The distribution coefficient KD = (Mg/ Fe)garnet/(Mg/Fe)biotite has been calculated for eight garnet-biotite pairs from rocks ranging from the garnet isograd to the staurolite zone along the Stonehaven coast of NE Scotland. This distribution coefficient depends upon metamorphic temperature and is not dependent upon the Mn and Ca content of garnet. Garnet-biotite geothermometer yields temperatures of about 500, 550, and 575–620°C for the garnet, chloritoid + staurolite, and staurolite zones respectively.

IT was not until 1966 that pseudorutile was first defined. Earlier, its X-ray diffraction spectrum had been confused with that of futile and, to a lesser degree, with those of hematite and ilmenite. Subsequent work has shown that pseudorutile has a world-wide distribution in detrital ilmenite-bearing heavy mineral deposits. The present work has confirmed its magnetic susceptibility and density. In addition pseudorutile is shown to be a magnetic spin glass with a peak susceptibility at 23 °K.

Altered ilmenites, in which pseudorutile occurs as a secondary alteration product, display a range of chemical composition and magnetic susceptibility. The most highly magnetic fractions are not necessarily those containing the least-altered ilmenite, and in material from Capel, Western Australia, the most highly magnetic fractions were those containing grains of ferrimagnetic ferrian ilmenite.

Quantitative X-ray diffraction has shown that West Australian altered ilmenite contains significant amounts of amorphous ilmenite, pseudorutile, and rutile. The magnetic susceptibility of paramagnetic fractions of altered ilmenite from Capel, Western Australia, can be calculated from normative compositions based on chemical analyses.

Camptonite sills of Permian age, intruded into Cambrosilurian sediments, are present in the northern part of the Oslo region, immediately to the north of Gran, Hadeland. They have been considered, along with associated dykes of camptonite and other minor intrusions of maenaite, to be related to the predominantly gabbroic rocks comprising the Oslo-essexite plugs (Brögger, 1894). Porphyritic varieties of camptonite are the most common, containing phenocrysts of concentrically zoned clinopyroxene and concentrically zoned brown amphibole; but, aphyric varieties with abundant brown amphibole also occur. A second amphibole (actinolite) forms thin, patchy, green overgrowths to some phenocryst or matrix brown amphiboles.

The pyroxenes range in composition from diopside to sahlite. They contain up to 0.7% Cr2O3 in the most Mg-rich zones and up to 0.6% Na2O in the most Fe-rich zones. TiO2 and Al2O3 are in the ranges 0.5–3.8% and 3.0–9.2% respectively. These compositions are similar to pyroxenes from other camptonites and alkali basaltic and alkali gabbroic rocks. ‘Normal’ zoned phenocrysts, in which an outer pinkish zone surrounds a colourless core in thin section, have more Fe, Ti, and Al, and less Mg and Si in the outer zone. In ‘reversed’ zoned crystals, the slightly pinkish core contains more Fe, Ti, and Al, and less Mg and Si. Oscillatory zoning, restricted to three concentric zones (pinkish-colourless-pinkish) show the same chemical changes with colour variation. In terms of end-member molecules, substitutions involving CaTiAl2O6, CaAl2SiO6 plus CaFeSi2O6 replacing CaMgSi2O6 are responsible for the zoning.

The majority of the zoned brown amphiboles are kaersutite following the criteria of Leake (1978). Those zones with Ti < 0.5 atoms per formula unit range from titanian pargasite and titanian ferroan pargasite to titanian ferro-pargasite. These compositions are similar to kaersutites crystallizing from other camptonites and as inclusions, or as megacrysts and phenocrysts in alkali basalts and derivative rocks. A simple concentric zoning with more Fe and less Mg in the outer zones occurs; but, an oscillatory zoning pattern, with an intermediate lighter brown zone (containing more Mg and less Fe) between two concentric darker zones, is more common. Other elements do not show consistent variations between zones, but Ti generally increases with increase in Aliv.

