact circumstances favorable for the formation of dolomite today can be constrained by observations at several modern sites of dolomite formation. Well-known localities include a few selected spots in the Bahama Banks, around the Arabian Gulf, on Bonaire Island in the Caribbean, and in the Coorange Lagoon of Australia. In all of these so-called primary dolomite localities, dolomite apparently forms by almost instantaneous reaction of newly precipitated calcite or aragonite crystals with seawater and/or percolating ground water. The geochemistry of the dolomitizing water is unusual. Salinity is higher than normal seawater (3.5%) and the ratio of dissolved magnesium to calcium is 3 to 4 times that of normal seawater. Such primary dolomite localities are either supratidal flat areas or hypersaline lagoons. The production of the dolomitizing fluids results from the evaporative conditions. Evaporation increases the salinity of seawater to the point where chemical precipitation of aragonite (CaCO3) and gypsum (CaSO4 2H2O) can occur. Formation of these minerals selectively removes calcium from seawater and raises the magnesium/calcium ratio above the critical threshold level necessary for a fluid with dolomitizing capability. The overall process is referred to as evaporative reflux because evaporation is the fundamental mechanism for producing the fluid and because the process itself is aided by the flushing (reflux) of the denser, more saline solution through previously deposited limestone.
Many ancient dolomites probably formed by the evaporative reflux mechanism, particularly the so-called bedded dolomites (for example, the Ordovician Knox Group and Beekmantown Dolomite of the Appalachians). These bedded dolomites are regionally extensive units of alternating beds of limestone, dolomite, and evaporite. Individual beds, each several centimeters to meters thick and continuing across broad regions, probably result from the dolomitization of previously deposited limestone almost immediately after their formation. The original aragonitic or calcitic sedimentary bands presumably reacted almost immediately with brines produced across supratidal flats, hence the ubiquitous and total replacement of thin bands. Various sedimentary structures indicative of supratidal and intertidal environments in which evaporation would likely occur (mudcracks, stromatolites, salt crystal imprints, interbedded evaporate seams, fossils indicative of hypersalinity) are commonly associated with such bedded dolomites.
Other ancient dolomites have a more problematical origin. Particularly puzzling are the patchlike masses of dolomite that occur sporadically within ancient limestone units, as random dolomite crystals scattered through a limestone framework, as dolomite cement lining cavities, and as a replacement or partial replacement of normally calcite or aragonite shelled organisms. A variety of mechanisms have been suggested to explain these kinds of deposits. Almost all invoke the reaction of preexisting limestone with a dolomitizing fluid long after the initial crystallization of the limestone. One popular mechanism is simply a time-delayed variation of evaporative reflux. Dolomitizing brines produced in a supratidal or intertidal setting (identical with the settings described above) are simply flushed down through deeper, underlying limestones. Such fluids would be propelled downward by their higher density, though some might follow fault zones, cracks, and fissures. Other mechanisms of dolomitization invoke complex chemical reactions between seawater and percolating meteoric water along the coastline. The exact mechanisms of dolomitization have obviously yet to be pinpointed!
The masses of CO2 sequestered during the considered process through dissolution (and speciation) in the aqueous solution (solubility trapping) and through incorporation in solid carbonates, mainly magnesite and subordinately siderite and dolomite (mineral fixation) are plotted against both the reaction progress and time inFig. 7.8. The total mass of sequestered CO2, corresponding to the sum of these two terms, is also shown.
Figure 7.8. High-pressure (Ptot = PCO2 = 100 bar) CO2 injection in a deep aquifer (T = 60C) hosted in serpentinitic rocks: masses of CO2 sequestered through dissolution in the aqueous solution (solubility trapping) and incorporation in solid carbonates (mineral fixation), and total mass of sequestered CO2 as a function of time and the reaction progress variable. Note that letters A to E mark the onset of the precipitation of different secondary minerals as shown in Figure 7.5.
Inspection ofFig. 7.8 shows that CO2 sequestration through dissolution in the aqueous solution is instantaneous, but relatively limited, as the maximum amount of CO2 which can be dissolved in 1 kg of water is 50 g, at the specified conditions of 100 bar Ptot = PCO2, and 60 C. Carbon dioxide incorporation in secondary solid carbonates begins later on, namely at points B (dolomite), C (magnesite) and E (siderite). The log-mass of CO2 incorporated in precipitating solid carbonates increases linearly with log , attaining values much greater than the mass of CO2 dissolved in the aqueous solution, but this process requires comparatively long time intervals. For instance, the contribution of mineral fixation attains the same level of solubility trapping after 5 years. However, mineral carbonation becomes more and more important afterwards, whereas solubility trapping does not change with time. Again, the sequestration capacity of the process is large and time is less than the residence times of high-pH waters in deep aquifers. In particular the value of 2,000 g kg1 water is attained in 350 years. The same sequestration capacity is attained in 1,000 years in the dunite model by Xu and coworkers, who used somewhat different figures for several parameters (seeSection 7.3.1).
Hydrocarbon reservoirs are porous and permeable lithological units or a set of units capable of holding the hydrocarbon reserves. It comprises of one or more subsurface rock formations of either sedimentary or carbonate origin. Reservoir rocks are characterized by good porosity and permeability and bounded by impermeable barriers, like shales, that trap the hydrocarbons. A cross-section of a simple reservoir is shown in Figure 5.1.
A hydrocarbon reservoir may contain liquid, gas, or both, and the vertical occurrence of fluids in the structure is governed by the gravitational segregation. Hydrocarbon reservoirs do have a seal known as cap rock, which is of low permeability that impedes the escape of hydrocarbons from the reservoir rock. Common seals include evaporites, chalks, and shales.
Main reservoir rocks are either sandstones or carbonates. More than 60% of the world oil reserves are found in sandstone reservoirs. Sandstone is a sedimentary rock composed mainly of sand-sized rock grains cemented together. Most of the sandstones are composed of quartz (SiO2) and/or feldspar because these are the most common minerals in the Earth's crust. Reservoirs that are primarily composed of sandstone allow percolation fluids and are porous enough to store large quantities of hydrocarbons. There are many types of sandstones like shaly, carbonate, etc., and one can refer to Pettijohn et al. (1987) and Folk (1965) for a detailed description and classification of sandstones.
Carbonate rocks are also a class of sedimentary rocks that are composed primarily of carbonate minerals. The two major types of carbonate rocks are limestone (CaCO3) and dolostone, primarily composed of the mineral dolomite (CaMg(CO3)2). Carbonate rocks can be of various origins like:
Oil reservoirs can be classified as saturated and undersaturated reservoirs. The degree of saturation in a gas-saturated reservoir is a function of reservoir pressure and temperature. The bubble point pressure (pressure at which gas begins to come out of the solution) in a reservoir is either equal or less than the reservoir pressure. If the bubble point is equal to reservoir pressure, oil in the reservoir is gas saturated. This means oil has dissolved in all of the gas it is capable of holding under given conditions. Oil in the reservoir is undersaturated if there is less gas present in the reservoir than the amount that may be dissolved in oil under given conditions. In the case of saturated oil, gas begins to come out of the solution as soon as the reservoir pressure begins to decrease, but in the case of unsaturated oil, the dissolved gas does not start coming out of solution until the reservoir pressure drops to the level of the bubble point. Presence of a gas cap in a reservoir always indicates saturated oil.
Undersaturated oil reservoirs are usually one phase. There is insufficient gas to saturate the oil; hence, there is no excess gas to form a gas cap. In an under saturated reservoir, single-phase hydrocarbon is produced by solution gas drive or expansion drive as the reservoir pressure is drawn down until the bubble point is reached. After the bubble point pressure is reached, the production mechanism will be gas cap driven, the same as the originally saturated reservoir.
