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Development of a Continental Volcanic Field: Petrogenesis of Pre-caldera Intermediate and Silicic Rocks and Origin of the Bandelier Magmas, Jemez Mountains (New Mexico, USA)
1 2 3 4 5


CA 95929, USA



The Miocene^Quaternary Jemez Mountains volcanic field (JMVF) is the site of the Valles caldera and associated Bandelier T Caldera formation was preceded by410 Myr of volcanism uff. dominated by intermediate composition rocks (57^70% SiO2) that contain components derived from the lithospheric mantle and Precambrian crust. Simple mixing between crust-dominated silicic melts and mantle-dominated mafic magmas, fractional crystallization, and assimilation accompanied by fractional crystallization are the principal mechanisms involved in the production of these intermediate lavas. A variety of isotopically distinct crustal sources were involved in magmatism between 13 and 6 Ma, but only one type (or two very similar types) of crust between 6 and 2 Ma. This long history constitutes a record of accommodation of mantle-derived magma in the crust by melting of country rock. The post-2 Ma Bandelier T and associated rhyolites were, in contrast, generated uff by melting of hybridized crust in the form of buried, warm intrusive rocks associated with pre-6 Ma activity. Major shifts in the location, style and geochemical character of magmatism in the JMVF occur within a few million years after volcanic maxima and may correspond to pooling of magma at a new location in the crust following

solidification of earlier magma chambers that acted as traps for basaltic replenishment.

KEY WORDS: crustal anatexis; fractional crystallization; Jemez Mountain Volcanic Field; Valles Caldera; radiogenic isotopes; trace elements

The Jemez Mountains region, situated on the western flank of the Rio Grande rift in northern New Mexico, contains a record of over 20 Myr of volcanism prior to the catastrophic eruptions of the rhyolitic Bandelier Tuff and formation of the Valles^T oledo caldera complex in two episodes at 1Á6 and 1Á2 Ma. The Jemez Mountains volcanic field (JMVF) is largely built of intermediate to silicic lavas, with dacites being especially prominent in the few million years preceding the rhyolitic caldera-forming eruptions. Prior studies have indicated a major role for crustal

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anatexis in the generation of the intermediate and silicic JMVF rocks (Gardner, 1985; Singer & Kudo, 1986; Ellisor et al., 1996). The need for continuous basaltic input into the crust to sustain long-lived silicic centers has long been recognized, and the JMVF, like many silicic centers, may be regarded as ‘fundamentally basaltic’ (Hildreth, 1981). However, identification of specific mantle and crustal components, the processes by which they are blended, and their changing roles during the evolution of any one center remain problematic. For example, DePaolo et al. (1992) and Perry et al. (1993) noted, on the basis of Nd isotope ratios, that the Precambrian crustal contribution to JMVF magmatism decreased with time, reaching a minimum in the caldera-forming rhyolitic Bandelier Tuff, and inferred that fractional crystallization of mafic magma, with little accompanying assimilation, must have played a major role in the generation of the caldera-related rhyolites. The difficulty with this model is that the Valles caldera is the site of a pronounced gravity low (Segar, 1974; Nowell, 1996) and, although other geophysical data indicate a significant role for mafic underplating beneath the volcanic field (Ankeny et al., 1986; Steck et al., 1998; Aprea et al., 2002), there is no evidence for a large volume of mafic cumulate residue beneath the JMVF such as must accompany generation of many hundred cubic kilometers of high-silica rhyolite by fractional crystallization of basaltic magma. Crustal hybridization (Johnson et al., 1990; Riciputi & Johnson, 1990; Riciputi et al., 1995) is an alternative model, in which the crust gradually takes on the isotopic signature of the mantle as a result of repeated intrusion of mantle-derived magma during the lifetime of a volcanic field. Crustal melting late in the history of the field therefore produces magmas that have isotopic affinities with mantle rather than crust. Recently, Wolff et al. (2005) showed that the lithospheric mantle has been a major source of basaltic magma throughout the history of the Jemez Mountains region and demonstrated the existence of two end-member parental magmas: (1) basanites and nephelinites derived from lowdegree partial melting of lithospheric mantle with residual amphibole; (2) olivine tholeiites, which could be derived either by higher-degree partial melting of the same source or from asthenospheric mantle. Crustal contamination modified the strongly silica-undersaturated compositions to weakly ne-normative alkali basalts, hawaiites, and derivative basaltic andesites and mugearites that erupted alongside tholeiites through most of the lifetime of the volcanic field. In this paper, we extend the work of Wolff et al. (2005) to consider the petrogenesis of the intermediate and silicic pre-caldera lavas and tuffs during construction of a moderate-sized ($2000 km3), long-lived volcanic field that ultimately hosted rhyolitic caldera-forming eruptions of catastrophic magnitude. T ogether, the two studies are intended to provide a basis for future, more detailed studies

