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See K. B. Krauskopf, Introduction to Geochemistry (1967); G. Faure, Principles and Applications of Geochemistry (1991).
the science of the chemical composition of the earth, of the laws governing the abundance and distribution of chemical elements, and of the methods of the association and migration of atoms in the course of natural processes. Geochemistry is part of cosmochemistry. The units of comparison in geochemistry are atoms and ions.
One of the most important tasks of geochemistry is to study the earth’s chemical evolution on the basis of the abundance of chemical elements and to attempt to give a chemical explanation for the origin and history of the earth and its differentiation into shells (geospheres). Geochemistry devotes maximum attention to problems of the abundance and distribution of chemical elements.
Abundance of chemical elements. The abundance of the various chemical elements depends on the synthesis of their nuclei by various thermonuclear reactions in the interiors of stars. The stage of evolution of the star (its temperature) determines the nature of this synthesis.
According to the prevailing cosmogonic hypotheses, on the formation of the sun from a contracting rotating plasma during the final stage of contraction, there separated from the center of contraction a considerable mass of hot plasma, which formed around it a “protoplanetary” cloud in the form of a disk. The cloud cooled rapidly and a spontaneous condensation of matter developed in it. As a result of multistage reactions (the condensation growth of nuclei and their coagulation, accretion, and agglomeration processes), the gaseous cloud transformed into a gas-dust cloud. Simultaneously, the cloud lost gases to cosmic space. Because of rotational instability, the cold gas-dust cloud broke up into a series of condensations—the protoplanets—which underwent adiabatic contraction. As a result of this process, the cold matter of the protoplanetary cloud formed earthlike planets and an asteroid belt with asteroids and meteorites. Finally, at the edge of the protoplanetary cloud, the condensation of the blown-off gases (H, He, NH3, CH4, and so forth) took place at very low absolute temperatures, resulting in the formation of the large planets—Jupiter, Saturn, Neptune, and Uranus.
It is not possible to determine directly the general composition of a planet. However, the availability of astronomical (spectral) data on the composition of the sun and data on the chemical composition of stony meteorites (of the most widely distributed ones, the chondrites) makes it possible to determine the abundance of chemical elements on the earth and other planets. It is evident from Table 1 that the abundance of elements in the sun and meteorites are the same. The most abundant elements (isotopes) have nuclei with an even number of protons and neutrons—126C, 168O, 2412Mg, 2814Si, 3216S, and many others. Elements with an even-odd number of protons or neutrons occupy a middle position. Elements having an odd number of protons and neutrons have a very low abundance, for example, 73Li, 115B, and 31T. Elements having even atomic numbers are more abundant than the neighboring elements with odd atomic numbers (see Figure 1). The light elements Li, Be, and B show a deficiency, because they “burn up” in reactions with protons. The nuclei of the elements at the end of the Mendeleev system have a large excess of neutrons and are therefore unstable. These elements undergo radioactive decomposition (U, Th, Ra, and other elements) and spontaneous fission (U, Th, and some actinides).
From data on the composition of the earth’s shells, it follows that the earth has a meteoritic composition. Meteorites are divided into stony (chondrites and more rarely achon-drites), iron (of Fe-Ni alloy) and mixed. Chondrites have lost all volatile substances other than those that have firmly combined with the solid substances of the meteorites—H2O, FeS, C, NH3, and others. Thus in abundance of elements, the solids correspond to those of the sun: Mg, Si, Fe, and O occupy first place (on the basis of the number of atoms, Si/Mg = 1), followed by S, Al, Ca, and other elements. The silicate phase of chondrites consists primarily of metasili-cates and orthosilicates—the pyroxenes (MgSiO3) and olivines [(Mg,Fe)2SiO4]—that is, it is a ternary system of MgO, SiO2, and FeO. Stony meteorites are multiphase systems; in addition to the main phases—silicate and metallic (Fe-Ni alloy)—they have sulfide, chromite, carbide, and phosphide phases. The ratio of the silicate phase to metallic phase varies with the meteorite. On the basis of an analogy with meteorites, many scientists believe that the earth-type planets also have a silicate phase and a metallic core, the ratio between these phases varying with each planet. According to this hypothesis, the earth has about 31 percent of the metallic phase, or about 40 percent Fe (including oxides).
