Attention is called to the value to be derived from application of diffraction analysis of modern grade to specific, familiar, approximately known crystallizations. Features consequent to ordering of coordination polyhedra of differing sizes in layer structures are discussed, and their bearing upon relations between lattice parameters and chemical compositions are outlined.

The problems of hydrous decomposition of alkali feldspars and the formation of the micaceous minerals from the decomposition products are discussed in terms of the surface chemistry and the Al-Si ordering of the feldspars, and a crystal growth model for the micas.

The principal end members of the micas believed to be common in sediments have been synthesized and some of their stability relations determined. The polymorphs of muscovite and paragonite, the principal dioctahedral end members, obtained were 1Md, 1M and 2M1, and those of phlogopite and annite, the principal trioctahedral end members, 1Md and 1M or 3T. The range of stability of each of the polymorphs could not be fixed accurately because of the slow rate of transformation; however, the transformations 1Md → 1M → 2M1 were effected for muscovite and paragonite and 1M → 1Md or 3T and 2M1 → 1M or 3T for phlogopite. The growth characteristics of these micas in the laboratory are believed to be analogous to the formation of micas in sediments.

Knowledge of the synthetic mica.s contributes greatly to an understanding of the natural materials called illite, hydromica, and high-silica sericite. The dioctahedral members of these materials and related minerals may be delineated accurately in the system muscovite—Al-celadonite—pyrophyllite and their iron analogues. The trioctahedral members of some of the same materials may be outlined in the system phlogopite— eastonite—tale and their iron analogues. The postulated substitution schemes in these systems are mainly MgSi ⇌ AlVIAlIV, KAl ⇌ Si, and H3O ⇌ K. In materials intermediate between these systems, such as most biotites and vormiculites, the substitution of 3Mg ⇌ 2AlVI is of major importance. The mixed-layer structures involving micas are elucidated.

Sepiolite is a valid mineral with the following chemical formula: $8MgO\, \cdot \,12Si{O_2}\, \cdot \,n{H_2}O.$$

Its unit cell characteristics are: $a\, = \,13.4{\AA};\,b\, = \,26.8{\AA};\,c\, = \,5.28{\AA};Z\, = \,2;\,D_{{2^h}}^6\, - \,P\frac{{{2_1}}}{n}\,\frac{2}{c}\,\frac{2}{n}.$

They were determined on the sepiolite from Ampandrandava (Madagascar). Its structure was determined by means of the “trial and error” method and the observed and calculated intensities rendered a good agreement.

The formula of the structure is: $\mathop 1\limits_\infty {[{H_2}O]_n} \cdot \mathop 1\limits_\infty \{ {(Mg,F{e^{3 + }},F{e^{2 + }},\,Al)_8}{({H_2}O)_4}{(O,OH)_4}[{(Si,Al)_{10}}{O_{30}}]\} \cdot n\, \leqslant \,8.$

The mineral contains zeolitic, crystal- and constitution-water. A good agreement was achieved between the chemical, thermal, and morphological qualities and the x-ray data. The mineral remains stable up to about 350°C and begins to change into an “anhydride” above this temperature.

The formula of the structure of sepiolite. “anhydride” is: $\mathop 1\limits_\infty \{ {(Mg,F{e^{3 + }},F{e^{2 + }},\,Al)_8}{(O,\,OH)_4}[{(Si,\,Al)_{10}}{O_{30}}]\}.$

Its unit coll characteristics are: $a\, = \,10.9{\AA};\,b\, = \,23.3{\AA};\,c = 5.28{\AA};\,a \approx \,{90^ \circ };\,Z = \,2;\,C_{{2^h}}^5\, - \,P_n^{{2_1}}\,11.$

The minerals xylotile, mountain wood and gunnbjarnite are isotypical with sepiolite.