of the temporal development of the JMVF, and to provide a foundation for understanding the petrogenesis of the climactic rhyolitic magmas.

The JMVF is perhaps best known for the work of R. L. Smith and associates (Smith, 1960, 1979; Ross & Smith, 1961; Smith & Bailey, 1966) on the caldera-forming Bandelier Tuff and associated rhyolitic lavas. Caldera formation, however, was preceded by some 23 Myr of volcanic activity in the area. The JMVF is built on the western shoulder of the Espanola basin, one of the north^ ‹ south-trending series of Cenozoic en echelon sedimentary basins that make up the Rio Grande rift (Fig. 1). The JMVF is located at the intersection of the rift and the Jemez lineament (Fig. 1). The lineament is a long-lived tectonic feature that originally formed during the Proterozoic assembly of North America by collision of the Yavapai and Mazatzal crustal provinces between 1Á6 and 1Á7 Ga, and is associated with a south-dipping lithospheric suture (Shaw & Karlstrom, 1999; Karlstrom et al., 2002; Magnani et al., 2004). An associated low-velocity region (Duecker et al., 2001) in the lithospheric mantle to the north of the suture extends between the Moho at 40^50 km depth and 120 km depth, the probable base of the lithosphere. This lithospheric mantle is the likely source of JMVF primitive mafic magmas [see the paper by Wolff et al. (2005), which also includes a more detailed description of regional geology and tectonics]. Rifting in northern New Mexico began at $30 Ma and has continued episodically to the present (Aldrich, 1986; Gardner et al., 1986; Morgan et al., 1986). Following a period of quiescence from at least 18 to 13 Ma, extension reinitiated and continued for several million years, during which time much of the JMVF was constructed. A second period of reduced rate of extension between 7 and 4 Ma (Gardner et al., 1986) that coincides with a lull in volcanic activity, ended with the development of the Pajarito fault zone (Fig. 1), which marks an eastward shift of the western boundary of the Espanola basin from the earlier Canada ‹ ‹ de Cochiti fault zone (Fig. 1). Currently, low-velocity mantle extends to the base of the crust beneath the rift axis, where any former lithospheric mantle is now presumably either geophysically indistinguishable from asthenosphere as a result of shearing and heating, or has been thermally eroded. A similar low-velocity zone exists at the base of the crust beneath the JMVF (Aprea et al., 2002; Steck et al., 1998). T emperatures may exceed 9008C in the lower crust beneath the rift (Baldridge et al., 1984; Clarkson & Reiter, 1984; Morgan & Golombek, 1984). This is sufficiently close to the range of crustal solidus temperatures that thermal perturbations caused by


