Journal of Petrology | Volume 44 | Number 4 | Pages 713-732 | 2003
© Oxford University Press 2003
Petrogenesis of Mesozoic, Peraluminous Granites in the Lamoille Canyon Area, Ruby Mountains, Nevada, USA
1 DEPARTMENT OF GEOSCIENCES, TEXAS TECH UNIVERSITY, LUBBOCK, TX 79409-1053, USA
2 DEPARTMENT OF GEOLOGY AND GEOPHYSICS, UNIVERSITY OF WYOMING, LARAMIE, WY 82071-3006, USA
3 US GEOLOGICAL SURVEY, MS 975, 345 MIDDLEFIELD ROAD, MENLO PARK, CA 94025, USA
E-mail: Cal.Barnes{at}ttu.edu.
RECEIVED DECEMBER 27, 2001; ACCEPTED OCTOBER 27, 2002
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONFOOTWALL OF THE CORE...THE INFRASTRUCTURE IN THE...DISCUSSIONPARTIAL MELTING AS THE...IMPLICATIONS FOR THE ORIGINS...SUPPLEMENTARY DATAREFERENCES |
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Two groups of closely associated, peraluminous, two-mica granitic gneiss were identified in the area. The older, sparsely distributed unit is equigranular (EG) with initial
Nd 
- 8·8 and initial 87Sr/86Sr 0·7098. Its age is uncertain. The younger unit is Late Cretaceous (80 Ma), pegmatitic, and sillimanite-bearing (KPG), with Nd from -15·8 to -17·3 and initial 87Sr/86Sr from 0·7157 to 0·7198. The concentrations of Fe, Mg, Na, Ca, Sr, V, Zr, Zn and Hf are higher, and K, Rb and Th are lower in the EG. Major- and trace-element models indicate that the KPG was derived by muscovite dehydration melting (<35 km depth) of Neoproterozoic metapelitic rocks that are widespread in the eastern Great Basin. The models are broadly consistent with anatexis of crust tectonically thickened during the Sevier orogeny; no mantle mass or heat contribution was necessary. As such, this unit represents one crustal end-member of regional Late Cretaceous peraluminous granites. The EG was produced by biotite dehydration melting at greater depths, with garnet stable in the residue. The source of the EG was probably Paleoproterozoic metagraywacke. Because EG magmatism probably pre-dated Late Cretaceous crustal thickening, it required heat input from the mantle or from mantle-derived magma. KEY WORDS: leucogranite; Late Cretaceous; crustal melting; Sevier orogeny; peraluminous
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONFOOTWALL OF THE CORE...THE INFRASTRUCTURE IN THE...DISCUSSIONPARTIAL MELTING AS THE...IMPLICATIONS FOR THE ORIGINS...SUPPLEMENTARY DATAREFERENCES |
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The Ruby Mountains–East Humboldt Range (RM–EHR) is a Cordilleran metamorphic core complex in the hinterland of the mid-Cretaceous to early Eocene Sevier fold-and-thrust belt (Royse, 1993
). Its present tectonic setting is within the highly extended Basin and Range province (Howard, 1980; Snoke, 1980; Snoke & Miller, 1988; Snoke et al., 1990b, 1997; Fig. 1). The RM–EHR is part of a discontinuous belt of core complexes that can be traced from Canada to Mexico (Crittenden et al., 1980; Armstrong, 1982). These complexes are products of large-magnitude crustal extension, which is commonly superposed on an earlier history of tectonic thickening related to crustal shortening. Like all core complexes, the RM–EHR can be divided into a footwall separated from a hanging wall by a normal-sense, plastic-to-brittle fault zone. Before Tertiary development of the core complex, abundant peraluminous two-mica granites (now granitic gneisses) were emplaced in the footwall (Howard, 1980, 1987; Kistler et al., 1981). Dated units are late Middle Jurassic and Late Cretaceous in age (Hudec & Wright, 1990; Wright & Snoke, 1993).
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On a regional basis, coeval Mesozoic magmatism ranged widely in elemental and isotopic compositions (e.g. Barton, 1996
). Because of this range of compositions, the magmas have variously been interpreted as the result of crustal melting owing to thermal relaxation following crustal thickening (Patiño Douce et al., 1990; Wright & Snoke, 1993; Camilleri & Chamberlain 1997; McGrew et al., 2000) or owing to mantle heating (Barton, 1990), or the result of melting, mixing, and assimilation caused by influx of mantle-derived magmas into the lower crust (Barton & Hanson, 1989; DePaolo et al., 1992). We present major- and trace-element, stable and radiogenic isotope, and mineral compositional data for two distinct, Mesozoic, peraluminous, two-mica granitic gneiss units in the northern Ruby Mountains. The younger, Late Cretaceous unit is widespread in the East Humboldt Range and Ruby Mountains. The older, undated unit is of limited geographic extent. It is enigmatic because it shows field relations that suggest a linkage to the Late Cretaceous granites but isotopic compositions that suggest Jurassic or Early Cretaceous emplacement. Whole-rock geochemical and isotopic compositions of the granites are consistent with an origin by crustal melting, with little evidence of mass contribution from mantle-derived magmas. As such, these granites provide an `end-member' with which more complex Mesozoic magmatic systems may be compared.
