Gneissic pegmatitic leucogranite forms a dominant component (>600 km3) of the midcrustal infrastructure of the Ruby Mountains–East Humboldt Range core complex (Nevada, USA), and was assembled and modified episodically into a batholithic volume by myriad small intrusions from ca. 92 to 29 Ma. This injection complex consists of deformed sheets and other bodies emplaced syntectonically into a stratigraphic framework of marble, calc-silicate rocks, quartzite, schist, and other granitoids. Bodies of pegmatitic granite coalesce around host-rock remnants, which preserve relict or ghost stratigraphy, thrusts, and fold nappes. Intrusion inflated but did not disrupt the host-rock structure. The pegmatitic granite increases proportionally downward from structurally high positions to the bottoms of 1-km-deep canyons where it constitutes 95%–100% of the rock. Zircon and monazite dated by U-Pb (sensitive high-resolution ion microprobe, SHRIMP) for this rock type cluster diffusely at ages near 92, 82(?), 69, 38, and 29 Ma, and indicate successive or rejuvenated igneous crystallization multiple times over long periods of the Late Cretaceous and the Paleogene. Initial partial melting of unexposed pelites may have generated granite forerunners, which were remobilized several times in partial melting events. Sources for the pegmatitic granite differed isotopically from sources of similar-aged interleaved equigranular granites. Dominant Late Cretaceous and fewer Paleogene ages recorded from some pegmatitic granite samples, and Paleogene-only ages from the two structurally deepest samples, together with varying zircon trace element contents, suggest several disparate ages of final emplacement or remobilization of various small bodies. Folded sills that merge with dikes that cut the same folds suggest that there may have been in situ partial remobilization. The pegmatitic granite intrusions represent prolonged and recurrent generation, assembly, and partial melting modification of a batholithic volume even while the regional tectonic environment varied dramatically from contractile thickening to extension and mafic underplating.
Investigations into the degree to which various plutonic bodies represent frozen big tanks of magma or, alternatively, incremental accumulations from many small intrusions have focused particularly on geochronologic and structural evidence (Glazner et al., 2004; Lipman, 2007; Miller, 2008). Geochronologic evidence of prolonged magma emplacement and accumulation over intervals >1 m.y. or even as long as 7 m.y. has been offered as evidence of gradual incremental assembly of some plutons (Coleman et al., 2004; Condon et al., 2004; Matzel et al., 2006). The recycling of plutonic zircon antecrysts and xenocrysts into younger melts both in young volcanic systems and plutons can pose difficulties in assessing crystallization ages (Bacon and Lowenstern, 2005; Matzel et al., 2007).
The presence of granitoid sheets around little-disturbed country-rock remnants in pluton border zones has been used as another line of evidence for incremental assembly for certain intrusions (Weinberg and Searle, 1998; Mahan et al., 2003; Glazner et al., 2004). Relict or ghost stratigraphy defined by stratigraphically and structurally coherent host-rock enclaves has long been known from injection complexes, migmatites, and within some mapped plutons (e.g., Cushing et al., 1910; Pitcher, 1952, 1970; Pitcher and Read, 1959; Pitcher and Berger, 1972; Blundy and Sparks, 1992; Mahan et al., 2003; Ciavarella and Wyld, 2008). Such structurally coherent relics are in contrast to xenoliths pictured as individually rotated even though in stratigraphic order within a pluton (Miller and Miller, 2002; Hawkins and Wiebe, 2004). Coherent relict stratigraphy and structure can result from intrusive wedging by a plexus of coalescing granite sheets and thus may offer insight into a type of incremental assembly of plutonic bodies, such as the Main Donegal Granite pluton of Ireland (Pitcher and Read, 1959; Hutton, 1982, 1992; Pitcher and Berger, 1972; Pitcher and Hutton, 2003).
We describe pegmatitic leucogranite that dominates a batholith-size (>600 km3) intrusive complex in the Ruby Mountains of Nevada (USA) in which relict stratigraphy is preserved as remnants of metacarbonate rock, quartzite, and subordinate schist (Figs. 1 and 2). The pegmatitic granite forms an injection complex grading downward into a plutonic massif and is at least partly syntectonic with ductile folding and metamorphism. Field evidence, combined with ion microprobe U-Pb geochronology and zircon geochemistry, suggests that the pegmatitic granite and intrusive forerunners were generated, assembled, and modified episodically over very long periods of both Late Cretaceous and Paleogene time from countless small intrusions. The combined field and geochronologic evidence thus offers new perspectives about processes and long time spans that can be involved in generation, assembly, and evolution of some plutonic complexes.
The Ruby Mountains in the Basin and Range province expose part of the Ruby–East Humboldt metamorphic core complex, in the hinterland of the Cretaceous Sevier fold and thrust belt and in a backarc position relative to the Sierra Nevada batholith (Fig. 1). Stratified and intrusive rocks in the core complex record Mesozoic tectonic thickening of a thick miogeoclinal section, metamorphism, and intrusion, followed by Paleogene magmatism, Miocene extensional unroofing, and continued relative uplift along normal faults into the Quaternary (Table 1).
Stratigraphic and Structural Framework: Ghost Stratigraphy
Metamorphosed miogeoclinal strata in the core complex form the host to abundant Jurassic to Paleogene granitoids (Table 1). Glaciated exposures in canyons 0.6–1.2 km deep reveal exceptional cross sections of fold nappes and relations of intrusions to host-rock stratigraphy and structure (Fig. 2). Nearby mountain ranges expose little-metamorphosed equivalents of the same miogeoclinal strata and a few Jurassic, Early Cretaceous, Late Cretaceous, and Eocene plutons (Coats, 1987; du Bray, 2007).
The metamorphosed strata in Lamoille Canyon (Ruby Mountains) consist of (1) the metamorphosed Prospect Mountain Quartzite (Neoproterozoic and Cambrian), comprising micaceous feldspathic quartzite (85% quartz) having semipelitic laminae; (2) the marble of Verdi Peak (Cambrian and Ordovician), comprising marble, diopside- bearing calc-silicate rocks, and minor amounts of plagioclase-quartz-biotite ± muscovite-sillimanite schist and plagioclase-biotite-hornblende schist; and (3) small amounts of Ordovician to Devonian quartzite and marble. Relative proportions of the metamorphosed stratal rock types are similar to those in unmetamorphosed or weakly metamorphosed protoliths to which they are correlated in the southern Ruby Mountains and nearby ranges (Howard, 1971, 2000; Coats, 1987). Elsewhere in the core complex are older Neoproterozoic strata of the McCoy Creek Group, and possible (disputed) Neoarchean ortho gneiss (Howard, 1971, 1980, 2000; Snoke, 1980; Lush et al., 1988; McGrew et al., 2000; Howard and MacCready, 2004; Premo et al., 2008, 2010; McGrew and Snoke, 2010).
At deep structural levels the stratigraphic framework consists of only scattered remnants among variably foliated granitoid intrusions, but it becomes increasingly better revealed at higher structural levels where proportions of granite decrease. Geologic maps have emphasized the stratigraphic continuity of metasedimentary remnants (Smith and Howard, 1977; Howard et al., 1979; Howard, 2000; Howard and MacCready, 2004; Figs. 2 and 3). Relict and ghost stratigraphy of the metasedimentary remnants defines thrust slices and large recumbent folds. Upper and lower plates of the premetamorphic Ogilvie thrust fault are folded together around the gently north-plunging Lamoille Canyon nappe (Fig. 3; Smith and Howard, 1977; Snoke and Howard, 1984; Howard, 1980, 1987; Howard et al., 1979). The eastward vergence of this nappe is consistent with upper-crustal Cretaceous overthrusting in the Sevier orogenic belt to the east (fold nappes elsewhere in the core complex have different vergence directions; Howard, 1980; McGrew et al., 2000). The Lamoille Canyon fold nappe doubles the preexisting thrust-duplicated stratigraphy, with the result that Lamoille Canyon's walls expose four stacked alternations of gently dipping metamorphosed Prospect Mountain Quartzite and the marble of Verdi Peak (Fig. 3). Consistently oriented fabrics in the metasedimentary rocks show that even where granitic rocks dominate, isolated remnants of metasedimentary rocks are part of a coherent framework rather than free-floating, rotated rafts (Fig. 4).
Mesozoic and Cenozoic Evolution of the Core Complex
The folded Ogilvie thrust fault and Jurassic intrusions in the core complex record early Mesozoic events that preceded the pegmatitic granite (Table 1). Maximum pressures determined from the metamorphic assemblages increase northward from ∼4 kbar in the central Ruby Mountains to ∼9 kbar (and 700–800 °C) in the northern East Humboldt Range, likely reflecting greater tectonic burial to the north by stacks of nappes (Hodges et al., 1992; Snoke, 1992; Hudec, 1992; Jones, 1999; McGrew et al., 2000; Howard, 2003). Early Cretaceous(?) assemblages in the East Humboldt Range, estimated to record to 8–10 kbar, were interpreted to record deep tectonic burial (Hodges et al., 1992; McGrew et al., 2000; Howard, 2003). This metamorphism was followed by intrusion of pegmatitic granite and thermal overprinting of kyanite assemblages by lower-pressure sillimanite assemblages in a clockwise pressure-temperature-time (P-T-t) path (Hodges et al., 1992; McGrew et al., 2000). Cretaceous crustal anatexis that produced pegmatitic granite and other leucogranites in the core complex may have been a response to regional crustal thickening and/or to subsequent collapse and decompression, or possibly to cryptic mafic underplating that left no rock record (Miller and Gans, 1989; Barton, 1990; Hodges and Walker, 1992; Lee et al., 2003; Whitney et al., 2004; Wells and Hoisch, 2008; Rey et al., 2009; Miller and Snoke, 2009).
