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Settling velocities of particulate systems, 1. Settling velocities of individual spherical particles

F. Concha E.R. Almendra 《International Journal of Mineral Processing》1979,5(4):349-367

Combining boundary-layer theory and experimental data for the pressure distribution and boundary-layer thickness over the surface of a sphere, the following expression was obtained for the drag coefficient:

C_{D} = 1+ 9.06 R_{e}^{12}^{2}

Using this formula as a basis, an expression was developed for relating the settling velocity of spherical particles to their diameter at any value of the Reynolds Number. The validity of such expressions discussed and compared with experimental data. 相似文献

Dissociation constants for alkali earth and sodium borate ion pairs from 10 to 50°C

E.J. Reardon 《Chemical Geology》1976,18(4):309-325

Potentiometric measurements in dilute sodium borate solutions with added alkali earth chlordie salts yield the following expressions for the dissociation constants of alkali earth borate ion pairs from 10 to 50°C:

p K(MgH_{2} BO_{3}^{+} =1.266+0.001204 T

p K(CaH_{2} BO_{3}^{+} =1.154+0.002170 T

p K(SrH_{2} BO_{3}^{+} =1.033+0.001738 T

p K(BaH_{2} BO_{3}^{+} =1.942+0.001850 T

where T is in °K. Enthalpies for the dissociation reactions at 25°C are less than 1 kcal./mole for all the alkali earth borate ion pairs.Values for pK(NaH₂BO₃°) from 5 to 55°C computed from the experimental data of Owen and King are in good agreement with those determined potentiometrically. The average value from both methods is 0.22 ± 0.1 at 25°C.Application to seawater of computed pK's for MgH₂BO₃⁺, CaH₂BO₃⁺ and NaH₂BO₃⁰ yields an apparent dissociation constant for boric acid of 8.73 vs. 8.70 measured by Lyman, 8.68 by Buch and 8.73 by Byrne and Kester. 相似文献

Thermodynamic stability studies of the basic copper carbonate mineral,malachite

James L Symes Dana R Kester 《Geochimica et cosmochimica acta》1984,48(11):2219-2229

Natural malachite is a well defined solid demonstrating reproducible solubility behavior over a wide range of pH. The following equilibrium constants associated with the malachite dissolution equilibrium at 25°C, 1 atm were determined:

K_{sp} = a^{2}_{cu^{2+}} a_{CO_{3}^{2?}} K^{2}_{w} a^{2}_{H+} = 3.5 ± 0.6 × 10^{?34}

(infinite dilution)

K^{1}_{sp} = [Cu^{2+}]^{2} [CO^{2?}_{3}]K^{2}_{w} a^{2}_{H+} = 10. ± 0.2 × 10^{?32}

(0.72 ionic strength)

K′_{sp} = m^{2}_{Cu^{2+}} m_{COsu2?}_{3} K^{2}_{w} a^{2}_{H+} = 1.3 ± 0.1 × 10^{?28}

(36.9‰ salinity seawater). The temperature dependence of a “mixed” equilibrium constant, K_sp⁺, of the form:

K^{2}_{sp} = [Cu^{2+}]^{2} m_{CO^{2?}_{3}} K^{2}_{w} a^{2}^{H+}

has been measured at I = 0.72, yielding the relationship:

log K^{2}_{sp} = (? 9.8 ± 0.03) × 10^{4} (1 T°K) + (1.52 ± 0.09)

within a 5–25°C temperature range. The effect of pressure on the solubility of malachite in water and seawater was estimated from partial molar volume and compressibility data. For 25 °C at infinite dilution

K'_{sp} (1000 bar) K'_{sp} (0) = 240

and in seawater

K′_{sp} (1000) K'_{sp} (0) = 44

.Comparison of stoichiometric and apparent malachite equilibrium constants has been used to estimate the extent of copper(II) ion interaction at the ionic strength of seawater. In dilute carbonate medium (total alkalinity, TA = 2.4 meq/kg H₂O, pH 8.3), 2.9% of total dissolved copper exists as the free copper(II) ion and in seawater (S = 36.9%., TA = 2.3 meq/kg H₂O, pH = 8.1),

[Cu^{2+}] T(Cu)

is 3.1%.Total dissolved copper levels of approximately 450–750 nMol/Kg are necessary to attain malachite saturation conditions in the open ocean. Observations of malachite particles suspended in seawater must be explained by precipitation or solid phase substitution reactions from localized environments rather than by direct precipitation from bulk seawater. 相似文献