It is postulated that the earliest pyroxene zones crystallized from an alkali basalt magma at approximately 20 km depth. Crystallization of the remainder of the pyroxene phenocrysts and the kaersutite took place at high temperatures (> 940 °C) and at probably greater depths (7–15 km) than crystallization of nearby exposed gabbroic plugs (Oslo essexites). A mechanism of alternate crystallization of pyroxene and kaersutite in response to changes in PH2O, with some partial resorption of the non-crystallizing phase, is suggested to account for the observed petrographic features and oscillatory zoning in both minerals. The actinolite is a product of late-stage metasomatic activity.

Analyses are given of four pairs of coexisting pyroxenes; their Kp ranges from 1.282 to 1.369, suggesting equilibration temperatures as high as 1200 °C and pressures of the order of 8 to 9 kb.

The meteorite was an oriented polyhedron, which broke up into at least two parts after entry into the Earth's atmosphere. It has undergone some degree of terrestrial weathering. A cut surface of the meteorite shows light-grey interior with sporadic rusty-brown patches and a distinct brown alteration zone close to the fusion crust. Weathering has resulted in preferential replacement of NiFe by limonite, and veining of minerals by goethite.

Rewari is an equilibrated chondrite with rare ghosts of chondrules and at least one lithic fragment. Composition of olivine, as indicated by microprobe analysis is Fa23, which agrees well with bulk wet chemical analysis; that indicated by d130 is Fa18–20. From the outer surface inwards, four petrographic zones can be distinguished in the meteorite: a skin, about 0.01 mm thick, a troilite-poor zone slightly thicker than the skin, a troilite-rich ‘soaking zone’, about 0.5–0.6 mm thick, and a relatively coarse-grained interior. These are described in detail.

The interior of the meteorite is composed of relatively coarse-grained crystalline silicates with disseminated metallic minerals including plessitic and zoned inter-growths of kamacite and taenite. The matrix shows a high degree of integration with the chondrules. The coarse texture and zonation of taenite may be the result of protracted heat treatment responsible for recrystallization. The constituent grains show considerable shock effects such as fracturing, comminution, veins of shock-melted pseudotachylite, pressure twinning, and undulose extinction. Chemical composition (mean of two wet chemical analysis) of the meteorite is: metallic Fe 7.475, Ni 0.975, Co 0.045; as sulphide Fe 3.200, Ni 0.090, Co < 0.01; SiO2 38.060, TiO2 0.10, Al2O3 2.34, Fe2O3 0.175, Cr2O3 0.485, FeO 13.950, MnO 0.210, NiO trace, CaO 1.875, MgO 26.265, Na2O 0.89, K2O 0.115, P2O5 0.285, H2O− 0.295, H2O+ 0.81, CO2 trace, S (total) 1.890, C (total) 0.19 per cent. The chemistry, mineralogy, and texture show that the Rewari meteorite is an L6 chondrite. Compared to average L-group chondrite it has a higher content of MgO and lower of SiO2, a little lower oxidation state, and tends to be enriched in siderophilic elements.

The mineralogy and major-element geochemistry of thirty-seven samples from a borehole in the Upper Cretaceous Chalk (Santonian) of Kent has been determined. The mean carbonate content is 97% and the carbonate is a low-Mg calcite (0.42 mol % MgCO3). Most of the silica is deduced to be biogenic chert. The total clay content probably varies through the sequence due to changes in the rate of carbonate sedimentation. The clay composition remains relatively constant with smectite more abundant than illite. The smectite and K-feldspar could be formed from volcanic debris whereas illite and quartz are thought to be detrital.

CALC-GRANULITES at Garividi are intruded by microcline granites and pegmatites. At the contacts fluorapatite, fluorian allanite, ferroan sahlite, and epidote occur; rarely, the epidote is associated with a partly metamict allanite, an analysis of which is given, also X-ray data for the unheated and heated mineral.