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FOOTWALL OF THE CORE COMPLEX |
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TOPABSTRACTINTRODUCTIONFOOTWALL OF THE CORE...THE INFRASTRUCTURE IN THE...DISCUSSIONPARTIAL MELTING AS THE...IMPLICATIONS FOR THE ORIGINS...SUPPLEMENTARY DATAREFERENCES |
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Rocks in the footwall of the core complex range from high-grade metasedimentary rocks complexly intruded by Mesozoic and Tertiary granitic rocks (northern Ruby Mountains and East Humboldt Range) to low-grade metasedimentary rocks (southern Ruby Mountains) and the late Eocene Harrison Pass pluton (see Fig. 1). In the RM–EHR, non-mylonitic igneous and metamorphic rocks in the footwall (i.e. infrastructure) are overlain by a kilometer-scale, down-to-the-west, normal-sense mylonitic shear zone. This mylonitic shear zone is a prominent feature from the East Humboldt Range to the vicinity of the Harrison Pass pluton (Fig. 1). South of the pluton, the footwall consists of low-grade metasedimentary rocks separated from an unmetamorphosed hanging wall by a brittle, low-angle normal fault (Howard et al., 1979
; Hudec, 1990; Snoke et al., 1990a; Burton, 1997).
The infrastructure experienced a complex Mesozoic igneous, metamorphic, and deformational history (Howard, 1980; Snoke & Miller, 1988; Snoke et al., 1990b, 1997; McGrew et al., 2000). In the central Ruby Mountains, polyphase metamorphism and deformation were synchronous with emplacement of a suite of Jurassic two-mica granites (153 Ma granite of Dawley Canyon) that intruded Neoproterozoic–Lower Paleozoic metasedimentary rocks (Hudec, 1990, 1992; Hudec & Wright, 1990; Jones, 1999). Jurassic granitic rocks have not been recognized elsewhere in the RM–EHR (but see below). In Late Cretaceous time, the infrastructure was again affected by intrusion, metamorphism, and deformation; these events were chiefly related to regional tectonic thickening in the hinterland of the Sevier orogenic belt (Snoke et al., 1992; Camilleri & Chamberlain, 1997; McGrew et al., 2000). At this time, the Cordilleran miogeoclinal sequence was imbricated by regional thrust faults (Camilleri & Chamberlain, 1997) and grossly thickened into an orogenic crustal welt that may have reached >50 km in overall structural thickness (Coney & Harms, 1984; Camilleri et al., 1997; McGrew et al., 2000).
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THE INFRASTRUCTURE IN THE LAMOILLE CANYON AREA |
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TOPABSTRACTINTRODUCTIONFOOTWALL OF THE CORE...THE INFRASTRUCTURE IN THE...DISCUSSIONPARTIAL MELTING AS THE...IMPLICATIONS FOR THE ORIGINS...SUPPLEMENTARY DATAREFERENCES |
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The infrastructure is well exposed in the Lamoille Canyon area (Fig. 2). The oldest unit is the Cambrian–Neoproterozoic Prospect Mountain Quartzite, which consists of micaceous, feldspathic quartzite and subordinate pelitic schist. It is stratigraphically overlain by the marble of Verdi Peak, a sequence of calcite marble, abundant calc-silicate paragneiss, and scarce metadolomite (Howard, 1971
, 2000) correlated with Ordovician and Cambrian sedimentary units exposed in the southern Ruby Mountains (Sharp, 1942; Hudec, 1990; Burton, 1997). Other distinctive and mappable metasedimentary units in the Lamoille Canyon area are the Ordovician Eureka metaquartzite, a mid-Paleozoic metadolomite sequence, and Upper Devonian graphitic calcite (±dolomite) marble (Howard, 1971, 2000). A large-scale recumbent fold with a core of metaplutonic orthogneisses (the Lamoille Canyon nappe; Howard, 1966, 1987, 2000) folds an inferred premetamorphic fault [the Ogilvie thrust of Smith & Howard (1977)], and two formations tectonically duplicated by the thrust: the Prospect Mountain Quartzite and marble of Verdi Peak (Howard, 1980, 2000).
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The Neoproterozoic–Paleozoic metasedimentary rocks are intruded by Mesozoic and Tertiary granitic rocks. Mesozoic granitic rocks are generally dikes, sills, sheets, or irregular bodies, now commonly with gneissic texture (Fig. 2a). Mesozoic granitic rocks are more abundant in deeper structural levels than at higher levels and locally account for >50% of the infrastructure volume (Howard, 2000
). The intrusive patterns of the Cretaceous granitic rocks are so complex and pervasive (Fig. 2a) that delineation of individual intrusive bodies is virtually impossible except at a very large scale (e.g. 1:1000; e.g. Lee & Barnes, 1997; Jeon, 1999).
Mesozoic granitic rocks
Introduction
The older of the two Mesozoic granitic gneisses is an equigranular, medium- to coarse-grained, muscovite–biotite gneiss of monzogranitic to leucomonzogranitic composition (EG). The younger is pegmatitic, sillimanite-bearing, two-mica granitic gneiss (KPG). The two units are commonly associated in the Lamoille Canyon area. However, the EG is much less abundant, and its outcrop area may be restricted to Lamoille Canyon. Where clear crosscutting relations were observed, a tectonic foliation in the EG is cut by the KPG.
