The electron spin resonance of some structural Fe3+ for montmorillonites having low Fe3+ content, is perturbed by electrostatic interaction between exchange cations and structural charge sites. The position of charge centers of organic and inorganic cations in the interlayer can thus be determined at various levels of solvation. Dielectric media between the silicate layers lower the electrostatic attraction between the silicate and the exchange cations. The silicate charge appears to be partially delocalized on structural oxygen atoms as shown by electron spin resonance and i.r. spectroscopy. There is also evidence that divalent exchange cations on dehydrated montmorillonites cause hydrolysis of water; the protons so produced migrate to structural charge sites.

Montmorillonite clay samples were coated with 16 m-equiv/g of clay or iron plus aluminum as hydrous oxides and aged 1 yr, in suspensions of pH 6 or 8. The magnesium exchange capacity (MgEC) decreased linearly with the amount of non-crystalline aluminum hydrous oxide associated with the clay. Eight to 16 m-equiv of iron per g of clay reduced the MgEC by 20 m-equiv/100g at pH 6, but did not affect the MgEC at pH 8. The quantity of non-crystalline aluminum associated with the clay depended on the suspension pH and aging time, and was unaffected by the coprecipitation of 8–16 m-equiv of iron hydrous oxide/g clay. The crystalline form of aluminum hydrous oxide depended on the suspension pH and was shown by X-ray diffraction to be gibbsite at pH 6 and bayerite at pH 8. Gibbsite and bayerite formed rapidly with a rate dependent on the suspension pH when excess non-crystalline aluminum hydrous oxides were present. The quantity of non-crystalline aluminum hydrous oxides remaining after one year in suspensions of iron hydrous oxides and montmorillonite varied from 2·3 m-equiv/g of montmorillonite at pH 8-4·0 m-equiv/g of montmorillonite at pH 6. Differential thermal analysis and MgEC measurements indicated some regular organization of the iron hydrous oxides, however, crystalline iron minerals were not detected by X-ray diffraction.

A closed form equation is derived for the calculation of the oriented diffraction pattern given by single particle size layer silicates having randomly interstratified interlamellar species. A general method for treating any particle size distribution is indicated and closed form results are presented for the Poisson, normal, gamma and binomial distributions. No restriction is placed on the number of interlayer types. The structure factors for these types are explicitly introduced. Graphs of two of the variables appearing in the equations applicable to particle size distributions provide a means of visualizing the effects of both interstratification and particle size on observed X-ray patterns.

Hexafluorotitanic acid (H2TiF6) selectively dissolves kaolinite and most other phyllosilicate minerals of soils and sediments, concentrating free crystalline (Ti,Fe)O2 minerals (partially substituted anatase and rutile) in the residue. A series of H2TiF6 reagents was standardized by analysis of the Ti content and by tests with pure anatase and commercial kaolinites. The Ti in the H2TiF6 solution selected (made from 49% HF + reagent TiO2) was 16·5% by weight as analyzed by the Tiron method. Treatment of pure anatase with the reagent H2TiF6 resulted in a 98% by weight recovery of TiO2 in the residue. The fraction of TiO2 recovered in the residue of commercial Georgia kaolinites was 88–101% after treatment with the selected H2TiF6 reagent. Isolates from nine Georgia kaolinite samples with varying amounts of TiO2 and Fe2O3 were examined by X-ray powder diffraction, scanning electron microscopy and elemental analysis. The main constituent of the (Ti,Fe)O2 isolates was anatase for all samples, with minor amounts of coarser rutile and mica from coarser kaolinite. The anatase and rutile isolates contained 74–93% (Ti,Fe)O2 with 0·5–3·1% Fe. The other constituents of the isolates were muscovite of mica (0·3–7%), quartz (0–9%) and amorphous relics of vermiculite and/or kaolinite (6–19%). Rutile, muscovite and quartz appear to be detrital but the anatase and relics are probably authigenic. Fine anatase appears to stick on the muscovite flakes as revealed by scanning electron microscopy and heavy liquid data for separation of these two minerals. The (Ti,Fe)O2 isolates from kaolinites which passed with the first magnetic concentrate of anatase were coarse, on the order of a few microns dia., as revealed by the scanning electron microscopy. Those passed with subsequent extensive magnetic concentrates from the same samples were finer. The anatase isolated from kaolinite purified by removal of as much of the impurities as possible by magnetic means was extremely fine, most of the particles being on the order of 0·1 µm dia. More than one third of the total Fe2O3 in kaolinites magnetically separated in the first pass was extracted by the citrate-bicarbonate-dithionite treatment after hot NaOH dissolution of 52–74% of the kaolinite, showing that the Fe2O3 had been mainly associated within the kaolinite. Only 2–6% of the total Fe2O3 was extracted from magnetically purified kaolinite after 40–50% of this kaolinite had been dissolved, indicating that most of the Fe is in the anatase and rutile fraction.