Distribution of chemical elements. The earth, like other planets of its type, and the moon have a shell structure, consisting of a series of geospheres: the core, mantle, lithosphere, hydrosphere, and atmosphere. The earth’s solid shells, the constituent rocks, the paragenetic mineral associations, and so forth as a rule are multicomponent silicate systems. The processes during which they form proceed at finite speeds and are irreversible. Nonequilibrium systems are encountered in geochemistry, which are characterized by mass, volume, entropy, pressure, temperature, and chemical potentials. To make use of thermodynamics in geochemistry it is necessary to know the behavior of the specific phases, components, and systems under geological conditions, especially
|Table 1. Abundance of chemical elements in the sun and stony meteorites (chondrites)|
|* ΣMg is the number of atoms of a given element per 106 atoms of Mg|
|1 H...............||10.64||4.4 × 1010||—||—|
|3 Li...............||<-0.46||<3.4 × 10-1||1.54||3.5 × 101|
|4 Be...............||0.98||9.55||-0.14||7.19 × 10-1|
|5 B...............||2.24||1.7 × 102||1.18||1.5 × 101|
|6 C...............||7.15||1.4 × 107||4.30||2.02 × 104|
|7 N...............||6.70||5.0 × 108||2.54||3.47 × 102|
|8 O...............||7.47||3.0 × 107||6.55||3.5 × 108|
|9 F...............||—||—||3.01||1.02 × 103|
|11 Na...............||4.94||8.7 × 104||4.69||4.93 × 4|
|12 Mg...............||6.00||1.0 × 106||6.00||1.00 × 106|
|13 Al...............||4.84||6.9 × 105||4.89||7.61 × 105|
|14 Si...............||6.34||2.2 × 106||6.01||1.04 × 106|
|5.88||7.6 × 105|
|15 P...............||3.98||9.6 × 103||3.72||5.23 × 103|
|16 S...............||5.94||8.7 × 105||5.00||1.01 × 105|
|17 Cl...............||—||—||2.50||3.20 × 102|
|19 K...............||3.34||2.2 × 103||3.55||3.52 × 103|
|20 Ca...............||4.68||4.8 × 104||4.75||5.66 × 104|
|21 Sc...............||1.49||3.1 × 101||1.46||2.88 × 101|
|22 Ti...............||3.45||2.8 × 103||3.34||2.20 × 103|
|3.27||1.9 × 103|
|23 V...............||2.81||6.5 × 102||2.35||2.23 × 102|
|24 Cr...............||3.76||5.8 × 103||—||—|
|3.65||4.5 × 103||3.97||9.35 × 103|
|25 Mn...............||3.49||3.1 × 103||3.87||7.37 × 103|
|26 Fe...............||5.44||2.8 × 105||5.84||6.96 × 105|
|27 Co...............||3.34||2.2 × 103||3.28||1.92 × 103|
|28 Ni...............||4.41||2.6 × 104||4.60||4.00 × 104|
|29 Cu...............||3.09||1.2 × 103||2.49||3.06 × 103|
|30 Zn...............||2.16||1.4 × 102||2.09||1.24 × 102|
|31 Ga...............||1.36||2.3 × 101||1.06||1.16 × 101|
|32 Ge...............||1.13||1.3 × 101||1.35||2.23 × 101|
|34 Sc...............||—||—||1.31||2.05 × 101|
|35 Br...............||—||—||1.78||6.08 × 10-1|
|37 Rb...............||1.12||1.3 × 101||0.75||5.69|
|38 Sr...............||1.66||4.6 × 101||1.27||1.85 × 101|
|39 Y...............||1.84||6.9 × 101||0.56||3.64|
|40 Zr...............||1.29||2.0 × 101||1.09||1.24 × 101|
|41 Nb...............||0.94||8.7||-0.28||5.23 × 10-1|
|45 Rh...............||0.01||1.0||-0.51||3.15 × 10-1|
|47 Ag...............||-0.61||2.4 × 10-1||-0.82||1.50 × 10-1|
|48 Cd...............||0.18||1.5||-1.14||7.21 × 10-2|
|49 In...............||0.09||1.2||-2.85||1.41 × 10-3|
|-0.94||1.1 × 10-1||0.88||1.33 × 10-1|
|53 I...............||—||—||-1.71||5.11 × 10-1|
|55 Cs...............||—||—||-0.91||1.22 × 10-1|
|57 La...............||0.67||4.7||-0.46||3.50 × 10-1|
|58 Ce...............||0.42||2.6||-0.24||5.78 × 10-1|
|59 Pr...............||0.09||1.2||-0.94||1.15 × 10-1|
|60 Nd...............||0.57||3.7||-0.17||6.74 × 10-1|
|62 Sm...............||0.26||1.8||-0.67||2.16 × 10-1|
|63 Eu...............||-0.40||4.0 × 10-1||-1.07||8.53 × 10-2|
|64 Gd...............||-0.23||5.9 × 10-1||-0.39||4.12 × 10-1|
|65 Tb...............||—||—||-1.29||5.10 × 10-2|
|66 Dy...............||-0.36||4.4 × 10-1||-0.46||3.49 × 10-1|
|67 Ho...............||—||—||-1.16||6.88 × 10-2|
|68 Er...............||—||—||-0.71||1.94 × 10-1|
|69 Tm...............||—||—||-0.42||3.84 × 10-2|
|70 Yb...............||0.17||1.5||-0.73||1.87 × 10-1|
|71 Lu...............||1.49||—||-1.49||3.24 × 10-2|
|72 Hf...............||—||—||-0.74||1.82 × 10-1|
|73 Ta...............||—||—||-0.75||1.79 × 10-1|
|74 W...............||—||—||-0.58||2.64 × 10-1|
|75 Re...............||—||—||-0.76||1.74 × 10-1|
|76 Os...............||—||—||-0.22||5.96 × 10-1|
|77 Ir...............||—||—||-0.38||4.22 × 10-1|
|79 Au...............||—||—||-0.79||1.65 × 10-1|
|80 Hg...............||—||—||-0.09||8.08 × 10-1|
|81 TI...............||—||—||-2.63||2.38 × 10-1|
|82 Pb...............||0.27||1.9||-0.81||1.56 × 10-1|
|83 Bi...............||—||—||-1.63||2.33 × 10-2|
|90 Th...............||—||—||-1.55||2.79 × 10-2|
|92 U...............||—||—||-1.99||1.02 × 10-2|
within the large range of temperature and pressure. Thus for example, the general concept of the direction of a geochemical process is given by the Le Châtelier-Braun principle, which states that in any system acted on by external forces, a change in any external factor causes a change that tends to offset the action of that factor. According to the law of mass action, a change in the activity of one of the components of a system shifts the equilibrium. For example, in the reaction
CaCl2 + Na2SO4 ⇄ CaSO4 + 2NaCl
the equilibrium shifts to the right as anhydrite precipitates out of the solution. In the reaction
CaCO3 + SiO2 ⇄ CaSiO3 + CO2
which gets under way above 350° C, the equilibrium shifts to the right, because at the same time the mineral wollastonite (CaSi03) precipitates out, carbon dioxide forms, escaping from the system. As the temperature increases in reactions involving a gaseous phase, the equilibrium shifts toward the side of smaller volume of the gaseous components. For example, in the reaction
SiO2 + 4HF ⇄ SiF4 + 2H2O
the equilibrium moves to the right. High (gas and lithostatic) pressure changes the direction and nature of the magma crystallization.