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Fig. 1. Location map of the Jemez Mountains Volcanic Field (JMVF) showing pre-caldera formations and the Valles caldera [see Wolff et al. (2005) for sources of information]. Blank areas represent formations lying beneath the base of the JMVF, and cover by Bandelier Tuff and associated caldera-related rhyolites. Locality abbreviations referred to in the text: CP, Cerro Pavo; MdM, Mesa del Medio; PP, Polvadera Peak; TM, Tschicoma Mountain; MdG, Mesa de la Gallina, outlined by dotted line; LC, Los Cerritos; TE, T oledo Embayment; CR, Cerro Rubio; PM, Pajarito Mt.; CG, Cerro Grande. Area abbreviations: LGP, La Grulla Plateau; CdSD, Canon de San Diego, EB, Espanola ‹ ‹ basin. PFZ, trace of the main portion of the Pajarito fault zone that forms the present western tectonic boundary of the Rio Grande rift in this area. CFZ, trace of Canada de Cochiti fault zone in the south‹ ern Jemez, and inferred to extend below Mesa del Medio in the north (Smith et al., 1970). Inset: the Rio Grande rift (RGR) and Jemez lineament (JL, dashed dark gray line) in Colorado and New Mexico, showing location of the study area. Light gray area indicates the extent of asthenospheric contributions to New Mexico Cenozoic volcanism (McMillan, 1998).

Fig. 2. T emporal ranges of JMVF geological units. SFG: Santa Fe Group; LGP: La Grulla Plateau. The ordering of units from left to right corresponds very approximately to areas of maximum exposure in southern, northern and eastern sectors of the JMVF. Santa Fe Group sediments underlie most of the JMVF, with a minimum age of $4 Ma in the east. Dotted lines indicate interbedded volcanic and sedimentary rocks. Modified after Goff & Gardner (2004).

succession is overlain by a veneer of Cenozoic sedimentary units, the uppermost of which is the eastward-thinning, rift-filling Santa Fe Group, which includes minor mafic lavas and tuffs (Bailey et al., 1969; Smith et al., 1970).

Summary of the petrogenesis of the mafic lavas
The Santa Fe Group lavas and tuffs are the oldest volcanic rocks exposed in the area (Fig. 2). They form part of a more widespread, small-volume, late Oligocene to middle Miocene volcanic episode in the Espanola basin ‹ (Gibson et al., 1993; Woldegabriel et al., 2003), coincident with the first extensional phase, which pre-dates construction of the JMVF (Gardner et al., 1986; Fig. 2). Santa Fe Group volcanic rocks fall into two compositional groups: strongly silica-undersaturated nephelinites and basanites, which include primitive compositions (up to 16% MgO), and tholeiites and derivative quartz-normative basaltic andesites. Marked depletions in K relative to other

intrusions of mantle-derived primitive basaltic magma into the sub-rift crust are very likely to induce partial melting of the Proterozoic crustal rocks. The JMVF is built on a substrate of Upper Paleozoic sedimentary strata that rest on the Proterozoic basement, which consists of granitoid and metavolcanic rocks (typically intermediate amphibolites) locally dated at 1Á62^1Á44 Ga (Eichelberger & Koch, 1979; Brookins & Laughlin, 1983; Laughlin et al., 1983). Metavolcanic and metasedimentary rocks are also exposed in uplifts around the borders of the Espanola basin. The Paleozoic ‹






