Contents Page Table 1 1. Average chemical analyses of certain types of plutonic igneous rocks (R. A. Daly) 2 2. Approximate mineral compositions of principal types of plutonic igneous rocks (E. S. Larsen) 3 3. Composition of stony meteorites (L. LaForge) 4 4. Composition of iron meteorites and meteoritic iron (L. LaForge) 5 R emarks : Each average composition was computed from single analyses of specimens to which the respective reporters gave the name printed at the top of the column. The average may thus be regarded as at the center of gravity of opinion as to the composition of rocks which may properly be described by this name. For many purposes, these average compositions may be taken as adequately descriptive of the materials falling into the same classifications in other tables; for other purposes, the analysis of the individual specimens on which measurements have been made is essential. Since most of the analysts neglected the effect of the adsorption of water by their specimens when pulverized, the proportion given for this oxide is in general somewhat too high. The analyses of significantly weathered specimens were excluded, insofar as possible, from the data finally used. On the whole, about three-fourths of the total water indicated was driven off at temperatures no higher than 105°C. It should be noted that the percentage of TiO 2 in many of the older analyses is too low. This error of itself would make the reported percentages of Al 2 O 3 slightly too high. The average composition of the . . .

Contents Page Table 2 1. Density of minerals (H. Berman) 8 2. Average densities of holocrystalline igneous rocks (R. A. Daly) 14 3. Densities of crystals and corresponding glasses (R. A. Daly) 15 4. Density of natural glasses (R. A. Daly) 16 5. Density of crystalline rock and corresponding glass (artificially prepared) (R. A. Daly) 16 6. Porosity and bulk density, dry and saturated, of sedimentary deposits (H. C. Spicer) 17 7. Loss of weight on drying at elevated temperatures (H. C. Spicer) 26 T able 2 – 1.—D ensity of M inerals (H. Berman) The densities of minerals here listed are selected from the complete list by Rosenholtz and Smith, Rensselaer Polytechnic Institute Eng. and Sci. Series, No. 34, Troy, 1931, who have compiled all the published values. The ranges given are, in most cases, not significant, since they arise from an uncritical compilation of the limiting values found in the literature. Some minerals have a considerable spread in specific gravity as a result of isomorphous substitution, and in these instances it is necessary to correlate a specific chemical analysis with the given value of density. In a number of instances, new values marked * have been substituted for those in Rosenholtz and Smith. Room temperature, 1 atm. Name Density gm./cm. 3 Name Density gm./cm. 3 Acanthite 7.2 – 7.33 Annite 2.8 – 3.4 Actinolite 2.9 – 3.2 Anorthite 2.70 – 2.76 Adamite 4.34 – 4.35 Anorthite, pure 2.76 Adularia 2.56 Anorthoclase 2.56 – 2.65 Aegirinaugite 3.5 Anthophyllite 2.86 – 3.2 Aegirite 3.5 – 3.56 Antigorite 2.55 – 2.62 Aegirite-hedenbergite 3.502 * Antimony

Contents Page Table 3 1. Thermal expansion of single crystals 30 1.1. Thermal expansion of quartz 35 2. Density at high temperature, liquid and crystalline states 36 3. Density and thermal expansion of a few commercial glasses at high temperature 37 4. Thermal expansion of rocks 37 The following tables contain the expansion of a selected list of naturally occurring substances, together with a few artificial ones like steel and silica glass for comparison. All the minerals for which high-temperature data are available are included, although some are of more interest as refractories than in geology. A number of other minerals whose expansions have been measured only to 100°C. or less are also included; a much more complete list of these latter will, however, be found in the International Critical Tables. In view of the many irregularities which have been found, extrapolation from these low temperatures to those of greater geologic interest is unsafe. The tables list the total expansion from 20°C., in per cent of the length or volume at 20°C. The nomenclature in regard to axes, angles, etc., follows Dana’s Textbook of mineralogy (John Wiley and Sons, New York, 1922). For isometric substances only the linear expansion is given; the volumetric expansion may be computed to a close approximation by multiplying the linear expansion by three. For the other crystallographic classes, both linear and volumetric values are quoted. With tetragonal and hexagonal minerals, a change of temperature distorts an original sphere into an ellipsoid of revolution so that . . .