In an under saturated reservoir, a produced gasoil ratio (GOR) is constant. Gas is only released from the oil when the pressure drops to the bubble point at the surface, but in a saturated reservoir, the bottom-hole pressure is always less than the bubble point pressure and the produced GOR is always higher than the GOR solution. Produced GOR will increase more with a decline in reservoir pressure. Thus, while taking oil samples for pressure volume and temperature (PVT) analysis, the reservoir condition should be close to initial state. While sampling, the flow should be single-phase oil and no excess gas (McAleese, 2000). The bubble point pressure of the oil generally decreases with increasing depth for a given reservoir, and thus the composition of oil varies vertically and we get lighter oil at the top (Coss, 1993). As the bubble point decreases, the dissolved gas comes out of the solution; this makes the oil more viscous and heavy and affects the recovery factor. Apart from affecting the recovery, dissolved gas has an important effect on the volume of production as the release of dissolved gas at the surface will shrink the oil volume produced. Generally, the physical properties of reservoir fluids are determined using samples in a laboratory or by using charts and graphs of empirically derived data.
Gas reservoirs can be classified as: (1) retrograde condensate gas reservoirs and (2) dry gas reservoirs (McAleese, 2000). In some reservoirs, retrograde condensation occurs in petroleum gases containing heavy hydrocarbons as single-phase fluids in deep reservoirs at high pressure and temperature. Reduction in pressure at a constant temperature may hit the dew-point curve, which leads to retrograde condensation. This means that production may cause rapid condensation of hydrocarbons. In this case, well-stream composition changes with depletion, which means the gas produced will be depleted of initial heavy hydrocarbons. Hence, for PVT analysis, samples should be collected at the initial reservoir condition. In case of dry gas reservoirs, temperature is higher than the critical condensation temperature of the reservoir fluid mixture and bottom-hole pressure is above the dew point. Dry gases consisting of pure methane and ethane are produced from these reservoirs.
Fluids present in the reservoir differ significantly in volume and quality when they reach the surface because of change in pressure and temperature. The light oils produce a lot of gas at the surface compared to the heavy oils. Dry gas reservoirs yield only gas at the surface; however, condensate reservoirs yield a lot of condensates.
The clastic medium granular sediments are represented by sands as unbound sediments, and sandstones as solid rocks. The sands and sandstones are sediments that are predominantly composed of detrite grains sizes of sand with grain in diameter ranging between 0.063 and 2.00mm. The rocks are characterized typically by dominance of sand size grains with minor share of powder size clay particles and of tiny gravel. The fundamental material include grains of sand derived from weathering and erosional component of any rock (Pettijohn et al., 1972).
The mineral composition of sands and sandstone can be very different and complex depending on the parent rocks, nature of weathering and erosion, transfer, and deposition. The clasts that makeup sand residue and sandstone include mineral grains and rock fragments of siliciclastic and carbonate composition, as well as fossil remains of the skeleton and shells of organisms, that is, fossil detritus.
The siliciclastic components include all grains of quartz, silicate minerals, and rock fragments containing quartz and silicate minerals all in the form of clasts, muddy, and clay matrix, and ingredients left after the physical and chemical weathering and erosion of silicate minerals and rocks. These grains are transferred to the precipitation area from land (terrigenous components). The carbonate components and carbonate detritus are carbonated grains, mostly fragments of limestone, dolomite, and fragments of calcite and dolomite minerals remaining from wear of carbonate rocks and minerals, primarily calcite, dolomite, and siderite veins. The carbonate detritus in its origin may be either of the following:
Intrabasinal belongs to ooids, oncoids, and pellets formed in the surrounding shallows and even intraclasts that originate from the destruction of carbonate rocks within the depositional area, and are nearly as old as the sand in which deposited. It is often the case in Badenian sediments in Pannonian basin, east-central Europe, in which siliciclastic material derived from weathering of older crystalline and lower Miocene rocks on mainland. The intrabasinal carbonate detritus from the destruction of reef Badenian limestone from coastal shallows, and underwater reefs. These are calcarenaceous sandstones (Section 220.127.116.11).
The fossil components or fossil detritus include the fossil remains of flora and fauna in the form of whole shells and/or skeletons or their fragments known as bioclasts. The fossil detritus in sandy sediments may originate from resedimentation from older rocks or carbonate detritus. It may be intrabasinal belonging to planktonic and benthic organisms residing within the depositional area. The redeposited fossil detritus from older Baden corallinaceabryozoa ridge rocks are often found in Sarmatian and Pannonian sandstones and intrabasinal fossil components (bioclasts of corallinacea, bryozoa, echinoderms, and mollusks) in Baden biocalcarenites sandstones of Pannonian basin.
The essential ingredients of sands and sandstones are quartz, feldspar, and rocks fragmentsmicas, carbonate, and clay minerals, and heavy minerals (density >2.85g/cm3). Certain types of sandstone can contain a substantial proportion of muddy matrix, fossil detritus, or glauconite.
Feldspars are particularly abundant ingredients of some sands and sandstone, especially molasse type, whose detritus derived from severe physical wear and rapid deposition at the foot of mountain massifs built from neutral and acidic igneous rocks and gneisses
Excerpts of quartz and feldspars originate from wear of mafic intrusive (plutonic), and intermediate extrusive (volcanic) igneous rocks, numerous sedimentary rocks (in particular, siltstone, sandstone, chert, limestone, and dolomite), and many metamorphic rocks (especially quartzite, phyllite, mica schist, and gneiss) are primary ingredients of many sands and sandstone.
Carbonates are predominantly salts of carbonic acid with Ca2+, Mg2+, and Fe2+ as the common cations. They are formed in aqueous environments (oceans) and deposited through sedimentary (chemical and biologically mediated) processes on the continental shelf (Scholle et al., 1983). The carbonate rocks formed in these sedimentary environments are mainly limestones and dolostones (dolomitic limestone). Limestones are predominantly made up of calcite (CaCO3) while dolostones or dolomitic limestones are made up of the mineral dolomite [CaMg(CO3)2]. Thick sequence of carbonates are found in all continents as a result of marine transgression and regression of shallow marine seas that covered the stable continental areas from time to time, during the late Precambrian, Paleozoic and Mesozoic eras (Scholle et al., 1983). Ancient carbonate rocks contain quartz (SiO2) as a common impurity but also minor amounts of calcite group minerals such as magnesite (MgCO3), rhodochrosite (MnCO3) and siderite (FeCO3) are also found in some carbonates. Modern carbonate sediments are composed almost entirely of aragonite (CaCO3) and Mg-rich calcite; both recrystallize during diagenesis to form calcite. In the oceans, the calcium carbonate precipitates out of ocean water to form layers of non-terrigenous sediment on the ocean floor. Carbonates make up the shells of diverse organisms and are often associated with chert and opal, which are precipitates of silica by different organisms. The carbonate and silica sediments accumulate on the ocean floor to form sedimentary deposits. Plate tectonic process leading to subduction, or an intrusive contact, both bring a mineralogical transformation consistent with the degree of thermal and pressure variables. This fluid-driven mineralogical transformation leading to recrystallization without melting is the process of metamorphism. After metamorphic transformation, limestone and dolomite develop into marble, while in the presence of silica they produce a suite of calc-silicate minerals coupled with decarbonation reactions.
Impure carbonate rocks are mainly composed of calcite, dolomite, and quartz. Recrystallized limestones containing dolomite and quartz are very useful indicators of metamorphic grade because of the considerable range of Ca-Mg silicate minerals formed during prograde metamorphism. The bulk composition of sedimentary carbonates can be represented in the CaO-MgO-SiO2 system. But these impure carbonate rocks also have H2O-rich pore fluid in addition to CO2 gas which is released during metamorphism. The carbonate rocks thus define the chemical system CaO-MgO-SiO2-H2O-CO2 (Fig. 1). Symbols for minerals in figures and text are given after Kretz (1983).
The reactions in this system depend on T and fluid compositionmole fraction of CO2=CO2CO2+H2O , such that the prograde metamorphic sequence of minerals is conventionally represented in a T-XCO2 diagram and the phase relations are displayed in the triangular phase diagram with SiO2-CaO-MgO at its three corners. Minor amounts of other components and accessory phases do not notably change the reactions among the Ca-Mg-silicates discussed below in light of the experimental investigations by Greenwood (1967), Eugster and Skippen (1967), Metz and Puhan (1970) and Skippen (1971, 1974).
