Alkali feldspars in plutonic igneous rocks vary both in their exsolution textures and in the structural state of their components. The primary factor leading to this diversity is the availability of hydrothermal fluids during their cooling history. In many plutons feldspar variation is related to degree of fractionation as indicated by rock chemistry. The variation reflects build-up of water with magmatic evolution and implies that fluids did not circulate freely in the intrusives in the temperature range of unmixing and ordering. Experimental work bearing on the role of fluids in subsolidus changes in feldspars is reviewed, and points of overlap between crystallographic studies of feldspars and stable isotope studies are discussed.

The nature of the upper mantle below the ancient cratonic areas can be deduced by study of the xenolith suites in kimberlites. Studies on the proportions of xenoliths, together with their mineralogy and chemistry, suggest an upper mantle containing an upper harzburgite zone and a lower lherzolite zone, with both of these zones containing chemical and mineralogical variants, together with minor rock types such as glimmerites, MARID-suite rocks, pyroxenites, and eclogites. Isotopic studies of the phases in xenoliths have yielded restricted ranges of values for the isotopic composition of hydrogen, carbon, oxygen, and sulphur which are tentatively identified as the true isotopic values for these elements in the upper mantle. In addition, recent discoveries suggest that diamond may be a primary uppermantle phase.

The textures and fabrics of the xenoliths indicate that plastic deformation has taken place in connection with the intrusive kimberlite event, and also in earlier events unconnected with the kimberlite event. In addition, brittle fracture has been observed, this fracturing often being accompanied by the filling of the ensuing veins and joints by fluids that have crystallized potassium-, titanium-, and water-rich phases; limited metasomatism of peridotite wall rocks accompanies this vein filling and more widespread pervasive metasomatism may also be present.

Although most upper-mantle rocks are now metamorphic, in some rare instances there are relics of earlier rock types that have not been completely obliterated by subsequent metamorphic events; most of these could be attributed to an igneous origin and in most cases during subsequent metamorphism the original rock types have been subjected to increasing pressures and/or lower temperatures.

Although the source of most materials now present at the earth's surface can be directly attributed to some identifiable source within upper-mantle rocks, no source has yet been identified for CO2, N, P, and the rare gases.

Metamorphic rocks of undoubted crustal origin have been described in recent years, principally from Mediterranean collision zones that have been subjected to PT conditions along very low geothermal gradients (∼ 7°C/km) and have reached pressures up to 30 kbar. MgAl-rich metapelites develop particularly diagnostic high-pressure minerals and mineral assemblages that have been and are being studied experimentally in model systems involving the components K2O, MgO, Al2O3, TiO2, SiO2, P2O5, and H2O up to pressures of 50 kbar and temperatures of 1000°C.

In the present review the following synthetic phases and phase assemblages are discussed, emphasizing their water-pressure-temperature stability fields (approximated in parentheses here), their reaction relationships, and their known or potential occurrences in metamorphic rocks. Sudoite (0 to ∼ 12 kbar, 150? to 380°C) occurs in very low-grade metapelites. Mg-carpholite (∼ 7 to ∼ 45 kbar, ∼ 200 to 600°C) is found in subducted metabauxites, metapelites, and related quartz veins. Mg-chloritoid (18 to 45 kbar?; 400 to 760°C) has not been found in nature as pure or nearly pure end-member; it requires silica-deficient environments. Yoderite, known in nature only from a single talc-kyanite schist occurrence, has only a small stability field (9 to 18 kbar?, 700 to 870°C?), cannot coexist with quartz, but may be stabilized by Fe3+. Pyrope (∼ 15 to at least 50 kbar, ∼ 700°C to melting), with or without relic coesite inclusions, occurs spectacularly in quartzites. Mg-staurolite (∼ 14 to some 90 kbar?, 700 to 1000°C), recently discovered as inclusions in pyrope, requires silica-deficiency. MgMgAl-pumpellyite is a new synthetic phase in which Mg totally replaces Ca of normal pumpellyite; because of its very high-pressure, low-temperature stability (∼ 37 to at least 55 kbar, < 400 to 780°C) it may not form within our globe. Ellenbergerite, the new high-pressure mineral forming inclusions in pyrope, apparently exhibits a rather composition-dependent stability with Ti-ellenbergerite, requiring higher pressures (> 20 kbar) than P-bearing, Ti-free members; a pure hydrous Mg-phosphate with ellenbergerite structure was synthesized at 10 kbar. Phengites, the widespread MgSi-substituted muscovites, require increasingly high water pressures (up to ∼ 20 kbar) for higher degrees of substitution, but the Al-celadonite end-member is not stable under any conditions; the compositions of phengites coexisting with limiting assemblages such as phlogopite, K-feldspar, and an SiO2 phase are useful geobarometers. The common assemblage Mg-chlorite + Al2SiO5 (mainly kyanite) has an extensive stability field ranging from near zero to 31 kbar at temperatures varying from some 320 to ∼ 760°C depending on pressure. The whiteschist assemblage talc + kyanite (6 to ∼ 45 kbar, 550 to 810°C) plays an important role in collision zone metamorphism as it forms from the greenschist assemblage chlorite + quartz at low grades but is also known to break down into pyrope + coesite at the highest grade observed thus far. The assemblage talc-phengite (11 to at least 35 kbar, 300? to 820°C depending on pressure), on the other hand, is well known from subducted metapelites. At pressures of 15–20 kbar and temperatures of 400–650°C a very K,Mg-rich, siliceous fluid forms as a consequence of the mutual reaction of the minerals K-feldspar and phlogopite (biotite) which are very common in crustal rocks including granites. Such fluids are bound to cause metasomatism in neighbouring mantle rocks which, upon subsequent increase of temperature, produce post-collisional ultrapotassic, lamproitic melts.