Contents Page Table 4 1. Compressibility of elements crystallizing in the cubic system 43 2. Compressibility of elements crystallizing in the hexagonal and tetragonal systems 46 3. Compression of a few miscellaneous elements 47 4. Compression and thermal expansion of the alkali metals 48 4.1. Effect of pressure on thermal expansion 48 5. Compression of elements to 50,000 kg/cm 2 49 6. Compression of compounds to 50,000 kg/cm 2 50 7. Compressibility of cubic compounds 52 8. Compressibility of hexagonal, trigonal, and tetragonal compounds 54 9. Compressibility of orthorhombic compounds 56 10. Compressibility of monoclinic and triclinic compounds 58 11. Change of compressibility on melting 59 12. Compressibility of several artificial and natural glasses 60 13. Compressibility of rocks at low pressures 61 14. Compressibility of rocks at high pressures 62 Density and specific volume are dependent not only upon the temperature but also upon the stress. A stress system consisting of uniform pressure in all directions is known as “ hydrostatic” pressure; the change of specific volume or of density for not too great changes of hydrostatic pressure may be described in terms of a single coefficient, the compressibility β, defined by β = 1 V 0 ( d V d P ) T = 1 ρ 0 ( d ρ d P ) T , where V 0 is the specific volume, ρ 0 the density at 1 atmosphere, and P the pressure. Since dV/dP is intrinsically negative, β is a positive number, with the dimensions of the reciprocal of a pressure or stress. In general, β depends upon the pressure and the temperature. Its reciprocal K is known as the bulk modulus. The compressibility . . .

Contents Page Table 5 1. Elasticity of single crystals 66 1.1. Extreme values of Young’s modulus and of rigidity in certain metal single crystals 69 1.2. Effect of temperature on the elasticity of quartz 70 2. Elasticity of polycrystalline metals at ordinary pressure and temperature; effect of pressure on rigidity 71 3. Effect of temperature on the elasticity of polycrystalline metals 72 4. Elastic constants of rocks at ordinary pressure and temperature 73 5. Effect of stress on Young’s modulus of rocks 78 6. Effect of stress on Young’s modulus and on σ in granite. 79 7. Effect of moisture upon velocity in Amherst sandstone 79 8. Elastic parameters of certain rocks at 4,000 kg.cm −2 and 30°C 80 9. Rigidity and velocity of shear waves, in rocks, as function of pressure 81 10. Effect of temperature on the velocity of shear waves in rocks at high pressure 83 11. Elasticity of glass and glassy rocks; effect of pressure upon rigidity 84 12. Effect of temperature on elasticity of glasses 85 13. Elasticity of rocks from seismic determinations 86 (For compressibility, see Section 4.) (For velocities determined in the field, see Section 7.) For sufficiently small stresses, two elastic “constants” are necessary and sufficient to determine the elastic behavior of a homogeneous, isotropic, elastic solid. Various pairs of constants are in common use; those usually measured in the laboratory are Young’s modulus E , the modulus of rigidity or the shear modulus G , Poisson’s ratio σ, and the compressibility β or its. . . .

Contents Page Table 6 1. Internal friction in single crystals 89 2. Internal friction in polycrystalline metals 90 3. Internal friction in glass 91 4. Internal friction in rocks 92 The vibrations of solid bodies are accompanied by dissipation of energy attributable to their “ internal friction”; this loss of energy is additional to whatever external losses may occur. There are various ways of specifying the internal friction, and few of the published researches have adopted a common terminology. In this section the results have been reduced to show a dimensionless quantity, 1/Q , which may be called the “dissipation function” or the “internal friction.” The logarithmic decrement Δ of free vibrations is related to this function by Δ = π/Q. If dE is the loss of energy per cycle, and E the total energy, then 1/Q = dE/(2πE), Thus the internal friction is small when 1/Q is small. Only in the last few years has the technique of measuring internal friction reached a point where the losses in single crystals can be studied [ 4 , 16 , 17 ]. The mechanism of internal dissipation of energy in single crystals probably involves plastic flow and strain hardening even for very small strains. The internal friction is sensitive to the condition of the surface and to annealing [ 16 ]. Unannealed single crystals may show losses approaching those of polycrystalline material. In the latter, internal friction arises from a number of distinct sources, including (1) losses within the individual crystals, (2) losses at the surfaces of the . . .