Ideal, ordered dolomite has a formula of CaMg(CO3)2 and consists of alternating layers of Ca2+CO32Mg2+CO32Ca2+, etc., perpendicular to the crystallographic c axis. Most natural dolomite contains up to a few per cent Ca surplus (and a corresponding Mg deficit), as well as less than ideal ordering. Protodolomite contains about 5560% Ca, is poorly ordered, i.e., the alternating cation layer structure is poorly developed, and is common as a metastable precursor of well-ordered, nearly stoichiometric dolomite in both laboratory experiments and in nature. Good arguments have been made to abandon the term protodolomite or to restrict it to laboratory products, yet the term is useful to describe metastable precursors of dolomite in nature. The term dolostone refers to a rock that consists largely (>75%) of the mineral dolomite. This term has been rejected by some, but has gained wide acceptance during the last 20 years. The term dolomites is the best term to use to refer to types of dolomite that vary in texture, composition, genesis, or a combination thereof.
Two types of dolomite formation are common, i.e., dolomitization, which is the replacement of CaCO3 by CaMg(CO3)2, and dolomite cementation, which is the precipitation of dolomite from aqueous solution as a cement in primary or secondary pore spaces. Dolomites and dolostones that originate via replacement of CaCO3 are called replacement dolomites or secondary dolomites, especially in the older literature. A third type of dolomite formation is direct precipitation from aqueous solution to form sedimentary deposits. Dolomites that form in this way may be called primary dolomites.
Genetically, all natural dolomites can be placed into two major families, i.e., penecontemporaneous dolomites and postdepositional dolomites. Penecontemporaneous dolomites may also be called syndepositional dolomites. They form while a carbonate sediment or limestone still resides in the original environment of deposition as a result of the geochemical conditions that are normal for that environment. Such dolomites are also called primary or early diagenetic, although these terms are not strictly synonymous with penecontemporaneous. True penecontemporaneous dolomites appear to be relatively rare. Most known cases are of Holocene age, and are restricted to certain evaporitic lagoonal and/or lacustrine settings. It is quite possible, however, that such dolomites are much more common in the geological record than presently known, but their presence is hard to prove because of later diagenetic overprinting.
Postdepositional dolomites may also be called postsedimentary. They form after a carbonate sediment has been deposited and removed from the active zone of sedimentation, which may happen via progradation of the sedimentary surface, burial and subsidence, uplift and emergence, eustatic sea-level fluctuations, or any combination of these. Such dolomites and dolostones are often called late diagenetic, although this term is not synonymous with postdepositional. Almost all known examples of massive, regionally extensive dolostones are postdepositional.
One aspect that transcends the above genetic grouping is that of hydrology. Whether syndepositional or postdepositional, the formation of large amounts of dolomite requires advection, i.e., fluid flow, because of chemical mass balance constraints. On the other hand, small amounts of dolomite can be formed without advection. In such cases, the Mg for dolomite formation is locally derived and redistributed, or supplied via diffusion. Examples include dolomite formed from Mg that was contained in (high-)Mg calcite, adsorbed to the surfaces of minerals, organic substances, or biogenic silica, or that was contained in older primary or secondary dolomites.
As a result of reactions (1)(5), the chemical character of water in caves is one rich in dissolved Ca2+ and HCO3. The pH of cave waters tends to be relatively high, because the protons released from carbonic acid dissociation (Eqs. 3 and 4) are consumed in the dissolution of calcite (Eq. 1). Eq. (1) demonstrates the molar relationships of Ca2+ and HCO3 in this system: two moles of HCO3 are generated for every mole of Ca2+. The typical groundwater collected from limestone terrain, then, has very large concentrations of dissolved HCO3 compared to groundwater in other geological settings (Table 2).
Unlike the hypothetical case of pure water in contact with pure calcite and a CO2-containing atmosphere (e.g., Table 1), Ca2+ and HCO3 are not the only dissolved chemicals in cave water or groundwater in limestone. Limestone bedrock, although predominantly made up of calcite, ordinarily has some other minerals present. Rainwater infiltrating through soil and groundwater migrating to greater depths comes in contact with these various solid phases that have some degree of solubility in water. As those minerals dissolve, other solutes enter aqueous solution.
In addition to the common occurrence of small amounts of dolomite in limestone bedrock, some caves are developed in dolostone, bedrock made up primarily of the mineral dolomite. Dissolved Mg2+ is, therefore, a major constituent in many cave waters (e.g., Table 2). Gypsum, CaSO42H2O, may be associated with calcite and dolomite in some sequences of sedimentary rock. If present, this highly soluble sulfate mineral can have a big impact on the chemical composition of cave waters
For instance, in the sample from Manatee County, FL, United States (Table 2), groundwater circulation brought water into contact with a gypsum layer at depth, and Eq. (7) proceeded to the right resulting in groundwater with higher dissolved SO42 concentration than HCO3 concentration.
The other common rock-forming elements sodium and potassium may enter solution as the ions Na+ and K+ from the dissolution of minerals or from ion-exchange reactions. Ion-exchange occurs among cations occupying the electrostatically charged surfaces of clay minerals in the limestone bedrock or in the sediments of cave streams according to
This mass-action expression (Eq. 8) illustrates the shift in the composition of the ion-exchange complex on the clay mineral surface with shifting concentrations of dissolved cations. When more dissolved Ca2+ is present from calcite dissolution, it tends to shift toward sorption onto the clay (i.e., Eq. 8 shifts to the right) with Ca2+ displacing Na+ from surface sites and Na+ entering aqueous solution. When more dissolved Na+ is present from saltwater intrusion, Na+ will move toward occupying more surface sites (i.e., Eq. 8 shifts to the left) releasing Ca2+ to aqueous solution.
The anions commonly present in cave waters may derive from any of several sources. Neither Cl nor NO3 are major rock-forming constituents. Both aqueous species have atmospheric sources, however, especially in regions where the atmospheric chemistry is impacted by industrial emissions that contribute to the chemical composition of rainfall. Cl may be introduced to groundwater in near-coastal settings by the dissolution of sea spray or marine aerosols into rainfall or by saltwater intrusion in limestone aquifers. Several samples in Table 2 have been influenced by the proximity of marine waters (Mallorca, Spain; Manatee County, FL, United States; Yucatan Peninsula, Mexico). Additionally, both solutes may enter solution as contaminants in waters influenced by animal fecal wastes, road salts, or agricultural chemicals.
Military applications of geology first became apparent in Napoleonic times. Indeed, the first general to take geologists as such on campaign was Bonaparte himself. The French army he led into Egypt in 1798 was accompanied by a civilian Commission of Sciences and Arts that included Dodat de Dolomieu (after whom the mineral dolomite was later named) (Figure 1) and several of Dolomieu's former geology students, recent graduates of the School of Mines in Paris notably Louis Cordier (later immortalized by the mineral name cordierite), Franois-Michel de Rozire, and Victor Dupuy. However, their role was that of mineralogists to support the army by exploring the geological resources of the country, rather than by contributing tactical or strategic advice. Karl von Raumer (17831865), from 1811 Professor of Mineralogy at the University of Breslau (now Wroclaw in Poland), served as a staff officer in Prussia's 181314 war of independence from Napoleonic domination, but was used to communicate dispatches rather than appraise geology. The distinguished Swiss mining geologist, Johann Samuel Gruner (17661824) (also known as von Grouner), having moved to Bavaria in 1803, became a captain in a volunteer rifle battalion when Bavaria joined the alliance against Napoleon late in 1813. Postwar in 1820, seemingly at the instigation of the Bavarian Bureau of Military Topography, he combined his military and geological experience to write a memorandum on the relationship between geology and military science. Published posthumously in 1826, this is the earliest known work in its field.