The equilibrium conditions are also governed by Gibbs’ phase rule, which states that the number of thermodynamic degrees of freedom of a system is f = k - n + 2, where n is the number of phases of a system and k is the number of components. Since the number of degrees of freedom in a closed system is f ≤ 2 (pressure and temperature), the number of phases is n ≥ k. This mineralogical phase rule, which V. M. Goldschmidt first used in geochemistry, has been verified for a variety of rocks.
The laws of the distribution of individual elements over numerous phases—minerals—depend mainly on the structures of the outer electron shells of the atoms. Consequently, in geochemistry there is extensive use of the laws established by crystal chemistry. Ions and atoms in crystal lattices have different radii Ri. The value of Ri is connected with the chemical element’s position in the Mendeleev system. In the vertical groups, Ri usually increases with increasing atomic mass and decreases with increasing ion valence within the period (see Figure 2).
In natural processes of separation, ions and molecules are separated according to size. The crystal lattices of the main rock-forming minerals accept some ions (or atoms) and not others, depending on such properties as size and charge. If ions have different valences but close Ri, values, then the ion with the larger charge most frequently enters the lattice. If the ions have the same valence and do not differ by more than 15 percent in size, they are often isomorphously placed in the crystal lattice. An atom replaces an atom, an ion replaces an ion, or a group of atoms replaces a group of atoms, depending on the type of lattice, dimensions Ri charge, and so forth. Isomorphous replacement plays a large role in the distribution of elements in various minerals. The use of Ri, in geochemistry explained the association of such dissimilar elements as U, Th, and the rare-earth elements (in such minerals as thorianite and yttrialite), as well as the constant association of rare-earth elements. Under the deformation of one ion by another in a compound with a low-radius cation and a large-radius anion, so-called polarization arises, destroying the physicochemical properties of the substance— for example, hardness and volatility. The ratio of the cation Ri to the anion Ri, determines the number of atoms surrounding the central atom in the compound—its coordination, that is, the coordination number. This in its turn indicates the nature and structure of the crystal lattice. The coordination number can vary with conditions of formation of a mineral. The structure of the crystal lattices of minerals vary from very simple and symmetrical structures of closely packed spheres to quite complex ones with little symmetry. During crystallization, atoms and ions strive to arrange themselves in the crystal lattice in such a way that the crystal lattice energy is minimal. From all these data the geochemical classification of the elements was compiled, based on the physicochemical properties of the chemical elements (Table 2).
With the discovery of isotopes, the geochemistry of isotopes began to develop. This is a study of the natural processes of the separation of the isotopes of chemical elements, particularly of such light elements as H, C, O, N, and S. In this manner it is often possible to establish the method and conditions of the separation of chemical elements and of the formation of individual minerals and ore deposits.
On the earth, geochemical processes of the separation of elements are primarily maintained by the heat generated by radioactive elements (radiogenic heat) and by gravitational energy. At the earth’s surface a considerable part is played by solar radiation energy, which in particular is converted by living matter into the chemical energy of petroleum and coal.
|Table 2. Geochemical classification of chemical elements|
|Siderophiles (iron)||Chalcophiles (sulfides)||Lithophiles (silicates and others)|
|Fe, Ni, Co,||S, Se, Te,||H, 0, N, Si, Ti,|
|Ru, Rh, Pd,||Cu, Zn, Cd,||Zr, Hf, F, Cl, Br, I,|
|Os, Ir, Pt,||Pb, Sn, Mo,||B, Al, Sc, Y,|
|(Mo), Au, Re,||Ge, As, Ga,||Li, Na, K, Rb, Cs,|
|(P), (As), (C),||Sb, Bi, Ag,||Be, Mg, Ca, Sr, Ba, Ra,|
|(Ge), (Ga), (Sn),||Hg, In, Tl,||V, Cr, Mn, W, Th,|
|(Sb), (Cu)||(Fe), (Ni), (Co)||Nb, Ta, U, Ac, Pa,|
|(S), (P), (Sn), (C), (Ga),|
|(Fe), (Ni), (Co),|
Geochemical processes. The primary separation of cold undifferentiated matter of the earth into shells took place under the action of the heat of the adiabatic contraction of the planet and radiogenic heat. At various depths in the earth’s mantle, particularly in the asthenosphere, numerous molten pockets arose. Separation into shells proceeded by way of zone melting, not requiring the complete fusion of the mantle. The silicate matter of the planet was divided into a refractory phase—the ultrabasic rocks of the upper mantle—and an easily fusible phase—the basic rocks (basalts) of the lithosphere. The easily fusible matter melted the roof of the magmatic pockets, and the refractory matter crystallized at the bottom of the pocket. Thus the easily fusible matter transported up to the earth’s surface. The metasilicates incongruently broke up into orthosilicates and silicic acid, enriched with chemical elements that lower the melting point, namely, the alkali metals, Si, Ca, Al, U, Th, Sr, and other rare lithophilic elements. Substances that raise the melting point (such as Mg, Fe, Ni, Co, and Cr) were retained mainly in the refractory phase, that is, remained in the earth’s mantle. Zone melting was accompanied by the degasification of the upper mantle.