KPG-type granitic gneiss is abundant and widespread; similar rocks crop out from the northern East Humboldt Range to the northern margin of the Harrison Pass pluton (Fig. 1). Monazite from an exposure of the KPG near the Terraces day-use area (UTM zone 11, 638000mE, 4499050mN) in Lamoille Canyon yielded a Late Cretaceous U–Pb age (80 Ma) (Wright & Snoke, 1993) and zircon from KPG in the East Humboldt Range yielded an age of 84·8 ± 2·8 Ma (McGrew et al., 2000). It is coarse to very coarse grained and commonly pegmatitic. The KPG forms networks of sills and dikes and/or irregular bodies. At structural levels where wall rocks are relatively abundant, the KPG occurs in discrete sheets subparallel to, and folded with compositional layering or foliation in the host metasedimentary rocks (Fig. 2a). In contrast, the deepest exposed parts of the infrastructure are underlain by extensive exposures of sheet-like KPG gneiss with enclaves of marble, calc-silicate paragneiss, pelitic schist, and metaquartzite. Along contacts with calc-silicate paragneiss or impure calcite marble, garnet–idocrase–diopside–scapolite skarns are locally developed (Jeon, 1999).
Petrography
Mineral compositions were determined on an automated JEOL JXA-8900 electron probe microanalyzer at the University of Wyoming, using a suite of natural and synthetic standards. Nominal operating conditions were 15 kV accelerating voltage, 20 nA beam current (10 nA for plagioclase), and 1 µm spot diameter; ZAF corrections were used. Representative analyses for biotite, muscovite, and plagioclase are given in the Electronic Appendix, Tables 1–3. The Electronic Appendix may be downloaded from the Journal of Petrology website, at http://www.petrology.oupjournals.org.
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Equigranular two-mica granitic gneiss (EG). This unit is characterized by tectonic foliation and local banding formed by variable proportions of feldspars and micas. Plagioclase is generally weakly zoned (An21–24) except for sample LEE10 (An12–29; Fig. 3a). Quartz is generally anhedral with undulose extinction. Alkali feldspar is interstitial or is graphically intergrown with quartz; some phenocrysts (2–5 mm) are present. Biotite and muscovite occur as inclusions in plagioclase and alkali feldspar, and in gneissic bands with granoblastic quartz and alkali feldspar. The proportions of the micas differ from sample to sample, but biotite is typically much more abundant than muscovite. Anhedral garnet is locally present as well as sparse sillimanite. Other common accessory minerals are apatite, monazite, and zircon.
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, rims. (b) Biotite Al contents (per 22 oxygens) vs Fe-number [atomic Fe/(Fe + Mg)]. (c) Muscovite TiO2 (wt %) vs Aliv + Alvi (per 22 oxygens). 
, data from sample UL1-60.Pegmatitic two-mica granitic gneiss (KPG). These rocks are very coarse grained, with a wide range of textural and mineralogical variations. For example, modes vary widely: 20–40% quartz, 10–60% alkali feldspar, 10–40% plagioclase, 1–3% biotite, and 0–2% muscovite (Howard, 1966
). Foliation, where present, is defined by the micas. Plagioclase core compositions are in the An15–24 range and rim compositions in the An12–19 range (Fig. 3a). In the coarsest rocks, alkali feldspar and plagioclase crystals reach at least 30 mm in diameter and many show graphic textures. Plagioclase is porphyroblastic and surrounded by finer (<2 mm), granoblastic quartz and alkali feldspar. Quartz is generally interstitial, with undulose extinction. Accessory minerals are apatite, allanite, monazite, zircon, and sparse Fe–Ti oxides. Sillimanite typically occurs as needles and fibrolitic mats in biotite, and less commonly in muscovite. Some deformed muscovite grains are partially replaced by sillimanite. We interpret the sillimanite as a product of metamorphic reactions during high-grade deformation of the gneiss and not as residue from the source.
Mica compositions
Major-element variability in biotite is chiefly in FeOt (total iron expressed as FeO) content (20·1–30·2 wt %) and MgO content (5·0–8·5 wt %). The relatively low MgO contents are similar to those of biotite from Late Cretaceous two-mica granites in the Snake Range of eastern Nevada (Lee & Van Loenen, 1970; Lee et al., 1981). Aluminum contents of biotite from the EG range from 2·7 to 3·9 cations per formula unit, but they show no correlation with Fe/(Fe + Mg) (Fig. 3b). In contrast, biotite from the KPG has a narrower range of Al contents over a range of Fe/(Fe + Mg) values similar to that of EG biotite (Fig. 3b).
The most distinctive feature of the muscovite analyses is the high TiO2 (Fig. 3c). Anderson & Rowley (1981) and Miller et al. (1981) noted that Ti contents of primary muscovite are markedly higher than that of secondary muscovite or sericite. The distinct Ti contents of sample UL1-60 suggest that both igneous and metamorphic muscovite are present in this sample.
Estimates of pressure and temperature
Although the granitic gneisses and semipelitic schists in the host rocks contain plagioclase, biotite, sillimanite, and garnet, the GASP assemblage was not observed. Furthermore, garnets from the central Ruby Mountains analyzed by Kistler et al. (1981) are typically enriched in the spessartine component and inappropriate for GASP thermobarometry (Anderson, 1996).
Hodges et al. (1992) reported P–T results for two pelitic schists collected about 14 km north of Lamoille Canyon. Garnet rim compositions indicated final equilibration at 580–650°C and 360–420 MPa (Hodges et al., 1992). However, Gibb's method analysis (Hodges et al., 1992) indicated higher P–T conditions before mid-Tertiary extension, with the maximum conditions estimated to be 775°C and 670 MPa.