Adsorption isotherms for water vapor, c-spacing and heat of immersion in water of Na- and Ca-montmorillonite were measured at 25°C at various r.h. The amount of water adsorbed as a function of the r.h. increased gradually, whereas the c-spacing increased, and the heat of immersion (per mole of adsorbed water) decreased in steps. There was good agreement between the calorimetric data, the heat calculated from the isotherms by use of BET equation, and the calculations from the ion-dipole model. A model is presented to describe the hydration of sodium and calcium montmorillonite.

Halloysite (metahalloysite) of various particle sizes has been altered with oxalic and EDTA acids, at room temperature and during different periods of time (5–90 days). The oxalic acid attack at first achieved only a recrystallization of halloysite. The recrystallization is much more significant the smaller the size of the treated halloysite particles. Later the material is destroyed. The EDTA treatment also has provoked during the first days a recrystallization of the halloysitic material which is destroyed again after about 20–25 days. Later kaolinite is formed. The kinetic curve of kaolinite formation is symmetrical with respect to that corresponding to the diminution of amorphous material in the sample. The influence of the halloysite particle size and the complexing effect of the acids in relation to the resulting products are discussed.

An electron-microscopic and Mössbauer spectroscopic study of a range of kaolinites has revealed three distinct types of iron contamination within these minerals: a) Ferric ion may substitute for aluminium and be evenly distributed throughout the lattice, b) Ferric ion may be present as a crystalline coating of goethite, as indicated by lattice-imaging studies, or c) as an amorphous coating. The distribution of the iron in the groups b and c is non-uniform and is highest at the flake surfaces. Ferrous ion, when detected, is thought to be evenly distributed throughout the lattice. The size of the contaminating goethite crystallites and the observed Mössbauer spectra of these samples suggest that such particles are super-paramagnetic. All kaolinites can be cleaned by acid treatment except those having iron substituting for Al3+.

The mineral in monomineralic glauconite pellets is an iron-rich mixed-layer illite-smectite (here called glauconite), often composed almost entirely of illite layers. The nature of the interlayering is closely analagous to that of aluminous illite-smectite and varies with the proportions of the layer types: >30 per cent smectite, randomly interstratified; 15–30 per cent smectite, allevardite-like ordering; < 15 per cent smectite, ‘IMII’ ordering.

Glauconite is analagous to aluminous illite-smectite chemically as well as structurally. A good correlation has been found between the number of potassium atoms per O10(OH)2 in structural formulas calculated from the chemical analyses and the proportion of illite layers as determined by X-ray powder diffraction methods. This relationship indicates a remarkably systematic increase in the potassium content of the illite layers with an increasing proportion of illite layers. This feature and the existence of ordered interlayering at high proportions of illite layers can be explained by crystal-chemical effects of illite layers on neighboring smectite layers. Glauconite differs from aluminous illite-smectite in that glauconite contains significantly less potassium per illite layer than does aluminous illite-smectite with the same proportion of illite layers except near the pure illite composition. The strength with which the interlayer potassium is held and the ease of conversion of smectite to illite layers in glauconite may be attributed to its 1M structure and, perhaps, to its high octahedral iron content, which lead to stronger bonding of potassium by allowing a higher tilt angle of the O-H axis of hydroxyls adjacent to the potassium ion.

The apparent octahedral cation occupancy in excess of two-thirds of the octahedral positions in many glauconites appears largely attributable to the presence of significant amounts of interlayer hydroxy-iron, aluminum and magnesium complexes in the smectite layers.

The formation of hematite from amorphous Fe(III)hydroxide in aqueous systems at pH 6 and 70°C, both with and without oxalate, was followed by kinetic measurements, electron microscopy, i.r. spectroscopy and thermal analysis.

In the absence of oxalate, small amorphous particles coalesce into aggregates which eventually become single crystals of hematite. When oxalate is present, crystal growth is much faster and does not proceed through the intermediate stage of aggregation. Aggregates, when formed, consist of groups of single crystals. It is suggested that oxalate accelerates the nucleation of hematite crystals by acting as a template, the Fe-Fe distance in Fe-oxalate ions being similar to that in hematite.