Figure 1. Dodat de Dolomieu (17501801), in 1798, by Andr Dutertre. A former cavalry officer and Knight of Malta, Dolomieu taught geology at the School of Mines, Paris, before serving as a senior scientist with Napoleon's expeditionary forces in Egypt. From portrait supplement to Saintine, X-B, Marcel JJ and Reybaud L (eds.) (18301836) Histoire Scientifique et Militaire de l'Expdition Franaise en gypte. Paris: Dnain & Delamare. By permission of the British Library, London (Shelfmark1311.h.2). Also from Rose EPF (2004) Napoleon Bonaparte's Egyptian campaign of 1798: the first military operation assisted by geologists? Geology Today 20: 2429, by permission of Blackwell Publishing Ltd.
In the UK, geological mapping was initially perceived as a military skill. From 1809 to 1814, military (Board of Ordnance) objectives and funding generated geological fieldwork by J. MacCulloch, and from 1814 to 1826 his geological mapping in Scotland. From 1826 to 1846, Royal Engineer officers (successively captains J.W. Pringle, J.E. Portlock, and H. James) pioneered government geological surveys in Ireland. The British Geological Survey was founded in 1835 and sustained until 1845 under military (Board of Ordnance) auspices, and its earliest directors (H.T. De la Beche until 1855, R.I. Murchison until 1871 (see FAMOUS GEOLOGISTS | Murchison)) were both men who had received a military rather than a university education. The world's oldest geological society, the Geological Society of London founded in 1807, included amongst its earliest influential members those also active in the reserve army (Lieutenant G.B. Greenough) or the regular army (T.F. Colby, J.W. Pringle, and J.E. Portlock), and a veteran of the Napoleonic wars in receipt of military half-pay throughout his geological career (Lieutenant W. Lonsdale).
Because of its evident practical applications, geology was soon introduced into the curriculum at many military colleges. It was taught at Addiscombe in Surrey to officer cadets of the East India Company's army, by J. MacCulloch from 1819 to 1835 and by D.T. Ansted from 1845 until college closure in 1861. For the British army, J. Tennant lectured on geology to Royal Engineer and Royal Artillery cadets at the Royal Military Academy, Woolwich, from 1848 to 1868; T. Rupert Jones to cadets of the infantry and cavalry at the Royal Military College, Sandhurst, from 1858 to 1870; Jones also to young officers of all arms at the Staff College, Camberley, until 1882; and A.H. Green at the School of Military Engineering, Chatham, from about 1888 to 1896. Jones (Figure 2) was employed full time as a military professor of geology, but the others were essentially visiting professors who concurrently held better known appointments elsewhere Ansted and Tennant for a while at King's College London, Green at Oxford. The United States Military Academy at West Point began some geological teaching in about 1823. Publications by R.B. Smith in 1849, F.W. Hutton in 1862, and articles within a massive three-volume, Aide-Mmoire to the Military Sciences, by G.G. Lewis and others (first edition, 184652; second edition, 185362), published primarily for use by the Royal and the East India Company's Engineers, demonstrated quite specifically the importance of geology to the British military profession; however, military teaching of the subject declined sharply towards the end of the nineteenth century as perception of its practical value waned.
Figure 2. Professors of the Staff College, Camberley, in 1874. The geologist, T. Rupert Jones, is seated at the left rear. Reproduced with permission from Rose EPF, Husler H, and Willig D (2000). In: Rose EPF and Nathanail CP (eds.) Comparison of British and German applications of geology in world war, pp. 107140. Geology and Warfare: Examples of the Influence of Terrain and Geologists on Military Operations. London: Geological Society. Courtesy of the Joint Services Command and Staff College, Swindon, and the Geological Society, London.
Professional geologists were not used operationally as such until World War I, and then primarily as a response to the near-static battlefield conditions on the Western Front in Belgium and northern France. The German army deployed the engineering geologist W. Kranz in 1914, and H. Philipp plus many more geologists in 1915. A German military geological service was constituted in 1916, and made use of some 250 geologists by the end of the war in 1918 when it was disbanded. In the British army, far fewer geologists were used: Lieutenant (later Captain) W.B.R. King from 1915 to guide the development of water supplies; Major (later Lieutenant-Colonel) T.W. Edgeworth David from 1916 to guide siting of mine tunnels and dugouts; and a few others, primarily in the Tunnelling Companies of the Engineer Corps. All returned to their former civilian life at the end of the war: King to the British Geological Survey and then the academic staff of the University of Cambridge and later London, and David as Professor of Geology and Physical Geography to the University of Sydney in Australia. Led by Lieutenant-Colonel A.H. Brooks, nine (of potentially 18) geologists were assigned for service with the American Expeditionary Force in 1918, primarily from the United States Geological Survey, but the war ended before their full operational deployment.
The German army was quick to reinstate a military geological organization prior to World War II, from 1937. Military geology textbooks were generated by J. Wilser in 1921, E. Wasmund in 1937, K. von Blow and others, C. Mordziol, and W. Kranz, all in 1938. Military geologists were deployed with troops invading Poland in 1939, and France and the Low Countries in 1940. By late 1941, there were 32 teams of military geologists providing support to the German army throughout its area of occupation, the number of terms increasing to 40 by November 1943. Making use eventually of some 400 geologists, largely recruited or conscripted from university staffs, the German army thus developed the largest organization ever to provide military applications of geoscience in wartime. Additionally, smaller but still significant numbers of geologists served with the German air force, navy, Waffen-SS, or the paramilitary construction agency Organisation Todt.
In contrast, the British army again used very few geologists as such: Major (later Lieutenant-Colonel) W.B.R. King from 1939 to 1943 and his Cambridge protg Captain (later Major) F.W. Shotton from 1940 to 1945; this group was increased to perhaps a dozen in various roles by 1944, largely in preparation for the Allied invasion of Normandy, and many of them formerly academic geologists from British universities. America entered the war in December 1941, its forces soon supported by a Military Geology Unit at the United States Geological Survey. With a wartime roster of 88 geologists, 11 soil scientists, 15 other specialists, and 43 support staff, it produced 313 studies including 140 major terrain folios, 42 other major reports, and 131 minor investigations in total containing about 5000 maps, 4000 photographs and figures, 2500 large tables, and 140 terrain diagrams.
After the war, the British army maintained continuity in geological expertise through a small group of officers in the reserve army (the Territorial Army or, from 1953 to 1967, the Army Emergency Reserve) for peace-time engineering projects and operational planning, and active service in times of crisis, whether occasioned by military conflict or humanitarian need. German capability in military geology was abolished at the end of the war, but later re-established within the Bundeswehr. By the end of the Cold War in 1990, the German army employed more than 20 full-time geologists as civilians but with commitment as reserve army officers. In the USA, geological support for the armed forces continued to be based on the United States Geological Survey which, between 1945 and 1972, briefly used about 150 scientists to compile terrain intelligence on a global scale.
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Most minerals have a range of atomic structures with variable amounts of impurities and substitutions. Not so with Dolomite which has an unusual constant ratio between its content of MgCo3 and of CaCo3.
Dolomites, Italian Alpi Dolomitiche, mountain group lying in the eastern section of the northern Italian Alps, bounded by the valleys of the Isarco (northwest), the Pusteria (north), the Piave (east and southeast), the Brenta (southwest), and the Adige (west). The range comprises a number of impressive peaks, 18 of which rise to more than 10,000 feet (3,050 metres). The highest point is the Marmolada (10,964 feet [3,342 metres]), the southern face of which consists of a precipice 2,000 feet (610 metres) high. The range and its characteristic rock take their name from the 18th-century French geologist Dieudonn Dolomieu, who made the first scientific study of the region and its geology. Geologically, the mountains are formed of light-coloured dolomitic limestone, which erosion has carved into grotesque shapes. The resulting landforms include jagged, saw-edged ridges, rocky pinnacles, screes (pebble deposits) of limestone debris, deep gorges, and numerous steep rock faces at relatively low levels. Glaciated features occur at higher levels; 41 glaciers lie in the region. Many of the lower and more gentle scree slopes were once forested; only patches of woodland remain, however, interspersed with grassy meadows.