Pertinent general properties of the planets are listed. The condensation of the solar nebula is set in the context of stellar evolution and meteorites with sections on astronomical observations, chemical composition of the solar nebula, physical properties of the solar nebula, and chemical and physical aspects of condensation, accretion, and planetary differentiation. A cool nebula is preferred to allow survival of pre-solar grains with isotopic anomalies. Equilibrium progressive condensation of the solar nebula is regarded as a useful theoretical boundary, but complex processes involving crystal-liquid differentiation in, and collisions between, planetesimals are used to interpret the properties of meteorites and terrestrial planets. Chemical differentiation in the nebula begins with condensation and aggregation of dust, which can yield oxidized and reduced products depending whether C/O is less or greater than unity. Simple models for direct accretion of condensed materials into planets are reviewed but not adopted. Physical interactions involving small bodies include collisional accretion of dust-covered bodies, and differentiation of silicate and metal from mechanical, magnetic, and electrostatic forces. Physical and chemical differentiation involving large bodies involves head-on and glancing collision of planetesimals, orbital deflection, and disintegration within the Roche limit, and collision with debris rings and moons. Planetary accretion: dynamics, time scale, and heat sources involves more rapid growth of a larger body than a smaller one with ultimate development of one planet in each feeding zone, which flares out and ultimately overlaps with adjacent zones. Mars is small, and a planet did not develop in the asteroid belt, because of perturbations from Jupiter. The giant planets deflected material into the inner solar system. Melting of early planetesimals is invoked to explain differentiated meteorites. Chemical differentiation inside planet esimals and planets describes the phase equilibria for metal, sulphide, and peridotite, either dry, wet, or containing CO2. A wet body could begin crystal-liquid differentiation near 1250 K with sinking of Fe,S-rich liquid and rising of basaltic melt. The peridotitic residuum might undergo a subsequent differentiation at higher temperature under volatile-free conditions. Mineralogical storage of H2O, CO2, S, Cl, F, and alkalies is discussed. Chemical differentiation in planetary atmospheres briefly mentions escape of light species.

For the Earth, the early history is constrained by Archaean rocks dating from −3.8 × 109 yr whose properties indicate a non-reducing atmosphere, and a mantle that yielded volcanic rocks mostly similar to recent ones. The upper mantle (above 200 km depth) contains peridotitic rocks attributable to crystal-liquid differentiation and metamorphism. Volatile elements exist in mica and other minerals, but are sparse. Abundances of siderophile and chalcophile elements are high enough to require late accretion of material rich in these elements, the presence of a barrier between upper mantle and core, and some extraction by sinking sulphide. The mantle (deeper than 200 km) and core are inaccessible to direct study but interpretation of seismic data coupled with high-pressure laboratory studies requires inversion to dense phases in the mantle (especially perovskite?) and presence of light elements in the core (mainly S?). The bulk composition is modelled by cosmochemical analogy constrained by geophysical and geochemical parameters. The early condensate may be augmented by 1.5 ± 0.5? over (Mg+Si), and metal by 1.2±0.1?, while alkalies are probably depleted six-fold. Radial heterogeneity from a reduced interior to an oxidized exterior is suggested. For the Moon, sections cover observations, petrologic interpretations, and bulk chemical composition. For the origin of the Earth and Moon, age constraints, chemical constraints, and dynamical and accretional constraints allow comparison of suggested origins, with the conclusion that the Moon formed by either fission or disintegrative capture during early growth of the Earth, followed by simultaneous accretion coupled with disintegrative capture of planetesimals.

Mercury must be Fe-rich, but the silicate may not be just early condensate. For Venus, reviews are given of the surface properties, atmosphere, speculations on bulk composition and speculations on surface mineralogy and atmospheric compositions. The CO2 is in the atmosphere; most H may have been lost with concomitant oxidation of rocks; K/U ratios suggest basaltic and granitic rocks, and the high-surface temperature (c.740 K) implies granulitic metamorphism. For Mars, reviews are given of surface morphology, atmosphere and volatiles, mineralogy and petrology, and geophysical and geochemical models. Prolonged emission of Fe-rich lavas is suggested. Volatiles were removed by mineralogical processes from the atmosphere to give ice caps and sediments affected by aeolian processes and oxidation from photochemically generated H2O2. Fe-rich layer silicates, maghemite, and Mg-sul-phate may dominate the sediments.