Free energies of formation of liebigite, Ca2UO2 (Co3)3·10H2O, swartzite, CaMgUO2(Co3)3·12H2O, bayleyite, Mg2UO2(CO3)3·18H2O, and andersonite, Na2CaUO2(CO3)3·6H2O have been determined from solution studies at various temperatures. ΔGf(298.2K)° values for the above minerals are −6226±12, −6607±8, −7924±8 and −5651±24 KJ mol−1 respectively. ΔH0f(298.2K) values respectively are −7037±24, −7535±20, −9192±20, and −5916±36 kJ mol−1. These results have been used to construct the stability diagram for the four minerals shown in fig. I. Andersonite can only form when the activity of the sodium ion, aNa+, is relatively high and aCa2+ and aMg2+ are small. The interconversions of the sodium-free species are defined. When aMg2+/aCa2+ < 0.32 liebigite is the stable phase. If aMg2+/aCa2+ > 7.94, bayleyite forms preferentially. One might therefore expect the bayleyite-andersonite association to be more common than liebigite-andersonite, but this is entirely dependent upon the relative concentrations of Mg2+(aq) and Ca2+(aq) in the solutions from which the minerals form.

Discontinuities occurring in the expansion behaviour of aluminosilicate-sodalites with large cavity anions such as I− and are thought to occur when the coordinate of the cavity cation becomes 0.25 (Henderson and Taylor, 1978). We also suggested that further thermal expansion work could explore how the disposition of sodalites in a diagram such as fig. 4 of Henderson and Taylor relates to the presence or absence of discontinuities.

We have now studied the expansion of three more sodalites which were expected to show discontinuities, namely one aluminogermanate-sodalite (Na8(Al6Ge6O24)I2) and two aluminate-sodalites (Sr8(Al12O24)(CrO4)2 and Sr2Ba6(Al12 O24)(SO4)2). Sr8(Al12O24)(CrO4)2 was thought to have an ∼ 9 Å cubic cell at room temperature but the occurrence of broadened and split reflections in our sample suggests that this is not the case. Stepwise heating experiments on our sample showed the presence of a reversible, presumably displacive, transformation at 40±2 °C above which the X-ray reflections were sharp.

All three sodalites showed significantly smaller expansion rates than for aluminosilicate-sodalites having similar degrees of structural collapse. The lower mean expansion coefficient (0–500 °C) for Na8(Al6Ge6O24)I2 of 12.5 × 10−6C−1 compared with that for Na8(Al6Si6O24)I2 of 15.1 × 10−6C−1 is particularly significant as the only chemical difference between these two phases is substitution of Ge for Si. In addition none of the three sodalites showed the discontinuities expected and this suggests that the low expansion rates do not allow the cavity cation to reach a coordinate of 0.25 over the temperature range investigated.

Our earlier concept of the mechanism of expansion of the sodalite structure assumed that the expansion of the cavity cation-cavity anion bond forced the cavity cations against and between the framework oxygens so untwisting the partially collapsed structure (Henderson and Taylor, 1978). It now appears that this concept was over-simplified and that the expansion characteristics of sodalites depend on the nature of the tetrahedrally coordinated framework cations present as well as on the cavity cations and anions.

A number of diorite complexes occur within the Channel Islands region, notably on Jersey, Alderney, and particularly Guernsey. Much of northern Guernsey is made up of the largest of these complexes (fig. S1), the Bordeaux diorite. In the north-western part of this diorite, around Chouet, a complicated association of plutonic rocks occurs. Although the field relationships in this area are sometimes difficult to interpret—this is often the case in diorite complexes—three separate groups of rocks may be distinguished within the association: a diorite group; a granodiorite group; and an inhomogeneous suite of rocks (fig. S2).