Peak metamorphic conditions in the rocks exposed in Lamoille Canyon accompanied intrusion of pegmatitic granite and resulted in muscovite–quartz–sillimanite–K-feldspar assemblages in semipelitic and granitic rocks. We infer peak pressures at ∼6–7 kbar (Hodges et al., 1992; Snoke et al., 2004). Deformation into fold nappes, including the Lamoille Canyon fold nappe (Fig. 3), likely occurred during the Late Cretaceous Sevier orogeny, although Eocene overprinting obscures the timing (MacCready et al., 1997; McGrew et al., 2000). Cretaceous regional contraction, possible extension, metamorphism, and intrusion of widely distributed leucogranites, including pegmatitic granite, were followed by quiescence until Eocene time.
Hornblende 40Ar/39Ar release patterns indicate either continued high temperatures into the Eocene or a renewed thermal spike to >500 °C, approximately coincident with Late Eocene gabbro to leucogranite intrusions, and possibly with some extensional unroofing (Wright and Snoke, 1993; McGrew and Snee, 1994; McGrew et al., 2000; Snoke et al., 2004; Premo et al., 2005). The compositionally expanded Eocene magmatism likely was driven by mafic underplating. Solid-state foliation and lineation parallel to fold nappes continued to develop during or after further granitic intrusion in Oligocene time (MacCready et al., 1997), possibly through deformation in response to crustal softening due to abundant melt (cf. Norlander et al., 2002).
The resulting gently dipping metamorphic-igneous infrastucture is deformed in a carapace hundreds of meters thick by a west-directed Tertiary mylonitic extensional shear zone that affects rocks as young as 29 Ma (Fig. 3). The Lamoille Canyon fold nappe and its limbs thin westward and upward into this mylonitic zone (Fig. 3). Cooling of the core complex progressed rapidly in Early and Middle Miocene time through fission-track annealing (zircon, titanite, apatite) and helium diffusion temperatures (250–50 °C), consistent with rapid extensional unroofing that placed fault slices of unmetamorphosed cover rocks onto the igneous-metamorphic complex and shed thick sediments into adjacent basins (Snoke, 1980; Dokka et al., 1986; Gifford et al., 2007; Colgan et al., 2010).
Jurassic, Cretaceous, Eocene, and Oligocene granitoids, largely gneissic granite, account for half or more of the rocks in the igneous- metamorphic complex (Table 1; Wright and Snoke, 1993; Hudec and Wright, 1991; MacCready et al., 1997; McGrew et al., 2000; Lee et al., 2003; Premo et al., 2005). Variably gneissic pegmatitic leucogranite of Late Cretaceous and Paleogene age, the subject of this report and here referred to as pegmatitic granite, forms the bulk of these intrusive rocks, occurring in countless sheets, dikes, and irregular bodies. In addition to this pegmatitic granite, the core complex exposes smaller amounts of two Jurassic peraluminous plutons (the granodiorite gneiss of Seitz Canyon [Fig. 2] and the muscovite granite of Dawley Canyon [Fig. 1]); widespread sheets of equigranular early Late Cretaceous peraluminous granite gneisses; Eocene sheets and dikes of gabbro, quartz diorite, granodiorite, and granite; the Eocene Harrison Pass pluton of granodiorite and granite (Fig. 1); and dispersed small bodies of Oligocene (29 Ma) biotite granite and associated late pegmatite and leucogranite dikes (Hudec and Wright, 1991; Wright and Snoke, 1993; Lee and Barnes, 1997; Snoke et al., 1997, 2004; McGrew et al., 2000; Barnes et al., 2001; Lee et al., 2003; Howard and MacCready, 2004; Premo et al., 2005). Thus in addition to the pegmatitic granite, other granitoids in Lamoille Canyon intruded ca. 161, 94–91, 40–37, and 29 Ma.
The Jurassic granodiorite gneiss of Seitz Canyon forms a deformed pluton as thick as 400 m in the core of the Lamoille Canyon fold nappe (Fig. 3). The pluton exhibits interlayering with bordering stratigraphic units and contains mapped enclave trains of quartzite and of calc-silicate rock suggestive of relict stratigraphy and structure (Howard, 1966, 2000).
Equigranular gneissic leucogranite bodies are ∼5–10 times less abundant than the pegmatitic granite. Even though the pegmatitic and equigranular gneissic leucogranites are commonly interleaved, locally appear gradational, and in each case include dated Late Cretaceous and Eocene members, the two textural rock types in Lamoille Canyon are isotopically different where measured (Wright and Snoke, 1993; Lee et al., 2003; Premo et al., 2005). For example, based on 21 total samples, values of epsilon Nd (εNd) are between –13 and –8 for Eocene and Cretaceous gneissic equigranular leucogranites, but between –15 and –18 for Eocene and Cretaceous gneissic pegmatitic leucogranites. Equigranular early Late Cretaceous two-mica gneissic leucogranites form small sheets in upper Lamoille Canyon (EG unit of Lee et al., 2003) and an extensive mapped sill, the granite gneiss of Thorpe Creek (both 91–94 Ma per J. Wright, 2002, written commun.; Premo et al., 2005; Figs. 2 and 3). The EG unit is cut by dikes of the gneissic pegmatitic leucogranite. The Thorpe Creek unit earlier yielded Eocene monazite U-Pb dates (Wright and Snoke, 1993; MacCready et al., 1997), but Premo et al. (2005) determined a 91 Ma zircon U-Pb age. Equigranular Eocene gneissic granite (Premo et al., 2005) forms additional leucogranite components in Lamoille Canyon. The various granite types in the area define subtly different and commonly overlapping compositional fields (e.g., Lee and Barnes, 1997).
The dominant gneissic pegmatitic leucogranite pervades the core complex. In Lee et al. (2003), this granite was described as the KPG unit. It is characteristically coarse to very coarse grained, foliated, and commonly contains sillimanite (Fig. 5). This foliated granite is distinguished from younger crosscutting pegmatite dikes that contain undeformed books of mica and cut cleanly across the gneissic foliation (Lee and Barnes, 1997; Howard and MacCready, 2004; Snoke et al., 2004). (The field distinction becomes uncertain for some coarse unfoliated leucogranite bodies that lack sillimanite.) The pegmatitic granite's very coarse texture distinguishes it from the less abundant equigranular leucogranite gneisses. Pegmatitic granite outcrops are nearly white in areas where metasedimentary remnants are calc-silicate rocks, and are slightly browner (from iron-oxide staining) in areas where small fractions (even <2%) of quartzite remnants are present. The pegmatitic granite typically consists of large (rarely to 60 cm) white K-feldspar crystals, oligoclase, and interstitial quartz (Fig. 5A). Some quartz is graphically intergrown with K-feldspar. Mafic minerals total 0%–4%. Low rock magnetic susceptibility averaging 0.02 × 10−3 SI (International System) units reflects a lack of magnetite. Biotite is more abundant than muscovite, and garnet is commonly an accessory phase. Garnet textural relationships are consistent with either an igneous or metamorphic origin. Metamorphic sillimanite occurs as needles and mats included within biotite and less commonly muscovite. Foliation is defined by oriented biotite and muscovite and by zones of granoblastic feldspar and quartz. Mica wraps around coarse feldspar and quartz crystals. Ductile deformation features in the infrastructure (the region below the mylonitic carapace) include annealed, irregular gneissic foliation (Fig. 5). In the mylonitic shear zone carapace, the granite shows various degrees of dynamic recrystallization and a strongly developed foliation and associated elongation lineation (H79Ruby-14 in Fig. 5A).
The pegmatitic granite intimately intrudes older granitoids and the sedimentary host rocks (Table 2), forming ∼40%–50% of the exposed rocks in the complex (Figs. 4 and 6). Outcrops of it can be found in nearly every square kilometer of exposed crystalline rock in the northern Ruby Mountains and East Humboldt Range. It decreases in abundance upward on the canyon walls and increases downward into canyon bottoms (Figs. 6B, 6C). On many ridge tops at high structural levels, pegmatitic granite mostly comprises less than one-third of the rock compared to >two-thirds host metasedimentary rocks; only occasional pods of pegmatitic granite occur at some structurally high exposed levels. At the deepest exposed levels, in 1-km-deep Lamoille Canyon and other deep canyons and to the south in the central Ruby Mountains, mapped gneissic pegmatitic granite occupies 90%–100% of the rocks, and metasedimentary remnants are rare or absent (Figs. 2 and 3; Howard et al., 1979). This resulting structurally deep pegmatitic granite massif exhibits biotite-rich banding in places, interlayered with subordinate finer grained gneissic granite, and the rare stratigraphic remnants (Fig. 7D). In calling this plutonic body a massif, we emphasize its indistinct upward transition into an injection complex and its ghost stratigraphy of metasedimentary relics, which together define a complete gradation into the overlying injection complex.