Dissimilar fabrics, observed by scan electron micrography (SEM) on freshly broken surfaces of kaolins from weathered, water-laid material vs hydrothermally altered material in Mexico are illustrated, contrasted, and tentatively related to their respective environments of genesis. Coarse, open-textured, well-preserved books and vermiforms of kaolinite characterize the kaolin from water-laid material at Jacal, Mexico. Hydrothermal kaolins tend to have more tightly compacted crystals of either kaolinite flakes (individuals or in packets) or elongates (halloysite morphology) or mixtures of both in the same deposit. Variations in porosity of parent rock whether sedimentary or igneous, and dissimilar permeation by meteoric water and by vapor and liquids of hydrothermal fluids are deemed to be significant in the processes producing distinctive fabrics. The question is raised as to the fundamental distinction between kaolinite and halloysite and their nomenclature.

The Si/Al ratio of an hydrothermal system plays an important role in kaolinite synthesis. If the atomic Si/Al ratio of a system is greater than 2-0, kaolinite will disappear at 345 + 5°C and 2 kbars water pressure according to the reaction kaolinite + 2 quartz → pyrophyllite + H2O. If the atomic Si/Al ratio is less than 2-0, however, kaolinite will persist until 405°C where it will react according to the equation 2 kaolinite → pyrophyllite + 2 boehmite + 2 H2O. The Si/Al ratio of the system and temperature are also factors in determining whether b-axis ordered or disordered kaolinite will crystallize. The ordered variety is favored by a lower Si/Al ratio and a higher temperature than is the disordered form.

Hydrothermal experiments also show that kaolinite can be synthesized at 150°C and 5 bars pressure in distilled water from amorphous starting materials. Previous investigators were unsuccessful in forming kaolinite under these conditions because their systems were contaminated with alkalis.

Attempts to synthesize halloysite and dickite failed, but halloysite was converted to kaolinite at 150°C, suggesting that halloysite can be synthesized only at low temperatures.

The distribution of clay minerals in Recent sediments can be explained by the relative stability of the clays. The rates of particle aggregation for three clays were determined in the laboratory in synthetic estuarine solutions; from the kinetic studies stability values were calculated. The results indicate that illite is more stable than kaolinite which is more stable than montmorillonite. The distribution of the clays in the Pamlico River Estuary can be explained on the basis of relative clay stability where kaolinite which aggregates rapidly (relatively unstable clay) is found upstream of illite.

Accelerated soil erosion from construction sites and the resulting increase in downstream sediment load constitute a significant environmental problem. Laboratory studies indicate that small percentages of hydrated lime or of Portland cement will stabilize clay soils against rainstorm erosion by preventing particle detachment. Coordinated measurements of the size distribution of water-stable aggregates, of pore size distribution by mercury porosimetry and of microstructure by scanning electron microscopy and energy dispersive X-ray spectrometry were used to clarify aspects of the mechanisms responsible for the development of erosion resistance. Attainment of such resistance was marked by aggregation of a significant part of the clay into water stable aggregations of the order of several mm in size and of minimal change in porosity and pore size distribution on exposure to the test rainstorms. At least some of the clay particles in the aggregations appeared to be partly converted to calcium-bearing reaction products and formation of the “reticulated network” variety of calcium silicate hydrate gel linking adjacent particles was demonstrated.

An attempt has been made to synthesize nitrogenous clay minerals hydrothermally from silica-alumina gels in the presence of amino acids, namely glycine and lysine. The products have been characterized by X-ray powder diffraction, by analyses for C and N contents, and by their infrared spectra.

Amino acid-montmorillonites have been prepared under hydrothermal conditions of 200–250°C and 1000 atm. Above 250°C the amino acids were degraded to ammonium ions, and ammonium-micas were obtained. Syntheses without the addition of amino acids to gels yielded kaolinite.

The role of organic compounds in the formation of clay minerals seems to be of considerable geochemical significance.

The exchange of [Co(NH3)6]3+ from its vermiculite complex has been carried out with a number of inorganic and alkyl quaternary ammonium ions of varying ionic sizes. The distribution and selectivity coefficients of the desorbing ions increase in the order: Li<Na<NH4<K<Rb<H<Cs for the monovalent, Ca<Mg for the bivalent and (C2H5)4N<(CH3)4N≪CTA (cetyl trimethyl ammonium) <CP (cetyl pyridinium) for the organic ions. The ΔG° values of some of the cation exchange reactions have been calculated by an equation of Kielland (1935).

