Abstract Since its discovery in 1891, the Cripple Creek district has produced more than 653 tonnes of Au(21 million oz) from high-grade gold-telluride veins. About99.5tonnes (3.2mi1lion ounces) of gold have been added to the district resource as a result of recent exploration (past 5 years) which has delineated low-grade, near surface disseminated gold deposits that occur as broad zones in permeable rocks adjacent to major structures (pontius, 1992;Harris et sl., 1993). The ore deposits of the district are localized within and adjacent to an elliptical northwest-trending diatreme complex that covers about 18 km 2 , Although some of the deposits occur within Proterozoic rocks surrounding the complex, most of the ore bodies are spatially associated with Tertiary alkaline igneous rocks and breccias within the complex. During the most productive mining years beginning in 1893, several comprehensive studies of the Cripple Creek district were conducted that included detailed descriptions of the mineralogy, texture, and field characteristics of each rock type and of the ores (Cross and Penrose, 1895;Lindgren and Ransome, 1906;Loughlin and Koschmann, 1935). Over the past 12 years, selected ore deposit studies have focused on the fluid inclusion characteristics, mineralogy, and alteration of the deposits (Thompson et el., 1985; Saunders, 1986; Nelson, 1989; Seibel, 1991; Burnett, 1995). The close spatial association between the deposits and alkaline igneous rocks, and relatively hot (average of 350±C), hypersaline (>40 wt. % NaCI equivalent) fluid inclusions in quartz from early stages of vein formation (Thompson et sl., 1985) have led many workers to suggest that the ore fluids were magmatically derived (Lindgren and Ransome, 1906; Thompson et sl., 1985; Pontius, 1992).

The alphabet soup of granite types and its association with certain ore deposit types is called into question. We now realize that granites do not retain a menu of ingredients inherited at the site of their generation, and that initial magma chemistry is only one of several influences on the potential for later ore formation in granite systems. Rather, processes leading to fluid release and crystal-melt and fluid-melt fractionation of critical elements, mainly metal species and volatile complexing agents, are central to the formation of ore. These processes are certainly influenced by tectonic setting, source mineralogy and chemistry, assimilation, contamination, and geothermal gradient, but are not strictly controlled by them. The delicate relations between a granite magma, its crystallizing phases, volatile content and species, and oxygen fugacity, plus the timing and mechanism of fluid release and the efficiency of metal extraction, ultimately control the formation of an ore deposit. As experimental data become more available, models for granite-related ore deposits must incorporate and emphasize the important role of crystal and volatile fractionation in the system. In other words, we need to model the process as well as the source.

Constraints may be placed on the functional dependence of the partitioning of ore and other trace, minor, and major elements in melt (liquid)-vapor-crystal systems by relating empirical Nernst-type partition coefficients to true mass action expressions and equilibrium constants. This process is herein termed stoichiometric analysis. The resulting expressions can be used to determine the functional dependence of partition coefficients on the bulk composition of the phases in question. In this chapter, stoichiometric constraints are used to place limits on the partitioning of europium between melt and vapor, to determine the functional dependence of the HCl concentration in the magmatic aqueous phase on the basis of the composition of the melt, and to suggest controls on the iron concentration in the magmatic aqueous phase. Based on this analysis and data available in the literature, it can be shown that the Nernst partition coefficient for europium is probably a function of the square of the chloride concentration in the magmatic aqueous phase; at the very least, it is limited to a polynomial dependence on the chloride concentration raised to the second and third power. The HCl concentration in the aqueous phase is shown to increase with the peraluminosity and chlorine concentration in the melt phase, and the aqueous iron concentration is shown to increase with increasing activity of HCl and decreasing f O 2 for a given activity of magnetite component in a magmatic system.

The distribution of F and Cl between aqueous fluids and melts of haplogranite and topaz rhyolite composition was determined experimentally at 0.5 to 5 kbar and 775° to 1,000°C. The distribution coefficients, D i , for F and Cl were calculated as the parts per million by weight (ppmw) of i in the fluid/parts per million by weight (ppmw) of i in the melt. D Cl ranges from 0.8 to 85; however, under typical geologic conditions, Cl partitions more strongly into an aqueous fluid relative to F-bearing granitic melts. Cl partitions increasingly in favor of the fluid as F in the fluid and magma decrease and as the X H 2 O fl (molar H 2 O/H 2 O + CO 2 in fluid), temperature, pressure, and Cl in the fluid and melt increase. Cl partitions in favor of haplogranite melts that contain >7 wt % F and ⩽1,200 ppm Cl at 1,000°C and 2 kbar. D F ranges from 0.2 to >1.0; however, F typically is concentrated in granitic melts relative to aqueous fluids. F concentrates more strongly into topaz rhyolite melt as pressure, F in the fluid, and melt, temperature, and the X H 2 O fl decrease. However, F partitions in favor of aqueous fluids relative to topaz rhyolite melts at 800°C and 2 kbar, if the melt contains ⩾7 wt % F. Computations indicate that extreme enrichments in F (>4 wt %) and Cl (>5,000 ppm) may occur in magmas and in associated magmatic hydrothermal fluids during the end stages of crystallization of topaz rhyolite magmas and magmas associated with Climax-type molybdenum deposits if the initial H 2 O, CaO, and ferromagnesian contents in the magma are low, and if the pressure at which water saturation occurs is high.