Contents Page Table 7 1. Field determinations of wave velocities at small depths 95 2. Velocity of compressional waves in soils, water, petroleum 96 3. Velocity in relation to depth and age 97 4. Seismic velocities and thickness of upper layers in various regions 99 5. Seismic velocities in the interior of the earth 101 Seismic prospecting or surveying constitutes one of the most important geophysical diagnostic methods; knowledge of the wave velocities in different materials is evidently a primary requirement, to which a great deal of effort has been devoted in the field and laboratory. Unfortunately for the sensitivity of the diagnosis, a given material is not characterized by a unique set of velocities, different from those of all other materials. For each material, the velocities depend upon a number of factors, such as alteration, present and perhaps former depth, and in some cases moisture, direction with respect to schistosity, bedding, and so on. Thus a range of values must exist even for a fairly well defined material, and this range frequently overlaps, to a large extent, the ranges of other materials. A velocity determined in the field for a more or less well-defined formation will represent an average for the varying conditions within this layer along the path of the wave; for different depths of penetration, different conditions and hence different velocities may be expected. The determinations which have been published are not adequate for the formulation of definite relations between velocity and the various other factors except . . .

Contents Page Table 8 1. Figure of the Earth’s surface 104 2. Gravity at the surface; mass, mean density, moment of inertia 104 3. Pressure, density, acceleration of gravity and ellipticity of layers of equal density in the Earth 105 4. The Earth’s core 106 1. Benfield, Z. f. Geophys. 13, 157, 1937. 2. Birch, Bull. Seism. Soc. Amer. 29, 463, 1939. 3. Bullen, M. N. R. A. S. Geophys. Supp. 3, 395, 1936. 4. Bullen, Trans. Roy. Soc. New Zealand, 67, 122, 1937. 5. Bullen, idem, 69, 188, 1939. 6. Bullen, Bull. Seism. Soc. Amer. 30, 235, 1940. 7. Gutenberg and Richter, Beitr. z. Geophys. 54, 94, 1939. 8. Gutenberg and Richter, M. N. R. A. S. Geophys. Supp. 4, 363, 1938. 9. Jeffreys, M. N. R. A. S. Geophys. Supp. 4, 537, 548 and 594, 1939. 10. Lambert, The Internal Constitution of the Earth, Chap. 13, McGraw-Hill, New York, 1939. 11. Olczak, Acta Astron. 3, 81, 1938. 12. Williamson and Adams, J. Washington Acad. Sci. 13, 413, 1923. T able 8–1.—F igure of the E arth’s S urface (Ref. 10) Semi-axes of International Ellipsoid of Reference (1924): Equatorial, a = 6378.388 km. Polar, c = 6356.912 km. Radius of sphere of equal volume, α 0 = 6371 km. Ellipticity of surface, ( a – c )/ a = 1/297 = 0.00337. (For ellipticity of internal levels, see Table 8–3.) Quadrant of a meridian = 10,002.288 km. Quadrant of the equator = 10,019.148 km. Area of surface = 5.101 × 10 18 cm. 2 Volume = 1.083 2

Contents Page Table 9 1. Standard crushing strengths of rocks 116 2. Critical data of translation gliding 117 3. Critical data of twin-gliding 120 4. Tests of unconsolidated sand 122 5. Strength of rocks confined in steel jackets 122 6. Short-time compressive strength of unjacketed materials with confining pressure of kerosene 123 7. Resistance to shearing under high confining pressure 126 8. Creep of certain materials 129 9. Creep tests of wet alabaster at different stresses 129 Illustrations Page Figure 1. Tresca’s apparatus 110 2. Adams’ apparatus 111 3. Schematic diagram of Griggs’ high-pressure apparatus 112 4. Bridgman’s shearing apparatus 113 5. Karman’s stress-strain diagram of jacketed marble tested in compression under confining pressure 124 6. Böker’s stress-strain diagrams of jacketed marble tested in tension under confining pressure 125 7. Griggs’ stress-strain diagrams for unjacketed Solenhofen limestone 130 8. Creep curves of alabaster 130 In normal laboratory tests or in building applications most rocks are brittle rather than plastic and possess a well-defined strength which is independent of time or temperature within the usual limits of observations. This behavior has led to common acceptance of the concepts of “elastic limit” and “strength” of rocks, and these terms have been used frequently in reference to the behavior of rocks deep in the earth’s crust with the tacit assumption that such concepts apply under conditions of earth deformation without great change in form. On the other hand, geological observations provide abundant evidence that rocks have exhibited a high degree of plasticity in . . .