The cross section (Fig. 3) illustrates the proportional distribution of pegmatitic granite along the walls of Lamoille Canyon. We estimate a volume of ∼600 km3 of pegmatitic granite projected through ridges and across canyons in the Ruby and East Humboldt ranges. Unexposed deeper volumes may be several times that. Pegmatitic granite 20 km to the south of the area of Figure 2 underlies a mapped area of ∼120 km2, projects under the area of Figure 2, and could correspond to a thickness of several kilometers (Howard et al., 1979). Pegmatitic granite is concentrated with depth and structurally updip to the south, even though geobarometry suggests that greater burial depths are encountered northward, probably under stacks of nappes (Howard, 2003).
The pegmatitic granite was previously considered Jurassic (Howard et al., 1979; Kistler et al., 1981), but a sample was dated as ca. 85 Ma by U-Pb on zircon in the East Humboldt Range (McGrew et al., 2000). An age of ca. 83 Ma in Lamoille Canyon, based on a preliminary monazite U-Pb (thermal ionization mass spectrometry, TIMS) age (Wright and Snoke, 1993), formed the geochronological framework for the interpretation in Lee et al. (2003) that this type of granite originated when Late Cretaceous tectonic thickening of the crust led to anatexis of underlying pelitic rocks. Subsequent U-Pb dating of monazite and zircon grains from the same outcrop by ion probe, as described in the following, shows a range of Late Cretaceous crystallization ages and a younger likely intrusive age. Our geochronology also expands the range of crystallization ages in pegmatitic granite into the Paleogene.
Composition and Source
The pegmatitic granite is peraluminous to strongly peraluminous, ilmenite series leucomonzogranite. Major element variation diagrams of the pegmatitic granite gneiss lack clear fractionation trends (Lee et al., 2003). Rare earth element (REE) and Sr/Y ratios typically are distinct from younger granitic rocks in the area and indicative of residual garnet in the source or left behind by fractional crystallization (Lee and Barnes, 1997; Lee et al., 2003; Batum, 1999). Trace element variations, particularly REE patterns, suggest local feldspar accumulation elsewhere (possibly residual) giving rise to fractionated residual magmas with negative Eu anomalies (Lee et al., 2003). Values of δ18O typically +10 to +12, εNd(t) values of –15 to –18, and initial 87Sr/86Sr values of 0.715–0.725 together suggest that the pegmatitic granite magmas were derived from partial melting of metasedimentary rocks lacking much residual feldspar (Kistler et al., 1981; Lee et al., 2003). The older EG equigranular two-mica gneissic granite of upper Lamoille Canyon was interpreted to have similar origins, but its higher εNd (∼–9) and lower initial 87Sr/86Sr (∼0.710) indicate a distinct crustal source (Lee et al., 2003). Lee and Barnes (1997) found that in comparison to younger, Oligocene crosscutting pegmatite dikes, the pegmatitic granite has higher Sr, lower Nb +Y, and overlapping but generally lower Rb, lower A/CNK [Al2O3/(CaO + Na2O + K2O)], and higher Zr values.
The exposed metacarbonate and quartzite stratigraphic hosts and their protoliths are inappropriate as major sources for granitic magmas. Small proportions of granitic fluid could have been derived from the impure quartzite host rock through dehydration reactions, as muscovite content decreases in this rock type with increased metamorphic grade, while sillimanite and K-feldspar contents increase (fig. 8 in Howard, 1966). Andesine-quartz-biotite ± muscovite-sillimanite schist, rarely containing garnet, forms only a few percent of the section and so is unsuitable as a major source rock. The isotopically distinct equigranular granites could not have been melt sources. Thus the near lack of suitable exposed host rocks and lack of residual mineral assemblages indicate that initial anatexis that produced pegmatitic granite melts mostly originated below the level of exposure (Lee et al., 2003)
In Lee et al. (2003), it was suggested that the pegmatitic granite is the result of muscovite-dehydration partial melting, on the basis of major element contents of the rocks and low estimated magmatic temperatures (620–775 °C estimated from the zircon saturation thermometer; Watson and Harrison, 1983). These low magmatic temperature estimates are consistent with the results of Ti-in-zircon thermometry (see following), which yielded temperature estimates <700 °C. The nonequigranular, coarse-grained texture of the pegmatitic granite may be the result of undercooling in H2O-rich environments (Nabelek et al., 2009).
Lee et al. (2003) suggested that the source rocks were pelitic parts of the Neoproterozoic McCoy Creek Group that stratigraphically underlies the strata exposed in the Lamoille Canyon area. Rocks that likely correlate to the McCoy Creep Group in the Ruby Mountains north of Lamoille Canyon consist of impure quartzite and local metacarbonate rocks; pelitic rocks are rarer than in protolith equivalent strata, so partial melting may have consumed them. In the East Humboldt Range, schist derived from a McCoy Creek Group protolith reached P-T conditions (∼9 kbar, 750 °C) at or above the muscovite stability field (McGrew et al., 2000; Lee et al., 2003). In contrast to the pegmatitic granite, the early Late Cretaceous equigranular two-mica granite gneiss of upper Lamoille Canyon was instead modeled as melted from even deeper levels at higher temperature by biotite-dehydration partial melting of rocks such as metagraywacke, that may be present in Paleoproterozoic basement (Lee et al., 2003).
Intrusion Style and Relation to Structure
Pegmatitic granite forms a migmatitic latticework of thousands of sheets, lenses, and discordant bodies intimately injected into marble, calc-silicate rocks, quartzite, older granite and granodiorite, and subordinate schist. Subhorizontal sheets, sheet-dike networks, and irregular bodies range from a few centimeters thick to sills as thick as 100 m (Figs. 4–8). Some sill-like sheets as thick as 50 m can be traced at least 2 km with little change in thickness. Some sills merge with crosscutting dikes (Figs. 6E, 6F).
Some of the pegmatitic granite bodies cut folds and foliation in the host rocks, other bodies are concordant to intricately folded layering in the metasedimentary hosts, and some folded sills merge with oblique dikes that cut across the same folds (Figs. 6E, 6F). Some folded sills are boudined, while others include ballooned parts where the timing of intrusion relative to folding may be ambiguous (Figs. 6E–6G). The pegmatitic granite in places exhibits folded gneissic foliation or it encloses folded dikes of older equigranular leucogranite. The axes of mesoscopic folds in the infrastructure in the Lamoille Canyon area trend consistently north-northeast parallel to well-mapped hinge lines of the Lamoille Canyon nappe (Howard, 1966, 1980). Except in the mylonitic shear zone, mineral lineation (Li of MacCready et al., 1997, where the subscript i refers to infrastructure) in both the metasedimentary host rocks and in the granite parallels these mostly northward-trending mesoscopic folds parasitic to the nappe (Howard, 1980, 2000; Howard and MacCready, 2004). Together these structural features indicate that intrusion of pegmatitic granite overlapped in time with folding parasitic to the nappe.
The proportion of pegmatitic granite increases downward across the Lamoille Canyon fold nappe, so that the upper limb exposes smaller amounts of pegmatitic granite than the lower limb (e.g., Fig. 6B). A similar downward increase of pegmatitic granite across the Winchell Lake fold nappe in the East Humboldt Range led to the interpretation that intrusion in part postdated that fold, and in part predated the fold because the granite there is involved in some of the folding (McGrew et al., 2000). These are the same relations as those shown by pegmatitic granite in the Lamoille Canyon fold nappe, and suggest protracted intrusion during folding.
Intrusive contacts of the pegmatitic granite against host metasedimentary rocks are mostly sharp. Metamorphic or magmatic fluids were present during emplacement, as shown by local development of grossularite-idocrase-diopside-scapolite-epidote skarns as much as a few meters across. In addition, Jeon (1999) demonstrated locally pervasive oxygen isotopic exchange between calc-silicate metasedimentary rocks and meter-scale pegmatitic granite sheets. This is in contrast to more limited isotopic exchange of ∼10 cm length scales between equigranular Paleogene monzogranites and their host rocks (Jeon, 1999). Metasomatism in host rocks is further suggested by the presence of feldspar and late-stage sodic scapolite in calc-silicate marbles, and greater content of diopside and feldspar in small isolated remnants of calc-silicate rocks than in larger calc-silicate masses. In deep canyon bottoms, gneissic pegmatitic granite volumetrically dominates (95%–99% of the rock). Contacts are sharp against rare calc-silicate relics, but the pegmatitic granite may be interlayered lit-par-lit with quartzite remnants on a scale of centimeters (Fig. 6C), or else surround isolated quartzite enclave layers a few centimeters thick or wispy biotite-rich pelitic seams a few millimeters thick.