Muscovite and K-feldspar (adularia) were dry-ground to — 200/inch particles, which were suspended in water. These suspensions were “titrated” with KCl, and the pH was recorded as a function of KCl concentration and temperature. The results indicate that muscovite and adularia react with water to produce a surface film in which H+ has displaced K+. The “titration” curves show some characteristics attributable to exchange reactions, others apparently related to equilibria among solids of fixed composition. Maximum release of K+ from adularia by reaction with water is much greater than that from mica. The interpretation is made that the first result of reaction of mica and adularia with water is a surface layer that grades from an outer portion that is structurally disrupted to an inner portion that retains the original silicate structure but with H+ substituted for K+. Addition of K+ as KCl to the suspending solution displaces H+ from the disrupted zone, but all H+ originally taken up by the solids was not returned to the solution by concentrations of KCl up to 1.0 M. Experiments were of a few hours duration; work by others has shown that the disrupted zone releases appreciable concentrations of silica and alumina to solution over longer time intervals.

These hydrolysis experiments indicate that at 25°C an H-feldspar or H-mica structure is favored over a K-feldspar or K-mica structure except in solutions in which the ratio of aK+ aH+ exceeds 109–10 or 107–8 respectively. These ratios decrease in the temperature range 25-65°C by a factor of about 100.7. These results, where considered in relation to the observed behavior of feldspar and mica under weathering conditions, indicate that the major energy change for the reactions, $2\,mica + 5{H_2}O = \,3\,kaolin + \,2{K^ + }\, + \,2O{H^ - }$$3K - feldspar + 2{H_2}O = \,K - mica + 6quartz + \,2O{H^ - }$ can be considered to result from H+-K+ exchange, and that the energy contribution from other changes is small.

The patterns of distribution of the mica and clay minerals resulting from alteration of feldspars are much less regular and well defined in the “porphyry copper” deposits than in the vein deposits such as those at Butte, Montana. There are, however, interpretative problems common to both. Chemically, the destruction of the K-feldspars may be interpreted in the light of laboratory investigations of the system potash-alumina-silica-water at various K+/H+ ratios. The experiments indicate that the K-mica (sericite) stability field lies between the K-feldspar and kaolinite fields. The K-mica field is approximately 1.5 pH unit wide, and its boundaries with K-feldspar and kaolinite are parallel below 500°C. The boundaries slope toward higher K+/H+ ratios at lower temporatures. Thus, considering the. K-feldspar—sericite—kaolinite system of minerals, one might expect field patterns showing a central zone of kaolinite bordered by a zone of K-mica which in turn grades outward into unaltered K-feidspar. Such a pattern is demonstrated by Butte’s inner diekite-kaolinite and sericite and possibly by the zoning at San Manuel described by Schwartz. Alteration of plagioclase introduces additional complications into the distribution patterns of clay and mica minerals, the chief of which is the outer argillic zone so characteristic of both vein and disseminated deposits.

Tobermorite [Ca4 (Si6O18H2)Ca·4H2O] is a hydrated calcium silicate mineral with a layer structure which in some respects resembles that of vermiculite. Its dehydration has been studied using single crystals from Ballycraigy, N. Ireland. The three most frequently encountered hydration states are characterized by basal spacings (d002) of 14.0, 11.3 and 9.35Å. Dehydration to the 9.35Å state is complete by 300°C and is accompanied by a stacking change so that the pseudo-cell (a 5.58, b 3.66, c 18.70Å) becomes A-face centered. The 9.35Å structure persists up to 700°C, by which temperature all the water has been expelled, and there is some evidence that interlayer Si-O-Si bonds are formed to an increasing extent as the temperature rises.

At about 800°C, the 9.35Å hydrate changes to β-CaSiO3 twinned in two orientations. The b-axis of the 9.35Å hydrate becomes b for both orientations of the product, and the (201) planes of the latter are formed parallel to the (101̄) and (101) planes of the original material. The mechanism of the change is discussed and is compared with some other transformations occurring under similar conditions. An orientation-determining step is suggested in which the principal effect is a migration of silicon atoms or ions, the calcium—oxygen skeleton remaining relatively undisturbed.