Contents Page Table 10 1. Viscosity of certain mineral and rock glasses 133 2. Viscosity in the system orthoclase-albite 136 3. Viscosity in the system diopside-albite-anorthite 136 4. Viscosity computed from field measurements 136 5. Effect of pressure upon viscosity 137 According to Newton’s law of fluid friction, the tangential or shearing stress in a liquid in motion is proportional to the rate of change, with time, of the angle of shear. Real liquids may be divided into two categories: “Newtonian” or “viscous” liquids, for which the factor of proportionality, the “viscosity,” is independent of the “rate of shear”; and “non-Newtonian” liquids, for which the factor depends upon this rate. The “Newtonian” behavior of many ordinary liquids, and of some glasses, has been verified for considerable ranges of shear rate. “Non-Newtonian” behavior has been observed for such materials as certain colloidal solutions, asphalt, and other bituminous products; no single value of “viscosity” is sufficient to describe the flow of such materials. ( See also Section 9.) If the stress is measured in dynes × cm. −2 , and the rate of shear in radians × sec. −1 , the viscosity will be given in dyne × sec. × cm. −2 or the equivalent unit, gram × sec. −1 × cm. −1 This unit is called the poise and is used in the tables in this section. The viscosity of a given material depends upon the pressure and temperature. Other factors of great importance for geological applications remain to be investigated, such as the effect of dissolved gases . . .

Contents Page Table 11 1. Elements 142 2. Oxides 143 3. Hydrous and hydrated oxides 146 4. Binary aluminates 148 5. Binary borates 149 6. Binary oxide systems 150 7. Three or more oxides (except SiO 2 ) 152 8. Binary silicates 153 9. Ternary and higher silicate systems 155 10. Aluminosilicates 158 11. Borosilicates 161 12. Miscellaneous systems containing silicates 161 13. Hydrothermal alteration of silicates and other minerals 162 14. Carbonates 163 15. Sulfates 164 16. Oxygen salts 166 17. Haloids 168 18. Sulfide-type minerals 170 19. Ternary sulfides 173 Scope of the tables .—These tables list the melting (congruent and incongruent) temperatures, boiling temperatures, and transition temperatures for the more important substances of interest to the geologist, the geochemist, and the petrologist. Eutectic and related data are given when the substance under discussion has been studied as a part of a binary or higher system. The aim in compiling the tables has been to describe the known thermal reactions of the compounds and the systems briefly, but in enough detail to be of use to an investigator in the field where extensive reference books are usually not available. Arrangement .—The entries are arranged alphabetically by chemical symbol (elements) or formula (compounds). Compounds composed of two or more oxides (except carbonates, sulfates, and other oxygen salts) are entered in the increasing alphabetical order of the constituent oxides. Mineralogical names are given following the nomenclature employed in Dana-Ford : Textbook of mineralogy (John Wiley and Sons Inc., New York).

Contents Page Table 12 1. Melting and solid-solid transitions of elements 177 1.1. Melting temperatures of A, H 2 , He, N 2 , Ne, and O 2 180 1.2. Formulas for melting curves 182 2. Melting and transitions of compounds 183 3. Liquid-vapor critical data 185 In systems containing a single component, the change Δ T of the temperature of equilibrium between two phases resulting from a change of hydrostatic pressure Δ P upon the two phases is given by Clapeyron’s equation, Δ T /Δ P = T Δ V /Δ H , where T is the absolute temperature, and Δ V and Δ H are the volume change and heat absorption, respectively, for a given mass of material in passing from one phase to the other. One phase may be considered the “high-temperature” phase, the other, the “low-temperature” phase. Heat is always absorbed in passing from the low-temperature phase to the high-temperature phase at constant pressure (the actual quantity may be negligibly small). In the same way, one of the phases will be the “low-pressure” phase, the other the “high-pressure” phase; in passing from the low-pressure phase to the high-pressure phase at constant temperature, the density always increases, or Δ V is negative. There is, however, no a priori way of knowing whether the high-temperature phase will be the high-pressure phase or the low-pressure phase. The liquid phase is “normally” the high-temperature, low-pressure phase with respect to solid phases; thus, in passing from a “normal” solid phase to the liquid phase, Δ V and Δ H are both positive, and Δ T /Δ P is positive (the melting temperature increases with . . .