The remelting and degassing of the mantle have a periodic character. After the transport of heat and matter from the depths of the earth to the surface, time was needed to again heat up the pocket. The entire rhythm of tectonic-magmatic and volcanic activities and of metamorphic transformations is connected with such a geochemical cycle. This process also occurred on the moon, and obviously on all earth-type planets. The earth’s chemical evolution is maintained and regulated by a continuous process of the remelting and degassing of substances in the mantle by the energy generated by radioactive decomposition.
The material in the earth’s mantle (peridotites, dunites, and other ultrabasic rocks) have a chemical composition approximating that of meteorites. (Table 3.) The high temperatures and pressures prevailing in the mantle result in polymorphic changes in the minerals, for example, in the formation of “stishovite,” a quartz with a density of 4,350 kg/m3 (at normal temperature and pressure). Owing to this, the material in the mantle is divided into zones of different density. The material of the upper mantle penetrates to the surface at continents in dunite belts rich in chromites, platinoids, and high temperature sulfides and in oceans, in rift valleys of the midocean ridges.
With reference to the presence of sulfide ores in the earth’s crust, geologists previously considered it probable that the mantle contained a sulfide shell. However, the determination of the isotopic composition of lead from various sulfide ores showed their different absolute ages; consequently, the rejection of sulfides by the rocks occurred at a different time, so that the hypothesis of a sulfide shell is not well founded. The process of the formation of the metallic Fe-Ni alloy, of which the earth’s core is composed, is studied the least. The core probably formed during the processes of agglomeration in the
|Table 3. Chemical composition of the earth’s rocks, lunar rocks, and meteorites|
|Oxides and elements||Stony meteorites (chondrites)||Ultrabasic rocks of earth||Primary basalts of earth (tholeiitic)||Eucrites (basaltic stony meteorites)||Rocks from lunar surface||Average composition of sedimentary rocks of earth||Earth’s granites|
|Crystalline (basalt)||Finely dispersed (regolith)|
|Apollo 12||Luna 16||Apollo 12||Luna 16|
|In percent by mass|
|FeO...............||12.45||9.84(+2.51 Fe2O3)||9.0(+2.88 Fe2O3)||17.6||21.3||19.35||17||16.64||1.95(+3.3 Fe2O3)||1.8(+ 1.6 Fe2O3)|
|10-40 percent by mass|
protoplanetary cloud and further during the adiabatic contraction of the earth, which continued over a long period of time.
Above the mantle is the lithosphere, which is separated from the material of the mantle by the Mohornviôić discontinuity. There are two types of lithospheres: the continental and the ocean. The thickness of the former is 35-40 km, of the latter 6-8 km. The primary (tholeiitic) basalts of the oceanic crust are a more complex system than the material of stony meteorites; they consist of at least four main components—MgO, SiO2, FeO, and A12O3. The Si/Mg ratio in these basalts equals 6.5, that is, the basalts are not of solar composition. Basalts of the lithosphere, lunar rocks (from the surfaces of the lunar “seas”), and eucrites (basaltic stony meteorites) have identical compositions and the same ophitic structure. In silicate and other systems, water and other volatiles that lower the system’s melting point play an exceptional part. The most significant effect on magmatic processes is exerted by water in a state close to the supercritical.
Seismological methods reveal pockets filled with liquid magma in the mantle beneath volcanoes. The outflow of basalts is accompanied by the liberation of water vapor— about 7 percent by mass (20 percent by volume) of the discharged basalt and acid fumes and gases (C02, HF, HCl, S, and SO2). Mainly CO2, HF, and HCl are liberated in the high temperature stage of the basalt cooling (600°-800° C). At medium temperatures (about 200° C), sulfur compounds are also liberated. CH4, NH4C1, H3BO3, CO2, and other gases are liberated at low temperatures and in the postvolcanic (fumarole) stage, as well as mineralized solutions. The formation of CO2, CO, and CH4 is a result of the reaction between carbon and H2O in the magma at various temperatures and pressures. This process is accompanied by the partial separation of carbon isotopes—by a weight increase of carbon (the 13C content) in CO2, diamonds, and carbonatites (CaCO3 of kimberlite pipes) compared with the carbon of other rocks. On cooling, the basalt lava undergoes fractional crystallization with the formation of various magmatic rocks, which have common properties. Liquation (for example, the separation of high temperature Cu-Ni-Fe sulfides from silicates) and gas transport are possible in the magmatic stage of differentiation. Magnetite and titanomagnetite can arise in an early stage of the fractional crystallization of magma as a result of oxidation in the magma Fe2+-» Fe3+. Magnetite does not dissolve in the silicate melt and carries with it Ti because of the closeness of the Ri, of Fe3+ (0.65) to Ri of Ti4+ (0.60). In the main crystallization stage, plagioclase is formed, from labradorite to oligoclase, and many other aluminosilicates. As the melt cools, the more easily fusible and volatile compounds accumulate in the melt, which at a certain stage react with previously separated higher temperature compounds (the Bowen reaction principle). In this partition mechanism, ions that did not enter into rock-forming minerals because of their large or very small Ri values are concentrated in the residual melt. The origin of pegmatites and other rocks, rich in rare elements, is connected with these residual melts.