Barshad’s method of calculating clay formation from the chemical analyses of the whole soil, the clay fraction, the nonelay fraction, and the mechanical analysis was applied to a large number of soils to determine the effect of climate, topography, parent material, vegetation and time on amount of clay formation. The amount of clay formed from 100 g of the nonclay fraction of the parent material was used as the yardstick for comparing clay formation in the different soils.

It was found that for a true evaluation of the effect of these factors on clay formation, it is necessary to determine clay formation in every horizon of a profile and that all soil profiles should be sectioned into horizons which are as nearly alike in thickness as possible. Maximum clay formation in a soil profile occurs in the horizon which is about 2–10 in. below the surface rather than at the immediate surface; as a rule, the amount of clay forming decreases below this maximum horizon. An increase in moisture and temperature enhances clay formation. The temperature factor mainly affects clay formation in the Great Soil Groups in the following order: Podzols < Gray Brown Podzols < Brown Earths < Chernozems < Prairies < Red and Yellow Earths (Latosols) < Laterites. Poor drainage enhances clay formation. Grass-type vegetation is more effective in clay formation than tree-type. Clay formation is greater from the finer grained and the more basic rocks. The more easily rocks fragment, the more clay is formed. The increase in age of a soil is believed to be expressed in the same manner as an increase in the intensity of clay formation, namely, as an increase in clay formation with depth and as an increase in the thickness of a soil profile. It was concluded that clay formation is relatively slow: clay forms at a rate ranging from 0.00001 g to 0.002 g of clay per year per 100 g parent material or from 0.6 to 120 lb of clay per acre foot of parent material per year.

The frequency distribution or relative abundance of minerals in soils varies with the five principal classes of factors that govern soil formation. The characteristics of the minerals of the parent material, the time factor, climatic factors, relief factors and biotic factors each can be shown to have important independent effects on clay mineralogy of soils under proper circumstances. The soil parent material exerts a control over the frequency distribution of minerals in soils by introduction of the clay minerals into the soil directly, by controlling the course of chemical weathering in the soil through the relative susceptibility of its minerals to weathering, by furnishing abundant divalent metallic cations, by impediment of drainage, or by acceleration of leaching when highly permeable. The time factor is conspicuous as long times give an advanced degree of weathering even in temperate climates. Climate is important, since highly weathered materials inevitably occur as a result of intense leaching in warm tropical and equatorial climates. Relief is important in concentrating leaching water and metallic cations, in affecting oxidation or reduction. The biotic factor affects minerals conspicuously where an A0 horizon develops and resulting cheluviation moves R2O3 out of the A2 horizon.

Inherited minerals such as illite, quartz, feldspars, ferro-magnesian minerals, carbonates and gypsum are most abundant in clays of little-weathered parent materials and soils of the zonal Desert, Brown, Chestnut and Tundra soils as well as intrazonal Mountain groups and azonal Regosols and Lithosols. Secondary layer silicate minerals such as vermiculite, secondary chlorite, montmorillonite, kaolinite and halloysite are most abundant in clays of moderately weathered parent materials and soils of the zonal Chernozem, Prairie, Gray—Brown Podzolic, Podzol, Red—Yellow Podzolic, and Low Humic Latosol groups as well as intrazonal Planosol, Rendzina, Dark Magnesium soil, and Wiesenboden groups. Secondary sesquioxide minerals such as hematite, goethite, allophane, gibbsite and anatase and residual resistant primary minerals such as ilmenite and magnetite predominate in the more highly weathered parent materials and soils of the zonal Ferruginous Humic Latosols, Hydrol Humic Latosols, Latosolic Brown, and Ando soils and Laterites, as well as the intrazonal Tropical Savannah and Ground Water Podzol ortstein soils.

Clay minerals undergo significant changes when exposed to weathering agents. Some of these changes were revealed by studying a sequence of samples collected from weathering profiles developed on shales. Illite and chlorite are modified by the development of expandable layers within the illite and chlorite structures. The suggested mechanism of alteration of chlorite is the oxidation and removal of octahedral iron. Chemical analyses indicate that interlayer potassium is removed from illites. Removal of potassium may be accomplished by an exchange reaction between potassium ions and hydronium ions in the environment.