Contents Page Table 13 1. Solubilities of different solids in water at atmospheric pressure 191 2. Effect of high pressures on the solubilities of certain salts in water at 25°C 193 3. Effect of pressure on the liquidus composition in the system dimethylaniline in monomethylaniline at 0°C 196 4. Effect of pressure on the eutectic temperature and composition in various systems 197 Illustrations Page Figure 1. Effect of pressure on the solubility of certain substances in water 190 2. Pressure-composition equilibrium diagram for the system sodium chloride-water at 25°C 191 3. Pressure-composition equilibrium diagram for the system potassium sulfate-water at 25°C 192 4. Pressure-composition equilibrium diagram for the system ammonium nitrate-water at 25°C 192 5. Pressure-temperature relations for the equilibrium: sodium chloride dihydrate-sodium chloride-saturated solution 194 6. Pressure-temperature relations for the equilibrium: sodium sulfate decahydrate-sodium sulfate-saturated solution 195 7. Effect of pressure on the solubility, at 30°C., of m-dinitrobenzene in ethylacetate, and of naphthalene in tetrachloroethane 196 8. Effect of pressure on the temperature-composition equilibrium curves in the binary system aniline-phenol 199 9. Effect of pressure on the temperature-composition equilibrium curves in the binary system m-cresol-o-cresol 200 10. Solidus and liquidus curves at different pressures in the system p-dichlorobenzene-p-dibromobenzene 201 11. Effect of pressure on the temperature-composition curves in the system aniline-cyclohexane 202 1. Adams, J. Am. Chem. Soc. 53, 3769 (1931). 2. Adams, J. Am. Chem. Soc. 54, 2229 (1932). 3. Adams and Gibson, J. Am. Chem. Soc. 54, 4520 (1932). 4. Cohen and Sinnige, Z. physik. Chem. 67, . . .

Contents Page Table 14 1. Triple points 205 2. Equilibrium between two solid phases 206 3. Melting curves 207 Specific volume of liquid water: 4.1. From 0° to 350°C.; 1 to 1000 bars 208 4.2. From −20° to 100°C.; 1 to 12,000 bars 208 4.3. At 25°C.; 1 to 12,000 bars 209 4.4. From 25° to 85°C.; 1, 500, and 1000 bars 209 5. Liquid-vapor saturation curve; heat of vaporization 210 Specific volume of water vapor: 6.1. From 100° to 650°C.; 1 to 2000 bars 211 6.2. From 700° to 1200°C.; 50 to 4000 bars 211 7. Heat capacity of air-free water 212 The data for specific volume are presented in separate tables rather than in a single composite because they are not all of equal accuracy. In order to express the results as a single composite, it would be necessary to assign statistical probabilities to each set of data, a procedure which at present would be arbitrary and unwarranted. Furthermore, no single set of data can be treated as self-consistent— i.e., as of uniform accuracy, and the change in experimental error with variations of temperature and pressure may at times have taken place in an unrecognized manner. Correction factors worked out for small overlaps cannot, therefore, be safely extrapolated over large areas. In this section temperature is expressed in degrees centigrade, pressure in bars, and volumes or changes in volume in cubic centimeters per gram. Although there seems to be no apparent uniformity of usage with respect to large . . .