The composition of comets and interplanetary dust may be inferrable from micrometeorites whose complex properties are suggestive of carbonaceous meteorites. Asteroids should be supplying at least many of the meteorites to Earth. A general description and review of remote-sensing studies culminate in a review of spatial descriptions and implications. The main belt is dominated by dark C-type asteroids assumed to be the primordial inhabitants produced by primary condensation and local accretion. The inner zone contains some brighter S-type asteroids interpreted as having undergone separation of metal from silicate as in stony-iron meteorites, as well as some E- and M-types perhaps matching enstatite-bearing meteorites and irons. Perhaps these differentiated(?) asteroids, as well as basaltic Vesta, are strays perturbed outwards from the inner solar system.

A review of meteorites covers carbonaceous meteorites, ordinary chondrites, enstatite chondrites and achondrites, reduced irons, forsterite-bearing meteorites and silicate inclusions in irons, irons, pallasites, chassignites and nakhlites, ureilites and lodranite, eucrites and shergottites, diogenites, howardites and mesosiderites, oxygen isotopes, and original location of meteorites. Emphasis is placed on mineralogical properties demonstrating crystal-liquid differentiation, brecciation, agglomeration, and even aqueous alteration and vapour transfer, though some evidence remains of direct condensation of gas to solid. The meteorites demonstrate the existence of many parent bodies (probably over 60), and most macrometeorites are ascribed to collision debris from the main-belt asteroids that has been deflected past Mars. The range of oxygen isotopes is explained by mixing of supernova debris rich in 16O with gas-solid differentiates of the nebula.

The culmination of the review is a suggested synthesis, which emphasizes a new model of heterogeneous accretion from planetesimals whose diverse compositions range from reduced material near Mercury to oxidized material from the asteroid zone outwards. Growth of a terrestrial planet begins from near-by slow planetesimals, and ends with distant fast planetesimals. Only the Earth and Venus are big enough to retain debris from late volatile-rich planetesimals deflected to high speed by the giant planets. Mercury, Moon, and Mars are volatilepoor. The Earth is zoned from a reduced interior, composed of crystal-liquid segregates of planetesimals captured early, to an oxidized exterior containing some material captured from outside the orbit of Mars. A mantle barrier hindered chemical equilibration. Hydrogen loss augmented oxidation. The mantle may be chemically zoned inwards from olivine-rich composition to pyroxene-rich composition, and the core should be reduced and contain substantial S and perhaps C and P. Mercury and Venus should have accreted reduced material, and only Venus should contain volatiles obtained from late planetesimals from outside the Martian orbit. Mars should have accreted mainly S-type asteroids, and have captured little volatile-rich material.

The history of the chemical analysis of minerals is shortly reviewed. Following rapid developments between 1770 and 1810, few important changes took place for close on 130 years; new procedures led to simpler and more convenient separations but not, on the whole, to greater accuracy. The older analyses are, in general, neither worse nor better than modern ones, and analyses can only be assessed on internal evidence, if available, or on the reputation of the analyst.

The development of microchemical and colorimetric techniques in the last 40 years has resulted in a great reduction in the amount of material required for a mineral analysis; 5 to 20 mg is usually sufficient, and in favourable cases a complete analysis may be possible on less than I mg. Some modern instrumental methods are considered, and it is concluded that they cannot wholly replace the classical wet-chemical methods in their modern developments.

W. F. Hillebrand (1894) deplored the tendency of some petrologists to demand large numbers of very incomplete rock analyses, and his remarks are considered apposite today.

From G. F. Becker's and L. V. Pirsson's early enunciations linking the dynamics of magma chambers to the rock records of sills and plutons to this day, two features stand at the centre of nearly every magmatic process: solidification fronts and phenocrysts. The structure and behaviour of the envisioned solidification front, however, has been mostly that akin to non-silicate, non-multiply-saturated systems, which has led to confusion in appreciating its role in magmatic evolution. The common habit of intruding magmas to carry significant amounts of phenocrysts, which can lead to efficient fractionation, layering, and interstitial melt flow within extensive mush piles, when coupled with solidification fronts, allows a broad understanding of the processes leading to the rock records of sills and lava lakes. These same processes are fundamental to understanding all magmas.

The spatial manifestation of the liquidus and solidus is the Solidification Front (SF); all magmas, stationary or in transit, are encased by SFs. In the ideal case of an initially crystal-free, cooling magma, crystallinity increases from nucleation on the leading liquidus edge to a holocrystalline rock at the trailing solidus. The package of SF isotherms advances inward, thickening with time and, depending on location — roof, floor, or walls — and the initial crystallinity of the magma, is instrumental in controlling magmatic evolution. Bimodal volcanism as well as much of the structure of the oceanic crust may arise from the behaviour of SFs.

In mafic magmas, somewhere near a crystallinity (N) of 55% (vol), depending on the phase assemblage, the SF changes from a viscous fluid (suspension (0<N<25) and mush (25<N<55%)) to an elastic crystalline network (rigid crust (55<N<100%)) of some strength containing interstitial residual melt. With thickening of the roofward SF of some mafic magmas, the weight of the leading, viscous portion repeatedly tears the crust near N ∼ 55–60%, efficiently segregating the local residual melt into zones of interdigitating silicic lenses. This is SF instability (SFI), a process of possible importance in continental crust initiation and evolution, in producing silicic segregations in oceanic crust, and in recording the inability of the viscous part of the upper SF ever to detach wholly in typical (<∼ 1 km) sheet-like magmas. These granophyric and pegmatitic segregations, individually reaching 1–2 m in thickness and 30–50 m in length, form thick (∼ 50–75 m) zones that can be misconstrued as sandwich horizons where the last liquids might have accumulated. In effectively splitting the magma chemically and spatially, SFI is, in essence, a form of chaos (i.e. silicic chaos).