In the East Humboldt Range, the sillimanite-bearing assemblages of the Neoproterozoic schists indicate peak pressure and temperature of >900 MPa and 800°C (McGrew, 1992; McGrew & Peters, 1997; McGrew et al., 2000). These P–T conditions were interpreted to be Late Cretaceous by McGrew et al. (2000). Camilleri & Chamberlain (1997) showed that in the Wood Hills, east of the East Humboldt Range, the Cretaceous metamorphic grade increases northward. This is consistent with the northward increase in P–T estimates observed in the RM–EHR and supports the idea that these estimates represent Late Cretaceous metamorphism. Therefore, conditions of emplacement of the Late Cretaceous granites exposed in Lamoille Canyon were probably at pressures slightly less than 670 MPa at temperatures <775°C.
Zircon-saturation temperatures (Harrison & Watson, 1983) were calculated for the EG (750–800°C) and the KPG (615–770°C; Electronic Appendix Table 4). The effect of inherited zircon, the presence of which is likely (Wright & Snoke, 1993), would be to raise the T estimate. However, the low calculated temperatures and the higher T estimates for EG are broadly consistent with melting models discussed in a later section.
Isotope data
Whole-rock and quartz oxygen isotope ratios were analyzed in the Department of Geosciences, Texas Tech University, using BrF5 as the fluorinating agent. Oxygen isotope values for quartz range from +12·3 to +15·5
and whole-rock oxygen isotope values from +11·7 to +12·8 (Table 1, Fig. 4). No clear difference between the EG and KPG was observed.
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Strontium and Nd isotopic compositions and Rb, Sr, Sm, and Nd concentrations were determined at the Department of Geology and Geophysics, University of Wyoming. The data and analytical details are given in Table 2. The EG is clearly distinct, with lower initial 87Sr/86Sr (0·70955–0·70998) and higher initial
Nd (-8·7 to -8·9) than the KPG (Fig. 5), which has initial 87Sr/86Sr from 0·7157 to 0·7198 and initial Nd values from -15·8 to -17·3.
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Nd, Sr, and Pb isotopic studies of Mesozoic and Tertiary granites in the RM–EHR suggested that isotopic compositions vary as a function of age, tectonic history, and geographic location (Wright & Wooden, 1991
; Wright & Snoke, 1993). Jurassic and Early Cretaceous granites have higher Nd and lower initial 87Sr/86Sr, whereas younger granites have lower Nd and generally higher initial 87Sr/86Sr (Fig. 5). These distinctions were thought to arise as a result of thrusting of the middle and upper crust eastward, away from the eugeoclinal sources of the older granites. However, on the basis of a larger dataset, Wooden et al. (1998) interpreted these isotopic differences to reflect variations in the depth of crustal source rocks. Wright & Wooden (1991) and Wright & Snoke (1993) also identified subdivisions within Late Cretaceous and Tertiary granites (Fig. 5). One group has lower Nd (‘lower array’) associated with Archean isotopic signatures; the other has higher Nd (‘upper array’) associated with Proterozoic isotopic signatures (Fig. 5). These distinct isotopic groups were thought to result from differences in the age of the source: Archean (and Proterozoic) crust beneath the East Humboldt Range and only Proterozoic crust beneath the Ruby Mountains (Fig. 1). Figure 5 shows that the EG is isotopically similar to Jurassic–Early Cretaceous plutonic rocks of the RM–EHR (Wright & Snoke, 1993), whereas the KPG is similar to the ‘upper array’.
Major- and trace-element geochemical data
Selected major- and trace-element variations in the EG and KPG are illustrated in Figs 6 and 7, and representative analyses are listed in Electronic Appendix Table 4. Both types of granitic gneiss are SiO2 rich (70–76 wt %) and both are peraluminous [A/CNK, molar Al2O3/(CaO + Na2O + K2O) > 1; Electronic Appendix Table 4]: from 1·04 to 1·17 among EG samples and from 1·03 to 1·33 among KPG ones. Major-element oxides in the EG show broad negative correlation with SiO2, except for Na2O, K2O, and P2O5, which exhibit no systematic correlation with silica content (Fig. 6). In contrast, among the KPG samples, SiO2 is uncorrelated or, at best, weakly correlated with all major oxides. This group has generally lower FeO, MgO, Na2O and CaO, and higher K2O than the EG (Fig. 6).
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The gneisses are characterized by widely variable Sr (1388–175 ppm) and Ba (1523–299 ppm), with the highest Sr values among the EG and the highest Ba values among the KPG (Fig. 7). Both element concentrations are negatively correlated with SiO2. Niobium and Y contents are low in both units (both <23 ppm). Zr is typically <100 ppm among the KPG samples and ranges between 170 and 80 ppm among the EG samples (Fig. 7). The contents of V, Zn, and Hf are also higher in the EG, whereas Rb and Th are lower (Fig. 7; Electronic Appendix Table 4).
Rare earth elements
Rare earth element (REE) patterns for the EG are relatively uniform, with steep slopes (LaN/LuN = 28–97) that result from depleted heavy REE (HREE) contents relative to light REE (LREE) (Fig. 8a). Eu anomalies are not apparent except for sample RD31, which has a weak negative Eu anomaly.