Contents Page Table 15 1. Composition and vapor pressure of saturated solutions 216 2. Properties of 1 gram of saturated solution (KCl + H 2 O) 220 3. Freezing-point curves for three-phase silicate-water systems 221 4. Solubility of water in silicate melts at high temperatures and pressures (two-phase equilibrium) 222 This section presents the limited amount of information, some of it still unpublished, regarding solubility and vapor pressure in aqueous solutions at temperatures above 100°C. No attempt is made to cover the wealth of data which exists for temperatures below 100°C., and which is exhaustively tabulated in other places (Int. Crit. Tables, Vol. IV; Handbook of Chemistry and Physics, Chemical Rubber Publishing Co.). Solubility determinations from the cryohydric point to the triple point of the salt have been carried out for sodium and potassium nitrates (m.p. 308°C. and 334°C., resp.) and for boric oxide (m.p. 450°C.), but for higher-melting salts only parts of certain binary systems with water as one component have been investigated. The experimental methods, in order of decreasing accuracy, are: (1) Heating water and excess salt in a closed thermostated system, with some device for sampling a portion of the saturated solution for analysis [ 2 , 3 ]; (2) Heating known weights of salt and water in a closed glass system, and observing the temperature at which the last portion of salt passes into solution [ 4 , 5 ]; (3) For glass-forming solutes, quenching charges of glass after treatment with excess water in a closed bomb [ 6 , 10 ]; (4) Observing the discontinuity of . . .

Contents Page Table 16 1. Heat capacity of minerals 228 2. Heat capacity of rocks 235 3. Heats of transformation and of fusion 237 4. Heat capacity of gases 241 Heat capacities of minerals .—In the tables (in this section), the minerals are classified according to the first elements of their chemical formulae; for example, albite (NaAlSi 8 O 8 ) will be found under sodium, and calcite (CaCO 3 ) under calcium. The heat data are given in terms of absolute joules per gram (one joule = 0.23895 gram calories at 15°C.). True or instantaneous heat capacities may be obtained by measuring the amount of heat needed to raise the temperature of the sample one degree or fraction thereof. This is the preferable type of data, available for many materials at low temperatures. At high temperatures this method is experimentally difficult because of large radiation losses so that ordinarily a different technique is employed—namely, the method of mixtures. ( See W. P. White, The Modern Calorimeter, N. Y., 1928.) This latter method yields data in the form of mean or interval heat capacities. The mean heat capacity is usually given for the temperature interval between 273.1°K. (0°C.) or room temperature and some higher temperature. The quantity measured is, however, the difference in heat content between some elevated temperature and that of the calorimeter. In order to convert these quantities to true heat capacities the mean heat capacity data (at constant pressure) are fitted to the equation 1 c ¯ p | T T o = a + b T o + T 2 − c T o T where c̄ p is the mean heat capacity at . . .

Contents Page Table 17 1. Thermal conductivity of cubic single crystals 245 2. Conductivity of crystal powders cemented by compression to 8000 atmospheres 247 3. Thermal conductivity of noncubic single crystals 248 4. Thermal conductivity of rocks 251 4.1. Effect of wetting and of simple compression on the thermal conductivity of certain rocks 258 5. Thermal conductivity of soil, snow, ice 259 6. Thermal conductivity of glass 260 7. Conductivity of a few common metals 262 8. Thermal conductivity of miscellaneous materials 263 9. Thermal conductivity of some common liquids, as function of pressure and temperature 265 10. Effect of hydrostatic pressure upon the thermal conductivity of solids 266 In an isotropic homogeneous material the conduction of heat depends upon a single “constant” of the material, known as the thermal conductivity; this “constant” is a function of temperature, pressure, and other variables. The quantity of heat dQ conducted in unit time across an element of surface dS is given by the fundamental relation, d Q = − K d T d n d S where K is the conductivity and dT/dn the gradient of the temperature in the direction of the normal to the surface element dS . Two “ c.g.s.” units of heat are in common use; the calorie (gram-calorie) and the joule; 1 calorie equals 4.18 5 joules. Using the centimeter, second, and Centigrade temperature scale, the corresponding units for thermal conductivity are the cal.·sec. −1 cm. −1 deg. −1 and the watt.·cm. −1 deg. −1 (1 watt equals 1 joule/sec.). 1 c a l . ⋅ sec . − 1 cm . − 1 deg . − 1 = 4.18 5 watt ⋅ cm . − 1 deg . − 1 1 watt ⋅ cm . − 1 deg . − 1 = 0.239 cal . ⋅ sec . − 1 cm . − 1 deg . − 1 A unit sometimes encountered in engineering work is the British thermal unit per square foot per hour for . . .