The origin of underclays is controversial. In an attempt to shed some light on this subject, three shale-underelay profiles were studied. The changes in clay minerals from shales to underelays are very similar to those in weathered shale profiles. This similarity lends support to the theory that some underelays are products of weathering.

Studies of the Recent indicate that, at the most, somewhat less than half the clay minerals are altered to any extent in a marine environment. Probably most of this alteration is in the form of cation adsorption or reconstitution of slightly weathered illites and chlorites to their original form. There appears to be little if any evidence that the detrital basic lattice is being altered to a measurable extent in Recent marine sediments. In the near-shore environments there is usually a coincidence of clay mineral suites and environments.

A major change in the clay mineral composition of sedimentary rock occurs within the Mississippian. Illite is the dominant clay mineral of the pre-Upper Mississippian sediments. Post-Lower Mississippian clay suites are more variable in composition; illite becomes less abundant and montmorillonite and kaolinite more abundant. This change is best related to a change in regional tectonics. The clay minerals seem to have no preferred lithologie associations, although owing to epigenetie alterations porous sandstones commonly have different clay mineral suites from those of adjacent shales and carbonate rocks.

In many instances clay mineral facies coincide with environmental facies. As the clay mineral criteria for distinguishing any given type environment are extremely variable, it is thought that segregation of clay mineral suites by sorting is usually more effective than by diagenesis.

Expanded clay minerals appear to be partially contracted by the time they have been buried to 10,000–15,000 ft. It remains to be proved whether this is caused by chemical modification of the basic lattice with burial or if the detrital clay lattice has the inherent ability to contract, without chemical rearrangement, when buried to a sufficient depth.

The refractory shales of the Purgatoire and Dakota formations in central Colorado were studied to determine the genesis of the kaolinite. The compositions of the shales were similar, with variation only in relative amounts and grain size of the component minerais. The shales consist primarily of kaolinite, subangular to subrounded quartz, and hydro-muscovite, with minor amounts of feldspar, montrcorillonite, mixed-layer clays, and occasional traces of gypsum and zeolites.

An inverse relation between the kaolinite and hydromuscovite contents suggests that hydromuscovite was an intermediate mineral in the formation of the kaolonite. Derivation of vermicular kaolinite from feldspar and mica was evident from examination of thin sections.. The thinly banded interfingered lenses of montmorillonite suggest a detrital origin rather than alteration in situ of a volcanic ash to bentonite.

Genesis of the kaolinite probably is similar to that proposed recently by Kesler for the Cretaceous clays of Georgia and South Carolina whereby feldspathic sands deposited in a near-shore environment were subsequently weathered to produce residual kaolinite. Winnowing of some of these clays and redeposition in downstream lagoons formed the fine-grained, carbonaceous, “flint” clays, whereas the plastic clays were formed by introduction of detrital montmorillonite.

In a previous paper (van Olphen, 1956) data were presented for the rheological and sedimentation behavior of sodium bentonite suspensions in water. From these data particle interaction energies were derived. Analogous data have been collected for calcium bentonite suspensions. It was observed that more clay is required to obtain a gel or a certain sediment volume with calcium bentonite than with sodium bentonite. This result can be quantitatively interpreted on the basis of a larger average particle thickness in the calcium bentonite suspensions. It is furthermore concluded that the interparticle forces are of the same order in both calcium and sodium bentonite suspensions.

A viacomotrio study of dilute montnioritlonite dispersions, using lor tho most part dilution-typo capillary viscometers supplemented by a falling-head capillary viscometer, has involved a determination of the concentration dependence of viscosity of such dispersions in both distilled and electrolyte-containing water. The resulta of the study have been interpreted in terms of a theory of incremental viscosity, and of the Solmlz—Rlaschke (1941) equation. Use of this equation yields good extrapolations to zero concentration for the determination of limiting viscosity number, and permits calculation of “interaction indices” which are considered, for reasons discussed in detail, to be a measure of paitiele-particle interaction. These indices are related to the cationie heat of hydration of the electrolyte used, and relationship is demonstrated between the yield stress of a 3 percent montmorillonite—eleetrolyte—water system and such interaction indices. The recent work of Paekter (1956, 1957) is discussed, and the relation of the results of this investigation to the data of van Olphen (1956, p. 204) is noted.