Differentiation of initially crystal-free, stationary magmas is limited to processes occurring within SFs, which operate in competition with the rate of inward advancement of solidification. Local processes operating on characteristic time scales longer than the time for the SF to advance a distance equal to its own thickness are suppressed. Enormous increases in viscosity outward within the viscous, leading portion of the SF efficiently partition the distribution of melt accessible to eruption. Eruptible melts lie essentially inward of the SF and are thus severely restricted in silica enrichment. The silica-enriched SFI melts are thus generally inaccessible to collection and eviction unless the host SF is reprocessed or “burned back” through, respectively, later regional magmatism or massive, late-stage re-injection. And because of large viscosity contrasts between SFI melts and host basalts, once freed, SFI melts are literally impossible to homogenize back into the system and may collect and compact against the roof to form large silicic masses. Unusually voluminous, bulbous masses of silicic granophyre present along, and sometimes warping, the roofs of large diabase sills may reflect collections of remobilized blobs of SFI melts. These bulbous masses may be later added to the continental crust through solid state creep.

In sheets made of phenocryst-rich, singly saturated magma, most phenocrysts are able through settling or floating to avoid capture by the advancing SFs. Significant differentiation is possible through extensive settling of initial phenocrysts and upward leakage of interstitial residual melt from the associated cumulate pile, which over-thickens the lower SF, greatly tipping the competitive edge against suppression of melt leakage by advancing solidification. Dense interstitial melts may similarly drain from roofward cumulates of light phenocrysts. The variation in crystal size and modal abundance in these cumulate piles are intimate records of prior crystallization, transport, and filling.

Magmas in transit erode SFs and thoroughly charge the magma with crystals, facilitating fractionation and differentiation, especially if the body occasionally comes to rest. The key to protracted differentiation through fractional crystallization is not crystallization in stationary, closed chambers, but the repeated transport and chambering of magma or the periodic resupply to chambers of phenocryst-rich magma. This is punctuated differentiation, which may be the general case. Close corollaries are that thick, closed sheets of initially crystal-free, multiply-saturated magma undergo precious little overall differentiation, and that deciphering the sequence and crystallinity, including in transit phenocryst entrainment, growth, and sorting, of the filling events is central to unravelling intrusive history.

Variations in temperature, whether on phase diagrams or in actual magmas, are intrinsically linked to commensurate variations in space and time in magmatic systems. The spectrum of all physical and chemical processes associated with magma is accordingly strongly partitioned in space and time.

The idea of a magma chamber as a vat of low crystallinity melt crystallizing everywhere within and differentiating through crystal settling is unrealistic. A magma chamber formed of any number of crystal-laden inputs, encased by inward-propagating, dynamic solidification fronts, and where significant differentiation is tied to the dynamics of late-stage, interstitial melt within extensive mush piles is more in accord with the rock record.

The geochemical evolution of tin-tungsten deposits and their associated sulphides can be discussed in terms of four sequential processes: acquisition of the ore-forming elements (OFEs) by the granitic magma, emplacement of these elements in minerals and residual melt of the crystallizing granite, release of the OFEs to the circulating hydrothermal fluids and transport to the depositional sites, and finally, deposition of ore minerals through interaction of these fluids with the wall rock. Based on their crystallographic behaviour, it is useful to distinguish three principal classes of OFEs, here identified as BOC, LHC, and ALC elements. BOC (bivalent octahedral cation) elements are similar to ferrous iron and here are represented mainly by Zn, Mn, and perhaps Cu. Li also belongs to this class, although it is monovalent. LHC (large highly charged cations) elements encompass As, Nb, Mo, Sn, Sb, Ta, and W and they are similar to ferric iron or titanium in their crystallographic role. ALC (alkali-like cations) are capable of occupying alkali positions and are represented mainly by Pb, Ag, and Hg.

LHCs are rejected from the polymerized silicate liquid network and become enriched in the roof of the acid magma chamber, where more non-bridging oxygens are available. Transport to the roof may be enhanced by the formation of hydrous complexes, as is the pronounced enrichment of Na and Li. BOCs, along with Cl, F, and B, fractionate strongly into the vapour phase during vesiculation. HCl in the ore fluid is crucial for the alteration process and can be produced during boiling by a hydrolysis reaction of the NaCl dissolved or immiscibly present in the silicate magma.

Considerable laboratory information is available concerning release mechanisms of the OFEs to hydrothermal fluids. We can distinguish congruent and incongruent dissolution, both in response to acid buildup, as well as congruent and incongruent exchange not involving HCl. Melt-fluid fractionation is also thought to be important, though the physical mechanisms are not well understood. Any of these release mechanisms may be coupled with reduction or oxidation reactions. LHC, BOC, and ALC elements respond differently to each of these mechanisms, and these differences may in part be responsible for the observed separation of ore minerals in space and time. It is suggested that LHC elements are released preferentially during acid, non-oxidizing conditions typical of early stages, while BOC elements respond more readily to later acid-oxidizing environments, as well as exchange reactions with or without oxidation.