The processes leading to the generation of pegmatite fabrics and rare-element ore deposits are evaluated in terms of pressure, temperature, and compositions of fluids through the combination of recent field, fluid inclusion, and experimental data. The results indicate that primary consolidation of miarolitic and massive rare-element (Li-Be-Ta) pegmatites occurs in the range of 700° to 450°C at 400 to 250 MPa. The model for pegmatite genesis proposed here involves disequilibrium crystallization of quartz and feldspars in a hydrous melt-crystal system that approaches chemical equilibrium but rarely attains it. Fluxing anionic components of boron, phosphorus, and fluorine are not essential to this model, but augment the zoning process by promoting expansion of the liquidus fields of quartz and K-silicates (feldspar and mica) and by driving residual melt toward alkaline, Na-rich compositions enriched in lithophile trace elements. Aqueous vapor, if present, may have a nominal affect on primary zonation and fabric development within pegmatites. Various lines of evidence suggest that rare-element pegmatites may not become vapor saturated until they approach solidus conditions, at which point internal retrograde recrystallization and wall-rock alteration occur. Recent experiments have succeeded in replicating most aspects of pegmatite geology, including fractionation trends, mineral fabrics, and spatial zonation, at vapor-undersaturated conditions.

Late-stage volatile-rich topaz granites occur widely but sparsely throughout the southwest England Sn-Cu-polymetallic mineralized S-type biotite granite batholith. New observations from the St. Austell area have clarified field relations, and demonstrate the importance of an aureole of tourmalinization affecting both granitic and sedimentary host rocks. Topaz granite contacts are often marked by pegmatitic zones showing undirectional solidification textures and carrying vugs of quartz-tourmaline; minor intrusive sheets have symmetrical haloes of tourmalinization within adjacent host rocks. The topaz granites are mineralogically complex, containing primary topaz, zinnwaldite, or lepidolite, amblygonite (and other phosphates), and various Nb-Ta rich accessory phases, as well as albite, orthoclase, and quartz. Fluorite is secondary. They are chemically distinct from the biotite granites, showing markedly higher concentrations of Li 2 O (as much as 0.5 percent), F (as much as 1.5 percent), P 2 O 5 (0.5 percent), Nb (as much as 65 ppm), Ta (as much as 30 ppm), Ga (as much as 50 ppm), and Rb (as much as 2,000 ppm), in particular. Late differentiates of the biotite granites include tourmaline granites, but this differentiation trend principally involves an increase in B with little change in Li, F, or P. It is considered that the topaz granites are unlikely to be derived by fractional crystallization of the southwest England biotite granite magma (they are intruded by rhyolite porphyry dikes, which belong to the biotite granite suite), and an origin involving limited partial melting of subbatholithic fusion residues during an episode of potassic basic magmatism is preferred. Comparison with other volatile-rich granitic rocks indicates that certain lithium pegmatites (e.g., Tanco), other topaz granites (Seward Peninsula, Erzgebirge, etc.) and volcanic glasses (e.g., Macusani) share important characteristics with the southwest England topaz granites. It is suggested that these rock types may represent a fundamentally similar volatile-rich granite magma type whose formation, while debatable, may be controlled by limited partial melting of lower crustal fusion residues that had previously generated more “normal” granite magmas.