Acid rocks—granites, granodiorite, and so forth—are widely distributed in the lithosphere. Some of them have a considerable amount of Ca (approximately 2.5 percent) and heavy metals and are low in alkalis and volatiles, while others are poor in Ca (about 0.5 percent) and heavy metals but rich in alkalis and volatiles. Most scientists associate the origin of granites with eutectic melting and with the process of the granitization (metamorphism and metasomatism) of sedimentary rocks at different levels in the lithosphere. The higher 18O content in the quartz of granites corresponds to the relatively low temperatures of mineral formation.
In the lithosphere of continents ore deposits are formed. These are beds of many chemical elements, in particular Fe, Cu, Ni, Co, Pb, Zn, Mo, Ag, and Hg, in the form of oxides, sulfides, and so forth. Their origin is connected with hydrothermal solutions, which also carry gases. Despite the known diversity of their composition, which depends on depth, temperature, and other conditions of formation, they all possess common properties, for example, Si02-Au or Pb-Zn-Cu associations in the form of sulfides, or SnO2. WO3-H3BO3-F associations in hydrothermal and greisen deposits. Hydrothermal formations and greisens are regarded as end products of a tectonic magmatic process or granitization. The sources of the ore material of thermal springs can be both subcrustal processes and processes within the lithosphere. The mode of transport of heavy elements is disputed. Gas transport of heavy metals, for example, in the form of fluorides, is not excluded, while fluorine often gives large diffuse halation in enclosing rocks. The equilibriums of fluorides, chlorides, and metals with H2O at various temperatures and pressures are unclear.
The chemical and physical conditions for ore formation is indicated by the composition of the gas-liquid inclusions in ore minerals, which contain solutions of NaCl, MgCl2, MgSO4, KCl, H2S, SiO2, and carbonates and traces of metals; often the CO2 pressure is high, up to 2,000 atmospheres. These solutions are almost neutral; the temperature at which they formed ranges from 50° to 550° C. The ordinary sulfides of such heavy metals as Pb, Zn, Cu, and Fe are not readily soluble in water, and changing the temperature and pressure does little to change their solubility. For example, it would be necessary to evaporate about 10 km3 of water to precipitate 1 ton of zinc from a solution of ZnS. The transport of sulfides as colloidal solutions, or sols, is also improbable. However, there are complex compounds of heavy-metal sulfides that are more soluble than simple sulfides, for example, the ions HZnS22- or HgS-. A large role in the transport process of heavy metals by hot solutions is played by the concentration of CO2 and probably that of other gases: O2, H2S, and PH3. For example, U forms complexes [UO2(CO3)3]4- that are readily soluble in H2O at a specific CO2 concentration. Decreasing the amount of CO2 in the solution destroys this complex and causes deposition of U compounds. Deposition of heavy metals is also controlled by the partial pressure of H2S, which determines the sequence of metal deposition in a sulfide body, as well as by the partial pressure of CO2, the oxidation potential, and so forth. The crystallization of sulfides, as for example of Pb and Zn, and the distribution in them of such rare elements as In, Ga, Ge, and Tl take place according to the laws of isomorphism. The process of deposition of sulfides is reflected in the isotopic ratio S32/S34 of minerals, which is of diagnostic importance.
Magmatic rocks on the earth’s surface are destroyed by the action of climatic and a number of other agents: organisms, water, carbon dioxide, and organic substances. This process depends on the concentration of hydrogen and oxygen ions, the ionic potential, and other factors. Weathering leads to complex changes in rocks. For example, feldspars are converted to kaolinite, carbonates, and quartz, and Na, Mg, and K in the form of chlorides, sulfates, and carbonates turn into solution and are carried away by streams to the ocean. The total volume of rocks increases (Figure 3) as a result of hydration and carbonation. The destruction of rocks involves many chemical processes, for example, the hydrolysis of aluminosilicates which leads to the formation of laterite, free hydrated oxides A12O3, and bauxites which are enriched with Ti, Nb, Sn, Be, and so forth. Oxidation to higher valences is often effected by microorganisms, for for example, Fe2+#x2192;Fe3+ and Mn2+→Mn4+. Sedimentary iron ores are enriched with phosphates, arsenates, and vanadates and manganese iron ores, with Ba, Ra, Co, and so forth. Limestones as well as dolomites, phosphates, and some other salts are formed with the participation of organisms, and they accumulate Sr, Mn, Pb, F, and the rare-earth elements.
Salt-bearing deposits arise as a result of the evaporation of water in isolated basins. The order of successive deposition of NaCl, MgSO4, and other salts follows the laws of halogenesis. In this process solid salts separate from a saturated solution—natural brine—which contains the most soluble salts of Na, K, Sr, Li, B, and Br. Similar solutions are also encountered in underground highly mineralized waters.