Depositional reactions have been formulated with respect to two contrasting types of country rocks: carbonates and schists. Differences are related to the process of neutralization of the HCl produced by ore deposition: carbonate dissolution on one hand and feldspar-muscovite or biotite-muscovite conversion on the other. In carbonate rocks, evaporite-related sulphates may provide the H2S necessary for sulphide precipitation, while in schists disseminated sulphides and organic matter may be important sulphur reservoirs in addition to the sulphur liberated from the granite. A variety of situations can be envisaged with respect to the sources of the OFEs and the sulphur species required for ore deposition, including granite and wall rocks. Chloride is recognized as the crucial anion for OFE release, transport, and deposition, although F and B play a role yet to be evaluated. Final HCl neutralization is an essential step in the reactions responsible for the deposition of ore minerals.

The ultimate sources of the OFEs must be related to the continental material involved in the process of melt production by partial melting. Oxidized sediments provide sources for LHC and ALC elements in the form of heavy minerals and clastic feldspars and micas. Organic-rich reduced sediments are hosts to BOC and LHC elements as sulphides and ALC elements in organic matter. Remelting of igneous and metamorphic rocks can enrich LHC, BOC, and ALC elements in the melt by extraction from opaques, Fe-Mg silicates, feldspars, and micas.

The transmission electron microscope is now used in a great variety of mineralogical and petrological contexts, The development of such applications over the past thirty years is illustrated by reference mainly to studies on serpentine and amphibole minerals.

Spectral measurements of sunlight reflected from planetary surfaces, when correlated with experimental visible-near-infrared spectra of rock-forming minerals, are being used to detect transition metal cations, to identify constituent minerals, and to determine modal mineralogies of regoliths on terrestrial planets. Such remote-sensed reflectance spectra measured through earth-based telescopes may have absorption bands in the one micron and two micron wavelength regions which originate from crystal field transitions within Fe2+ ions. Pyroxenes with Fe2+ in M2 positions dominate the spectra, and the resulting 1 μm versus 2 µm spectral determinative curve is used to identify compositions and structure-types of pyroxenes on surfaces of the Moon, Mercury, and asteroids, after correcting for experimentally-determined temperature-shifts of peak positions. Olivines and Fe2+-bearing plagioclase feldspars also give diagnostic peaks in the 1 µm region, while tetrahedral Fe2+ in glasses absorb in the 2 µm region as well. Opaque ilmenite, spinel and metallic iron phases mask all of these Fe2+ spectral features. Laboratory studies of mixed-mineral assemblages enable coexisting Fe2+ phases to be identified in remote-sensed reflectance spectra of regoliths. Thus, noritic rocks in the lunar highlands, troctolites in central peaks of impact craters such as Copernicus, and high-Ti and low-Ti mare basalts have been mapped on the Moon's surface by telescopic reflectance spectroscopy. The Venusian atmosphere prevents remote-sensed spectral measurements of its surface mineralogy, while atmospheric CO2 and ferric-bearing materials in the regolith on Mars interfere with pyroxene characterization in bright- and dark-region spectra. Reflectance spectral measurements of several meteorite types, including specimens from Antarctica, are consistent with a lunar highland origin for achondrite ALHA 81005 and a martian origin for shergottite EETA 79001, although source regions may not be outermost surfaces of the Moon and Mars. Correlations with asteroid reflectance spectra suggest that Vesta is the source of basaltic achondrites, while wide ranges of olivine/pyroxene ratios are inconsistent with an ordinary-chondrite surface composition of many asteroids. Visible-near-infrared spectrometers are destined for instrument payloads in future spacecraft missions to neighbouring solar system bodies.

Most countries intend to dispose of their high-level radioactive wastes by converting them into a solidified wasteform which is to be buried within the earth. SYNROC is a titanate ceramic wasteform which has been designed for this purpose on the basis of geochemical principles. It comprises essentially rutile TiO2, ‘hollandite’ Ba(Al,Ti)Ti6O16, zirconolite CaZrTi2O7, and perovskite CaTiO3. The latter three phases have the capacity to accept the great majority of radioactive elements occurring in high-level wastes into their crystal lattice sites. These minerals (or their close relatives) also occur in nature, where they have demonstrated their capacity to survive for many millions of years in a wide range of geological environments. The properties of SYNROC and the crystal chemistry of its constituent minerals are reviewed in some detail and current formulations of SYNROC are summarized. A notable property of SYNROC it its extremely high resistance to leaching by groundwaters, particularly above 100°C. In addition, it can be shown that the capacity of SYNROC minerals to immobilize high-level waste elements is not markedly impaired by high levels of radiation damage. Current investigations are focused on developing a satisfactory production technology for SYNROC and progress towards this objective is described. The high leach resistance of SYNROC at elevated temperatures increases the range of geological environments in which the waste may be finally interred; in particular, SYNROC is well adapted for disposal in deep drill-holes, both in continental and marine environments. The fact that SYNROC is comprised of minerals which have demonstrated long-term geological stability is significant in establishing public confidence in the ability of the nuclear industry to immobilize high-level wastes for the very long periods required.