The main valleys provide relatively easy access to most parts of the Dolomites. The main northsouth road uses the Campolongo Pass (6,152 feet [1,875 metres]). The eastwest roads cross the well-known passes of Pordoi (7,346 feet [2,239 metres]), Falzarego (6,906 feet [2,105 metres]), Tre Croci (5,935 feet [1,809 metres]), Sella (7,404 feet [2,257 metres]), and Gardena (6,959 feet [2,121 metres]). The main centre of this tourist and mountain-climbing region is Cortina dAmpezzo. Other resorts are Auronzo, San Martino di Castrozza, and Ortisei, with its narrow-gauge railway. On the western and southeastern margins, respectively, are located the larger towns of Bolzano and Belluno.
Most of the main peaks were first climbed in the 1860s and 70s by English mountaineers. Landslides after heavy rainstorms in the southern Dolomites twice caused the Vaiont Dam (on a tributary of the Piave River) to overspill and drown the village of Longarone, causing the loss of more than 2,500 lives in 1963 and the destruction of houses and communications in 1966. In 2009 the Dolomites were inscribed on UNESCOs World Heritage List.
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Boning Tang, Chuanqing Zhu, Nansheng Qiu, Yue Cui, Sasa Guo, Xin Luo, Baoshou Zhang, Kunyu Li, Wenzheng Li, Xiaodong Fu, "Analyzing and Estimating Thermal Conductivity of Sedimentary Rocks from Mineral Composition and Pore Property", Geofluids, vol. 2021, Article ID 6665027, 19 pages, 2021. https://doi.org/10.1155/2021/6665027
In this study, thermal conductivities of 128 rock samples located in the Xiongan New Area and Tarim Basin were measured using the optical scanning and transient plane source methods. The thermal conductivities of the Xiongan New Area samples range from 1.14 to 6.69W/(mK), in which the mean thermal conductivities of dolomite and sandstone are and , respectively. In the Tarim Basin, sandstone samples have thermal conductivities ranging from 1.21 to 3.56W/(mK) with a mean value of . The results can provide helpful reference data for studies of geothermics and petroleum geology. Calculation correction and water-saturated measurements were conducted to acquire in situ rock thermal conductivity, and good consistency was found between both. Compaction diagenesis enhances bulk thermal conductivity of sedimentary rocks, particularly sandstones, by decreasing the rock porosity and mineral particle size. Finally, correction factors with respect to mineral grains were proposed to correct the thermal resistance of intergrain contacts and degree of intactness of crystals, and an optimized formula was adopted to calculate the thermal conductivity of sedimentary rock based on rock structure and mineral constituents.
Thermal conductivity directly characterizes rocks heat conduction capacity and plays an important role in the present thermal conditions including surface heat flow, geothermal field distribution, deep thermal structure, and thermal simulation, and it is also a significant parameter in engineering projects such as the construction of high-level radioactive waste repositories, buried heat exchangers of the ground source heat pump, tunnels construction, and oil, gas, and geothermal energy exploration . Until recently, it still lacks completed measurement methods to get the in situ thermal conductivity, so laboratory-measured results have been used to make corrections. Several factors, such as mineral composition, structure, porosity, moisture content, temperature, and pressure influence the rock thermal conductivity [7, 9, 10]. Therefore, it is crucial to understand each factor during the investigation of thermal conductivity.
Rock samples used for measurements of thermal conductivity were from the Xiongan New Area and Tarim Basin. These two sampling areas have different geological conditions. The Xiongan New Area belongs to rift basin background and the Tarim Basin belongs to craton basin background. In previous work, our research group did a lot of researches on both areas , and corresponding supporting projects allowed authors to obtain core samples. By studying the features of samples measurements from two different regions, it can help to avoid the impact of regional factors from the same region and find common laws suitable for most situations.
Located in the northwest of Bohai Bay Basin in the middle east of the North China Craton, the Xiongan New Area featured by rich geothermal resources is one of the hot spots for geothermal energy study and utilization in China. It stands geographically in a triangle with Beijing and Tianjin cities, and was established as a state-level new area by Chinese government in 2017 which has important significance in politics and economy. Because of the shallowly burial depth, large storage, and complete reinjection rate of reservoirs, the exploration and exploitation of geothermal resources in Xiongan New Area has already made great progress and has been built as a demonstration area for geothermal system utilization in China. The Tarim Basin located in northwest China is the largest petroliferous basin in China. Experiencing multiperiod tectonic movements, the Tarim Basin stores large quantities of oil and gas. Abundant gas resources make it become the main source area of the west-east natural gas transmission project which is a significantly strategic project set up by Chinese government. The thermal conductivities obtained in both areas could be used as important references for studies of regional heat flow and present-day thermal regime, and furthermore, they are nonnegligible parameters in thermal history simulation and geothermal, oil, and gas resource utilization.
There were several researches aiming to develop thermal conductivity estimation methods based on different factors. Woodside and Messmer  studied the variations in the thermal conductivity of consolidated rocks based on porosity, pore medium, and pressure. They found that the two-phase geometric mean (matrix and pore) agrees well with the measured results. Horai and Baldridge  estimated thermal conductivities of 19 igneous rocks based on their mineral and chemical compositions. Pribnow and Umsonst  built a layered model for crystalline rocks and investigated the effects of structure and anisotropy on the thermal conductivity. Furthermore, Fjeldskaar et al.  developed a model to estimate thermal conductivity of sedimentary rock. He et al.  used multiple models to estimate soil thermal conductivities. Based on these studies, it can be concluded that the mineral composition and interior structure of rocks are the main factors influencing the estimated thermal conductivity results. When discussing the relation between the thermal conductivity of rock and mineral composition, most previous researches have focused mainly on the igneous and metamorphic rocks comprising intact crystals with simple mineral compositions rather than sedimentary rocks with more complicated compositions and structures.
In this study, 128 samples of diverse lithologies, mostly in sedimentary rocks, were measured using the optical scanning (OS) method and transient plane source (TPS) technology. By identifying the porosity and composition of rock and combining both the estimation methods, factors affecting the thermal conductivity were comprehensively studied.
There are several approaches for measuring rock thermal conductivity, such as the line heat source, divided bar, OS method, and TPS method [5, 1921], with the latter two methods being commonly used in the laboratory. The OS method is widely used in the geologic field owing to its rapid measurement and convenience [22, 23]. The TPS method is used to determine thermal properties of various materials, such as metals, minerals, and alloys, and is widely applied in engineering and production activities owing to its high precision and accuracy . This study adopted OS and TPS methods which have different working procedures and handling methods to measure the thermal conductivity. Two methods can make cross-check analysis, ensure data integrity, and determine if there is a difference between both. By comparing characteristics of both methods, it is helpful to extend the testing conditions and make the measurement of thermal conductivity more efficient and high-quality.
The OS method has reached maturity owing to the efforts of Professor Popov [4, 20], and this method has been applied in a thermal conductivity scanner (TCS) with an accuracy of 2% for thermal conductivities ranging 0.2045W/(mK). The rock sample surfaces plane variation must not exceed 0.5mm during measurement. The core element comprises three or four temperature sensors and a heat source. The sample remains immobile on the countertop during measurement. The test is conducted by moving the scanning unit, with the front temperature sensors recording the temperature before heating, middle heat source heating the sample, and latter temperature sensors recording the maximum temperature increase after heating, as shown in Figure 1.
The temperature sensors and heat source can heat and record multiple samples lined up end to end as long as samples maintain a consistent distance and move at the same speed throughout the measurement process. The relational expression of a samples maximum temperature increase () is as follows : where is the samples maximum temperature increase, is the source power, is the distance between the heat source and sensor, and is the samples thermal conductivity.
If the maximum temperatures of standard samples are known, rock thermal conductivity can then be acquired as follows : where and are the standard samples thermal conductivity and maximum temperature increase, respectively.