Land organic matter when buried leads to the formation of coal, while the organic matter of benthic deposits of modern and ancient seas (mainly plankton) leads to the formation of petroleum and fuel gases. Isotopic 12C/13C analysis of individual petroleum fractions shows that they were formed at temperatures not exceeding 200-250° C. The appearance of coal and petroleum in the lithosphere altered the migration and distribution of a number of elements. Thus for example, U, V, and Ge are usually concentrated in sedimentary iron
ores. With the appearance of coal, their compounds began to accumulate in both coals and bitumens, often forming deposits of those elements. When rocks disintegrate, the most stable minerals—monazite, thorite, gold, magnetite, quartz, zircon, rutile, and cassiterite, among others—accumulate in coastal waters and form placer deposits in marine shelf zones.
Sedimentary rocks of continents reach a thickness of 20 km in some places, and on the average measure over 1 km thick. Table 4 shows the total amount of sedimentary rocks on the earth. Most of the rocks consist of clays and slates (about 55 percent), carbonate rocks (about 25 percent), and sands and sandstone (about 20 percent).
|Table 4. Amounts of sedimentary rocks on the earth (in kg)|
|Deep-sea regions...............||2.17 × 1020|
|Bathyal regions...............||1.0 × 1021|
|Ancient platform shields...............||1.4 × 1020|
|Young platforms...............||3.4 × 1020|
|Total...............||1.7 × 1021|
All magmatic and sedimentary rocks undergo varying degrees of metamorphism. The different processes in the solid matter of rocks proceed either without the loss and addition of material from outside (metamorphism proper) or with the loss and addition of material (metasomatism). There are alkaline metasomatism (sodium and potassium metasomatism) and magnesium, calcium, iron, sulfur (beresitisation of granites), phosphate, borate, and other metasomatisms. Clays are converted to slates, limestones to marbles, and so forth. Under the action of high temperature, deep-seated rocks can undergo remelting (palingenesis, granitization). All changes connected with metamorphism are directed at chemical equilibrium and recrystallization with a decrease in volume. High-density minerals are formed as are rocks of more or less similar mineral composition containing quartz, feldspar, and mica (systems of minimum free energy). Because of the complexity and diversity of metamorphic processes, the basis for classifying metamorphic rocks are their mineral associations (mineral facies), because they act as indexes of the conditions of formation of these rocks.
The regional unequal distribution of individual chemical elements makes it necessary to differentiate geochemical provinces on the earth. The study of the territorial distribution of chemical elements in connection with the geology of a region constitutes the problem of regional geochemistry, the ultimate aim of which is the compilation of geochemical maps of a territory on the basis of general geological data.
Geochemical processes in the hydrosphere, atmosphere, and biosphere. The earth’s water blanket, the hydrosphere, arose as a consequence of the outflow of basalts and the resulting effluence of water, CO2, and other gases. The world’s oceans and seas occupy about 71 percent of the earth’s surface and have a total volume of 1.37 × 1018 m3. The structure of the bottom of the oceans is a result of immense magmatogenic processes. Bottom sediments total about 1.2 × 1021 kg. Readily soluble materials enrich the aqueous solution; those of low solubility accumulate in the bottom sediments. The ratio of dissolved salts remains constant. Table 5 shows the main ions in ocean water. In the water, complex equilibria are established between organic matter, salts, gases, and other substances of the ocean solution and the chemical composition of the bottom sediments. All the waters of the continents (which are derived ocean waters) amount to 3 percent of the mass of the total ocean water. The main ions of the waters of rivers and fresh-water lakes are (in order of decreasing content) Ca2+, Na+, Mg2+, CO32-, SO42-, and Cl-.
|Table 5. The main ions of ocean water (per 1 kg of ocean water with salinity of 35.00 parts per thousand and chlorine content of 19.375 parts per thousand)|
|g/kg||g = equivalentl kg|
|The ocean’s present Sr content is taken as 8 × 10-4 percent|
About 500,000 km3 of water evaporates annually from the surface of the oceans; part of this water precipitates over continents, percolates through layers of sedimentary rocks, and forms groundwater. Buried waters of former marine silts form interstratal waters. The exchange between interstratal waters and rocks and the temperature of the strata are determining factors in the composition of groundwaters. Groundwaters of petroleum-bearing regions, rich in I and Br and sometimes in B, are known, as are waters containing calcium chloride (for example, in the Devonian rocks of the East European Platform) without sulfates and rich in Ra, waters containing hydrogen sulfide usually formed by reduction of SO42-, and waters rich in Li (in the Irkutsk amphitheater). In regions of ancient volcanic activity, the springs are cold and devoid of CO2. In regions of recent volcanic activity, hot springs of varied salt composition appear. Their classification has been worked out. The earth’s ancient gaseous shell was not very thick and consisted of CO2, H2O, possibly of CH4, and other gases. The present atmosphere arose subsequently, with the appearance on earth of free oxygen as a result of the photosynthetic activity of plants. After this, products of the volcanic exhalation of S, H2S, NH3, H2, and CH4 were oxidized and left the atmosphere, leaving the modern nitrogen-oxygen shell of the earth.