Studies of the inclusions contained in natural diamonds have shown the occurrence of minerals which must have formed at depths below the lithosphere and which may be closely matched with the silicate mineral assemblages determined by high pressure and temperature experimental studies for depths of 300 to 800 km in the Earth's mantle. The inclusions come principally from two main depth zones: (1) the lower asthenosphere and upper transition zone; (2) the Upper Mantle/Lower Mantle (UM/LM) boundary region and the uppermost LM. The inclusions from zone 1 are very largely majoritic garnets (with or without clinopyroxene) which indicate bulk compositions of eclogitic/metabasic affinity. The minerals from zone 2 include Ca-Si and Mg-Si perovskites and ferropericlase and are dominantly of metaperidotitic bulk composition, but include some possible metabasite assemblages. In many of these natural assemblages, the tetragonal almandine pyrope phase occurs rather than the garnet found in experiments.

As natural diamonds are believed to crystallize in fluids/melts, the hypothesis is developed that the restriction of diamonds and inclusions of particular compositions to the above two depth intervals is because they are controlled by loci of fluid/melt occurrence. Attention is focused on subduction zones because both suites of inclusions show some evidence of subducted protoliths. The lower zone (600–800 km) coincides with the region where dehydration may be expected for hydrous ringwoodite and dense hydrous Mg-silicates formed in subducted peridotites. The dehydration of lawsonite in subducted metabasites provides a particular location for melt formation and the inclusion of the shallower (~ 300 km) majoritic inclusions. For the deeper majoritic inclusions in the region of the upper transition zone, melt development may occur as a consequence of the hydrous wadsleyite-to-olivine transformation, and such melt may then interact with the upper crustal portion of a subducting slab. These suggestions offer an explanation of the depth restrictions and the compositional restrictions of the inclusions. The differences in δ13C values in the host diamonds for the two suites of inclusions may also be explained on this basis.

Data on the distribution of the precious metals (Au, Ag and the platinum=group elements Ru, Rh, Pd, Os, Ir, Pt) are rapidly accumulating as a consequence of the application of microbeam analytical techniques such as the electron microprobe, the proton microprobe, and the ion microprobe. The new data obtained with these techniques build on knowledge accumulated over the last 140 years using bulk analytical methods (e.g. spectrographic). The nature of the occurrence, the concentrations, and the variation in amount s of precious metals in minerals and mineral products are reviewed with comments as to the significance of the results with respect to industrial and environmental concerns. Future directions in mineral-specific analysis of trace precious-metal concentrations and applications outside of process mineralogy are discussed.

The need in geology for in situ U-Pb age determinations of minerals is illustrated by two examples: the internal age dispersion developed within the zircon SL13 shortly after original crystallization, and the occurrence within minerals of old cores and later overgrowths. SL13 contains rare μm-sized patches of unsupported radiogenic Pb and a mainly bimodal distribution of 206Pb/238U ages otherwise. Both observations are consistent with original crystallization at 580 Ma and Pb loss at 565 Ma. Age precision is controlled by the ions counted for radiogenic Pb, and varies with instrumental sensitivity, age and U contents of the target. Laser-ablation ICPMS has similar spatial resolution and sensitivity to SIMS but consumes more sample because of much greater hole-depth in practice. Like SIMS, the measured Pb+/U+ is biased and also changes with depth so comparison with a standard mineral is necessary. Analyses of reference zircons reported here indicate that the reproducibility of Pb/U ages by ICPMS is limited by residual bias, not ion counting errors. For multipurpose ICPMS at least, the Hg background at mass 204 prohibits the measurement of 204Pb for common Pb estimation. A third micro-analytical method, ‘CHIME’, and future developments in SIMS and ICPMS are discussed briefly.