The pegmatitic granite may form concordant and bulbous sills intruded into metacarbonate strata yet locally form dike networks into an adjacent quartzite host (Fig. 6D). This contrast in style suggests that at the time of intrusion the quartzite was stiffer and sometimes was injected by the granite through brittle cracking or hydrofracturing in contrast to more ductile metacarbonate host rocks, in which the granite commonly intruded as ballooning sills along lithologic layering (cf. Weinberg and Searle, 1998). Concordant sheets may merge with or be truncated at adjoining dikes of the gneissic pegmatitic granite with no discernable lithologic contact between them. Gneissic pegmatitic granite is interlayered with and locally crosscuts Jurassic granodiorite gneiss of Seitz Canyon (Howard, 1966), as well as small bodies of early Late Cretaceous equigranular gneissic two-mica granite (EG unit in Lee et al., 2003).
Except at the deepest structural levels where the granite greatly dominates over metasedimentary rocks, higher proportions of pegmatitic granite commonly are present in calcareous host-rock units compared to adjacent quartzite host rocks. For example, networks of sills and thick dikes of pegmatitic granite may constitute 60% of a mapped metacarbonate formation at intermediate structural levels but only 30% of adjacent mapped Prospect Mountain Quartzite. The concentration of pegmatitic granite in the metacarbonate formation host raises the question whether some attribute of the metacarbonate rocks, such as CO2 flux or ductile rheology, may have enhanced granite intrusion.
Interpretation of Intrusive Sequence
The abundance of small granite bodies together with the preservation of mapped relict stratigraphy, structure, and coherent fabric patterns in the host-rock relics and the granite suggest that assembly of the pegmatitic granite network was by incremental intrusion of many granitic sheets. The evidence of intrusion before, during, and after ductile nappe-related folding indicates that intrusion overlapped the time of folding. The spatial progression from networks of sparse granite toward granite networks having a higher proportion of granite and coalescing granite bodies can be interpreted to record successive stages of assembly of small granite bodies into the final network and massif of batholithic proportions. We infer that early stages of intrusion networks resembled the networks of relatively sparse granite seen at high structural levels, and that progressively more advanced stages in which more granite bodies had accumulated resembled denser networks of granite seen at lower levels.
Lack of internal contacts in granites associated with relict country-rock structure has led to varied interpretations of intrusion (Pitcher and Berger, 1972,;Mahan et al., 2003). Among other possibilities, a lack of obvious contacts could be consistent with the presence of mush zones of partial melt for sustained or recurring intervals during assembly of the many sheets (cf. Matzel et al., 2006).
The field evidence that folded sills merge elsewhere into dikes that cut the same folds could be explained if emplaced magma underwent simultaneous ductile and brittle deformation, or if deformation resulted in removal of residual magma to form a crosscutting dike. The field evidence alternatively is consistent with the possibility that the pegmatitic granites reached or maintained partial melting conditions and intrusive mobilization more than once.
Making Room for the Granite
The network of pegmatitic granite implies that the intrusions wedged into and inflated the host rocks in this part of the crustal column. A possible analog for this kind of inflation may be the splaying of ghost stratigraphy defined by enclave trains in the Main Donegal Granite pluton into wider zones than in the host rocks, as mapped (but not then so interpreted) by Pitcher and Read (1959). The intrusion complex represented by the Harney Peak Granite and associated pegmatites in South Dakota is another analog (Nabelek et al., 1999; Redden and DeWitt, 2008). Inflation of the host-rock structure in the Ruby Mountains by injection of primarily subhorizontal sheets may have been accompanied by downward displacement of the lower part of the framework as partial melt was transferred from deeper crustal levels (cf. Cruden, 1998; Wiebe and Collins, 1998; Brown and McClelland, 2000). Intrusion overlapped the timing of fold nappe development, and probably also southward flow parallel to the fold nappe axis, and possibly westward structural shear and thinning related to Paleogene crustal extension (MacCready et al., 1997).
Pegmatitic granite was sampled for SHRIMP-RG (sensitive high-resolution ion microprobe–reverse geometry) geochronology at locations shown in Figures 2 and 3 and listed in Table 3. The geologic context of each sample is summarized in Figure 7A. Samples, described in Appendix 1, were 2–5 kg, and appeared homogeneous. High-resolution U-Pb geochronology on separates of zircon and monazite from the pegmatitic granite used the SHRIMP-RG at the Stanford University–U.S. Geological Survey laboratory.
Zircon or monazite grains separated from the rock samples were mounted in epoxy and polished. Zircons were imaged by cathodoluminescence (CL) and monazites by backscattered secondary electrons (BSE). Probe spots (∼15–30 μm across) were selected based on CL zoning, and avoided internal boundaries. In addition to radiogenic daughter products of the U-Pb system, on two samples we collected data on REEs and certain other trace elements on zircon spots separate from those studied geochronologically. SHRIMP-RG data reduction used the SQUID 1 program and Isoplot developed by K. Ludwig (http://www.bgc.org/isoplot_etc/software.html, accessed 2006 July). Isotope ratios were standardized against zircon standard R33 (419 Ma; Black et al., 2004) and monazite standard Wendall (301 Ma; P. Holden, 2000, written commun.) for age, U, and Th. The trace element concentrations (for samples H04Rby113 and H04Rby107, Supplemental Table 11) were standardized against gem-quality zircon standard CZ3. The CZ3 trace element concentrations were determined by comparison to doped synthetic zircons analyzed by electron probe (Mazdab and Wooden, 2006).
U-Pb isotopic ratios are listed in Table 4 and dates are summarized in Figure 9. Ages presented are corrected for common Pb using the measured 207Pb/206Pb ratio (the Tera-Wasserburg method). High U contents led to high precision for most results, with 1ó errors ranging from 1.7 to 0.1 m.y. and most being <0.5 m.y.
Pegmatitic Granite Gneiss in Upper Lamoille Canyon (WRP02-7, RM-4)
Gneissic pegmatitic leucogranite samples WRP02-7 and RM-4 were collected separately within 1–2 m of each other from a roadside outcrop near The Terraces picnic area. The granite here contains mesoscopically visible sillimanite and monazite in addition to muscovite and biotite. The pegmatitic granite sheet, several meters thick, obliquely cuts layered and lineated calc-silicate rocks and fine- to medium-grained biotite-muscovite leucogranite gneiss. A steep pegmatite dike 20 m away cuts the dated granite (Fig. 7A). Nd, Sr, and Pb isotopic values for sample WRP02–7 are essentially identical to the value reported for RM-4 by Wright and Snoke (1993), confirming that the same rock was sampled. Zircons separated from WRP02–7 and monazites separated from RM-4 were investigated using the SHRIMP-RG.
The CL images of the zircons are shown in Figure 10. In transmitted light, the grains are subhedral to anhedral, look clear to cloudy, and have distinct areas of possible inherited cores and overgrowths. The CL images show that most prismatic grains consist of an inner, mottled, in places convoluted, core surrounded by a concentrically zoned outer region typical of magmatic zoning.
SHRIMP zircon data scatter along concordia with most 206Pb/238U ages from ca. 90 to 67 Ma (Fig. 11A). A concentration of 10 dates ca. 69 Ma (including 2 associated with U > 5000 ppm; Fig. 11B) yield a weighted mean age of 69.1 ± 0.6 Ma (mean square of weighted deviates = 1.6; Fig. 11C). Most of the zircon analyses with 206Pb/238U dates of ca. 69 Ma are from concentrically zoned outer regions (Fig. 10). A discordant analysis is at 37 Ma. U concentrations of non-core regions range from ∼2000 to nearly 14,000 ppm, and Th concentrations are low, all but one <100 ppm (Table 4). The zircon dated as older than 70 Ma mostly has higher U content than younger zircon. Some dates, including many older than 80 Ma, are associated with U contents >5000 ppm, for which the instrument is not well calibrated and the ages uncertain and maximums. All data exhibit Th/U values below 0.02 (Fig. 11D).
Earlier preliminary U-Pb dating of monazite in RM-4 using TIMS indicated a Pb-U age of ca. 83 Ma (Wright and Snoke, 1993). Figure 12A presents data on monazites from RM-4 investigated using the SHRIMP-RG. The monazite data are above concordia with 206Pb/238U dates ranging between 80 and 90 Ma, centered at 83–84 Ma. The spread of monazite dates resembles part of the spread of pre–70 Ma zircon dates in the adjacent sample WRP02–7. Monazite may be less or more refractory to age resetting than zircon, and in some cases yields ages that apparently represent neither the time of crystallization nor of cooling (e.g., Carr, 1992; Kapp et al., 2002).
Folded and Crosscutting Pegmatitic Granite in Central Lamoille Canyon (Rby113 and Rby 111)
Two samples of gneissic pegmatitic leucogranite were collected from bodies 60 m apart at the base of a cliff 300 m above the floor of central Lamoille Canyon, 4.5 km northwest of WRP02–7 (Fig. 6D). The granites intrude Cambrian calc-silicate rocks and marble near their folded contact against structurally inverted older metamorphosed Prospect Mountain Quartzite on the lower limb of the Lamoille Canyon fold nappe. Mesoscopic folds Fi in the metasedimentary rocks have axes that trend parallel to this nappe, and their S shapes are parasitic to the nappe. The two sample sites were chosen to compare folded granite with crosscutting granite.