Sodium and potassium coprecipitation in aragonite

Art F White 《Geochimica et cosmochimica acta》1977,41(5):613-625

The coprecipitation of Na and K was experimentally investigated in aragonite. The distribution functions were determined at pH 6.8 and 8.8 over aqueous Na and K concentrations of between 5 × 10^?4and 2.0 M and temperatures of between 25 and 75°C.The mole fractions of Na and K in aragonite are related to the aqueous ratios of Na and Ca by a function of the form

log X_{Na}_{2} CO_{3},K_{2} CO_{3} = C_{0} + C_{1} log a^{2}_{Na} ? ,K^{?} a_{Ca}^{2+}

where C₀ and C₁ are constants at a given temperature. This equation was derived by a statistical model assuming a heterogeneous energy distribution for the sites of incorporation. The independence of the coprecipitation process from aqueous anion activities suggests that carbonate is the only anionic species in the solid solution. 相似文献

Stability of ethylxanthate ion in neutral and weakly acidic media. Part 1: Influence of pH

M. Maillot J.-L. Cecile R. Bloise 《International Journal of Mineral Processing》1984,13(3):193-210

Xanthates are used in the flotation of sulfide ores although their aqueous solutions are not stable under certain conditions. Their stability in acidic and weakly acidic aqueous solutions was therefore investigated, as these media are required for some processes.The peak absorbances of ethylxanthate ion and carbon disulfide were first determined in aqueous solution. The decomposition of ethylxanthate ion was analyzed by measuring variations in absorbance (at 301 nm) and pH with respect to time. A pH regulation system was then used while measuring variations in absorbance and productions of protons caused by xanthate decomposition.The results concerning xanthate half-lives show good agreement with the literature, but the kinetic results deviate substantially. The following relation was obtained for half-life:

T_{12} =9.67×10^{?6} (pH)^{11};4?7;T_{12} in seconds

We established that ethylxanthate decomposition at pH 4 is a first order reaction with respect to ethylxanthate concentration, and postulating this order to the other pH values, the following kinetic relation was found:

v= ?(1.22×10^{4} [H^{+}]?1.36×10^{?2})([EtX^{?}]_{∞}) (4?pH?7)

where v is the rate of decomposition (mol l^?1 min^?1), and [EtX^?]_∞ is the ethylxanthate concentration when the decomposition equilibria are reached (mol l^?1). The better concentration was found to obey the law:

[EtX^{?}]_{∞} =3.142×10^{?5} pH ? 1.255 × 10^{?4} (4?pH?6)

相似文献

The effect of pressure on the solubility of minerals in water and seawater

Frank J Millero 《Geochimica et cosmochimica acta》1982,46(1):11-22

The effect of presure on the solubility of minerals in water and seawater can be estimated from In

(K^{P}_{sp} K^{0}_{sp}) + (?ΔVP + 0.5ΔKP^{2}) RT

where the volume (ΔV) and compressibility (ΔK) changes at atmospheric pressure (P = 0) are given by

ΔV = V ? (M^{+}, X^{?}) ? V ? [MX(s)]ΔK = K ? (M^{+}, X^{?}) ? K ? [MX(s)]

Values of the partial molal volume (

V ?

) and compressibilty (

K ?

) in water and seawater have been tabulated for some ions from 0 to 50°C. The compressibility change is quite large (~10 × 10^?3 cm³ bar^?1 mol^?1) for the solubility of most minerals. This large compressibility change accounts for the large differences observed between values of ΔV obtained from linear plots of In K_sp versus P and molal volume data (Macdonald and North, 1974; North, 1974). Calculated values of

K^{P}_{sp} K^{o}_{sp}

for the solubility of CaCO₃, SrSO₄ and CaF₂ in water were found to be in good agreement with direct measurements (Macdonald and North, 1974). Similar calculations for the solubility of minerals in seawater are also in good agreement with direct measurements (Ingle, 1975) providing that the surface of the solid phase is not appreciably altered. 相似文献

The effect of ionic interactions on the oxidation of metals in natural waters

Frank J Millero 《Geochimica et cosmochimica acta》1985,49(2):547-553

The effect of ionic interactions of the major components of natural waters on the oxidation of Cu(I) and Fe(II) has been examined. The various ion pairs of these metals have been shown to have different rates of oxidation. For Fe(II), the chloride and sulfate ion pairs are not easily oxidized. The measured decrease in the rate constant at a fixed pH in chloride and sulfate solutions agrees very well with the values predicted. The effect of pH (6 to 8) on the oxidation of Fe(II) in water and seawater have been shown to follow the rate equation