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Samples of the KPG show a range of REE patterns (Fig. 8b). In general, the HREE are strongly depleted, but the LREE concentrations vary widely. Most samples either lack an Eu anomaly or have a weak negative anomaly. However, the three samples with lowest REE abundances (LEE2, LEE5, and B150) show positive Eu anomalies. These three samples are sill-like sheets in host metasedimentary rocks, but are otherwise geologically and geochemically unremarkable.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONFOOTWALL OF THE CORE...THE INFRASTRUCTURE IN THE...DISCUSSIONPARTIAL MELTING AS THE...IMPLICATIONS FOR THE ORIGINS...SUPPLEMENTARY DATAREFERENCES |
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Although the EG is geochemically and isotopically distinct from the KPG, and the age of the EG is uncertain, the close spatial association of the two units prompted us to model them as contrasting peraluminous granitic suites. The variation of Sr, Ba, Na, K, and Ca contents within a narrow range of SiO2 in both units suggests that feldspar fractionation and/or accumulation played a role in the variation of the two units. These processes were modeled with major-element mass-balance calculations and then tested for consistency with trace-element abundances.
The mass-balance calculations (Bryan et al., 1969; Electronic Appendix Table 5) show that removal of some combination of plagioclase, K-feldspar, quartz, and micas can explain major-element variations in the EG and KPG units. Successful models for the EG (Model 1) require removal of high proportions of plagioclase relative to K-feldspar and biotite removal is required. Although these models have good to very good statistical fits (r2 < 0·1, where r2 is the sum of squares of residuals), they are not consistent with the variation of trace elements. For example, because Model 1 for the EG requires that plagioclase was the dominant fractionating phase, REE patterns of EG samples should show increasingly large negative Eu anomalies; these are not observed (Fig. 8). An additional test of the major-element models is shown in Fig. 9, in which variations in Rb, Sr, and Ba are predicted on the basis of the major-element model results. The trends for EG show superficially good fits to the data; however, the predicted value of F (fraction of melt remaining) is not in agreement with the fractional crystallization model (Electronic Appendix Table 5).
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Successful mass-balance results were also obtained for the KPG (Models 2–4, Electronic Appendix Table 5), yet even larger problems arise when trace-element tests are applied. The principal of these is that the sequence of samples in the calculated liquid line of descent is not the same as the sequence inferred from Fig. 9. Similar inconsistencies are observed among the REE data. Furthermore, values of F determined in models 3 and 4 are not in agreement with the trace-element tests (compare Electronic Appendix Table 5 and Fig. 9). This lack of consistent variation among trace element abundances and REE patterns shows that the apparent good agreement of major-element models with the Rb, Sr, and Ba data is fortuitous. We conclude that fractional crystallization does not adequately explain the compositional variations within either unit. This is not to say that accumulation and fractional crystallization had no effect on the KPG. Rare earth element patterns that show low abundances and positive Eu anomalies suggest accumulation of feldspars, and this implies at least local fractional crystallization. However, a fractionation scheme for the entire suite was not successfully modeled.
For similar reasons, we discount the effects of unmixing of residual solids on the geochemical variations in the two-mica granitic gneiss units. ‘Restitic’ textures (e.g. clots of mafic crystals, calcic cores of plagioclase, refractory metasedimentary enclaves) are absent (sillimanite in these samples is of metamorphic origin). This is in contrast to the late Eocene two-mica monzogranite of Green Mountain Creek in the Harrison Pass pluton (26 km south of Lamoille Canyon), which contains residual surmicaceous enclaves and massive quartz interpreted by Barnes et al. (2001) to be refractory material from the source.
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PARTIAL MELTING AS THE ORIGIN OF THE GRANITIC MAGMAS |
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TOPABSTRACTINTRODUCTIONFOOTWALL OF THE CORE...THE INFRASTRUCTURE IN THE...DISCUSSIONPARTIAL MELTING AS THE...IMPLICATIONS FOR THE ORIGINS...SUPPLEMENTARY DATAREFERENCES |
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The isotopic and elemental compositions of the Mesozoic RM–EHR granites are consistent with metasedimentary source rocks (e.g. Kistler et al., 1981
; Farmer & DePaolo, 1984). Our larger dataset can now be used to model such sources in some detail. Before that consideration, it is useful to evaluate possible contributions of mafic magmas to these granites.
If mafic magmas provided significant mass to the Cretaceous granites of the RM–EHR, one would expect greater variability of elemental compositions, anticorrelation of elements such as Fe, Ti, Sc, and V with SiO2, and wider ranges of isotope ratios. Such variation is particularly apparent among the Tertiary igneous suites of this part of Nevada (e.g. Grunder, 1992, 1995; Barnes et al., 2001), which show the effects of interaction of mafic magmas with the crust. Therefore, we consider the isotopic data for the EG and KPG to reflect a lack of mass input from mafic magmas. This conclusion does not rule out a thermal input from such magmas, but requires isolation of thermal effects from any geochemical ones. If this conclusion is correct, then the isotopic and elemental compositions of the granites should reflect their source compositions and residual mineral phases during melting.
Constraints that can be applied to melting models are the peraluminous compositions of the granites, their high 
18O, and uniform, yet distinct initial 87Sr/86Sr and Nd values (Fig. 2). In addition, the lack of Eu anomalies and the high Sr contents of the EG suggest that residual plagioclase was of minor importance, whereas the depletion of HREE is consistent with residual garnet. Similar constraints can be applied to the KPG, with the additional need to explain its high K2O/Na2O. Finally, detailed mapping and sampling of the infrastructure in the northern EHR (McGrew, 1992; Batum, 1999; McGrew et al., 2000) failed to identify granitic gneisses equivalent to the EG, whereas KPG-type granitic gneisses are widespread. This suggests, but does not prove, that source rocks of KPG underlie the entire RM–EHR, whereas source rocks for EG are restricted to the basement of the Ruby Mountains, south of the inferred southern margin of the Archean Wyoming province (see Fig. 1).