In this chapter we describe the petrogenesis of aplitic segregations in the fluorine-rich Proterozoic Butler Hill and Graniteville granites of the St. Francois Mountains volcano-plutonic terrane, southeastern Missouri. Both plutons contain an early coarse-grained type of granite that grades into or is crosscut by fine-grained aplitic segregations. The aplitic segregations are generally enriched in fluorine and alkalis, have more pronounced negative Eu anomalies in their rare earth element (REE) patterns and higher concentrations of heavy REEs compared to their coarse-grained high-silica counterparts. In the Butler Hill granite, the distinction in fluorine concentrations between the two rock types was obliterated in part by subsolidus hydrothermal alteration, which is indicated by complete chloritization of biotite, sericitization of feldspar, and recrystallization or reequilibration of muscovite with water-rich fluids. These chemical relations suggest that the aplitic segregations in both the Butler Hill and Graniteville plutons are the products of crystal-liquid fractionation in which fluorine played a significant role. Initially, the slightly peraluminous to metaluminous compositions of the fluorine-containing magmas were near the composition of a minimum melt in the Q-Ab-Or system. However, early crystallization of quartz and feldspars resulted in enrichment of the remaining melt in fluorine, causing the pseudoternary minimum to move away from quartz. As a consequence of the enlargement of the quartz field, quartz became the sole crystallizing phase, yielding the silica-enriched composition of the coarse-grained granites. The aplitic rocks crystallized from the relatively alkalic residual melt, which separated from the crystalline assemblage as the melt’s viscosity decreased due to an increase in the fluorine and water content. These results are in accord with published experimental data that show that the effect of fluorine is to decrease the silica content of residual liquids, contrary to normally observed fractionation trends in igneous rocks.

The quartz phenocrysts of rhyolitic rock that have erupted over the last 40 m.y. in the Thomas Range, west-central Utah, preserve glass inclusions that may contain trapped liquidus phases. These glass inclusions, analyzed by electron microprobe, show the melts of the Joy Tuff (crystal tuff member), Dell Tuff, and Topaz Mountain Rhyolite to be peraluminous potassic rhyolites, the latter enriched in fluorine (as much as ~1.9 percent). The older units (Joy and Dell Tuffs), tuffaceous calc-alkaline rhyolites, contain trapped Mn-rich aluminous clinopyroxene, zircon, fluorapatite, feldspar(?), titanite(?), and quartz. The anorogenic Topaz Mountain rhyolite preserves evidence of a ternary feldspar and magnetite (as well as quartz) on the liquidus. The potassic and peraluminous nature of the melts is consistent with anatexis of a crustal sequence. The distribution of fluorine, unlike that of chlorine, is heterogeneous in the Topaz Mountain Rhyolite. In this F-rich unit, primary β -quartz gave way to primary α -quartz. Both Ca and Fe were added to the bulk rocks following their devitrification.

Peraluminous rhyolites that are chemically somewhat similar to topaz rhyolites and anorogenic granites occur in an orogenic setting near Sierra Blanca in the Tertiary Trans-Pecos magmatic province. The Sierra Blanca rhyolites are even more enriched in most incompatible trace elements than are topaz rhyolites. Some of the extreme enrichments may in part be the result of chemical modification by crystallization from an F-rich vapor phase. The rhyolites were intruded as laccoliths at 36 Ma, during the main phase of Trans-Pecos igneous activity, which is characterized by ash-flow eruptions from numerous calderas and widespread mafic, intermediate, and silicic intrusions. A dominant east-northeast orientation of dikes and veins throughout the region indicates mild compression that was residual from Laramide deformation. This compressive tectonic setting, coupled with concurrent volcanism in Mexico and the east-northeast change in magma chemistry from calc-alkalic in western Mexico through alkali-calcic to alkalic in Texas, suggests that the rhyolites were emplaced in a continental arc. Extension did not begin in Trans-Pecos Texas until after 32 Ma; 31- to 17-Ma dikes are dominantly oriented north-northwest, perpendicular to the direction of extension during early Basin and Range deformation. Thus, the tectonic setting of the Sierra Blanca rhyolites contrasts with that of typical topaz rhyolites, most of which were emplaced during periods of crustal extension. The Sierra Blanca rhyolites are chemical and mineralogic oddities for the region, where most rhyolites are peralkaline or metaluminous. The rhyolites are depleted in the same elements as topaz rhyolites (Mg, Ca, Ti, Sr, Ba) but are more highly evolved than topaz rhyolites. Extreme trace-element enrichments (Li, F, Zn, Rb, Y, Zr, Nb, Sn, Ta, Pb, HREE, Th, U) are accommodated in Li-rich white mica, Zn-rich biotite, Rb-rich feldspars, and numerous trace minerals, including cassiterite, changbaiite, columbite, thorite, xenotime, yttrium- and REE–rich fluorides, and zircon. The rhyolites are large-tonnage, low-grade resources of several rare metals. Also enriched in Be (as much as 180 ppm), the rhyolites are the sources of Be and F in beryllium deposits in fluoritized limestones along the contacts with the laccoliths. Interaction with the limestones probably locally elevated the Ca, Mg, and Sr contents of the rhyolites. Vapor-phase crystallization has modified the original magmatic chemistry of the rocks. Evidence of vapor-phase crystallization includes the presence of minerals typical of pegmatites: cryolite (from 0 to 3 volume percent), alkali feldspars with nearly end-member compositions, polylithionite-zinnwaldite mica, and rutilated quartz, plus fluid inclusions defining quartz overgrowths on magmatic grains. Extreme HREE enrichments (Yb to 72 ppm; chondrite-normalized REE patterns with positive slopes) may also be the result of vapor-phase crystallization.