The widespread diorite group consists predominantly of an even-grained diorite, which is relatively homogeneous but which occasionally grades into an acicular diorite, the latter often containing pods and veins of appinite. The granodiorite group is the least common, occurring as bodies which are interpreted as intrusive sheets and bosses within the even-grained diorite, but occurring as angular blocks within the inhomogeneous suite of rocks. The granodiorite invariably contains rounded diorite xenoliths. The inhomogeneous suite consists of a variety of rocks from patchy, dark diorite, through quartz diorite to tonalite. Commonly, these rock types are intimately associated, often showing gradational contacts with each other and frequently with the more basic portions occurring as ‘xenolithic’ material within the more acidic portions. At contacts between the inhomogeneous suite and the even-grained diorite certain features (e.g. lobate margins and pipe-like structures) indicate that the diorite must have been close to its solidus temperature at the time of emplacement of the inhomogeneous suite. The field relationships between the three groups are interpreted as indicating that the diorite group was emplaced first, followed by the granodiorite group, with both of these clearly pre-dating the inhomogeneous suite.

Fifty-nine specimens, chosen to give a representative sample of each of the three rock groups, have been analysed for major, minor, and a selection of trace elements and thirteen of these specimens have been analysed for REE. The chemistry of the analysed rocks confirms the division into three groups, with each group showing distinctive characteristics. Furthermore, chemical plots (e.g. Al2O3, P2O5, Cr, and Ni v. SiO2) show discontinuities and areas of overlap between each group which cannot be explained within the constraints of a single genetic model relating the three groups to each other. This argument is particularly strong for the relationship between the diorite group, which spans the range 50 to 59% SiO2, and the inhomogeneous suite, spanning the range 53 to 68% SiO2. In the area of overlap (53 to 59% SiO2) the two groups are geochemically different. Therefore, for a variety of reasons, including emplacement order, the geochemical characteristics of the groups and the lithological inhomogeneity which is associated only with the chemically intermediate members of the association (the inhomogeneous suite), three quite different and genetically unrelated liquids are required to generate the three groups of rocks.

The even-grained diorite shows chemical variation (e.g. with increasing SiO2, decreasing Al2O3, MgO, CaO, Sc, V, Cr, and Ni, and increasing Na2O, La, Nd, and Y) consistent with amphibole + plagioclase fractionation up to 55% SiO2. At 55% SiO2 several elements show a change of slope (e.g. FeO + Fe2O3, TiO2, Rb, Ba, and Zr) indicating the introduction of biotite as a fractionating phase. Increasing total REE content with increasing SiO2 throughout the even-grained diorite supports the contention that amphibole is an important fractionating phase. The higher TiO2, P2O5, Sr, La, Ce, Nd, and Y contents and negligible Cr and Ni contents of the acicular diorite suggest an origin by delayed crystallization of volatile-enriched portions of the diorite group magma.

The granodiorite group shows little geochemical variation. Members of this group contain detectable amounts of Cr and Ni, unlike virtually all members of the inhomogeneous suite. For this reason, and because of the field relationships, the granodiorite is considered to be genetically unrelated to members of the inhomogeneous suite and a separate liquid is thus required for its genesis. This liquid may have been the fractionated derivative of some other magma (though if this is so the ‘parent’ is entirely unrepresented at the present erosion level) or it may represent a direct crustal melt. Diorite xenoliths within the granodiorite are chemically similar to the even-grained diorite.

Despite the lithological complexity of the inhomogeneous suite, its geochemical unity is clearly established in that, for instance, virtually none of the members of the suite (including even the most SiO2-poor) contain detectable Cr and Ni. Moreover, geochemical variation within the group is rational (with the possible exceptions of Sr, Zr, and Ba) and may be explained in terms of a crystal fractionation model. However, the fractionation must have acted on a liquid itself unrelated to either the diorite or granodiorite group magmas. An additional complication is that later derivative liquids intrude into and partly digest earlier-formed semi-solids of the suite to produce much of the observed inhomogeneity. The phases which have controlled fractionation within the suite include plagioclase (established petrographically as well as geochemically) and hornblende. The role of apatite is uncertain. The fractionation of hornblende is particularly useful in explaining the change in REE contents within the inhomogeneous suite. Total REE contents increase from the dark diorite to the quartz diorite, but decrease from the quartz diorite to the tonalite with concomitant relative HREE depletion. This is taken to be a reflection of the changing hornblende/liquid partition coefficients for REE with increasing SiO2, which are less than one for liquids of basaltic and andesitic composition but greater than one for liquids of dacitic composition.