The size and shape of clay particles in aqueous suspension have been determined by means of five experimental techniques—electro-optical birefringence, ultracentrifugation, viscometry, optical transmission and electron microscopy. Five clay minerals—montmorillonite, illite, hectorite, nontronite and attapulgite—each subdivided into five size fractions, were studied. The average particle dimensions in each fraction are tabulated.

An evaluation of the experimental techniques is presented. It is concluded that for a preliminary survey of the particle size and shape of clay minerals the technique of electron microscopy is best. However, once a model for the particles is obtained, a combination of the techniques of electro-optical birefringence and viscometry becomes more useful in characterizing the size and shape of the particles in suspension.

Electron microscope studies of chrysotile show that tubes are present in bulk specimens and that these tubes commonly have fuzzy, amorphous-looking material on both the inside and outside. Similar material is associated with synthetic chrysotile and has been noted previously in halloysite specimens. The existence of such material between and within the tubes, together with apparent irregularities in size, shape and packing of tubes, explains the apparent discrepancy between the measured density of bulk samples and the calculated density of a hypothetical sample consisting of close-packed, regular, hollow capillaries.

Replicas of fractured surfaces of halloysite (2H2O) from various localities reveal that the particles occur as curved to flat laths commonly possessing “hexagonal” terminations and surface features indicative of a higher degree of crystallinity than tubes of halloysite (4H2O).

It is suggested that a complete morphological series from plates through laths to tubes exists both in platy to fibrous serpentine and in kaolinite to halloysite (4H2O). In each series a number of structural varieties are to be expected between the morphologically distinct “end members.”

The distinctive response of clay minerals to water indicated that it should be possible to determine the water sensitivities of sedimentary rock samples directly by water- vapor adsorption measurements. Therefore, the adsorption of water-vapor by standard clay samples and by core samples from oil-producing formations has been investigated.

Water-vapor adsorption isotherms of the montmorillonite and kaolinite samples are characteristic of the clay type and serve as an aid in their identification. The illite isotherm was intermediate between those of the other two types.

Water-vapor adsorption by core samples from oil-producing formations varied with the contents of swelling clays and with water-sensitivities as indicated by previous x-ray analyses and permeability studies. Core samples containing montmorillonite clay adsorbed water vapor strongly and exhibited typical adsorption—desorption hysteresis. Correlations of nitrogen and water-vapor adsorption studies with clay analyses and permeability measurements show that the water-vapor adsorption method is a promising means for the direct measurement of the water sensitivities of sedimentary formations.

Detailed water-vapor sorption measurements were made on Li, ISTa, K, Cs and Mg kaolinite at 25°C. When desorption is started from P/P0 near saturation, all ionic modifications show reversible hysteresis. The size and P/P0 range of the hysteresis loop varies with the nature of the exchangeable ion and sample history.

No hysteresis occurs on Li and K kaolinite when desorption is started at P/P0 < 0.8 which along with other data indicates that Li and K exchangeable ions are not truly hydrated below P/P0 = 0.8. The hysteresis for all ionic modifications is traced to hydration of the exchangeable ions. The apparent anomalous position of Li and K ionic hydration is attributed to control of hydration of exchangeable ions by the difference between the hydration energy of the ion minus the specific adsorption energy of the ion for the clay. This difference must be positive before true ionic hydration occurs.

A molecular model for sorption is postulated in which sorption is assumed to take place at specific sorption sites upon the clay surface. The energy of sorption at different sites is believed to be different and also to be a function of the hydration state of the exchangeable ion. By the use of the proposed model and the detailed sorption data, hydration of the different exchangeable ions as a function of the relative vapor pressure can be obtained. The sorption model also predicts phase changes in the sorbed film at certain relative pressures. These calculated phase changes agree very well with phase changes indicated from Jura-Harkins’ plots of the experimental data.