Although the thermal evolution of the mantle before c. 3.0 Ga remains unclear, since c. 3.0 Ga secular cooling has dominated over heat production—this is time's arrow. By contrast, the thermal history of the crust, which is preserved in the record of metamorphism, is more complex. Heat to drive metamorphism is generated by radioactive decay and viscous dissipation, and is augmented by the influx of heat from the mantle. Notwithstanding that reliable data are sparse before the Neoarchean, we use a dataset of temperature (T), pressure (P) and thermobaric ratio (T/P at the metamorphic ‘peak’), and age of metamorphism (t, the timing of the metamorphic ‘peak’) for rocks from 564 localities ranging in age from the Cenozoic to the Eoarchean eras to interrogate the crustal record of metamorphism as a proxy for the heat budget of the crust through time. On the basis of T/P, metamorphic rocks are classified into three natural groups: high T/P type (T/P >775°C/GPa, mean T/P ~1105°C/GPa), including common and ultrahigh-temperature granulites, intermediate T/P type (T/P between 775 and 375°C/GPa, mean T/P ~575°C/GPa), including high-pressure granulites and medium- and high-temperature eclogites, and low T/P type (T/P <375°C/GPa, mean T/P ~255°C/GPa), including blueschists, low-temperature eclogites and ultrahigh-pressure metamorphic rocks. A monotonic increase in the P of intermediate T/P metamorphism from the Neoarchean to the Neoproterozoic reflects strengthening of the lithosphere during secular cooling of the mantle—this is also time's arrow. However, temporal variation in the P of intermediate T/P metamorphism and in the moving means of T and T/P of high T/P metamorphism, combined with the clustered age distribution, demonstrate the cyclicity of collisional orogenesis and cyclic variations in the heat budget of the crust superimposed on secular cooling since c. 3.0 Ga—this is time's cycle. A first cycle began with the widespread appearance/survival of intermediate T/P and high T/P metamorphism in the Neoarchean rock record coeval with amalgamation of dispersed blocks of lithosphere to form protocontinents. This cycle was terminated by the fragmentation of the protocontinents into cratons in the early Paleoproterozoic, which signalled the start of a new cycle. The second cycle continued with the progressive amalgamation of the cratons into the supercontinent Columbia and extended until the breakup of the supercontinent Rodinia in the Neoproterozoic. This cycle represented a period of relative tectonic and environmental stability, and perhaps reduced subduction during at least part of the cycle. During most of the Proterozoic the moving means for both T and T/P of high T/P metamorphism exceeded the arithmetic means, reflecting insulation of the mantle beneath the quasi-integrated lithosphere of Columbia and, after a limited reorganisation, Rodinia. The third cycle began with the steep decline in thermobaric ratios of high T/P metamorphism to their lowest value, synchronous with the breakup of Rodinia and the formation of Pannotia, and the widespread appearance/preservation of low T/P metamorphism in the rock record. The thermobaric ratios for high T/P metamorphism rise to another peak associated with the Pan-African event, again reflecting insulation of the mantle. The subsequent steep decline in thermobaric ratios of high T/P metamorphism associated with the breakup of Pangea at c. 0.175 Ga may indicate the start of a fourth cycle. The limited occurrence of high and intermediate T/P metamorphism before the Neoarchean suggests either that suitable tectonic environments to generate these types of metamorphism were not widely available before then or that the rate of survival was low. We interpret the first cycle to record stabilisation of subduction and the emergence of a network of plate boundaries in a plate tectonics regime once the balance between heat production and heat loss changed in favour of secular cooling, possibly as early as c. 3.0 Ga in some areas. This is inferred to have been a globally linked system by the early Paleoproterozoic, but whether it remained continuous to the present is unclear. The second cycle was characterised by stability from the formation of Columbia to the breakup of Rodinia, generating higher than average T and T/P of high T/P metamorphism. The third cycle reflects colder collisional orogenesis and deep subduction of the continental crust, features that are characteristic of modern plate tectonics, which became possible once the average temperature of the asthenospheric mantle had declined to <100°C warmer than the present day after c. 1.0 Ga.

Oxidation zones and mine wastes are metal-rich, near-surface environments, natural and man-made critical zones of ore deposits, respectively. They contain a number of minerals which, despite their metastability, occur consistently and in abundance. Field studies, presented as examples in this work, show that metastable minerals form not only directly from aqueous solutions, but also from more complex precursors, such as nanoparticles, gels, X-ray amorphous solids, or clusters. Initial precipitation of metastable phases and their conversion to stable phases is described by the Ostwald's step rule. Thermodynamic data show that there is a tendency, but no rule, that structurally more complex phases are also thermodynamically more stable. The Ostwald's step rule could then state that the initial metastable phases are structurally simple and easily assembled from aqueous solutions, nanoparticles, gels, disordered solids, or clusters. The structural similarity of the precursor and the forming phase is a kinetic factor favouring the crystallisation of the new phase. Calculation of saturation indices for mine drainage solutions show that they are mostly supersaturated with respect to the stable phases and the aqueous concentrations are sufficient to precipitate metastable minerals. In our fieldwork, we often encounter gelatinous substances with copper, manganese or tungsten that slowly convert to metastable oxysalt minerals. Another possibility is the crystallisation of various metastable minerals from solid, homogeneous ‘resins’ that are X-ray amorphous. Minerals typical for near-surface environments may be stabilised by their surface energy at high specific surface areas. For example, ferrihydrite is often described as a metastable phase but can be shown to be stable with respect to nanosised hematite.

Biominerals are a subset of the mineral kingdom, those created by living creatures. In spite of usually fine grain size and intimate association with organic materials, biominerals are readily identified as common mineral species. Iron hydroxides and oxyhydroxides, calcium carbonates and calcium phosphates from uni- and multi-cellular species are presented as examples of biominerals, and biomineralization processes. Their special morphological, and crystal chemical, characteristics provide unique structural contributions to the life forms that create them. Investigations of novel habitats should present opportunities to expand the number of biominerals and their potential for industrial applications.