Contents Page Table 18 1. Heat flow to the earth’s surface 269 2. Radioactive content and heat generation in rocks 270 3. Ultimate products of naturally occurring disintegration processes 271 4. Isotopic constitution of radioactive elements and their end products 272 5. Variations in the isotopic composition of common lead 273 6. Lead age ratios of selected radioactive minerals 274 7. Age results by lead isotope ratios 275 8. Helium age ratios of various magnetite deposits 276 THE EARTH’S HEAT A. Loss of heat .—Only a limited amount of data is available on the heat flux transferred to the surface from within the earth. These observations are summarized in Table 18–1 and show a mean value of 1.3 ± 0.1 × 10 −6 cal./cm. 2 /sec., when the necessary correction is made for the effect of the recent ice age. Considering the heat conductivity of the outer granitic layer as 0.006 cal./sec./cm.2/°C./cm., the corresponding surface geothermal gradient is about 2.0 × 10 −4 °C./cm., or 20°C./km. This heat is partially transferred to the atmosphere by convection and radiation with subsequent reradiation by the atmosphere into space. The remainder of the heat is radiated directly through the blanketing atmosphere. The proportion of heat loss attributable to each of these processes is still a matter of speculation. B. Sources of heat .—Aside from the residuum of heat associated with the earth at the time of its formation, the primary sources of heat in the earth are radioactivity, solar radiation, and exothermal chemical changes. 1. R adioactivity : Because . . .

Contents Page Table 19 1. Earth temperatures in the United States 281 2. Temperatures in foreign wells and mines 286 3. Temperature gradient, conductivity, and heat flow in South Africa 291 4. Heat flow in Great Britain 292 5. Maximum observed temperatures of lava at volcanic vents 292 The tables in this section are based on those of Van Orstrand [1]; data which have appeared since the preparation of Van Orstrand’s tables are inserted where possible [ 2 , 3 , 4 ], or are presented as separate tables [ 5 , 6 ], The observations of temperature in the mines, deep wells, and bore holes selected by Van Orstrand have been reduced by the method of least squares to equations of the form, T = a + b x where T is the temperature (°C.) at the depth x kilometers. For observations made in the United States, this reduction has been made for three ranges of depth: from 30.5 to 305 meters (100 to 1000 feet), from 30.5 meters to the deepest observation, and from 610 meters (2000 feet) to the deepest observation. These intervals will only accidentally bear any relation to the geological structure or stratigraphy. The constant a is the extrapolated surface soil temperature and is generally within a few degrees of the mean annual air temperature. The tabulated values of a are for the range, 30.5 to 305 meters; the values of a for the other ranges of depth, except in a few instances, are within less than 5 degrees of this value. The coefficient 6 is the mean . . .

Contents Page Table 20 1. Magnetic susceptibility of magnetite 295 2. Susceptibility of magnetic minerals 296 3. Range of magnetic susceptibility in major rock types 296 4. Susceptibility and remanent magnetization of rocks 297 Magnetic susceptibility of magnetite and of ferromagnetic minerals. —Magnetite, by reason of its relatively high magnetic susceptibility and wide distribution, is easily the most important ferromagnetic mineral. Indeed, the value of the magnetic susceptibility of a rock is generally determined solely by its contained magnetite. The other ferromagnetic minerals are few in number and of rarer occurrence. The most important ones are pyrrhotite, ilmenite, specularite, franklinite, and cubanite. In Table 20–1 are listed values of the magnetic susceptibility of magnetite. These values spread over a wide range and are dependent not only upon the specimen itself but also upon the intensity of the magnetizing field. In natural, large-scale occurrences of magnetite, such as ore bodies, the extremely high values of susceptibility found by Weiss [ 25 ] for single crystals and by Steinmetz [ 26 ] are rarely, if ever, observed. A value of about .3 is usually appropriate, but exceptions must be expected, since the susceptibility of magnetite in common with that of other ferromagnetic substances is sensitive to small changes in composition or structure. In Table 20–2 are listed measured susceptibilities for other ferromagnetic minerals. It will be noted that the susceptibility of magnetite is at least ten times that of other minerals. Magnetic susceptibility of rocks. —Published measurements of the magnetic susceptibility of rocks are . . .