Possible source rocks
Pre-Mesozoic stratigraphy in the eastern Great Basin is fairly well understood (Misch & Hazzard, 1962; Stewart & Poole, 1974; Stewart, 1980), and many of the characteristics of the Precambrian basement can be inferred from regional isotopic studies (e.g. Farmer & DePaolo, 1983, 1984; Wright & Wooden, 1991; Wright & Snoke, 1993; Wooden et al., 1998). As discussed above, Wright & Snoke (1993) used Sr, Nd, and Pb isotope data to infer that the RM–EHR is underlain by two contrasting basement provinces: in the north by the Archean Wyoming province and in the south by accreted Paleoproterozoic basement. The boundary between these two provinces strikes across the core complex (Fig. 1); it was interpreted to be a continuation of the Cheyenne belt, a Paleoproterozoic suture zone associated with a continental margin–oceanic arc collision (e.g. Hills & Houston, 1979; Houston et al., 1979; Karlstrom & Houston, 1984; Duebendorfer & Houston, 1987; Geist et al., 1989; Chamberlain, 1998).
Archean basement rocks are exposed in the core of the Winchell Lake fold–nappe in the northern EHR. These rocks are primarily gray biotite monzogranitic augen orthogneiss (2·52 Ga; Lush et al., 1988), with subordinate intercalated paragneiss (i.e. wall rocks of the orthogneiss). Within the fold–nappe the Archean rocks are encased in a heterogeneous, chiefly metasedimentary sequence that includes probable Paleoproterozoic quartzite, paragneiss, pelitic schist, and amphibolite (McGrew, 1992; Snoke, 1992).
No Precambrian basement rocks (either Archean or Paleoproterozoic) are exposed south of the closure of the Winchell Lake fold–nappe. As a result, the nature of the Proterozoic basement beneath the RM–EHR is uncertain. On the basis of Nd isotopic data for the late Eocene Harrison Pass pluton, Barnes et al. (2001) suggested that this basement is the northward extension of the Mojave province, a distinctive Precambrian isotopic province in southern Nevada and eastern California (Bennett & DePaolo, 1987; Wooden & Miller, 1990).
A miogeoclinal wedge rests on both the Archean and Proterozoic basement terranes; it is composed of metamorphosed and unmetamorphosed carbonate and siliciclastic strata and was at least 12–15 km thick before Mesozoic contractional deformation (Miller et al., 1988; Camilleri & Chamberlain, 1997). Parts of this wedge are well exposed in the eastern Great Basin (e.g. Misch & Hazzard, 1962; Stewart, 1980). The Paleozoic and Neoproterozoic carbonate- and quartzite-rich parts of the miogeoclinal sequence do not provide fertile sources for magma generation. In contrast, the lower part of this succession consists of thick pelitic to quartzofeldspathic units (McCoy Creek Group; Misch & Hazzard, 1962) parts of which are a potential source for silicic magmas. The McCoy Creek Group was probably undeformed until Mesozoic contraction and metamorphism (Miller et al., 1988), at which time it may have provided a relatively wet, fertile source for magma generation.
Source characteristics inferred from major-element data
Our understanding of the Precambrian basement rocks of the region suggests that possible source rocks include granitic orthogneiss (EHR only), meta-arenites, metapsammites, and amphibolites. Amphibolitic sources can be ruled out because partial melting of such rocks results in tonalitic magmas (e.g. Beard & Lofgren, 1991; Rapp et al., 1991; Winther & Newton, 1991; Wolf & Wyllie, 1991, 1993). Orthogneiss and mica-poor metapsammites were probably not fertile at the low temperatures implied by zircon saturation geothermometry (<800°C; e.g. Beard & Lofgren, 1991). However, mica-rich metapsammites (metagraywacke) and metapelites are appropriate source rocks and would result in the peraluminous to strongly peraluminous compositions observed.
If the compositional differences between the EG and KPG magmas were produced by partial melting of distinct source rocks, it should be possible to characterize these differences by reference to experimental studies. Patiño Douce (1999; Fig. 10) showed that major element concentrations of experimentally produced glasses vary systematically according to starting composition. In Fig. 10, samples of the EG have lower molar Al2O3/(MgO + FeOt) and (Na2O + K2O)/(FeO + MgO + TiO2) ratios than those of the KPG. Comparison with the compositions of glasses obtained from partial melting experiments suggests that the EG magmas are similar to melts produced from a metagraywacke source, whereas the KPG magmas are similar to melts from muscovite-rich metapelitic rocks (Fig. 10).
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Trace-element partial melting modeling for the origin of the KPG magmas
Magmas derived from metapelitic and metagraywacke sources generally form by dehydration melting of muscovite and/or biotite; such reactions are well constrained by experimental studies (e.g. Le Breton & Thompson, 1988
; Vielzeuf & Holloway, 1988; Holtz & Johannes, 1991; Patiño Douce & Johnston, 1991; Skjerlie & Johnston, 1992; Skjerlie et al., 1993; Vielzeuf & Montel, 1994). Consequently, the residual mineral proportions and the degree of melting (F) can be determined from the modal composition of a protolith by mass balance (e.g. Hertogen & Gijbels, 1976; Benito-García & López-Ruiz, 1992). Residual accessory minerals may strongly influence REE patterns and abundances (e.g. Hogan & Sinha, 1991; Harris & Inger, 1992). Therefore, we focused on variations of Rb, Sr, and Ba, which are controlled by major silicate phases that participate in partial melting (e.g. Harris & Inger, 1992; Harris et al., 1995).