Through the action of volcanoes, the light gases He4, He3, H, and D (helium respiration) are released from the rocks and pass into the atmosphere. They are not held by the earth’s gravitational field and are dispersed in cosmic space. Volcanoes are also sources of CO2 as well as of traces of HF, HCl, and so forth. The CO2 content of the atmosphere is influenced by the ocean, which absorbs CO2 at cold latitudes and liberates it at the equator. Hence, the partial pressure of CO2 in the atmosphere is somewhat higher at the equator. The 40Ar isotope of argon accumulates in the air through nuclear reaction 40K → 40Ar (K capture). Other inert gases—Ne, Kr, and Xe—are of primary origin. The atmosphere plays a large part in the transport of many easily volatile compounds, halogens, organic substances, and so forth. The gases of the atmosphere, for example, O2 and CO2, participate in the geochemical weathering of rocks. Nitrogen is fixed by bluish-green algae and some other plants. After the destruction of the latter by metamorphism, potassium nitrate is formed from their residues.
The subsurface gases, which fill porous rocks, vary in composition and are formed in different ways. Atmospheric gases can be captured by sedimentary rocks. This is characterized by a 40Ar/N2 ratio of about 1 percent. Nitrogen streamers without 40Ar are the result of the metamorphism of organic substances (biogenic gases). Subsurface atmospheres of CO2 are known, as are CO2 streamers in regions of volcanic activity and the petroleum gases CH4, C2H6, C3H8 and other hydrocarbons in petroleum-bearing regions, hydrogen sulfide, and such radiogenic gases as He and Rn.
The biosphere is a region at the boundary of the solid, liquid, and gaseous shells of the earth, occupied by living matter, a totality of organisms. The biosphere arose about 3.5 × 109 years ago. Owing to the small thickness of the primary atmosphere, cosmic radiation penetrated to the earth. Under the action of this radiation, the abiogenic synthesis of many complex optically inactive compounds of carbon with symmetrical molecules took place from volcanic fumes and such gases as H2O, CO, CO2, HF, HCl, CH4, S, H2S, S2, NH3, and H3B03. Against this background the biogenic synthesis of asymmetrical optically active molecules of living matter has arisen. After the emergence of the nitrogen-oxygen atmosphere as a result of photosynthesis, an ozone screen formed over it. Because of this, cosmic radiation practically ceased to penetrate to the earth’s surface, and the abiogenic synthesis of organic compounds ceased. Organisms not only altered the composition of the atmosphere, but directly or indirectly participated in numerous geochemical processes.
History of individual elements of the lithosphere. The geochemistry of the individual elements and their behavior in different natural processes constitute a special branch of general geochemistry and are often of considerable economic interest. Regular parageneses (associations of elements) are encountered in various natural processes, but subsequently the separation of elements can also take place. For example, all halogens in the form of HF, HCl, HBr, and HI come to the earth’s surface with volcanic emanations. Further, I- compounds under the influence of oxidizing-reducing reactions (and solar radiation) are more easily oxidized than other halogens, that is, transform into I2, which is transported through the atmosphere and completes its cycle on the earth’s surface (Figure 4). The HF of volcanic gases is immediately fixed by the matrix rocks, particularly by the P2O5 molecule, forming a stable molecule that is the basis of fluoroapatite. Salts of HCl and HBr pass into aqueous solutions and migrate together. For them, the separation process is mainly a process of the precipitation of salts by evaporationof solutions in isolated basins. NaCl precipitates, while Br salts remain in the brine of lakes. In the ocean the CI/Br ratio is close to 300, about the same as in lakes, rivers, and the like. But in halite deposits the CI/Br ratio is approximately 10,000 or more, while in brine (or the Dead Sea) it is about 50. Thus, the origins of mineral solutions can be established by this CI/Br ratio.
Another example: S, Se, and Te are ejected by volcanoes. In hydrothermal ore deposits and in the sulfides of heavy metals they occur together, but on the earth’s surface they separate: S is easily oxidized into SO42- and is discarded into the sea. When sea waters evaporate, precipitates are formed of calcium sulfate—gypsum and anhydrite. Se oxidizes with difficulty, and in the form of insoluble aqueous salts (Fe and others) of selenious acid forms occlusions. Te disperses on oxidation. The migration of Ca, Sr, Ba, and Ra has many general stages. However when Ba encounters SO42-, it gives the insoluble compound BaSO4. At the same time and place, RaSO4 also accumulates. Ca and Sr bicarbonates are discharged into the ocean in the form of aqueous solutions. In this case, owing to the high solubility of Sr2+ salts, this ion does not pass into the carbonate deposits but accumulates in the solutions. Even more complex separation processes take place during the formation of sulfide hydrothermal deposits and in many other cases. The migration of individual elements from one thermodynamic system to another is part of the general exchange or migration of material on the earth.
Geochemistry and other sciences: history. Geochemistry lies at the juncture of the geological, physical, and chemical sciences and is linked through biogeochemistry with the biological sciences. Geochemistry is most closely connected with the geological sciences—mineralogy and petrology— particularly in problems related to the origins of minerals, rocks, and geological processes. Regional geochemical investigations are carried out in close conjunction with geotectonic structures. Geochemistry uses modern physical and chemical methods of investigating matter and processes over a wide range of temperatures and pressures: spectral, mass spectral, resonance, and nuclear methods; mathematical methods are also used. The study of the behavior of a substance at high temperatures and pressures ties geochemistry to geophysics. The evaluation of absolute time, on which historical geology is based, and a number of other problems of the earth’s history, can be solved only by the precise methods of geochemical and radiochemical investigations. In paleontology, when solving problems of the formation of the solid skeletal parts of organisms and their evolution, it is important to know the geochemical conditions under which the organisms lived. The study of fossil organic matter reveals the processes of formation of mineral fuels. Geochemical ideas play a very large role in the development of soil science. They are directed at solving a number of important problems in agricultural chemistry and agronomy. The geochemical study of the soil mantle is very important in the geochemical prospecting of minerals. A geochemical trend is also developing in geography—landscape geochemistry. The study of geochemical processes connected with flora and fauna is of great importance in agriculture and medicine.