The OS method has been widely used for measuring thermal conductivity, thermal diffusivity, and volumetric heat capacity and has given high-quality measurements. Thousands of samples including diverse rocks types and minerals from various areas have been measured by optical scanning techniques [4, 2731]. Furthermore, this method has been used for researches of deep scientific boreholes in many parts of the world [27, 3234].
The OS method has several advantages, such as fast measurement, convenient movement, continuous operation, and nondestructive test process. And it also allows to measure thermal conductivity components and determine anisotropy coefficient easily. However, it still possesses some limitations. For example, measured results are easily influenced by environmental conditions, such as wind or heat, which might change the measured values to some extent. Additionally, to enable samples absorb enough heat, sample surfaces must be painted black, which can cause differences in impurities. In principle, the OS method is only supposed to measure a dry sample, but if applied to saturated test samples, evaporation from the saturated samples during heating may cause deviations if the water was not removed from the rock sample surface or the heating power was chosen too high.
The TPS method can simultaneously measure thermal conductivity, thermal diffusivity, and specific heat. It uses a hot disk thermal constant analyzer. The core component of the instrument is a thin disk-shaped temperature-dependent sensor comprising a thin sheet with a double-helix structure etched by the conductive nickel inside and a double-layer Kapton protective layer outside, which can heat the samples and simultaneously record temperatures, as shown in Figure 2.
The instrument uses a resistance thermometer, wherein temperature dynamic change in a sample is reflected by the variations in the resistance of the sensor. The sensor is fixed on the sample surface within a specimen holder during testing, and a fixed power current is applied to the sensor. An increase in the temperature changes the sensor resistance (), giving rise to voltage variations at both ends of the sensor. There are different changes in voltage due to varying heat losses in different samples. By recording voltage variations over a limited period, rock thermophysical property can be obtained, as shown in Figure 3. The electrical bridge in the sensor is shown as follows :
The relational expression between temperature variation and voltage can be expressed as follows: where is the sensor resistance at time , is the initial sensor resistance, is the initial current through the sensor, is the resistance of a resistor in series with the sensor, is the resistance of sensor leads, is the temperature coefficient of sensor resistance, is the mean voltage variation at time , and is the mean temperature increase at time . and play important roles in keeping sensor power stable during measurements.
If the sensors heat source is located in an infinite sample surface, the variation in the temperature increases in the sensor combined with samples can be expressed as follows [35, 36]: where and are the temperature increase in the Kapton protective layer and sample surface, respectively.
From the viewpoint of the sensors geometric characteristics, i.e., concentric ring sources, the mean space temperature increase can be defined as follows : where is the output power of the sensor, is the sample thermal conductivity, is the dimensionless specific time function representing sensor geometric characteristics (a time-dependent equation is independent of size), is the integral variable of , and are sum variables not more than the number of total rings in the double-helix structure, and is the zero-order modified Bessel function. is a function varying with , which is defined as follows : where is the specific time depending on sensor and sample conditions, is the outermost radius of the double-helix structure, and is the thermal diffusivity.
Only a smoothly flat surface is acceptable for measurement using the TPS method. Furthermore, the incircle diameter of a sample surface must be four times larger than the sensor radius and thickness must be no less than samples radius. The TPS instrument can be equipped with multiple sensors, each having different radii. The measurement accuracy is 2% for 0.0051800W/(mK) range.
Deviations in TPS method measurements occur mainly because the heat conduction equation is based on a sensor on an infinite surface and thus not controlling measurement time to keep heat away from samples edge. There is a thermal contact resistance between the sample and sensor, leading to an additional 200 points of increasing temperature; however, removing some of these initial points should eliminate the influence of this thermal contact resistance.
There are different sizes of sensors measuring samples with various sizes. And the hot disk has multiple modules to conduct analyses of the anisotropy and specific heat capacity and measurements of the thin film, slab, powder, and liquid materials. The TPS method possesses several advantages, such as high measurement accuracy and repeatability, simple specimen requirements, nondestructive measurement capabilities, and the ability to measure at high temperatures. It has been used in engineering domains to test various materials, such as metals, rock, ceramic, powder, film, plastic, and liquids [24, 36, 3840]. In recent years, TPS has entered the geological domain. However, it is still less used than the OS method. TPS method has been used in a few recent studies. Di Sipio et al.  discussed the necessary parameters for establishing a geothermal model using thermophysical data measured by the TPS method. Aurangzeb and Maqsood  measured the thermal conductivity of olivenite at different temperatures to study their relation with each other. Li and Liang  analyzed factors affecting thermal conductivities of methane hydrate-bearing sediments using the TPS method.
Herein, 128 samples were collected from the Xiongan New Area and Tarim Basin located in North and Northwest China, respectively. Sampling positions are shown in Figures 4 and 5. In the Xiongan New Area, we collected 96 samples from 34 boreholes from depths of 7006000m, including Archaean gneiss and leptynite, middle Proterozoic dolomite rock, and Paleogene and Neogene sandstone and mudstone. Samples from the Tarim Basin comprised 32 sandstones that were obtained from 13 boreholes over a wide area with sampling depths of 11206320m, covering the main sedimentary layers from the Ordovician to Cretaceous.
Thermal conductivities of the rock samples were measured using the OS and TPS methods at room temperature (21C) and normal atmosphere. All samples were quite complete. Although the thermal conductivity is an anisotropic parameter, for many sedimentary rocks, the testing difference caused by anisotropy is small and close to the measurement accuracy [20, 44], so the anisotropy of sedimentary rocks is not discussed here. The anisotropy of gneiss is obvious and nonnegligible , but this article aims at thermal characteristics of sedimentary rocks, the thermal conductivity of gneiss is mainly used for increasing the thermal data.
Descriptive results of thermal conductivity are shown in Table 1 and Figure 6. The thermal conductivities of Xiongan New Area samples range from 1.14 to 6.69W/(mK). Dolomite rock values range from 1.44 to 6.69W/(mK) but are mainly distributed in 4.007.00W/(mK) with a mean value of (). For sandstones, thermal conductivity results range from 1.00 to 3.00W/(mK) with a mean value of (). Leptynites and gneisses have similar values at 1.603.20W/(mK) with mean values of () and (), respectively. There are only few mudstone and breccia samples, with the former and latter having the lowest and highest mean values of () and (), respectively. The sandstones from the Tarim Basin range 1.004.00W/(mK) evenly with a mean value of ().
Figures 7 and 8 show the varying thermal conductivities of the main rocks with depth. For rocks in the Xiongan New Area, the thermal conductivity of sandstones increases obviously with depth. However, this tendency is not present in dolomite rock, which shows hardly any variation with depth. There are no obvious relations between depth and thermal conductivity for other rocks in the Xiongan New Area since the number of samples is less. In the Tarim Basin, sandstone thermal conductivity slightly increases with depth.
Both OS and TPS methods have own unique characteristics for measuring thermal conductivity (Table 2) and have successfully achieved measurements. Upon comparing the two methods, the TPS method is able to measure more materials and smaller size than the OS method, and the TPS can directly obtain the specific heat capacity by adjusting the special module. The OS method should be conducted in an enclosed space because it is more easily affected by the environment such as the wind, and the TPS method could remove environmental impacts by being covered with a specific lid. The testing efficiency of the OS is faster than that of TPS because the OS can test many samples simultaneously. When testing the same sample several times, variations in the results obtained by the TPS method could be limited to two or three decimal places, whereas those obtained by the OS method changed one or two decimal places under the same conditions.
Five different OS standards whose thermal conductivity had been already known were measured by the TPS. All the relative deviations of results are below 1% (Table 3); therefore, it is available to measure the thermal conductivity by the TPS.
The TPS and OS methods measured thermal conductivities of dry samples. And both method values are consistent with each other, as shown in Figure 9. Their overall deviations are between 11.1% and 16.54%, and mean deviations range from 3.5% to 5.97% with an integration of 2.59%. This demonstrates that the results for both methods should be valid because their relative deviations are less than 4%, which is the sum of the test accuracy of both methods at 2%. Hence, the OS and TPS methods could be combined in different measurement conditions based on their various features.