Melt model for KPG
The mass-balance equation for incongruent melting was formulated by Hertogen & Gijbels (1976) and extended to the more general case by Benito-García & López-Ruiz (1992). Calculation of the residual mineral proportions and melt fraction during incongruent melting used the equation of Benito-García & López-Ruiz (1992; Table 3) and their program ANATEX.BAS. The calculations assume equilibrium during melting and that the dehydration melting reactions used are applicable to partial melting of the miogeoclinal metapelites of the region. Parameters such as PA (fractional contribution of a phase to the melting) and t (fractional proportion of a phase forming a new phase) in the equation were calculated from mass-balanced reactions for mica dehydration melting (Patiño Douce & Johnston, 1991).
According to Patiño Douce & Johnston (1991), vapor-present melting in metapelites can be represented by the following reaction [abbreviations after Kretz (1983) and Als for aluminosilicate]:
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Under vapor-absent conditions, the reaction is
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The miogeoclinal metapelites of the McCoy Creek Group show wide variations in muscovite and biotite proportions (Misch & Hazzard, 1962). It is therefore not appropriate to assume a single homogeneous source composition in the modeling. Instead, we used four compositions that cover the observed variations of mica content (Electronic Appendix Table 6) along with elemental data for miogeoclinal metapelites from the structurally deepest parts of Lamoille Cany on (Electronic Appendix Table 4) and the East Humboldt Range (Batum, 1999). Miogeoclinal metapelites that lack muscovite were not considered because at the low temperatures of melting they were probably infertile and because the major-element compositions of the KPG are consistent with muscovite-dominated melting (Fig. 10).
The concentration of a trace element in the melt is calculated using the batch-melting equation (partition coefficients given in Electronic Appendix Table 7), and the calculated residual mineral proportions and melt fraction from the previous steps. The batch-melting calculation assumes continuous re-equilibration of melt with residual solid phases until removal of the melt.
Vapor-present melting. Melting will proceed until one of the reactants (quartz, plagioclase, or muscovite) is exhausted. The calculated residual mineral proportions and melt fractions for the given source mineral proportions are shown in Electronic Appendix Table 6. The calculated melt fraction (F) is correlated with the modal proportion of muscovite in the source. Depletion of muscovite leaves plagioclase in the residue; however, the amount of residual plagioclase becomes minor as the muscovite proportion in the source increases. Consequently, the calculated melt has higher Sr contents than is observed (Fig. 11). Also, all of the models result in high modal proportions of biotite relative to plagioclase. This yields modeled Sr/Rb ratios higher than is observed (Fig. 11).
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Vapor-absent melting. Under vapor-absent conditions, alkali feldspar is produced by muscovite breakdown and contributes to the solid residue. As in the vapor-present melting calculations, muscovite was consumed in all models. However, vapor-absent incongruent melting of muscovite results in lower melt fractions (F) and a residue with larger proportions of feldspar relative to biotite. Consequently, batch-melting calculations from a given source indicate enrichment in Rb and depletion in Sr and Ba relative to those from vapor-present conditions (Fig. 12). The resultant Sr/Rb and Ba/Sr ratios are all within the ranges of observed values among the KPG.
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Melt model for the EG
Major-element characteristics of the EG are consistent with partial melts of metagraywackes rather than metapelites (Fig. 10). Biotite dehydration melting of metagraywacke is controlled by the Fe/Mg in biotite, plagioclase composition, relative modal proportions of plagioclase and biotite, oxygen fugacity, and pressure (e.g. Vielzeuf & Montel, 1994
; Patiño Douce & Beard, 1996). At relatively low pressures, garnet is unstable and relatively large amounts of plagioclase are left in the residue (e.g. Vielzeuf & Montel, 1994; Patiño Douce & Beard, 1996). However, at pressures >500 MPa, garnet becomes stable relative to plagioclase (Vielzeuf & Montel, 1994; Patiño Douce & Beard, 1996; Patiño Douce & McCarthy, 1998). Biotite is generally present in the residue unless melting temperature is higher than 900°C (e.g. Vielzeuf & Montel, 1994; Patiño Douce & Beard, 1996).
Under these conditions, partial melts of metagraywackes in the Paleoproterozoic basement in this region were probably in equilibrium with a garnet-rich, plagioclase-depleted residue (e.g. Patiño Douce & Beard, 1996). The resultant partial melts would have high Sr, lack Eu anomalies, and have depleted HREE. Because of the lack of elemental and modal compositions for the possible source (Paleoproterozoic basement metagraywackes), a semi-quantitative iterative method was applied for the purposes of demonstration. The model was tested using a range of source compositions compiled for Precambrian graywacke from Wyoming (Condie, 1969) and average Archean graywacke (Condie, 1981). The residual mineral phases and proportions were initially based on the experimental study of Patiño Douce & Beard (1996), after which the mineral proportions were adjusted to improve the results. The observed large ion lithophile element (LILE) abundances and REE patterns are best explained by partial melting of a source in equilibrium with residual quartz (44%), garnet (52%) and biotite (4%), and a melt fraction (F) of 0·5 (Fig. 13). This model explains the unusually high Sr contents, the highly differentiated REE patterns, and the typical lack of Eu anomalies in EG samples. The model slightly underestimates HREE abundances. However, considering uncertainties in source composition and the possible influence of accessory minerals, further modeling of REE abundances is at present not justified.