Late Cretaceous to Early Tertiary granitic plutons associated with W skarn or Sn greisen-skarn occur interspersed in a belt 70 × 200-km-long just northeast of Fairbanks, Alaska. All plutons intrude the late Precambrian–early Paleozoic Yukon-Tanana terrane and are similar in major-element compositions (dominantly granodiorite to monzogranite), initial Sr isotopic ratios (0.710 to 0.719), and Pb isotopic signatures ( 206 Pb/ 204 Pb = 19.17 to 19.37). Biotite compositions and opaque mineral abundances indicate both types of plutons crystallized along a buffered path intermediate between nickel–nickel oxide and quartz-magnetite-fayalite. Both suites contain multiple igneous units, with younger, usually equigranular, units spatially related to mineralized zones. Isotopic, trace-element, and mineralogical data suggest an “I-type,” “ilmenite-series” classification for both pluton suites. Because the W and Sn plutons appear to represent magmas with similar origins and source materials, differences in observed metallogeny are thought to be related to differences in environment of crystallization and vapor loss. Such differences include: age (102 to 87 Ma for W plutons, 73 to 50 Ma for Sn plutons), crystallization pressure (1 to 2 kbar for W plutons, <0.5 kbar for Sn plutons), vapor loss history (late for the W plutons and early + late for the Sn plutons), and fluorine trends (decreasing F with increasing differentiation for the W plutons and increasing F for the Sn plutons). Differences in confining pressure (depth) and vapor loss history are associated with differences in age: the younger (Sn) plutons are shallower, and the older (W) plutons are deeper. Trace-element patterns (e.g., Rb, B, Be, W, Sn, Li) are similar for least differentiated units of both pluton types, increasing modestly with increasing differentiation for the W plutons and increasing strongly for the Sn plutons. Data are most compatible with 80 to 95 percent fractionation (crystal-liquid) followed by vapor loss for the W plutons and 80 to 90 percnt fractionation (crystal-liquid) for the Sn plutons, with early vapor loss followed by (liquid-liquid?) “ultrafractionation.” Ultrafractionation and subsequent ore element enrichment occurs in the Sn plutons by early vapor loss and subsequent F enrichment in the residual magma. The data suggest that metallogeny differences for W vs. Sn plutons in our study area are not a function of differences in initial metal contents of the magmas but are more likely due to differences in magmatic evolution.

The amount and character of the tin resources in the Seward Peninsula tin granite belt in western Alaska are directly related to the depth of erosion of the plutonic system. Plutons that have little or no outcrop (little erosion), such as Kougarok or Lost River, are the site of lode tin deposits, whereas placer tin deposits are associated with somewhat eroded plutons (Cape Mountain, Ear Mountain) with a modest outcrop area. The largest pluton in the tin belt, the Oonatut Complex, is a deeply eroded pluton with little tin in either lode or placer deposits. Textural units in the granite plutons also vary with depth in the plutons. Late-stage, fractionated (Differentiation Index = sum of normative quartz + albite + orthoclase = 91.5 to 96.6) equigranular biotite granite is found at high levels in the plutons, and earlier, less fractionated (Differentiation Index = 71.8 to 92.5), seriate and porphyritic biotite granite is more abundant at greater depth. Fluids associated with the porphyritic and seriate granites were H 2 O-NaCl–rich and produced metal-poor, idocrase-scapolite-diopside–rich skarns in the carbonate host rocks. Fluids evolved during the late-stage crystallization of the equigranular granite were enriched in incompatible components and produced tin-bearing greisen in granite and hedenbergitic pyroxene-garnet-tourmaline-axinite-cassiterite skarns in marble. Mineral assemblages and chemistry reflect magma fractionation and record a two-stage evolution for the granite and associated skarns in the tin belt. Biotite compositions show progressive Fe enrichment (Fe/(Fe + Mg) = 0.69 to 0.99) with increasing fractionation of the enclosing granite. Biotite-bearing pegmatite and aplite dikes are associated with the early-crystallizing seriate and porphyritic units of some plutons. Biotite compositions in these dikes are very similar to those in the seriate and porphyritic granites, indicating the H 2 O-rich fluids associated with the dikes were magmatic in origin. Tourmaline is found as an interstitial phase in the equigranular granites; it is Fe-rich, similar to the coexisting biotite, suggesting a magmatic origin for this tourmaline. Later, secondary tourmaline is Ca-Mg–rich, and is associated with white mica alteration of biotite in granite and with sulfide deposition in skarns.