Sample Rby113 is from a bulbous sill of gneissic pegmatitic granite intruded into and folded with its host rocks of layered calc-silicate rocks and marble (Figs. 7A, 7B). Rby111 is from a body of gneissic pegmatitic granite that cuts the folded contact of the same calc-silicate-marble unit against metamorphosed Prospect Mountain Quartzite (Figs. 7A, 7C), and merges higher on the cliff face with sills concordant to folded layers in the host calc-silicate unit. Zircon grains from Rby113 and Rby111 were dated.
The grains in both samples mostly show euhedral terminations and aspect ratios typical of igneous morphology (Figs. 13A, 13B). CL images show oscillatory zoned rims and tips to the zircons typical of magmatic crystal growth, and cores that are mottled or speckled. Some crystal interiors also show oscillatory zoned chevrons over a mottled core and separated from the zoned rim also by a mottled zone. A few grains have slightly irregular external shapes and may be corroded.
Spot analyses by SHRIMP-RG on zoned rims and tips of zircons from the folded granite Rby113 (Fig. 13A) yielded concordant 206Pb/238U dates that range from 103 to 35 Ma (Figs. 14A, 14B). U contents are high (Fig. 14C). Excluding ages queried here associated with U contents >5000 ppm, a cluster of 6 dates occurs ca. 92–93 Ma and a cluster of 4 dates ca. 36 Ma. Mottled zones of grains yielded 206Pb/238U ages of 35, 37, 75, 81, 93, and 101(?) Ma. One zoned grain yielded a Cretaceous date, and another zoned grain yielded a Paleogene date. Some grains yielded Cretaceous dates from both a mottled core and zoned rim. Zoned rims yielded 206Pb/238U ages of 36, 37, 91, 92, 98, 98(?), 101(?), and 103(?) Ma. The Th/U ratios are all <0.02 for the spots that yielded Cretaceous dates, but range to more than twice that value for the spots that yielded Paleogene dates (Fig. 14D). Some uncertain dates associated with high U exceed 100 Ma. Ignoring these, Rby113 is notably polymodal with dominant date clusters ca. 92 Ma and 37 Ma.
Spot analyses by SHRIMP-RG on zoned rims and tips of a smaller set of zircons selected from the crosscutting granite sample Rby111 also show a wide range of Late Cretaceous and Paleogene 206Pb/238U dates from 97 to 32 Ma, exclusive of 2 questionable data for which U > 5000 ppm (Figs. 9, 13B, 14E, and 14F). A cluster of 5 dates at a mean of 69 Ma resembles the 69 Ma cluster in WRP02–7. The Rby111 data thus resemble the data in WRP02–7, which also cuts host-rock layers and fabric.
Pegmatitic Granite Gneiss in Lower Lamoille Canyon (Rby107 and RM-27)
Pegmatitic leucogranite gneiss was sampled at two sites downstream in Lamoille Canyon (RM-27 and Rby107). These sites are in structurally low levels where gneissic pegmatitic leucogranite constitutes >90% of the rocks, and the geologic map portrays pegmatitic granite rather than stratigraphy (Fig. 3; Howard, 2000). Rare thin quartzite enclaves, relics of the Prospect Mountain Quartzite, constitute <1% of the rock near both sample localities (Fig. 7A).
The pegmatitic granite here lithologically resembles the other dated pegmatitic granites, even though it is geochronologically distinct, as shown in the following. Using geochemical parameters defined for granitoids in the Ruby Mountains in Lee and Barnes (1997), the evolved composition (75%–76% SiO2) overlaps geochemically with other members of the pegmatitic granite suite and with evolved granites of the Eocene Harrison Pass pluton (Lee and Barnes, 1997; Barnes et al., 2001). It differs chemically from most Paleogene biotite monzogranite bodies in Lamoille Canyon such as those studied by Wright and Snoke (1993) and Snoke et al. (2004).
Sample Rby107 is from a horizontal sheet of locally sillimanite-bearing, gneissic pegmatitic leucogranite (Fig. 7D). The gneissic pegmatitic granite is intruded 3 m above the sample site by a concordant sill of biotite monzogranite 5 m thick that closely resembles rocks dated elsewhere in the complex as 29 Ma (Wright and Snoke, 1993; MacCready et al., 1997). That sill in turn is cut by small dikes of leucogranite. As in the other samples, zircon grains in Rby107 have euhedral shapes that are typical of magmatic grains, and contain mottled zones and oscillatory zoned rims and tips (Fig. 15).
Zircons in Rby107 analyzed by SHRIMP-RG (Fig. 15) show a range in 206Pb/238U dates from 46 to 27 Ma (Figs. 16A, 16B). Three of the four oldest dates and one other are unreliable as the spots contain >5000 ppm U (Fig. 9). Sample Rby107 zircon dates are widely spread but roughly may be grouped into an older population of 9 dates scattered around ca. 38–39 Ma, and a younger population of 10 dates scattered around ca. 29 Ma. The Th/U ratios in the zircon increase with Paleogene age, not unlike an increase with Cretaceous age in WRP02–7 (Fig. 11D). The pattern of increasing Th/U with age is opposite to the pattern in Rby113 where older grains, dated as Cretaceous, have lower Th/U ratios than younger grains, dated as Paleogene.
Sample RM-27 is sillimanite-muscovite-biotite pegmatitic leucogranite gneiss from a roadcut 3 km to the southeast of Rby107. Subordinate equigranular finer-grained gneissic granite interlayers at this site had been dated by U-Pb as 29 Ma (monazite sample RM-15a of Wright and Snoke, 1993; gneissic biotite monzogranite), and 32–40 Ma (zircon sample WRP02–6 of Premo et al., 2005; garnet- muscovite-biotite gneissic leucogranite). Monazite grains separated from RM-27 analyzed by SHRIMP-RG yielded 206Pb/238U ages of 41–37 Ma (Fig. 12B), for which we calculate a mean age of 39.4 ± 0.9 Ma. The ages intersect concordia at 37.1 ± 0.5–36.8 ± 0.8 Ma, depending on which data are included. The discordant older 206Pb/238U ages project distantly toward upper intercepts ages ca. 1.9 or 2.7 Ga, suggestive of possible inheritance from monazite comparable in age to Paleoproterozoic and Neoarchean zircon xenocrysts found in nonpegmatitic granitoids in the core complex (Wright and Snoke, 1993; Premo et al., 2005, 2008, 2010). Note that unlike the other dated members of the pegmatitic granite suite, samples Rby107 (zircon) and RM-27 (monazite) record Paleogene U-Pb dates and yielded no analyzed Cretaceous zircon or monazite.
Zircon dates show considerable scatter (Fig. 9), even though individual age determinations have high precision (Table 4). The data in combination suggest at least four age peaks. Cores and rims give similar dates in several samples, and zircons that yielded various Cretaceous dates show no characteristic difference in morphology compared to those that yielded various Paleogene dates. Dates peak ca. 92 Ma, 68 Ma, 38–36 Ma, and 29 Ma. Three of these peaks (ca. 92, 38–36, and 29 Ma) approximate the zircon crystallization ages reported for other granitoids in the core complex (Fig. 9). The dates for monazite in RM-27 are similar to the 38–36 Ma zircon age group. Intermediate Paleogene and Late Cretaceous dates are also present in our data. Dates are nearly absent from a gap between 65 and 46 Ma.
An age of ca. 69 Ma for much of the zircon in WRP02–7 is matched in the more sparsely sampled Rby111. Both of these samples contain a spread of (possibly inherited) older Late Cretaceous zircon dates and evidence of Paleogene crystallization or disturbance. Pre–69 Ma Late Cretaceous monazite in the WRP02–7 and RM-4 outcrop also suggests possible inheritance from an early Late Cretaceous event. Note bimodal zircon populations in Rby113, with peaks ca. 92 Ma and 37 Ma, and in Rby107, with age peaks centered ca. 38–39 and 29 Ma. The great range in Late Cretaceous to Paleogene U-Pb dates in the pegmatitic granite suite both at single samples and between different sites suggests recurrent and perhaps prolonged conditions of zircon growth or disturbance.
Late Cretaceous and Paleogene zircon and monazite age clusters from the pegmatitic granite mirror the combined dates from other, less voluminous granitoids in the core complex (summarized in Fig. 9). In contrast to evidence of inherited Proterozoic and Archean zircon in other granitoids in the Ruby–East Humboldt range core complex, we found no pre- Cretaceous zircon in the pegmatitic granite.
Zircon Geochemical Data
Geochemical data for the zircons (listed in Supplemental Table 1 [see footnote 1]) exhibit several characteristics that are unusual (Table 5) compared to most magmatic zircons (including zircons in Jurassic granodiorite and Eocene gabbro in Lamoille Canyon). Zircon samples Rby113 and Rby107 were the most completely analyzed in the pegmatitic granite suite.
Low Th/U values in zircon commonly are associated with metamorphic disturbance or crystallization from anatectic melt. The very low Th/U values, below 0.05, for the dated zircons accompany unusually high U contents (Fig. 17A). The Th/U values show systematic differences between age groups found within individual samples (Fig. 17A).
The Ti content of zircon allows crystallization temperatures to be estimated (Watson and Harrison, 2005; Watson et al., 2006; Hayden and Watson, 2007; Ferry and Watson, 2007). Temperatures (uncorrected for TiO2 and SiO2 activities) calculated using this Ti geothermometer for zircons in both Rby113 and Rby107 are mostly between 700 and 600 °C (Fig. 17B). The Hf values are very high, 13,000–26,000 ppm (Fig. 17B).