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IMPLICATIONS FOR THE ORIGINS OF RM–EHR GRANITES |
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Late Cretaceous pegmatitic granitic gneiss (KPG)
The origins of Late Cretaceous peraluminous magmatism in the Great Basin area have been explained by crustal thickening (Coney & Harms, 1984
; Miller et al., 1988), magmatic underplating or intraplating in the lower crust (Haxel et al., 1984; Speed et al., 1988), and progressive conductive heating of the crust by upwelling mantle (Barton, 1990). In the RM–EHR, Late Cretaceous granites have been proposed as products of crustal thickening followed by high-grade metamorphism and anatexis (Wright & Wooden, 1991; McGrew, 1992; Snoke, 1994; Camilleri & Chamberlain, 1997; McGrew et al., 2000), which is consistent with thermobarometric evidence for deep burial (e.g. Hodges et al., 1992; McGrew, 1992; Camilleri & Chamberlain, 1997; McGrew et al., 2000).
Trace-element models indicate that the crustal source rocks of the KPG contained abundant muscovite. In addition, the presence of KPG-type granites in the RM and EHR suggests a source common to both sides of the inferred local boundary between Archean and Proterozoic basement rocks. Neoproterozoic miogeoclinal strata of the McCoy Creek Group (Misch & Hazzard, 1962; Stewart, 1980) fit this geometric requirement. In addition, depleted mantle model ages for the McCoy Creek Group (Farmer & Ball, 1997) are in the same range as those of KPG. Furthermore, Farmer & Ball (1997) indicated that sediments in the McCoy Creek Group were locally derived. This could explain the distinctions in Nd between RM and EHR pegmatitic granites described by Wright & Snoke (1993).
The P–T–t path of these Neoproterozoic source rocks can be estimated from the results of McGrew et al. (2000), who determined P–T–t relations for the structurally deepest exposed part of the region, the East Humboldt Range, where peak pressure and temperature reached 860 MPa and 780°C (McGrew, 1992; McGrew et al., 2000). Before metamorphism, the Neoproterozoic rocks were probably at a stratigraphic depth of 11 km (P 300 MPa). Peak metamorphism accompanied a pressure increase of 500 MPa (Camilleri & Chamberlain, 1997). If the thickness of the source layer was 2·6 km (e.g. Misch & Hazzard, 1962), then decompression melting probably occurred at less than 35 km depth (940 MPa). The P–T–t path for the northern EHR (McGrew et al., 2000) and various muscovite and biotite dehydration melting solidi (Patiño Douce & Harris, 1998) are shown in Fig. 14. The P–T–t path intersects the muscovite dehydration solidus but does not intersect biotite dehydration solidi. This suggests that decompression melting by muscovite dehydration was possible and further suggests that added heat from mafic magmas was not necessary to produce the widespread KPG in the Ruby Mountains and East Humboldt Range.
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Equigranular granitic gneiss (EG)
In contrast to the KPG, equigranular granitic gneiss in the Lamoille Canyon area requires biotite dehydration melting at temperatures above 800°C and pressures likely to produce garnet in the solid residue. Such temperatures were probably not attained even at the base of the Precambrian miogeoclinal section during Late Cretaceous time (Fig. 14). Furthermore, because the age of the EG is uncertain, a petrogenetic link between it and Late Cretaceous crustal thickening cannot be established. In fact, the Sr and Nd isotope characteristics of the EG are most similar to those of Jurassic and Early Cretaceous granites of the northern Great Basin (Fig. 5).
Therefore, if the EG is Late Cretaceous, then its isotopic composition is unusual compared with the main volume of Late Cretaceous granites (KPG). Melting models for EG are most consistent with a metagraywacke source; such rock types are known to be components of the Mojave basement province (e.g. Bennett & DePaolo, 1987; geometry illustrated in Fig. 15). However, this scenario requires that EG melting (relatively high-T biotite dehydration melting) preceded KPG melting (lower-T muscovite dehydration melting).
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We prefer the interpretation that the EG is Early Cretaceous or Jurassic in age. A Jurassic age is more likely because Early Cretaceous peraluminous plutons are uncommon in the region (M. Barton, personal communication, 2002). The isotopic characteristics reflect a source that is distinct from KPG sources and was probably of limited geographic extent. This interpretation places the EG in a back-arc setting, which would provide the thermal flux necessary for biotite dehydration melting. Although the EG lacks chemical evidence of a mafic component, the high temperatures necessary for biotite dehydration melting probably require transfer of heat from mafic back-arc magmas at depth (e.g. Barton, 1990
; Thompson & Connolly, 1995; Patiño Douce, 1999). |
SUPPLEMENTARY DATA |
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TOPABSTRACTINTRODUCTIONFOOTWALL OF THE CORE...THE INFRASTRUCTURE IN THE...DISCUSSIONPARTIAL MELTING AS THE...IMPLICATIONS FOR THE ORIGINS...SUPPLEMENTARY DATAREFERENCES |
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Supplementary data for this paper are available on Bioinformatics online.
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ACKNOWLEDGEMENTS |
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We thank M. Jeon, M. Batum, and A. Strike for their assistance in the field and for generously sharing their ideas and data, and M. Barnes, J. Browning, and C. Beville for assistance in the laboratory. Additional thanks are due to M. Barton, Y. Janou
ek, and R. Evart for thorough, insightful reviews. This research was supported by NSF grant EAR-9627814 to C.G.B. and EAR-9627958 to A.W.S. Neutron activation analyses were obtained through the DOE Reactor Sharing Program at Oregon State University. |
REFERENCES |
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