The Devonian batholiths of western Tasmania represent a diverse assemblage of highly fractionated intrusions (70 to 77 percent SiO 2 ) that are the products of different source materials. The Housetop batholith exhibits compositional affinities to a fluorine-rich I-type magma. The Meredith batholith also has characteristics indicative of I-type source materials. The Heemskirk batholith is composite, and consists of a volatile (F, B, H 2 O)–rich S-type granite underlying an I-type granite. The Three Hummock Island, Interview River, Sandy Cape, and Conical Rocks plutons probably have an S-type source and are grouped together as the Sandy Cape Suite. Rapakivi texture is common in the Housetop, Meredith, and Heemkirk batholiths. Quartz-tourmaline nodules are found in the Conical Rocks pluton and the S-type portion of the Heemskirk batholith. The Conical Rocks and Interview River plutons yield high initial Sr isotopic ratios of 0.74242 and 0.76009, respectively. The Housetop and Meredith batholiths yield the lowest initial Sr isotopic ratios of 0.71041 and 0.71445, respectively. The S-type portion of the Heemskirk batholith has an initial Sr isotopic ratio of 0.76387. The 40 Ar/ 39 Ar release spectrum and Rb/Sr mineral isochron analyses corroborate previously reported Devonian to Carboniferous age estimates for these batholiths. A relatively low-temperature thermal event (<200°C) caused argon loss from the K-feldspars at about 105 Ma. This heating event is probably related to the continental breakup of Australia from Antarctica. Major-element compositions of the western Tasmanian granites are very similar. The highly fractionated Sandy Cape Suite leucogranites exhibit high Ga/Al ratios typical of A-type granites, but not their extreme Zr, Y, or Ce enrichments. A distinctive feature of the Sandy Cape Suite is the increase in P 2 O 5 concentration with fractionation. The increase in P 2 O 5 with fractionation is apparently due to extremely low Ca activity, which precludes the formation of apatite, thus allowing P 2 O 5 to behave incompatibly in the melt. All of the granitoids have LREE–enriched chondrite-normalized rare earth element patterns. REE fractionation within the individual granitoids can be summarized by two trends: those with LREE >> HREE depletion (Housetop, Meredith, and Heemskirk batholiths), and those with LREE = HREE depletion (Sandy Cape Suite). The first trend is caused by the initial undersaturation of accessory mineral assemblage that resulted from high concentrations of volatiles and/or alkali complexes. The second trend is caused by early saturation of accessory phases and/or refractory accessory phases.

The East Kemptville greisen-hosted tin deposit (58 million tonnes of 0.165 percent Sn), southwestern Nova Scotia, Canada, occurs beneath undulations in the contact between granitic rocks of the Davis Lake complex (DLC) and Meguma Group metawacke. Cassiterite-topaz ore precipitated from a F- and Sn-rich fluid derived from the East Kemptville leucomonzogranite. Controls on tin mineralization include the unusual primary Sn and F abundances of the DLC magma, the vertical chemical zonation of the magma prior to crystallization, generation of an aqueous phase, and associated Sn-Cl complexing before the separation of F from the magma and the flat-lying granite-metawacke contact. The Carboniferous DLC is composed of biotite-bearing monzogranites, leucomonzogranite, and high-F, low-B topaz greisen. The chemical and isotopic signatures of the monzogranites reflect the unusual source of this highly evolved pluton. Magmatic evolution was focused toward the granite-metawacke contact and culminated, at the current erosional level, in the East Kemptville deposit. Chemical variation is attributed to the vertical zonation of the magma prior to the fractionation of biotite, K-feldspar, ilmenite, and REE-bearing phosphate minerals. The remarkably high F and P contents resulted in unusual Al and P distributions and concentration of metals beneath the contact as a result of increased magmatic depolymerization and diffusion. Pegmatitic segregations record the evolution of a F- and Sn-rich aqueous fluid. This aqueous phase scavenged Cl, alkali elements, P, Sn, and other metals from the magma, concentrating them beneath the contact. In contrast to porphyry-style deposits, this fluid was not expelled from the granite. When crystallization was complete, more than 10,000 tonnes of F partitioned into this aqueous fluid, forming a hydrofluoric fluid that was neutralized by reaction with the leucomonzogranite. Quartz-topaz rock is either a direct precipitate from the F-rich fluid or extensively replaced leucomonzogranite. Quartz-mica greisen and incomplete greisen formed as the F/OH ratio in the fluid decreased. Veins and fractures emanating from massive greisen zones have alteration envelopes, indicating these structures were conduits for the F-rich fluid. Their orthogonal orientation suggests these veins are related to cooling, not hydraulic fracturing. Whole-rock Rb-Sr isotopic data from leucomonzogranite and quartz-topaz rocks yield dates and initial ratios statistically identical to the biotite monzogranite. High mean square of weighted deviations (MSWDs) resulting from the former data reflect contamination of the late-magmatic fluids by Sr derived from the metawacke and later thermal overprinting. Deformation, probably related to regional tectonothermal events, affected all rocks of the DLC, but postdates ore formation.