Factors affecting the rock thermal conductivity can be divided into two classes: internal factors, which depend on aspects of the rock itself, such as mineral composition, structure, intergrain contact, and porosity; and external factors, which are determined by outer conditions, such as temperature, pressure, pore medium, and medium saturation. Both dry and in situ samples have the same internal factors, but external conditions change significantly. To obtain in situ thermal conductivity from dry samples in the laboratory, external factors must thus be corrected.
Thermal conductivity increases with an increase in the pressure and decreases with an increase in the temperature, suggesting that the two effects might offset each other under certain conditions [29, 45, 46]. Hence, for the ease of convenience, the correction for temperature and pressure was not used here, and only the effect of pores on thermal conductivity was discussed. Pore characteristics arise from three aspects, namely, the porosity, pore medium, and medium saturation.
All the samples in situ were under the water table and had pores filled with water. The measured thermal conductivity value was the integration of the thermal conductivities of the rock matrix and air filled in pores. Therefore, it is necessary to make water saturation corrections to attain in situ thermal conductivity values reflecting real heat conduction processes.
Variations in the porosities of the main rocks with depth are shown in Figure 10. All porosities were tested by the helium porosimeter named CAT113 manufactured by the American Core Lab Company. The dolomite rock porosity varies irregularly with depth having a low range of 0%6%. Sandstones in the Xiongan New Area and Tarim Basin both have obviously higher porosity than dolomite with porosities of 2%25% and 1%6%, respectively. It is noticeable that the porosity decreases with an increase in the depth in the Xiongan New Area, which agrees with compaction diagenesis. The compaction effect enhances gradually as depth increases, making the rock structure more compact and decreasing porosity. However, this trend does not exist in the Tarim Basin because its proportion, which is 40 millionkm2, significantly larger than that of the Xiongan New Area (2000km2). In a small scale, samples present a continuous variation with depth due to being in a similar tectonic unit. On the contrary, samples in a wide scale belong to different tectonic units and are far apart. Therefore, no continuous change among porosities exists.
Two methods were adopted to perform water saturation corrections for thermal conductivity. The first method was to calculate the geometric mean as follows : where is the rock bulk thermal conductivity, is the matrix thermal conductivity, is the thermal conductivity of medium (such as water or air) in the pores, and is the porosity. According to Equation (7), the matrix thermal conductivity was obtained based on thermal conductivities of dry sample and air and porosity firstly; then, the water-saturated thermal conductivity was acquired by combining the water thermal conductivity.
The second method was to measure saturated samples simulating the true underground situation using the TPS method by saturating them with water under vacuum for more than 72h until the pores were full of water.
The correlation of results between these two methods is shown in Figure 11. Their relative deviations are between 10.53% and 10.16%, with a mean range of 3.24%4.32%. No obvious difference can be noted, so either method could be chosen to make corrections. For convenient analysis, all corrected thermal conductivities in the rest of this paper adopted the geometric mean.
Analysis and comparison of thermal conductivities before and after correction (Figures 12(a)12(c)) find that the correction of the dolomite rock changes little, with a mean of 5.20%, which is closely related to its low porosity at 1.64%. The accuracy of the TPS method is close to the correction difference, so water corrections have no significance in the dolomite rock. The mean correction difference of sandstones in the Xiongan New Area and Tarim Basin with a mean porosity of 7.8% and 7.31% is 29.3% and 27.2%, respectively. The correction differences and porosities of both areas are higher than those of the dolomite rock. As we know, the higher the porosity is, the higher the correction difference is . The sandstones thermal conductivity shows a rising tendency with depth before correction, but after correction, this trend weakens, and because of the separated distribution in the Tarim Basin, the positive correlation is also weaker than in the Xiongan New Area.
Variation in thermal conductivity with porosity (Figures 13(a)13(c)) shows that the dolomite rock has no evident regular changes with porosity, which is lower than 6%. Thermal conductivity of sandstone in the Xiongan New Area and Tarim Basin decreases as porosity increases without corrections, whereas after correction, this tendency weakens. This indicates that when pores are filled by media with low thermal conductivities, the rock bulk conductivity could show a negative correlation with porosity. However, once medium conductivity rises, the change in relation will be inconspicuous.
Mineral compositions and structure are key factors affecting rock thermal conductivity. Different lithologies comprising varying mineral compositions have different thermal conductivities. Additionally, the thermal conductivity of the same rock type depends mainly on the mineral type and content.
Previous researches have studied the conductivity of various minerals, debrises, and interstitial materials in detail (Table 4) [17, 45, 4850]. Among the main rock-forming minerals, thermal conductivity is obviously higher for quartz than others, i.e., 7.69W/(mK). On the contrary, clay minerals are apparently lower than others, at generally less than 1.0W/(mK). Accounting for mineral anisotropy, the values listed in Table 4 are integral results synthesizing every direction.
To clarify the relation between thermal conductivity and mineral type, 13 dolomite rocks and 8 sandstones from the Xiongan New Area were made into slices to determine their mineral compositions and grain sizes under the microscope (Tables 58). Table 6 shows that dolomite rock is characterized by dolomites, clays, and silicas with a small amount of pyrites and sands. The dolomite content ranges 45%97% with a mean value of , the clay content is relatively less at 0%15% with a mean value of , and silica exists in some of these samples, with content varying significantly from 0% to 55%. The dolomite rock has a mean grain size of 0.0241.606mm, which was calculated using the arithmetic mean according to particle diameter distributions.
The sandstone is characterized by quartz, feldspar (alkali feldspar and plagioclase), and debris (acidic magma, andesite-basic magma, and carbonate sediments) in terrigenous clastic, and calcite, dolomite, and clay in the interstitial material (Table 7). Among these, quartz has the highest ratio at 24%55.24% with a mean value of . Alkali feldspar is the second most common at 14.45%36.80% with a mean value of 22.96%. In the debris, acidic magma accounts for the highest proportion at 7.26%32.04% with a mean value of 15.91%. The highest content in the interstitial material is the clay ranging 1%20% with a mean of 9.50%. The remaining compositions are less than 3%. The sandstones are mainly fine, medium, and silt sandstones with grain size of 0.0750.250mm and mean value of .
Correlation analysis was employed to discuss the relation between the rock thermal conductivity and mineral composition. Keeping aside the low porosity of the dolomite rock, the measured dry thermal conductivities are directly applied to compare with mineral compositions (Table 9). Thermal conductivity negatively correlates with clay mineral content with a correction coefficient of 0.702 and reaches a statistically significant level (), indicating that dolomite rocks thermal conductivity decreases significantly with increasing clay content (Figure 14). The thermal conductivity correction coefficient of dolomite is and there is no correlation between both. The coefficient for silica is 0.402, but obviously exceeds the level of statistical significance at . There is no obvious correlation between dolomite rock conductivity and grain size with a correlation coefficient of 0.210.
Owing to the effect of the high porosity of sandstones on rock bulk thermal conductivity, the matrix thermal conductivity computed using the geometric mean was used to analyze the relations between thermal conductivity and mineral composition (Table 10). Results show that sandstones have thermal conductivity that positively correlates with quartz content with a correlation coefficient of 0.759 and a level of statistical significance of (Figure 15). The coefficient of acidic sediments and thermal conductivity is 0.632, but the statistical significance level is . Other compositions are not evidently correlated with thermal conductivity because the absolute value of their coefficient is less than 0.4 and significant level is more than 0.3. There is also a negative relation between thermal conductivity and grain size with a coefficient of 0.764 (Figure 15).
Pearson correlation analysis studies the relation between two variables; however, sometimes the correlation coefficient cannot reflect the real relation due to the influence of a third variable. Therefore, it is necessary to conduct a partial correlation analysis. Previously, it is found that there is a strong negative relation between dolomites and silicas with a coefficient of 0.933 (Table 9), but the two have nothing to do with thermal conductivity of dolomite rock. Considering the dolomite and silica as a mutual controlling variable, the correlation between them and thermal conductivity is again analyzed employing partial correlation analysis (Table 11). The partial correlation coefficient between thermal conductivity and dolomite is 0.755 with a significance level of 0.003 when silica is taken as a controlling variable, indicating an obviously positive correlation between them. When dolomite is the controlling variable, thermal conductivity increases with silica content with a partial coefficient of 0.812, which is an identical trend as the dolomite.