The Yb/Gd has a negative correlation with Th/U in a suite of granitic zircons where compositional change is related to melt fractionation at the time of zircon growth (Claiborne et al., 2010). The Yb/Gd values for zircons in the pegmatitic granite are high, indicative of relatively depleted middle–heavy REEs (Fig. 16C); Yb/Gd plotted versus Th/U (Fig. 17C) from Paleogene rims and grains in Rby113 cluster with Paleogene grains from Rby107, as opposed to higher Yb/Gd and lower Th/U in Cretaceous cores in Rby113. The age-dependent clustering suggests that two events of zircon crystallization in each of these samples involved distinctly different melts or fluids.
Metamorphic zircons typically exhibit stoichiometric balance among +3 REE cations and (+5) P, whereas magmatic zircons typically exhibit excess REEs in this measure, indicating charge balance by other elements (Hoskin and Schaltegger, 2003). In Figure 17D, a plot of molar +3 cations (REEs and Sc) divided by (+5) P clearly distinguishes Tertiary (younger than 55 Ma) rims and grains in Rby113 as >1, not stoichiometrically balanced. Cretaceous grains in the same sample, and Paleogene grains in Rby107 (all grains), instead show a value near 1, indicating charge balance.
The Eu contents are low in zircons from Rby113 and Rby107 compared to other REEs (Fig. 17E). The depth of the Eu anomaly tends to be slightly greater for Rby113 than for Rby107.
Lack of Ancient Inherited Zircon Cores
The lack of dated inherited pre–Late Cretaceous cores in zircon in the pegmatitic granite contrasts with nonpegmatitic Paleogene and Cretaceous leucogranites and most other dated granitoids in the Ruby–East Humboldt Range complex, which typically include Proterozoic and/or Archean zircon cores or U-Pb evidence of inherited old zircon (Wright and Snoke, 1993; Premo et al., 2005, 2008). A sampling bias cannot explain the lack in the pegmatitic granite of pre-Cretaceous zircon cores either observed or age measured in our 73 geochronologic analyses and age estimated in another 49 probed zircon spots (Supplemental Table 1 [see footnote 1]). A lack of ancient inheritance could relate to different melt sources compared to the other granites, to some method of rise through the crust that filtered or did not transport zircon, or to Cretaceous melting conditions in which older zircon did not survive. Pegmatites typically lack xenocrysts, due perhaps to ultraefficient extraction of pegmatite-forming melts from their source regions (London, 2005). However, that process would not predict the presence of Cretaceous and Paleogene zircon antecrysts and/or xenocrysts in our pegmatitic granite. If muscovite-bearing pelitic rocks were the initial anatectic source for the pegmatitic granite suite (Lee et al., 2003), and for Cretaceous and Paleogene granite forerunners, a likely protolith for the melt source would be a clay-rich mudstone, in which any clastic zircon may have been too fine grained to survive melting. In contrast to zircon, the projection of monazite data for RM-27 toward Paleoproterozoic and Neoarchean upper intercepts (Fig. 12B) suggests a possible inheritance signal in the monazite.
Zircon Crystallization Conditions and Compositional Constraints on Source
Grain morphology, oscillatory zoning, and geochemical evidence discussed in the following suggest that most of the dated Cretaceous and Paleogene zircons are low- temperature magmatic crystals, even though many features of the zircons are unusual for magmatic zircon (Table 5). The unusual mottling in many grains we speculate may relate to magmatic growth of high-U zircon in low-temperature, water-rich melts. Low calculated crystallization temperatures, mostly <700 °C (Fig. 17B), are consistent with the highly evolved bulk compositions of the leucogranites, and with the low calculated zircon saturation temperatures for the pegmatitic granite reported in Lee et al. (2003).
Paleogene zircons in Rby113 have a typically magmatic signature of variably high values of the ratio of molar +3 cations to +5 cations, expressed as (REE + Sc)/P (Fig. 17C). Stoichiometrically balanced values near 1 for Cretaceous zircons in Rby113 and Paleogene zircons in Rby107 are unusual for magmatic zircons, and we speculate that they may relate to crystallization in water-rich environments.
Systematic differences in Th content and in Th/U between different groups of analyses dependent on age in individual samples imply real compositional differences of liquid in equilibrium with the zircon at different times of crystallization. These differences distinguish ca. 69 Ma zircon rims from older Cretaceous cores in WRP02–7 (Fig. 11D), Cretaceous zircon from Paleogene zircon in Rby113 (Fig. 14D), and Oligocene from older Eocene zircon in Rby107.
Low Th/U and high Yb/Gd ratios in the zircons are consistent with crystallization from highly fractionated, low-temperature, hydrous, oxidized magmas. The low zircon crystallization temperatures, very high zircon U (Fig. 17B), bulk leucogranite composition, and pegmatitic textures are also consistent with H2O-rich, highly evolved, low-temperature magmas. High Hf values in the zircons (Figs. 17C, 17E), indicative of highly evolved melt compositions, suggest that the zircons were in equilibrium with highly fractionated magmas, and that residual zircon may have been left in the source. The very low values for Eu/Eu* in zircon (Fig. 17E) probably reflect strong negative Eu anomalies in the melt. The Eu anomalies suggest that feldspar crystallized before zircon, or that residual feldspar was left in the source.
The unusual characteristics of the zircons in the pegmatitic granite offer numerous opportunities for future research that could advance the understanding of melt extraction and crystallization (Table 5). The high and variable U contents in the zircon, the unusual mottled character that many zircon grains exhibit, and the lack of Precambrian zircon xenocrysts all deserve further study.
Ages of Crystallization
The geochronologic data suggest a history of melt generation and crystallization of pegmatitic granite multiple times in the Late Cretaceous and Paleogene. Before discussing the timing of emplacement of the granite bodies ages, we first address times of zircon and monazite crystallization.
Crystallization of igneous-appearing zircon occurred in at least four episodes (ca. 92, 69, 38, and 29 Ma), and three of these match other dated igneous events in the core complex (Fig. 9). Early Late Cretaceous ages of ca. 92 Ma (Rby113) approximately match the age of granite gneiss of Thorpe Creek (Premo et al., 2005) and equigranular granite gneiss in upper Lamoille Canyon (J. Wright, 2002, written commun.). Monazite in RM-4 and some zircon dates near 83 Ma may either be mixed ages (from beam overlap of more than one crystal zone) or suggest another possible episode of melt generation and crystallization ca. 83 Ma, an age of granite intrusion also reported from U-Pb zircon dating in the East Humboldt Range (McGrew et al., 2000; Premo et al., 2008, 2010). An age of 69 Ma for WRP02–7 and Rby111 does not match previously known events. Eocene crystallization ages of 39–36 Ma (Rby107, RM-27, Rby113) coincide in age with (1) other widespread intrusions in the core complex ranging in composition from gabbro to granite, (2) the reheating suggested by 40Ar/39Ar thermochronology, and (3) regional volcanism in northeastern Nevada. The youngest cluster of zircon ages in a pegmatitic granite sample (ca. 29 Ma in Rby107) matches the time of intrusion of small biotite monzogranite bodies in the core complex (Wright and Snoke, 1993). The data indicate crystallization of leucogranite magmas into pegmatitic granite, or into igneous forerunners that were sources of entrained zircons, at least four times during Late Cretaceous and Paleogene crustal melting events. Large temporal spreads of zircon dates between these ages may reflect episodes or long intervals of magmatic crystallization in intermittently reinvigorated igneous systems.
Ages of Intrusion
Emplacement ages can be interpreted in more than one way for the pegmatitic granite bodies that exhibit bimodal and scattered zircon age patterns. A wide range of intrusion ages for the pegmatitic granite would be consistent with the core complex's long history of peraluminous granite intrusion from Jurassic until Oligocene time (Table 1). We cannot exclude the possibilities that the sampled granite bodies are all Paleogene (requiring that inherited zircon grains dominate most of the samples), or that they are all early Late Cretaceous (requiring abundant to dominant metamorphic or age-reset grains). However, we favor the interpretation that the sampled bodies were intruded at a variety of different times, reflected in dominant U-Pb crystal ages.
We interpret an emplacement age of 69 Ma for the WRP02–7 and RM-4 site and the Rby111 site, on the assumption that the few Eocene and Paleocene zircon dates at those two sites represent resetting during Paleogene reheating, and that abundant 70–86 Ma and older zircon and 80–90 Ma monazite are inherited. Rby113 can be interpreted as intruded at 36 Ma with a large population of inherited 92 Ma zircon, or alternatively intruded at 92 Ma, with some new zircon growth from magmatic crystallization in situ at 36 Ma. The latter interpretation seems more consistent with the field relations suggesting that at least parts of the Rby113 granite body predate Fi folding that in turn is cut by nearby (69 Ma) Rby111. The monazite lower intercept age of 37 Ma in RM-27 suggests its approximate time of emplacement. Rby107 can be interpreted either as intruded at 29 Ma with a large component of inherited (ca. 38 Ma) zircon, or alternatively intruded ca.38 Ma with younger zircon crystallization at 29 Ma related to heating by intrusion of a nearby 29 Ma biotite monzogranite sill. The age data for all the samples together imply multiple intrusion episodes of the pegmatitic granite in Late Cretaceous and Paleogene time.