In this chapter we report on the use and limitations of oxygen and halogen fugacities to characterize granitoids associated with tungsten deposits. Tungsten-related granitoids generally occupy an intermediate geochemical position, less evolved than tin granites, but more differentiated than porphyry copper and molybdenum granodiorites. Microprobe analyses of the different populations of biotite in the CanTung system (in granite, aplite, lamprophyre, vein, and skarn) show differences in major-element chemistry. Fluorine-chlorine intercept values, however, indicate that the halogens in the biotites reequilibrated with one fluid. This reequilibration probably was at relatively low temperatures (300° to 350°C), but might have been at higher temperatures (500°C) if the fluid did not change its halogen fugacity during cooling. In both cases, this implies that the initial halogen content of the magmatic biotites is overprinted by the hydrothermal fluid. The observation that an aplite, away from the mineralization, has a lower intercept value (higher fluorine fugacity) may indicate that the mineralizing fluid evolved from the main biotite granite body and not from the leucocratic phase. High ferrous/ferric ratios in biotites, the absence of magnetite and titanite in the granite, and the skarn mineralogy indicate reducing conditions of the magmatic-hydrothermal fluids of the CanTung and some other high-grade deposits. We propose that the oxidation state of the magma, the amount of initial water, and the depth of emplacement play an important role in the formation of large tungsten deposits.

Wolframite-bearing quartz veins flanked by greisen alteration occur at and near the Black Pearl mine, Yavapai County, Arizona. The veins are genetically related to a small albitite stock, and cut a series of Proterozoic metasedimentary and intrusive rocks. The largest vein, the only one mined, is located at the apex of the stock. Field relations imply that this stock is a late-stage differentiate of the 1.4-Ga anorogenic Lawler Peak batholith, which crops out about 3 km to the south. Other, similar, albitites occur locally. Sharp contacts, relatively unaltered xenoliths, and uniform mineralogy indicate that the albitites are of igneous origin and have suffered only minor deuteric alteration. A thin (1 to 2 m) pegmatite unit (“stockscheider”) occurs at the contact of the Black Pearl Albitite stock with the country rocks. Directional indicators and other evidence suggest that the pegmatite was formed in the presence of a volatile-rich fluid phase close to the time of magma emplacement. The sudden change from coarse-grained microcline-rich pegmatite to fine-grained, albite-rich albitite suggests pressure quenching, possibly due to escape of fluids up the Black Pearl vein. Descriptions of other tungsten and tin deposits worldwide indicate that stockscheider-like textures typically occur near the apical contacts of productive plutons. The presence or absence of this texture is a useful guide in prospecting for lithophile metal deposits.