The crystallization history of four trondhjemite samples from the Scourian complex, NW Scotland, has been investigated using composite ilmenite-magnetite grains. A variety of compositions are present both as large- and small-scale exsolution lamellae, which can be used to unravel the complex cooling history of these rocks. The samples were collected near Upper Badcall, Sutherland, where intrusive trondhjemite sheets 1–2 m thick cut banded gabbro. The trondhjemites have a complex history that includes four stages: magmatic intrusions, granulite facies metamorphism, hydration and retrogression to amphibolite facies, and slow cooling with uplift.

Ilmenite-magnetite grains in samples HR. 49, 53, 86 display a complex exsolution pattern (fig. 1A). An original titanomagnetite exsolved into large-scale (up to 50 µm wide) ilmenite-magnetite lamellae from which have subsequently exsolved small-scale lamellae (c.4 µm wide) parallel to the earlier lamellae. The ilmenite-magnetite pairs form subhedral grains in a granoblastic aggregate of plagioclase and quartz. A little biotite overgrows some oxide grains. Ilmenite-magnetite grains in sample HR. 56 are composed of broad-zoned lamellae (fig. 1B); small-scale exsolution lamellae are absent. Silicate-grain boundaries are irregular and lower-temperature minerals (chlorite and carbonate) are more common.

The experimental results of Buddington and Lindsley (1964) allow the equilibration temperature and oxygen fugacity of coexisting ilmenite and magnetite to be determined from their chemical composition. Subsequent workers have shown that it is possible to determine liquidus temperatures and oxygen fugacity for volcanic rocks (Carmichael, 1967; Anderson, 1968a). Slowly cooled igneous and metamorphic rocks, however, have continued to equilibrate below their solidus and show a range of temperatures and oxygen-fugacity conditions (Anderson, 1968b; Duchesne, 1972; Oliver, 1978; Bowles, 1976, 1977).

This paper presents 42 new pairs of analyses made by electron-probe microanalysis, from 13 composite ilmenite-magnetite grains (Table I). Mole % ulvöspinel and R2O3 values have been calculated using the method of Carmichael (1967) and used to determine temperature and oxygen fugacity at equilibration, from the experimental data of Buddington and Lindsley (1964).

By using a scanning electron beam it is possible to obtain the average composition of a broad lamella that contains smaller exsolution lamellae in order to estimate its composition prior to exsolution. A −log10ƒo2ν.T °C plot of lamellae whose original composition has been determined in this way shows that they lie on a curve slightly above the Ni-NiO buffer between 1010 and 850 °C (fig. 2). Temperatures of the order of 1000 °C are probably magmatic temperatures since they are higher than is normally recorded for granulite-facies metamorphism; 850 °C is interpreted as the blocking temperature below which diffusion was unable to occur to form large-scale lamellae. A comparison may be made between this oxygen-fugacity curve and the curves determined by Carmichael (1967) (fig. 2) for acid lavas coexisting with different phenocryst phases. If an adjustment is made for the differences in bulk composition and pressure, some correspondence between the analysed points and the curve for hydrous silicates would be expected since hornblende is the earliest Fe-bearing silicate seen in the trondhjemite. However, correspondence is not found, implying that amphibole did not control the oxygen fugacity, either because it was not the main Fe-bearing phase at magmatic temperatures, or because the oxygen fugacity was externally controlled.

After the formation of broad high-temperature lamellae Ti diffusion continued on a smaller scale (2–3 µm) so that the lower-temperature history of these grains can be considered in terms of many independent microsystems. Limited diffusion continued across the boundaries of and within early magnetite and ilmenite lamellae. The compositions of small-scale exsolution lamellae in ilmenite and magnetite hosts have been determined. Lamellae of ilmenite in magnetite from different grains define separate log ƒo2-T curves for different grains. Lamellae of magnetite in ilmenite equilibrated at lower temperatures and oxygen fugacities. Individual microsystems have equilibrated at different temperatures and oxygen fugacities within the same grain and similar microsystems in different grains have equilibrated at different temperatures and oxygen fugacities, suggesting that the rock itself has become a series of independent closed systems.