Whether the crystallization and the emplacement events were short or long lasting is not resolved. Concordant 238U-206Pb ages that are between age peaks at 92, 68, 42, 37, and 29 Ma may include igneous ages, metamorphic ages, composite ages across more than one crystal zone, and/or apparent ages caused by Pb loss. Many of the dates come from oscillatory zoned zircon considered to be typical of igneous crystallization rather than of metamorphic overprinting. A paucity of dates in the age gap between ca. 65 and 45 Ma argues that few or none of the dates are composite mixes between Paleogene and Cretaceous zircon or monazite. Pb loss or metamorphic overprinting may explain a few younger dates, but the crystal zoning and the geochemistry discussed here suggest that metamorphic overprinting is unlikely to be a major cause of the large spreads of determined ages. Low Th/U ratios in zircon commonly are attributed to metamorphism, but we interpret the low ratios in our samples instead to reflect mainly the very high uranium contents. The high U contents possibly caused radiation damage that allowed Pb loss. Higher Th/U ratios in Rby113 associated with Paleogene ages compared to Cretaceous ones suggest crystallization from distinct magmas at two distinct times, presumably owing to a remelting event.
Metamorphic or igneous overprinting during Cretaceous and Eocene thermal events is expectable, given what is known of the thermal history of the core complex. Metamorphism at upper amphibolite facies in both Late Cretaceous and Eocene time is suggested by the presence and fabric of metamorphic sillimanite in ortho gneisses of both ages and by evidence from the East Humboldt Range (McGrew et al., 2000). McGrew et al. (2000) concluded that slowly falling temperatures after Late Cretaceous metamorphism and intrusion of pegmatitic granite in the core complex were interrupted by Eocene reheating to upper amphibolite facies conditions and temperatures >500 °C (as dated by 40Ar/39Ar hornblende thermochronology). Pb loss related to the Eocene reheating may explain some low apparent dates (e.g., Fig. 11).
We conclude that the pegmatitic granite intruded its host rocks ca. 69 and ca. 38 Ma, possibly also ca. 92 and 29 Ma, and perhaps also at intermediate times. Some monazite and zircon survived as inherited xenocrysts. Most zircon grains likely reflect crystallization or recrystallization under magmatic conditions. The monazite in RM-4 (80–90 Ma) is older than the zircon-based 69 Ma intrusion age established for adjacent WRP02–7, and is here considered as inherited. This inherited monazite contrasts with the granite gneiss of Thorpe Creek, from which Cretaceous zircon has been reported (Premo et al., 2005) even though monazite (in three samples) yielded much younger (ca. 39 Ma) U-Th ages (Wright and Snoke, 1993; MacCready et al., 1997).
Survival of inherited monazite grains and low-temperature magmatic zircon grains in several samples implies that at least some of the pegmatitic granite was sourced from partial remelting or remobilization of antecedent Cretaceous (and Paleogene?) granites. Some of the dated zircons, such as grains dated 70–90 Ma in WRP02–7 and Rby111, possibly represent antecrysts formed earlier in the same magmatic system, although how long igneous systems can be maintained is uncertain. The inferred forerunner granites may have been derived from the Proterozoic metapelite source for the pegmatitic granite proposed in Lee et al. (2003). We infer that fluxes of new magma or of hydrous fluids from deeper levels raised granite temperatures above the solidus. The similarity of the granite products suggests that these fluxes all reflect a similar source, but the dramatic changes in tectonic setting from early Late Cretaceous to Oligocene time suggest that the driving mechanisms varied. The fluxes could have been driven at different times by mechanisms as varied as tectonic thickening, strain heating, decompression melting, and magmatic underplating. The crustal levels are unknown where granitic forerunners were partially melted or remobilized, possibly in a deep anatectic zone near the ultimate pelitic source, or possibly even at the level currently exposed. Pegmatitic granite likely cannibalized itself through partial melting and remobilization either at deeper depths or in situ at the sites where now exposed.
Among the currently exposed rock types, muscovite-bearing leucogranites (dominantly pegmatitic granite) would have been the most susceptible to anatexis, more easily fusible than quartzite, calcareous rocks, and scarce biotite schist. The granites could have been susceptible to small-volume partial remelting by dehydration and reheating above their solidus, especially if accompanied by rising hydrous fluids. The pegmatitic granite melts did not mix much with equigranular Late Cretaceous and Paleogene leucogranites, because they are isotopically distinct. Sillimanite present in the pegmatitic granite and in the metamorphosed Prospect Mountain Quartzite may reflect muscovite-dehydration reactions that produced small amounts of melt both from granite and the impure quartzite (cf. Cheney and Guidotti, 1979; Patiño Douce et al., 1990).
For pegmatitic granites to become remobilized, either in situ or at depth, a melt fraction >50%, or perhaps less if water content is high, likely would be needed to lower the viscosity enough for bulk intrusive mobility (cf. Bagdassarov and Dorfman, 1998). Melt could escape by percolation at much smaller melt fractions. We infer that remobilization of granite forerunners required the addition of heat carried by fluids. Growth of the new zircon, compositional differences between older and younger zircon in single samples, and lack of much evidence of corrosion of older grains, imply that zirconium was added. Whether the granites were cannibalized at depth or near the now-exposed section, the zircon dates suggest the possibility that some intrusions may have been near or above their solidus temperature in a midcrustal environment for tens of millions of years in the Late Cretaceous and again in the Paleogene. Previously emplaced granite underwent intermittent partial remelting sufficient to cause some remobilization and then new zircon crystallization during solidification.
Intrusion episodes of the pegmatitic granite overlapped the time of development of the Lamoille Canyon nappe and its parallel parasitic folds and lineation. Some folding parasitic to the nappe is cut by 69 Ma pegmatitic granite (Rby111 and WRP02–7) and may postdate 92 Ma (Rby113), so the main nappe folding may relate to Late Cretaceous regional plutonism and eastward-vergent contractile tectonics during the Sevier orogeny. Some nappe-parallel lineation also overprints Oligocene biotite monzogranite suggestive of continued or reworked fabric formation (MacCready et al., 1997), and metamorphic sillimanite is along a foliation plane in the Rby107 body. Whether the Lamoille Canyon nappe formed mainly in the Late Cretaceous, with an overprint of Paleogene fabric enhancement, remains to be further tested. The presence of abundant partial melt likely promoted ductile deformation (cf. Norlander et al., 2002). Intrusion and deformation of pegmatitic granite during folding and ductile deformation are consistent with models in which deformation pumping helps drive buoyant magmas upward to sites of emplacement (e.g., Weinberg and Searle, 1998). Some have proposed that such magma migration during deformation must be at slow rates compatible with tectonic deformation (Weinberg and Searle, 1998). Thus, extended periods of millions of years for magma intrusion may be expectable.
Direct evidence for a Cretaceous mafic heat source is lacking, but Cretaceous melting may relate to tectonic overthickening and/or to possible Late Cretaceous extensional collapse of the overthickened crust (McGrew et al., 2000; Wells and Hoisch, 2008). Heat sources for the multiple Cretaceous melting episodes, as well for as multiple Paleogene ones 20–30 m.y. younger, will need to be incorporated into future tectonic models. The Eocene intrusive pulse in the core complex included mafic melts, suggesting that mafic underplating provided a new flux of heat, partial melting, and fluids that helped drive geothermal systems at higher crustal levels expressed in the nearby world-class Carlin Trend gold resource (Henry and Boden, 1998; Hofstra et al., 1999; Henry and Ressel, 2000; Howard, 2003).
The lack in the pegmatitic granite of inherited Archean and Proterozoic zircon recycled from basement rocks could possibly record an early thermal event that dissolved all relict zircon, but the highly peraluminous nature of the magmas argues against much zircon resorption (Watson, 1979; Watson and Harrison, 1983). Alternatively, we propose that the lack of ancient inherited grains reflects a very fine-grained pelitic source for the initial magmas, different from isotopically distinct and likely coarser basement sources of other Mesozoic and Paleogene granitoids in the area that contain inherited Precambrian zircon.
Nearly identical pegmatitic granite magmas were injected or remobilized repeatedly through tectonic regimes varying from contraction to extension, and ultimately to a time of mafic underplating. Whether inherited zircon within single samples of the pegmatitic granite was carried from depth with new magma or inherited from granite bodies remobilized in situ, the geochronologic evidence that the pegmatitic granites and inferred intrusive forerunners underwent partial remelting at several times over a span of tens of millions of years suggests a new time dimension for protracted evolution of large igneous systems (cf. Hildreth, 1981). The repeated intrusion resulted in coalescence with pegmatitic granite bodies already present. The resulting injection complex shows both spatial and temporal evidence of being assembled piecemeal over extended periods. The presence of ghost stratigraphy in the plutonic massif of >95% pegmatitic granite at deep levels of exposure suggests that it was assembled incrementally and is not simply a parental pluton to the higher injection complex.