Tin mineralization in the Black Range is present in a series of Oligocene rhyolite domes that cover an area of approximately 200 km 2 in southwestern New Mexico. These rhyolites are typically high silica (76 to 78 percent), peraluminous, and topaz bearing. The δ 18 O values (~7.0 ‰) of unaltered rocks are typical of normal I-type granites. Locally along dome margins, the glassy groundmass of the rhyolite is pervasively altered to smectite, although sanidine phenocrysts are usually fresh, even in contact with tin-bearing veinlets. Tin occurs (1) rarely as cassiterite in miarolitic cavities near tin-bearing veinlets, (2) within altered zones in widely spaced cross-cutting veinlets of cassiterite ± wood tin (colloform cryptocrystalline cassiterite), and (3) most abundantly as widespread placer accumulations of wood tin ± cassiterite in drainages within or marginal to the domes. Veinlet minerals include early quartz, K-rich sanidine and topaz, followed by hematite, cassiterite ± wood tin, and cristobalite. Late chalcedony, fluorite, durangite, clays, zeolites, and a complex suite of unusual minerals are locally present in and near veinlets. The δ 18 O values of most individual minerals are remarkably uniform throughout the area (cassiterite, 4.3 to 3.2 ‰; wood tin, 3.0 to 1.2 ‰; hematite, 3.5 to 2.6 ‰). Within these narrow ranges the δ 18 O values for wood tin and hematite are distinctive in each of the three types of tin occurrences and they generally decrease paragenetically in individual veinlet and placer samples. The δ 18 O values of temporally overlapping cristobalite in veinlets, however, range from 9.7 to 11.8 ‰ and increase paragenetically. The δ 18 O values for quartz in miarolitic cavities containing cassiterite are nearly the same as those for adjacent rhyolite quartz phenocrysts (7.5 ‰). A consistent Δ 18 O of 1.0 ‰ for quartz-sanidine phenocrysts indicates isotope equilibration temperatures of ~700°C for rhyolite during dome emplacement. Temperatures based on quartz-hematite 18 O fractionations range from ~800°C (miarolitic cavities) to ~485°C (wood tin–bearing veinlets). Late cristobalite in the veinlets precipitated at a maximum temperature of only ~230°C. Cassiterite-hematite Δ 18 O values are uniform (0.7 ± 0.1 ‰), suggesting that the cassiterite-water curve is parallel to the hematite-water curve between 800° and 400°C. An empirical 18 O fractionation curve for cassiterite-water suggests that most wood tin precipitated above ~350°C. The δ 18 O and δD values of smectite from pervasive argillic alteration range from 10.9 to 15.1 ‰ and −73 to −107 ‰, respectively, and along with whole-rock data on altered rocks, imply alteration temperatures of ⩽235°C. The δ 18 O values for chalcedony range from 15.9 to 34.4 ‰ and imply temperatures generally ⩽160°C for local silicification of rhyolite. The δ 18 O H 2 O of the tin-bearing fluids was 8.0 ±0.5 ‰, indicating that the fluids equilibrated with a magma or high-temperature rhyolite throughout most of the time-space milieu of mineralization. However, limited data indicate that the δD H 2 O of the fluids ranged from −60 to at least as low as −101 ‰ and may have decreased with successive stages of mineralization, even though there is no evidence of mixing of the fluids with unexchanged meteoric water. Isotopically similar fluids may have been responsible for the pervasive alteration of the host rock. Integrated stable-isotope, fluid-inclusion, and petrographic and geologic data suggest that tin mineralization resulted from NaCl-saturated fluids that derived from shallow magmas or that equilibrated with high-temperature rhyolite. Tin mineralization resulted largely from rapid temperature decrease and decompression of vapor-rich fluids, which produced cassiterite in fractures at deeper levels and wood tin in the carapace of the domes. At shallow levels, HCl from the hydrolysis of NaCl permeated the wall rock adjacent to the widely spaced veinlets and contributed to the pervasive alteration of the glassy groundmass of the rhyolite. This model suggests that the tin resource potential of the area has already been realized with the mining of the placer wood tin.

The Taylor Creek Rhyolite, a group of Tertiary high-silica-rhyolite lava domes and flows in southwestern New Mexico, contains cassiterite-bearing veins whose tin was derived from the host rhyolite as it degassed, cooled, and devitrified immediately after emplacement. Theoretical considerations and studies of fumarolic deposits at many volcanoes worldwide indicate that tin is highly mobile in a vapor phase, probably as halogen complexes, thus favoring the occurrence of such auto-mineralization in a cooling-lava environment. Mass-balance calculations for the New Mexico situation indicate that much of the tin evolved during devitrification of the rhyolite cannot be accounted for in the mineralized deposits. Some of this “missing” tin almost certainly was dispersed into alluvium during erosion of mineralized parts of the lavas, and some may have been transferred to the atmosphere around fumaroles rooted in the cooling lavas. In addition, tin may have been lost to the atmosphere from Taylor Creek Rhyolite magma that was erupted in fountains. The recent recognition of fountain-fed fallback in the New Mexico rhyolite field suggests that this third means of moving tin out of erupting magma may indeed have contributed to the overall tin history in the Taylor Creek Rhyolite magma system. Fountain-fed flows of silicic lavas are not well known, whereas mafic counterparts are known to be common as a result of observations of many eruptions of basaltic magma. Characteristic properties of silicic magmas that collectively tend to result in relatively high viscosity inhibit the occurrence of eruption columns whose fallback is hot enough to thoroughly weld and perhaps totally rehomogenize into a melt that subsequently feeds lava flows. However, high volumetric rates of eruption, high magmatic temperature (relative to solidus temperature), and any other conditions that help to reduce viscosity, in concert with factors that result in relatively brief periods of trajectory for lava clots, favor the formation of fountain-fed silicic lava flows. Fluorine-rich magma of about 830°C produced such silicic lavas in the Taylor Creek Rhyolite, and rocks with a similar mode of emplacement in other volcanic areas are herein hypothesized to be far more abundant than presently recognized. Possible examples elsewhere include large-volume sheets of silicic lavas in southwestern Idaho (eruption temperature of 950° to 1,100°C) and in Trans-Pecos Texas, where lava-flow and welded pyroclastic textures intermixed within individual eruptive units have led to confusion and difficulty in interpreting the mechanism of emplacement. Documentation of fountain-fed silicic lavas is rare; eruptions of silicic magmas are infrequent relative to the average human lifespan, and very few have occurred during historic time. Moreover, evidence of a lava-fountain origin may be only weakly preserved in the rocks so formed; the evidence also may be entirely lacking, as is commonly the case for the closely observed mafic examples.