Three sepiolite clays studied showed evidence for the presence of structural hydroxyl groups in three to five different environments depending on the composition of the clay. A 3720 cm−1 i.r. frequency is shown to be characteristic of SiOH at crystal edges which are very abundant in sepiolites. This band has not been seen by most workers because the Nujol, fluorolube or KBr used in sample preparation perturb it sufficiently to obscure it under other OH stretching bands. The 3680 cm−1 band is confirmed as being from the (Mg)3OH and evidence of a very small band near 3640 cm−1 is suggested to arise from limited trioctahedral substitution. The very crystalline Ampandrandava sepiolite shows only the above three bands. The intermediately crystalline Vallecas shows a 3620 cm−1 band in addition which is characteristic of dioctahedral systems and is due to either some vacancy sites or to the presence of attapulgite. This dioctahedral band is greater in the less crystalline Salinelles sepiolite; in addition, it has a smaller 3585 cm−1 band. Mg-Al-vacancy and Mg-Fe‴-vacancy are suggested as the source of the 3620 cm−1 and 3585 cm−1 bands.

Electron microscope microprobe analysis (EMMA) has been applied to the determination of the elemental compositions of the kaolinite particles (Mg, Al, Si, K, Ti and Fe) contained in the 0·2-0·3 µm e.s.d., the 0·9-1·0 µm e.s.d. and the 1·9-2·0 µm e.s.d. fractions of an English kaolin and an American kaolin. Particles with masses as small as 10-13 g were analysed. An EMMA-4 instrument (A.E.I. Ltd.) equipped with linear fully focussing spectrometers was used. The ratio method of analysis was employed. The operating procedures used to obtain the required high experimental precision in the measurement of Al:Si atom ratio are discussed.

Statistical analysis of the results gives the estimated mean and spread of the Al:Si atom ratios. In the English kaolin the mean Al:Si atom ratio differs from the ideal 1:1 at the 0·05 significance level. There is evidence for a variation in composition from kaolinite particle to kaolinite particle in the 1·9-2·0 µm fraction of each kaolin. In the 0·9-1·0 µm fractions, the mean Fe:Si atom ratio was close to 0·002 showing the presence of iron in the kaolinite structure. The mean K:Si ratio was about 0·002 which would be equivalent to 1 unit muscovite layer associated with a 0·175 µm thick kaolinite particle. In the American clay the Ti:Si atom ratio was 0·002 suggesting that some 12 per cent of the 'titania' found by conventional chemical analysis was associated with the kaolinite particles either as titania itself or as an isomorphous substituent.

Quartz isolated from soils by the pyrosulfate-H2SiF6 method and chert samples of various origins were examined with the scanning electron microscope. Quartz isolates of the 20–50μm from the A2 and B21t horizons of the Baxter soil (AR), with quartz δ18O of 18·2 and 19·0‰, respectively, showed a mixture of detrital quartz particles and chert clusters of aggregates of fine euhedral quartz crystals. The 2–5 μm fractions of both horizons consisted mainly of euhedral quartz particles. The 20–50 μm fraction from the underlying chert, with a δ18O = 29·6‰, consisted of aggregates of euhedral quartz particles 1–10 μm dia and of subhedral particles 0·1–0·5 μm dia. In the soil fractions, the size and shape of the quartz particles as well as oxygen isotope data indicated that the aggregates were from the underlying chert but that irregular, unaggregated grains were detritally admixed loess, particularly in the medium and coarse silt fractions. This mixing of chert (of low temperature origin and heavy oxygen) with detrital quartz (of high temperature origin and light oxygen) gave rise to the intermediate δ18O values in the soil quartz. The SEM of cherts of different geological ages showed different morphologies. Prismatic, polyhedral microcrystalline quartz of 1–10 μm size as well as submicron, euhedral particles were observed in cavities. Submicron, subhedral particles and interlocking quartz grains were characteristic of Precambrian chert. Quartz grains more than 100 μm in size isolated by HCl from Ordovician dolomite (WI) had large (many microns) subhedral overgrowths and attached clusters of 0·1–0·5 μm microcrystalline quartz. A Danish flint formed in chalk had calcite-lined cavities (x-ray emission determined) in which spheroidal fibrous chalcedony occurred.

Fibrous sepiolite crystals derive much of their commercial value from their molecular size channels and grooves. The crystals fold upon drying and these channels and grooves are lost. A model for the folding and unfolding of the crystals is presented. Extensive i.r., X-ray and thermogravimetric evidence shows that folding occurs when approximately half of the water of hydration, which is coordinated to the edge magnesium atoms inside of the channels, is removed. This occurs near 175°C under vacuum and near 300°C in air. When the crystals fold, all remaining water molecules enter a new environment, that of the hexagonal holes of the neighboring silica surface. A true anhydride is produced at about 500°C under vacuum when the final water is lost, but this final dehydration produces no important structural change. Rehydration of the anhydride to the normal hydrated sepiolite does not occur at room temperatures in 100% r.h. However, above, 60°C rehydration does occur.