Melting and crystallization events that led to assembly and modification of the batholith-sized volume of pegmatitic granite occurred repeatedly over long periods of Late Cretaceous and Paleogene time during dramatically changing tectonic settings. The injection complex, including its massif of >95% pegmatitic granite at deep structural levels, was assembled episodically by injection of a series of small, now aggregated and partly coalesced intrusions. The ghost stratigraphy and structure of metasedimentary remnants and their unrotated fabrics indicate that synkinematic intrusion inflated, but by itself did not disrupt, the evolving and deforming host-rock architecture. Instead, the joining of granite bodies around metasedimentary remnants and the downward proportional increase of the granite suggest incremental stages in construction of the batholithic volume of pegmatitic granite (>600 km3). The lowest, most granite-rich parts of the resulting injection complex may have been assembled in a manner analogous to construction of the Harney Peak Granite (South Dakota; Redden and DeWitt, 2008), or of the Main Donegal Granite pluton around its ghost stratigraphy of enclosed remnants (Hutton, 1992), but apparently over a much longer period of time.
Zircon U-Pb ages that cluster ca. 92 Ma, 69 Ma, 38 Ma, and 29 Ma, with morphologic and geochemical evidence of magmatic origins, indicate recurrent and possibly prolonged igneous zircon growth that was rejuvenated several times. Inherited Cretaceous and younger grains suggest that granites, thought to have formed initially by muscovite dehydration melting of very fine-grained metapelitic rocks, were remobilized multiple times by partial melting events, each accompanied by new zircon growth. Whether the remobilization occurred at depth or near the present level of exposure is unknown, but evidence that the intrusions predate and postdate folding, blend together without visible contacts in many cases, and isotopically resemble each other but not other granites in the Lamoille Canyon area is consistent with local remobilization. Some small folded sills may have been partly remobilized as crosscutting dikes. The pegmatitic granite intrusions represent recurrent additions to and possibly reassembly of a batholith-size massif and injection complex. These processes are interpreted to have overlapped the development of fold nappes during the Late Cretaceous Sevier orogeny and an Eocene–Oligocene time of compositionally expanded magmatism and continued ductile deformation.
The range of Cretaceous and Paleogene zircon and monazite U-Pb ages suggests repeated periods of crystallization of pegmatitic granite or its forerunners. Recurrent zircon growth suggests that melts were present or recharged many times in the Late Cretaceous and Paleogene. The batholith-sized injection complex of pegmatitic granite was constructed as a network of interconnected sheets and dikes of varying age of mobility. The intrusive style and U-Pb dating of pegmatitic leucogranite in the Ruby Mountains suggest that small granitic bodies intruded, and possibly recurrently mobilized, as they aggregated incrementally into larger granite bodies over long periods of time. The combined field relations and geochronology thus suggest very lengthy plutonic assembly and modification and offer new perspectives to the time spans that may be involved in generation, construction, and reconstruction of some plutonic complexes.
APPENDIX 1: ROCKS SAMPLED
Gneissic Pegmatitic Leucogranite in Upper Lamoille Canyon (WRP02-7 and RM-4)
Biotite-bearing gneissic pegmatitic leucogranite is from the north end of a roadcut on the Lamoille Canyon road (elevation 2565 m) overlooked by The Terraces picnic ground, 40 m from Lamoille Creek. The granite obliquely cuts large enclaves of layered diopside-rich calc-silicate rock. The granite contains sillimanite and monazite visible in hand specimen. The granite exhibits a north-south lineation, in part defined by sillimanite bundles. It also contains scarce winged feldspar porphyroblasts that indicate a northward sense of shear. Texture at this outcrop shown in Figure 5A (sample H80Ruby-4). Zircon in the granite is green. A preliminary monazite U-Pb age of ca. 83 Ma was reported (Wright and Snoke, 1993) from sample RM-4 (collected within 2 m of the WRP02-7 site). The site is identified on a map of the Verdi Peak quadrangle (Howard and McCready, 2004). Essentially identical Nd and Pb isotopic ratios in RM-4 (Wright and Snoke, 1993) and WRP02-7 (Premo et al., 2005) help to establish that RM-4 and WRP02-7 sampled the same rock.
Folded Pegmatitic Leucogranite in Central Lamoille Canyon (Rby113)
Gneissic pegmatitic leucogranite at the base of a cliff (elevation 2575 m) on the north wall of Lamoille Canyon (Fig. 6D) overlooking Thomas Canyon Campground intrudes as sills, sheets, and dikes into layered calc-silicate rocks and structurally overlying quartzite. The granite forms bodies ranging from centimeters to many meters thick, some of which are folded or boudined, and others cut across folds. Dated sample Rby113 is from the thick part of a sill, which we interpret as folded (Fig. 7B), that follows a fold in its host rocks (calc-silicate rocks and marble). Although layers of host rock thicken in the nose of a fold, the sill does not, which seems difficult to explain if the granite wedged apart already folded host rocks, and more easily explained if the sill was folded as a layer stiffer than its host rocks. Elsewhere the folded sill varies 0.4–2 m in thickness. The very coarse-grained muscovite-biotite leucogranite exhibits gneissic biotite folia in the outcrop, and locally contains massive quartz. The analyzed rock contained 75% SiO2.
Crosscutting Pegmatitic Leucogranite in Central Lamoille Canyon (Rby111)
This Pegmatitic leucogranite is at the base of the same cliff (elevation 2560 m), 50 m northwest of sample Rby113 (Fig. 6D). The granite forms bodies ranging from centimeters to many meters thick, some of which are folded or boudined; others cut across folds, and others are folded in one place, yet all grade into parts that cut across the folds. The folds are parasitic to the Lamoille Canyon nappe. Dated sample Rby111 is from the narrow (1.1 m) part of an irregular body (commonly 2-4 m thick) of pegmatitic leucogneiss where it discordantly cuts layering in both calc-silicate rocks and in quartzite and the inverted stratigraphic contact between these two rock types (Fig. 7C). The roughly planar layering in the stratigraphic remnants represents parts of larger fold limbs. The analyzed biotite-bearing granite rock contained 68% SiO2.
Gneissic Pegmatitic Leucogranite in Lower Lamoille Canyon (Rby107)
Pegmatitic leucogranite gneiss forms most of the exposed rock at deep levels in Lamoille Canyon and its Right Fork. Rby107 is pegmatitic leucogranite exposed in a roadcut at elevation 2035 m (Fig. 7D). The sampled rock is from a subhorizontal gneissic sheet >3 m thick, representative of most of the rock at this structural level. A nebulous patch ∼0.4 m across was sampled that is fully gradational to and slightly less structured than adjacent slightly more gneissic similar rock containing finer-grained muscovite and wispy biotite folia within a few centimeters of the sample site. The sample contains biotite and less muscovite. The texture is partly recrystallized. Vermiculite is present. The analyzed rock contained 76% SiO2. Sillimanite-biotite folia are present in this leucogranite sheet 10 m northwest of the sample site. A relict feldspathic quartzite layer 8 cm thick surrounded by the gneissic pegmatitic granite occurs a few centimeters below the sample site. A horizontal sill 5 m thick of biotite monzogranite intrudes the gneissic pegmatitic leucogranite 3 m above the sampled granite; this sill contains rare enclaves of pegmatitic leucogranite gneiss, is cut by rare leucogranite and pegmatite dikes, and closely resembles rock dated elsewhere as 29 Ma (Wright and Snoke, 1993; MacCready et al., 1997).
Pegmatitic Leucogranite Gneiss in Lower Lamoille Canyon (RM-27)
RM-27 is sillimanite-biotite leucogranite gneiss from a roadcut overlooking Camp Lamoille. The gneiss is interlayered with finer-grained equigranular biotite granite and leucogranite gneiss sheets. The leucogranite gneiss envelops rare thin layers of feldspathic quartzite, associated with biotite-sillimanite folia in the granite gneiss that are likely inherited from thin semipelitic interbeds in the quartzite. The rock has a deformed hypidiomorphic granular texture, with significant subgrain development in quartz and K-feldspar. Sparse brown biotite and fibrous to bladed muscovite occupy interstices. Sillimanite is present as inclusions in quartz and feldspars, and some fibrous muscovite is interpreted as relict sillimanite. The rock was analyzed as containing 75.3% SiO2 and 7.2% total alkalies. Subordinate fine-grained gneissic rocks are interleaved with the pegmatitic granite gneiss in the outcrop. A sheet of gneissic biotite monzogranite from this outcrop was dated by U-Pb on monazite as 29 Ma (Wright and Snoke, 1993), and a gneissic fine-grained garnet–two-mica granite from this outcrop was dated by U-Pb on zircon as 32 Ma (Premo et al., 2005).
We thank Wes Hildreth, Charlie Bacon, and two Geosphere reviewers for helpful and perceptive critiques of the manuscript. Many geologists have visited Lamoille Canyon with us on field trips and offered valuable comments on the field and geochronologic data. We thank Pedro Castiñeiras for help in the SHRIMP lab. Jim Wright performed U-Pb analyses on the monazites. This research was supported by the U.S. Geological Survey Mineral Resources Program and a Bradley Scholarship, and by National Science Foundation grants EAR-9627814 to Barnes and EAR-9627958 to Snoke.