Taylor Creek high-silica tin-bearing rhyolites are found in the northern Black Range of the Mogollon-Datil volcanic field in southwestern New Mexico, occurring near the stratigraphic top of a thick mid-Tertiary volcanic section. Initial ɛ Nd values for the high-silica rhyolite lavas range from −5.0 to −6.2, which are similar to those of the Garcia Camp tuff, a pyroclastic phase of the Taylor Creek Rhyolite. The older Kneeling Nun tuff, which crops out in the same area, also has a similar ɛ Nd 1 value, which indicates that the high-silica rhyolites and spatially associated silicic tuffs were derived from an isotopically similar source. Comparison with data from lower crustal xenoliths and data bearing on the isotopic compositions of the lower crust suggest that the melts were derived from 80 to 50 percent lower crustal sources. The Poverty Creek basaltic andesite and Bearwallow Mountain Formation andesite, stratigraphically below and above Taylor Creek Rhyolite, respectively, have more positive ɛ Nd 1 values of −4.7 and −2.3, respectively, indicating a greater mantle component. Initial 87 Sr/ 86 Sr ratios vary from 0.7046 to 0.7131 for the Taylor Creek Rhyolite. There is a broad positive correlation between initial 87 Sr/ 86 Sr and Sr content and a negative correlation with Rb, Ta, and Th content. These variations may be explained by late-stage upper crustal assimilation of radiogenic and relatively Sr-rich wall and roof rocks. Whole-rock Sr contents of the least radiogenic rocks as low as 3 ppm indicate that little assimilation would be required to affect the original Sr isotopic signature of the Taylor Creek Rhyolite magma. The Nd isotopes, however, were not measurably affected by the upper crustal processes.

The rapakivi granites (1.7 to 1.55 Ga) of southern Finland occur as epizonal batholiths (e.g., the Wiborg, Åland or Ahvenanmaa, Vehmaa, and Laitila batholiths) and stocks cutting the medium- to high-grade metamorphic Svecofennian (1.9 to 1.8 Ga) crust. Emplacement of the granites was associated with faulting and the intrusion of coeval sets of mainly west-northwest- (some north-northeast) trending diabase and quartz porphyry dikes, indicating an extensional continental tectonic regime. The rapakivi granite batholiths and stocks are multiple intrusions, several of which also contain minor anorthositic and gabbroic bodies. Granites of the early intrusive phases commonly crystallized from water-deficient magmas and contain biotite and hornblende (± fayalite) as dark constituents. The younger intrusive bodies contain biotite as the only ferromagnesian silicate, whereas the youngest, water-saturated intrusive phases are topaz-bearing granites, in which the dark mica is lithium-bearing siderophyllite. Fluorite, zircon, allanite, apatite, anatase, magnetite, and ilmenite are typical accessory minerals in the granites of the early and main intrusive phases. The biotite granites contain monazite instead of allanite, and the late-stage granites contain topaz, monazite, ilmenite, Nb- and Ta-rich cassiterite, and columbite as common accessory minerals. Topaz-bearing quartz porphyry dikes and greisen-type tin-polymetallic mineralization are often associated with the last intrusive phases. The rapakivi granites are metaluminous to slightly peraluminous rocks characterized normally by high K, K/Na, Fe/Mg, F, Ga, Rb, Zr, Hf, Th, U, and REE. The early and main intrusive phases are enriched in LREE and show deep Eu anomalies. The last minor intrusive phases show flattened normalized REE patterns with still deeper Eu minima. As a result of extreme differentiation and superimposed alteration, they are anomalously enriched in F, Ga, Rb, Sn, and Nb, and are impoverished in Ti, Ba, Sr, and Zr. The rapakivi granites exhibit geochemical characteristics of subalkaline A-type granites and within-plate granites. Nd isotopic studies from the northern part of the Wiborg rapakivi area indicate that the rapakivi-age diabase dikes [∊ Nd ( T ) values +1.6 to −1.0] crystallized from mantle-derived magmas that had experienced variable degrees of crustal contamination. The rapakivi granite-quartz porphyry magmas [∊ Nd ( T ) values −0.8 to −1.9] most probably originated by partial melting of the Svecofennian crust formed 0.2 to 0.3 b.y. earlier. Heat flow from the mantle-derived magmas contributed to the partial melting.