-d in [Fe(II)]/dt = k_{1} β_{1} α_{Fe} /[H^{+}] + k_{2} β_{2} α_{Fe} /[H^{+}]^{2}

where k₁ and k₂ are the pseudo first order rate constants, β₁ and β₂ are the hydrolysis constants for Fe(OH)⁺ and Fe(OH)⁰. The value of α_FE is the fraction of free Fe²⁺. The value of k₁ (2.0 ±0.5 min^?1) in water and seawater are similar within experimental error. The value of k₂ (1.2 × 10⁵ min^?1) in seawater is 28% of its value in water in reasonable agreement with predictions using an ion pairing model.For the oxidation of Cu(I) a rate equation of the form

?d ln [Cu(I)]/dt = k_{0} α_{Cu} + k_{1} β_{1} α_{Cu} [Cl]

was found where k₀ (14.1 sec^?1) and k₁ (3.9 sec^?1) are the pseudo first order rate constants for the oxidation of Cu⁺ and CuCl⁰, β₁ is the formation constant for CuCl⁰ and α_Cu is the fraction of free Cu⁺. Thus, unlike the results for Fe(II), Cu(I) chloride complexes have measurable rates of oxidation. 相似文献

Prediction of Gibbs energies of formation of compounds from the elements—II. Monovalent and divalent metal silicates

Yves Tardy Robert M. Garrels 《Geochimica et cosmochimica acta》1977,41(1):87-92

A parameter ΔO^2?, defined as the difference between the Gibbs energy of formation of a given oxide and its aqueous cation, was used to obtain linear relationships among Gibbs energies of formation from the elements of hydroxides, oxides and aqueous metallic ions (Tardy and Garrels, 1976). Use of this parameter has now been extended to meta- and orthosilicates for which the Gibbs energies of formation of silicates from their oxides are shown to be linear functions of the ΔO^2? values of their constituent cations. The function obtained for metasilicates is:

ΔG^{o}_{?} silicate ? ∑ΔG^{o}_{?} oxides = ? 23 (ΔO^{2?} cation ? ΔO^{2?} silicon

and that for orthosilicates is:

ΔG^{o}_{?} silicate ? ∑ΔG^{o}_{?} oxides = ? 44 (ΔO^{2?} cation ? ΔO^{2?} silicon

in which Δ^o_? silicate is the Gibbs energy of formation from the elements of a silicate of a given cation and ∑ΔG^o_? oxides is the sum of the Gibbs energies of formation from the elements of the constituent oxides of the silicate considered.These functions can be used to test for consistency within and between various sources of thermodynamic data and to estimate free energy of formation values for previously unstudied species. 相似文献

An experimental study of alkali metal distributions in feldspars and micas

A.E. Beswick 《Geochimica et cosmochimica acta》1973,37(2):183-208

K and Rb distributions between aqueous alkali chloride vapour phase (0.7 molar) and coexisting phlogopites and sanidines have been investigated in the range 500 to 800°C at 2000 kg/cm² total pressure.Complete solid solution of RbMg₃AlSi₃O₁₀(OH)₂ in KMg₃AlSi₃O₁₀(OH)₂ exists at and above 700°C. At 500°C a possible miscibility gap between approximately 0.2 and 0.6 mole fraction of the Rb end-member is indicated.Only limited solid solution of Rb AlSi₃O₈ in KAlSi₃O₈ has been found at all temperatures investigated.Distribution coefficients, expressed as

K_{d} = (Rb/K)

in solid/(Rb/K) in vapour, are appreciably temperature-dependent but at each temperature are independent of composition for low Rb end-member mole fractions in the solids. The determined

K_{D}

values and their approximate Rb end-member mole fraction (

X_{RM}

) ranges of constancy are summarized as follows: (°C)

T

K_{D}^{Phlog/Vap.}

X_{RM}

K_{D}^{Sandi/Vap.}

X_{rm}

相似文献

10.