One of the major obstacles to our understanding of the growth of continental crust is that of estimating the balance between extraction rate of continental crust from the mantle and its recycling rate back into the mantle. As a first step it is important to learn more about how and when juvenile crust is preserved in orogens. The most abundant petrotectonic assemblage preserved in orogens (both collisional and accretionary) is the continental arc, whereas oceanic terranes (arcs, crust, mélange, Large Igneous Provinces, etc.) comprise <10%; the remainder comprises older, reworked crust. Most of the juvenile crust in orogens is found in continental arc assemblages. Our studies indicate that most juvenile crust preserved in orogens was produced during the ocean-basin closing stage and not during the collision. However, the duration of ocean-basin closing is not a major control on the fraction of juvenile crust preserved in orogens; regardless of the duration of subduction, the fraction of juvenile crust preserved reaches a maximum of ∼50%. Hafnium and Nd isotopic data indicate that reworking dominates in external orogens during supercontinent breakup, whereas during supercontinent assembly, external orogens change to retreating modes where greater amounts of juvenile crust are produced. The most remarkable feature of εNd (sedimentary rocks and granitoids) and εHf (detrital zircons) distributions through time is how well they agree with each other. The ratio of positive to negative εNd and eHf does not increase during supercontinent assembly (coincident with zircon age peaks), which suggests that supercontinent assembly is not accompanied by enhanced crustal production. Rather, the zircon age peaks probably result from enhanced preservation of juvenile crust. Valleys between zircon age peaks probably reflect recycling of continental crust into the mantle during supercontinent breakup. Hafnium isotopic data from zircons that have mantle sources, Nd isotopic data from detrital sedimentary rocks and granitoids and whole-rock Re depletion ages of mantle xenoliths collectively suggest that ≥70% of the continental crust was extracted from the mantle between 3500 and 2500 Ma.

The widespread occurrence of pyrope garnet in Archean lithospheric mantle remains one of the 'holy grails' of mantle petrology. Most garnets found in peridotitic mantle equilibrated with incompatible-trace-element enriched melts or fluids and are the products of metasomatism. Less common are macroscopic intergrowths of pyrope garnet formed by exsolution from orthopyroxene. Spectacular examples of these are preserved in both mantle xenoliths and large, isolated crystals (megacrysts) from the Kaapvaal craton of southern Africa, and provide direct evidence that some garnet inthe sub-continental lithospheric mantle formed initially by isochemical rather than metasomatic processes. The orthopyroxene hosts are enstatites and fully equilibrated with their exsolved phases (low-Cr pyrope garnet ± Cr-diopside). Significantly, P-T estimates of the postexsolution orthopyroxenes plot along an unperturbed conductive Kaapvaal craton geotherm and reveal that they were entrained from a large continuous depth interval (85 to 175 km). They therefore represent snapshots of processes operating throughout almost the entire thickness of the sub-cratonic lithosphericmantle.

New rare-earth element (REE) analyses show that the exsolved garnets occupy the full spectrum recorded by garnets in mantle peridotites and also diamond inclusions. A key finding is that a few low-temperature exsolved garnets, derived from depths of ∼90 km, are more depleted in light rare-earth elements (LREEs) than previously observed in any other mantle sample. Importantly, the REE patterns of these strongly LREE-depleted garnets resemble the hypothetical composition proposed for pre-metasomatic garnets that are thought to pre-date major enrichment events in the sub-continental lithospheric mantle, including those associated with diamond formation. The recalculated compositions of pre-exsolution orthopyroxenes have higher Al2O3 and CaO contents than their post-exsolution counterparts and most probably formed as shallow residues of large amounts of adiabatic decompression melting in the spinel-stability field. It is inferred that exsolution of garnet from Kaapvaal orthopyroxenes may have been widespread, and perhaps accompanied cratonization at ∼2.9 to 2.75 Ga. Such a process would considerably increase the density and stability of the continental lithosphere.

During the past 70 years, more than 2000 metric tonnes of Pu, and substantial quantities of the ‘minor’ actinides such as Np, Am and Cm, have been generated in nuclear reactors. Some of these transuranium elements can be a source of energy in fission reactions (e.g. 239Pu), a source of fissile material for nuclear weapons (e.g. 239Pu and 2Np), and of environmental concern because of their long half-lives and radiotoxicity (e.g. 239Pu and 237Np). There are two basic strategies for the disposition of these transuranium elements: (1) to ‘burn’ or fission the actinides using nuclear reactors or accelerators; (2) to dispose of the actinides directly as spent nuclear fuel or to ‘sequester’ the actinides in chemically durable, radiation-resistant materials that are also suitable for geological disposal. For the latter strategy, there has been substantial interest in the use of actinide-bearing minerals, especially isometric pyrochlore, A2B2Oi (A = rare earths; B = Ti, Zr, Sn, Hf), for the immobilization of actinides, particularly plutonium, both as inert matrix fuels and nuclear waste forms. Systematic studies of rare-earth pyrochlores have led to the discovery that certain compositions (B = Zr, Hf) are stable to very high doses of α-decay event damage. Recent developments in the understanding of the properties of actinide-bearing solids have opened up new possibilities for the design of advanced nuclear materials that can be used as fuels and waste forms. As an example, the amount of radiation damage that accumulates over time can be controlled by the selection of an appropriate composition for the pyrochlore and a consideration of the thermal environment of disposal. In the case of deep borehole disposal (3—5 km), the natural geothermal gradient may provide enough heat to reduce the amount of accumulated radiation damage by thermal annealing.