Above all, the dolomite rock has increasing variance with higher dolomite and silica contents, decreasing variance with higher clay content, and thermal conductivity variance has no association with grain size. For the sandstone, thermal conductivity increases with an increase in the quartz content and decreases with an increase in the grain size. Based on the mineral thermal conductivity shown in Table 4, dolomite rock and sandstone thermal conductivities are obviously correlated with dolomite (5.51W/(mK)), silica (4.53W/(mK)), and quartz (7.69W/(mK)), which have thermal conductivities significantly higher than other common rock-forming minerals, ranging 24W/(mK) . The reason why there is no direct relation between thermal conductivity variance and dolomite and silica contents is that they have close thermal conductivity values, thus playing the same role in heat conduction, and can be considered the same unit. The reason for the decreasing tendency seen in dolomite rock with clay might be spilled into two ways. On the one hand, shaly clay with low thermal conductivity (generally <1W/(mK))  correlates negatively with thermal conductivity in dolomite rock. On the other hand, the dolomite rock mainly comprises dolomite, silica, and clay, and as dolomite and silica contents fall, clay content naturally rises, making it difficult to find which mineral is a leading factor.
Since quartz has a significantly higher thermal conductivity than the other minerals, all types of minerals in the sandstone could be categorized into two groups, namely, quartz and others. The mean ratio of quartz is 35% and that of dolomite and silica in the dolomite rock is at 95%. The proportion of the mineral with high thermal conductivity is lower in the sandstone than in the dolomite rock, which may be one of the reasons for sandstones mean thermal conductivity being less than that of dolomite in general.
There is a decreasing trend in sandstone thermal conductivity with grain size. This is probably due to grain size in connection with compaction diagenesis. Higher pressure on the rock yields a greater degree of compaction and thus a smaller grain size . As the structure within rock becomes more compact, thermal conductivity increases. The dolomite rock has hardly any grain size variance because it is formed by secondary replacement with the same dolomite particles in contact with each other within the rock, rather than the different contact relations seen among the different particles in sandstone.
In comparison with the thermal conductivities of the common minerals, a mineral with a significantly higher or lower thermal conductivity may have greater effect on rock bulk thermal conductivity. However, estimates of rock bulk thermal conductivity should not ignore the contribution of other minerals.
Based on the contribution of each mineral, the geometric mean (estimation method 1) (Equation (8)) was the method most commonly used to estimate matrix thermal conductivity based on their thermal conductivities and proportions : where is the matrix thermal conductivity, is the thermal conductivity of each mineral, and is the volume fraction of each mineral. The theory is the same for both estimation method 1 and water correction formula, where the matrix and pore medium are considered separately and compared with the numerous mineral constituents.
Additionally, Horai and Baldridge  employed another formula to estimate the thermal conductivity of 19 igneous rocks and found that its results were better than those from the geometric mean. The formula (estimation method 2) (Equations (915)) is as follows: where and are the maximum and minimum thermal conductivities among all minerals, respectively, and and are the upper and lower limits of the matrix thermal conductivity, respectively.
Both methods mainly focused on estimating thermal conductivity in igneous rocks, and their use with sedimentary rocks has been rare in previous works. Hence, herein, the two methods were used to estimate matrix thermal conductivity of sandstones. The parameter values used in this estimation process are shown in Table 4.
Our results demonstrate that estimated values are abnormally higher than the measured values (Figure 16) and significant differences exist. The reason why previous studies obtained good results is that they focused on igneous or metamorphic rocks [5, 15]. The values listed in Table 4 were measured on single crystals or monomineralic aggregates, and igneous rock, whose crystals are in a basically intact condition, differs from sandstone with fragmented crystals. Thermal conductivity of rock integrated with intact crystals must disagree with that from fragmented crystals even with identical compositions. Additionally, igneous rocks have relatively simpler compositions and fewer impurities than sandstones. Therefore, it is simpler to attain good results from estimations in igneous rocks. However, sandstone has very complicated compositions with different minerals touching each other and raising thermal contact resistance, affecting thermal conductivity to some extent . Based on these points, this study puts forward a matched simple correction factor to correct for the degree of intactness and contact relations among particles. The correction factor is multiplied by each compositions thermal conductivity to reduce the original values in Table 4, bringing them closer to the real situation in sandstone.
According to the ability of different compositions to resist deformation and damage, the more undamaged they can be, the higher their correction factors can be. In the meantime, in consideration of the thermal contact resistance could reduce the thermal conductivity of the bulk rock to a certain extent, this affection was considered by lowering suitably correction factors. Therefore, correction factors were adjusted continuously to make estimated values approach the measured values. Based on experimental results from multiple parameter adjustments, the following empirical correction factors were determined: 0.85 for quartz due to its strong weathering resistance of weathering, 0.8 for feldspar due to its correspondingly weaker resistance, and 0.7 for debrises and interstitial materials because debris is a broken product from the mother rock that was influenced easily by weathering and alteration and interstitial material is marked by clay with very low thermal conductivity. Results after adopting these factors are shown in Figure 17. The mean relative deviations of estimation methods 1 and 2 range 8.22%16.45% and 6.81%14.73% with an integrated mean value of 1.03% and 6.65%, respectively.
Estimation method 1 results are lower than those of method 2, but it is not wise to affirm which one is better because only a few samples were collected herein. Currently, both methods could attain good results, and the advantages and disadvantages among them may be discussed in the future after collecting more data. Additionally, the rock-forming minerals were formed in different environments, so their thermal conductivities would not be completely in accordance with the values listed in Table 4.
Employing these two estimation methods can almost obtain matrix thermal conductivity, verifying that each mineral has an essential contribution on bulk thermal conductivity. In comparison with igneous and metamorphic rocks, sandstone has complex compositions, broken crystals, and intricate contact relations, resulting in a difference in heat conduction.
(1)The cross-check revealed that the results of TPS and OS methods were in correspondence with each other and had high precision and efficiency. The mean thermal conductivities of dolomites and sandstones in Xiongan New Area and sandstones in Tarim Basin were , , and , respectively; these measured values can contribute to basin-related study. Since both the methods have their own unique characteristics, combining them would be suitable for measurements under various conditions(2)Water saturation correction demonstrated that there was little difference between the geometric mean and actual saturated measured results. There is no need to perform a water saturation correction for low porosity dolomite rock. But corrections must be used for high porosity sandstone. The thermal conductivity of sandstone increases as the depth increases or porosity decreases, and the tendency weakens after correction, while these relationships do not exist in dolomite(3)Correlation analysis showed that the dolomite rocks thermal conductivity was negatively correlated with shaly clay and positively correlated with dolomite and silica, and the sandstone positively correlated to the quartz content. When the mineral composition meets the requirement of significantly above (such as dolomite and quartz) or below (such as clay) the general range of thermal conductivity of 24W/(mK) and containing certain amount, there would be a significant correlation between matrix thermal conductivity and that of the constituents. Compaction diagenesis can make the grain size and porosity of sandstones smaller, and further make their thermal conductivities greater(4)Based on the contribution of mineral compositions on bulk thermal conductivity, considering the degree of intactness of crystals and complex relation of particle contact, empirical correction factors for quartz, feldspar, and debrises and interstitial materials ranked from high to low were developed to estimate the sandstones thermal conductivity as results from experimental parameter adjustments
This study was supported by the National Science and Technology Major Project of China (Grant No. 2017ZX05008004), the National Natural Science Foundation of China (Grant No. 41772248), and the National Key R&D Program of China (Grant No. 2018YFC0604302).
Copyright 2021 Boning Tang et al. This is an open access article distributed under the Creative Commons Attribution License, which permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly cited.
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