Oxygen and sulfur isotope equilibria in the BaSO4HSO4−H2O system from 110 to 350°C and applications

Minoru Kusakabe Brian W. Robinson 《Geochimica et cosmochimica acta》1977,41(8):1033-1040

Oxygen isotope exchange between BaSO₄ and H₂O from 110 to 350°C was studied using 1 m H₂SO₄-1 m NaCl and 1 m NaCl solutions to recrystallize the barite. The slow exchange rate (only 7% exchange after 1 yr at 110°C and 91% exchange after 22 days at 350°C in 1 m NaCl solution) prompted the use of the partial equilibrium technique. However, runs at 300 and 350°C were checked by complete exchange experiments. The temperature calibration curve for the isotope exchange is calculated giving most weight to the high temperature runs where the partial equilibrium technique can be tested. Oxygen isotope fractionation factors (α) in 1 m NaCl solution (110–350°C), assuming a value of 1.0407 for α_{CO₂H₂O} at 25°C, are:

10^{3} 1 n α_{BaSO_{4}?1 mNaCl} = 2.64 (10^{6} T^{2}) ? 5.3 ± 0.3

.These data, when corrected for ion hydration effects in solution (Truesdell, 1974), give the fractionation factors in pure water:

10^{3} 1 n α_{BaSO_{4}H_{2}O} = 3.01 (10^{6} /T^{2}) ?7.3 ± 0.1

.In the 1 m H₂SO₄-1 m NaCl runs, sulfur isotope fractionation between HSO^?₄ and BaSO₄ is less than the detection limit of 0.4%. A barite-sulfide geothermometer is obtained by combining HSO^?₄H₂S and sulfide-H₂S calibration data.Barite in the Derbyshire ore field, U.K., appears to have precipitated in isotopic equilibrium with water and sulfur in the ore fluid at temperatures less than 150°C. At the Tui Mine, New Zealand, the barite-water geothermometer indicates temperatures of late stage mineralization in the range 100–200°C. A temperature of 350 ± 20°C is obtained from the barite-pyrite geothermometer at the Yauricocha copper deposit, Peru, and oxygen isotope analyses of the barite are consistent with a magmatic origin for the ore fluids. 相似文献

11.

The kinetics of silica-water reactions

J.D. Rimstidt H.L. Barnes 《Geochimica et cosmochimica acta》1980,44(11):1683-1699

A differential rate equation for silica-water reactions from 0–300°C has been derived based on stoichiometry and activities of the reactants in the reaction SiO₂(s) + 2H₂O(l) = H₄SiO₄(aq)

(?a_{H_{4}SiO_{4}} ?t)_{P.T.M.} = (A M)(γ_{H_{4}SiO_{4}})(k+a_{SiO_{2}} a^{2}_{H_{2}O} ? k_a_{H_{4}SiO_{4}})

where (

A M

) = (the relative interfacial area between the solid and aqueous phases/the relative mass of water in the system), and k₊ and k_? are the rate constants for, respectively, dissolution and precipitation. The rate constant for precipitation of all silica phases is

log k_{?} = ? 0.707 ? 2598 T (T, K)

and E_act for this reaction is 49.8 kJ mol^?1. Corresponding equilibrium constants for this reaction with quartz, cristobalite, or amorphous silica were expressed as

log K = a + bT + c T

. Using

K = k_{+} k_{?}

, k was expressed as

log k_{+} = a + bT + c T

and a corresponding activation energy calculated:

(°C)T	$K_{D}^{Phlog/Vap.}$	$X_{RM}$	$K_{D}^{Sanid/Vap.}$	$X_{RM}$
500	$0.64 ± 0.11$	0–0.2	$0.17 ± 0.04$	0–0.07
700	$1.11 ± 0.11$	0–0.2	$0.33 ± 0.04$	0–0.1
800	$1.28 ± 0.03$	0–0.2	$0.45 ± 0.06$	0–0.1

相似文献

12.

Rates of reaction of pyrite and marcasite with ferric iron at pH 2

C.L Wiersma J.D Rimstidt 《Geochimica et cosmochimica acta》1984,48(1):85-92

The relative reactivities of pulverized samples (100–200 mesh) of 3 marcasite and 7 pyrite specimens from various sources were determined at 25°C and pH 2.0 in ferric chloride solutions with initial ferric iron concentrations of 10^?3 molal. The rate of the reaction:

FeS_{2} + 14 Fe^{3+} + 8 H_{2} O = 15 Fe^{2+} + 2 SO^{2?}_{4} + 16 H^{+}

was determined by calculating the rate of reduction of aqueous ferric ion from measured oxidation-reduction potentials. The reaction follows the rate law:

?dm_{Fe^{3+}} dt = k(A M)m_{Fe^{3+}}

where m_Fe³⁺ is the molal concentration of uncomplexed ferric iron, k is the rate constant and