The compositions of phases either side of early high-temperature lamellar boundaries have been measured. Analyses from different rocks yield different oxygen-fugacity curves in the same temperature range (765 to 610 °C). Higher temperatures were obtained for grains 49/3 and 86/4, which have exsolved into broad-zoned lamellae with no small-scale exsolution. The sense of the zoning is such that R2O3 in ilmenite decreases as it approaches magnetite. The equilibration temperature and ƒo, at the grain boundary increases from the centre of the grain to the edge.

In HR. 56 ilmenite magnetite grains show broad-zoned lamellae with no late small exsolution lamellae. The sense of zoning is such that magnetite grains increase in ulvöspinel content towards ilmenite and ilmenite decreases in R2O3 towards magnetite (fig. 1B). Even though the grains are in disequilibrium, it is assumed that equilibrium was at least established close to the boundary between lamellae. Equilibration temperatures thus obtained are between 410 and 430 °C at an ƒo2 between the Ni-NiO and QFM buffers. Ilmenite-magnetite grains coexist with a biotite richer in Ti and a hornblende depleted in Fe relative to those in samples yielding higher oxide temperatures suggesting that there was continuous Fe-Ti exchange between oxides and silicates as well as between ilmenite and magnetite.

In Lower Proterozoic rocks of the Olary Province of South Australia margarite occurs in conjunction with sodic muscovite, pseudomorphing chiastolitic andalusite in graphite-rich schist, and as a pseudomorphic phase in rocks consisting of coarse muscovite together with lesser amounts of chloritoid, sillimanite, staurolite, and corundum. In the latter occurrence the assemblage (muscovite-sillimanite - chloritoid ± staurolite ± corundum) replaces massive andalusite. The margarite and sodic muscovite within the pseudomorphs lack a preferred orientation which suggests that their development was associated with the waning stages of the Palaeozoic Delamerian Orogeny (cf. Glen et al., 1977) which was the last major tectonothermal event in the region.

Within the chiastolitic andalusite pseudo-morphs, equal proportions of margarite and sodic muscovite are separated from the rock matrix by a rim of coarser-grained muscovite, which has developed at the pseudomorph margins. A suggested reaction for the development of margarite is 10Al2SiO5+1.7Ca2++1.2K++Na++10H2O⇄(Ca1.6Na0.4)Al4(Si4Al4O20)(OH)3.6+(K1.2Na0.6Ca0.1)Al4(Si6Al20O20)(OH)4+6Al3++12.4OH−.

The excess aluminium and hydroxyl ions from the above reaction leave the pseudomorph system but probably react with free quartz and more K+ ions to produce the muscovite fringe. 6SiO2+2K++6Al3++12OH−⇄K2Al4(Si6Al2O20)(OH)4+8H+.

The above reactions are pertinent only for margarite and sodic muscovite produced by the pseudomorphism of chiastolitic andalusite.

Margarite and sodic muscovite in the muscovite- chloritoid - sillimanite ± staurolite ± corundum rocks replace coarser-grained (up to 1 mm) muscovite, sillimanite, corundum, and chloritoid; staurolite is unaffected. The coarse muscovite of the earlier pseudomorphic assemblage contains 7–16 mole % paragonite in solid solution whereas the sodic muscovite assocated with margarite contains 22–33 mole% paragonite in solid solution. Margarite contains negligible muscovite but substantial (17–26 mole%) paragonite in solid solution. The phases plagioclase and paragonite were not detected in any of the investigated samples. Staurolite has an Mg/(Mg+Fe) value of 0.21 and associated chloritoid a value of 0.20.

It is possible that much of the previously described sericitic alteration of andalusite (cf. D'arcy, 1977) within schists of the north-eastern Willyama Complex (north of Broken Hill) involves the production of margarite.