The ideas of geochemistry penetrate into astrophysics, atomic physics, chemistry, physical chemistry, chemical engineering, and metallurgy (especially of the rare metals). Geochemistry successfully develops and puts into practice the geochemical prospecting of mineral deposits and solves problems of the comprehensive use of mineral raw material. It plays an active part in the vast work in progress in the Soviet Union to introduce chemical processes into the national economy, particularly into agriculture.
Geochemistry arose on the basis of the study of atoms. Its roots can be found in past geological-mineralogical knowledge. Geochemical ideas had already appeared at the end of the 18th century. The German geologist K. G. Bischof and the French geologist Élie de Beaumont, among others, accumulated geochemical facts about the composition of matter and its migration in aqueous solutions and in magmatic and volcanic processes. In the first half of the 19th century the Swedish chemist and mineralogist J. J. Berzelius studied the chemical compositions of a large number of minerals and was the first to propose a chemical classification of minerals. The chemical analysis of minerals and rocks, investigations of the chemical compositions of natural gases and waters, and the chemical study of minerals made it possible, in the middle of the 19th century, to lay the foundations of geochemistry. In 1838 the Swiss chemist C. F. Schonbein first used the term “geochemistry.” By the end of the 19th century and the beginning of the 20th, much geochemical information had been obtained. It was the American geochemist F. W. Clarke who in 1882 made the first comprehensive survey of the data. It was the Soviet academicians V. I. Vernadskii and A. E. Fersman and the Norwegian geochemist V. M. Goldschmidt who formulated the principal tasks of geochemistry. N. S. Kurnakov and his schools contributed substantially to geochemistry with their work that laid the foundations of the geochemistry of halogenesis as well as of the physicochemical analysis of natural salt systems. The period after the Great October Socialist Revolution was particularly conducive to the development of the ideas of Vernadskii and Fersman. In the USSR their students—A. P. Vinogradov, D. I. Shcherbakov, P. N. Chirvinskii, N. V. Belov, A. G. Betekhtin, N. M. Strakhov, V. S. Sobolev, K. A. Nenad-kevich, V. G. Khlopin, A. A. Saukov, K. A. Vlasov, V. V. Shcherbina, V. I. Gerasimovskii, and N. I. Khitarov, to name a few—worked and continue to work on both general and individual problems of geochemistry.
In the second half of the 20th century studies of the radioactivity of rocks and minerals were intensified, and isotope geochemistry developed. In addition, work on the determination of the absolute ages of rocks began to develop extensively. In the USSR geochemical investigations have been carried out not only in scientific research institutes but also in numerous manufacturing enterprises. Geochemistry is taught at universities and other educational institutions. A number of geochemical institutes and divisions have been created, among them a biogeochemical laboratory, later reorganized into the V. I. Vernadskii Institute of Geochemistry and Analytical Chemistry. The journal Geokhimiia began publication in 1956.
REFERENCESVernadskii, V. I. Ocherki geokhimii, 4th ed. Moscow-Leningrad, 1934.
Fersman, A. E. Geokhimiia, vols. 1-4. Leningrad, 1933-39.
Fersman, A. E. Pegmatitty, 3rd ed., vol. 1. Moscow-Leningrad, 1940.
Vinogradov, A. P. Geokhimiia redkikh i rasseiannykh khimicheskikh elementov v pochvakh, 2nd ed. Moscow, 1957.
Vinogradov, A. P. Vvedenie v geokhimiiu okeana. Moscow, 1967.
Vinogradov, A. P. “Predvaritel’nye dannye o lunnom grunte, dostavlennom avtomaticheskoi stantsiei ‘Luna-16’.” Geokhimiia, 1971, no. 3.
Vinogradov, A. P. The Elementary Chemical Composition of Marine Organisms. New Haven, 1953.
Saukov, A. A. Geokhimiia [3rd ed.]. Moscow, 1966.
Clarke, F. W. The Data of Geochemistry, 5th ed. Washington, D. C, 1924.
Goldschmidt, V. W. Geochemistry. Oxford, 1954.
Rankama, K. Progress in Isotope Geology. New York-London, 1963.
Krauskopf, K. B. Introduction to Geochemistry. New York-London, 1967.
Handbook of Geochemistry, vols. 1-2. Edited by K. H. Wedepohl. Berlin, 1969.
Mason, Br. Principles of Geochemistry, 3rd ed. New York-London-Sydney, 1970.
Slater, J. C. “Atomic Radii in Crystals.” Journal of Chemical Physics, 1964, vol. 41, no. 10, pp. 3199-3204.
Ahrens, L. H. “The Use of Ionization Potentials: part 1—Ionic Radii of the Elements.” Geochimica et cosmochimica Acta, 1952, vol.2, no. 3.
A. P. VINOGRADOV