A M

is the surface area of reacting solid to mass of solution ratio. The measured rate constants, k, range from 1.0 × 10^?4 to 2.7 × 10^?4 sec^?1 ± 5%, with lower-temperature/early diagenetic pyrite having the smallest rate constants, marcasite intermediate, and pyrite of higher-temperature hydrothermal and metamorphic origin having the greatest rate constants. Geologically, these small relative differences between the rate constants are not significant, so the fundamental reactivities of marcasite and pyrite are not appreciably different.The activation energy of the reaction for a hydrothermal pyrite in the temperature interval of 25 to 50°C is 92 kJ mol^?1. This relatively high activation energy indicates that a surface reaction controls the rate over this temperature range. The BET-measured specific surface area for lower-temperature/early diagenetic pyrite is an order of magnitude greater than that for pyrite of higher-temperature origin. Consequently, since the lower-temperature types have a much greater

A M

ratio, they appear to be more reactive per unit mass than the higher temperature types. 相似文献

13.

Complexing of silica by iron(III) in natural waters

E.J. Reardon 《Chemical Geology》1979,25(4):339-345

Determination of amorphous silica solubility in acidified ferric nitrate solutions confirms the presence of ferric silicate complexing. A dissociation constant for the reaction:

FeH_{3} SiO_{4}^{2+} → Fe^{3+} + H_{3} SiO_{4}^{?}

of 10^?9.8 ± 0.3 pK units at room temperature (22 ± 3°C) is obtained, in close agreement with reported values at 25°C corrected to zero ionic strength of 10^?9.9 by Weber and Stumm and 10^?9.5 by Olson and O'Melia. Iron-silicate complexing may be of significance to the mobilization of silica in acid waters associated with oxidizing sulphide deposits and coal strip mining and the precipitation of secondary silicate mineral phases. 相似文献

14.

Extension of the empirical Alyavdin equation for representing batch grinding data

L.G Austin K Shoji V.K Bhatia F.F Aplan 《International Journal of Mineral Processing》1974,1(2):107-123

The Alyavdin equation for batch grinding data is:

1 ? P(χ, t) = [1 ? P(χ, 0)] exp ? c(x)t^{p}]

where P(χ,t) is the weight fraction less than size χ after grinding time t, c (χ) is constant with t and p is a constant close to one. It is shown that this equation is illogical (except for a single size of feed) unless c (χ) varies with P(χ,0), which makes the equation of little utility. A new empirical equation is developed for finite size intervals:

1 ? P(χ_{i+1}, t) = exp ? tK_{i}^{1γ} + ln 1 1 ? P(χ_{i+1},0)^{1γ_{i}}^{γ_{i}}

which reduces to the Alyavdin equation for a single size of feed, and which gives consistent computations for any feed size distribution. Techniques are given for determining K_i, γ values from sets of batch grinding data. The values are then used to predict size distributions for other times and other feed size distributions. The equation was quite successful in predicting size distributions in batch milling: (a) providing the feed size distribution was not un-natural, that is, not truncated or (b) if a truncated feed was used, the values of K_i and γ are determined from size distributions of grinding of the same type of feed. Thus, K_i, γ are not, unfortunately, completely independent of the starting feed size distribution. 相似文献

15.

Phase relations in the system NaCl-KCl-H₂O. Part I: Differential thermal analysis of the NaCl-KCl liquidas at 1 atmosphere and 500, 1000, 1500, and 2000 bars

I-Ming Chou 《Geochimica et cosmochimica acta》1982,46(10):1957-1962

A simple differential thermal analysis (DTA) technique has been developed to study phase relations of various chemical systems at elevated pressures and temperatures. The DTA system has been calibrated against known melting temperatures in the system NaCl-KCl. Isobaric sections of the liquidus in the system NaCl-KCl have been determined at pressures of 1 atmosphere and 500, 1000, 1500, and 2000 bars. Using the least-squares method, the following equation was used to fit the experimental data:

T(°C)= ∑ i=0 6 a_{i} X^{i}_{KCl}

where T is the liquidus temperature, X_KCl is mole fraction of KCl, and a_i (listed below) are the derived empirical constants.

	a	b	c	E_act(kJ mol ^-1)
Quarts	1.174	-2.028 x 10³	-4158	67.4–76.6
α-Cristobalite	-0.739	0	-3586	68.7
β-Cristobalite	-0.936	0	-3392	65.0
Amorphous silica	-0.369	-7.890 x 10^-4	3438	60.9–64.9

相似文献

16.

The surface chemistry of δMnO2 in major ion sea water

L.S Balistrieri J.W Murray 《Geochimica et cosmochimica acta》1982,46(6):1041-1052

相似文献

17.

Aqueous oxidation-reduction kinetics associated with coupled electron-cation transfer from iron-containing silicates at 25°C

Art F. White Andy Yee 《Geochimica et cosmochimica acta》1985,49(5):1263-1275

Mechanisms and kinetics of aqueous

Fe^{+2} Fe^{+3}

oxidation-reduction and dissolved O₂ interaction in the presence of augite, biotite and hornblende were studied in oxic and anoxic solutions at pH 1–9 at 25°C. Oxidation of surface iron on the minerals coincided with both surface release of Fe⁺² and by reduction of Fe⁺³ in solution. Reaction with iron silicates consumed dissolved oxygen at a rate that increased with decreasing pH. Both Fe⁺³ and O₂ consumption were shown to be controlled by coupled electron-cation transfer reactions of the form;

[Fe^{+2}, 1 zM^{+z}]_{silicate} + Fe^{+3} → [Fe^{+3}]_{silicate} + Fe^{+2} + 1 zM^{+z}

and

[Fe^{+2}, 1 zM^{+z}]_{silicate} + H^{+} + 14 O_{2} → [Fe^{+3}]_{silicate} + 1 zM^{+z} + 12 H_{2} O

where M is a cation of charge +z. The spontaneous reduction of aqueous Fe⁺³in the presence of precipitated Fe(OH)₃bracketed the surface oxidation standard half cell between +0.33 and +0.52 volts. Concurrent hydrolysis reactions involving cation release from the iron silicates were suppressed by the above reactions. Calculated oxidation depths in the minerals varied between 12 and 80Å and were apparently controlled by rates of solid-state cation diffusion. 相似文献

18.

Studies in diffusion—III. Oxygen in feldspars: an ion microprobe determination

B.J Giletti M.P Semet R.A Yund 《Geochimica et cosmochimica acta》1978,42(1):45-57

Self-diffusion of oxygen in adularia, anorthite, albite, oligoclase and labradorite has been measured by isotope exchange of oxygen between natural feldspars and hydrothermal water enriched in ¹⁸O. The analysis consisted of measuring the ¹⁸O/¹⁶O gradient inward from the feldspar surface using an ion microprobe, and fitting a solution of the diffusion equation to the data. Depth of the sputtered hole was measured with an optical interferometer. Linear Arrhenius plots were obtained:

P (bars)	$a_{o}$	$a_{1}$	$a_{2}$	$a_{3}$	$a_{4}$	$a_{5}$	$a_{6}$
1 atm.	800.1	?334.2	781.6	?6490.3	17553.1	?17638.4	6098.3
500	813.5	?354.9	743.3	?6011.7	16406.4	?16516.3	5702.8
1000	824.5	?406.7	1446.8	?8818.4	21253.5	?20343.7	6839.4
1500	838.6	?418.7	1434.7	?8819.0	21557.9	?20908.4	7123.1
2000	848.5	?381.5	1246.9	?8605.0	21785.8	?21449.1	7375.8

相似文献

19.

Diffusion of ions in sea water and in deep-sea sediments 总被引：3，自引：0，他引：3

Sandra Gregory 《Geochimica et cosmochimica acta》1974,38(5):703-714

The tracer-diffusion coefficient of ions in water, D_j⁰, and in sea water,

D_{j}^{1}

, differ by no more than zero to 8 per cent. When sea water diffuses into a dilute solution of water, in order to maintain the electro-neutrality, the average diffusion coefficients of major cations become greater but of major anions smaller than their respective

D_{j}^{1}

or D_j⁰ values. The tracer diffusion coefficients of ions in deep-sea sediments, D_j,sed., can be related to

D_{j}^{1}

D_{j,sed.} = D_{j}^{1} · α θ^{2}

, where θ is the tortuosity of the bulk sediment and a a constant close to one. 相似文献

20.

Isotopic anomalies of noble gases in meteorites and their origins—VI. Presolar components in the Murchison C2 chondrite

Leo Alaerts Roy S. Lewis Jun-ichi Matsuda Edward Anders 《Geochimica et cosmochimica acta》1980,44(2):189-209

相似文献

	$d_{0} (cm^{2} /sec)$	$Q (kcal/g-atom O)$	$T(°C)$
Adularia (Or₉₈)	$4.51 × 10^{?8}$	25.6	350–700
Albite (Ab₉₇, Ab₉₉)	$2.31 × 10^{?9}$	21.3	350–800
Anorthite (An₉₆)	$1.39 × 10^{?7}$	26.2	350–800