incompatible elements among the silica-undersaturated lavas were attributed by Wolff et al. (2005) to residual amphibole in the lithospheric mantle source during partial melting. Although an asthenospheric source for the tholeiites cannot be ruled out, Wolff et al. (2005) argued that the tholeiitic parental magma is more probably derived from the same lithospheric mantle source material via a higher degree of partial melting involving the complete consumption of amphibole. A solely lithospheric source region for all JMVF parental mafic magmas is consistent with existing data [as reviewed by McMillan (1998)] indicating that the northern limit of asthenospheric contributions to magmatism in New Mexico lies south of the JMVF (Fig. 1). The actual source for JMVF mafic magmas is envisaged as ancient oceanic lithosphere associated with the Proterozoic suture zone (Wolff et al., 2005), which partially melted in association with rift extension. This is consistent with the geophysical evidence for shallow low-velocity mantle beneath the area at the present day (Duecker et al., 2001). Many mafic lavas of the JMVF (Paliza Canyon Formation, Cerros del Rio and El Alto basalts, and a few Lobato basalts: Figs 1 and 2) have distinctive incompatible trace element characteristics, with Th/(Nb,Ta) and La/(Nb,Ta) greater than Bulk Earth, but K/(Nb,Ta) similar to Bulk Earth, that do not correspond to any common globally recognized basalt type. Wolff et al. (2000, 2005) showed that these compositions can be produced by contamination of K-depleted, silica-undersaturated magma with crustal melts having high K/(Nb,Ta) and proposed such an origin for the JMVF mafic rocks, with the implicit assumption that mantle partial melts produced during the early extensional phase are representative of those produced since $13 Ma. In this study, these rocks, which span several categories of formal classification (Fig. 3), are for convenience referred to as Type I mafic lavas. Tholeiites and derivative basaltic andesites are also crustally contaminated (Duncker et al., 1991; Wolff et al., 2000, 2005) and are referred to as Type II mafic lavas. We emphasize that all JMVF mafic magmas carry a component derived from continental crust. Type I and Type II JMVF mafic lavas contain variable proportions of Proterozoic granitoid crust (Wolff et al., 2005) and consequently exhibit significant variations in Pb and Nd isotope ratios (206Pb/204Pb ¼17Á20 ^ 18Á31; 143 Nd/144Nd ¼ 0Á51244 ^ 0Á51273). Sr isotope ratios, however, are low and relatively uniform (87Sr/86Sr ¼ 0Á7041 ^ 0Á7048), despite high 87Sr/86Sr among potential contaminating lithologies. Wolff et al. (2005) compared the effects of contamination by low-87Sr/86Sr crust with assimilation of high-87Sr/86Sr granitoids by partial melting, with Sr retained in a feldspathic residue. Both models reproduced the isotopic characteristics of the rocks, but the lack of measurable Eu anomalies among most JMVF mafic

lavas is inconsistent with a major role for plagioclase during petrogenesis.

The exact date of inception of volcanism in the JMVF proper is a largely semantic issue (Gardner et al., 1986; Goff & Gardner, 2004); at some locations, there is a sharp contact between rift-filling Santa Fe Group sediments and a continuous volcanic section at 12^13 Ma (Goff et al., 1990), but elsewhere lavas dated as old as 14 Ma to as young as 9 Ma are interbedded with the sediments (Aldrich & Dethier, 1990). Overall, construction of the volcanic field was under way by 13 Ma, coincident with the onset of an episode of crustal extension in the adjacent Espanola ‹ basin (Gardner et al., 1986). The earliest rocks are basalts of the Paliza Canyon and Lobato Formations, and tuffs of the Canovas Canyon Rhyolite (Fig. 2). The stratigraphy, geology, geochronology and petrography of the pre-caldera intermediate and silicic rocks are briefly described below; details of rocks with 557% SiO2 have been given by Wolff et al. (2005). Revisions to the original JMVF stratigraphic scheme of Bailey et al. (1969) and Smith et al. (1970) have been made as new mapping and geochronological data have become available. Here we follow the formation stratigraphy of Bailey et al. (1969) and Smith et al. (1970) as modified by Gardner et al. (1986), but find the broader division into Keres and Polvadera Groups, made largely on geographical grounds, to be unhelpful. In particular, it is now clear that the Polvadera Group, as originally defined, includes rocks of widely disparate ages and origins. In contrast, the formations for the most part exhibit geological, geochronological, and petrological coherence. The Lobato Basalt, dominated byType II tholeiites, and the El Alto basalt with both Type I and II mafic compositions, are omitted from the descriptions below because there are no samples from either with 457% SiO2 (see Wolff et al., 2005).

Paliza Canyon Formation
Exposures of the Paliza Canyon Formation are largely confined to the southern JMVF (Fig. 1), but on the basis of well data (Hulen et al., 1991), mapping (Smith et al., 1970; Gardner & Goff, 1996), and abundant lithic fragments in the Bandelier Tuff, it underlies the caldera and Tschicoma Formation in the north^central Jemez Mountains and has a total volume of 1000 km3. The Paliza Canyon Formation is volumetrically dominated by trachyandesites, trachydacites and dacites, with subordinate Type I and II mafic lavas, andesites and low-silica rhyolites (Fig. 3). Paliza Canyon flows, domes, tuffs and minor intrusives


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Fig. 3. T otal alkali^silica (TAS) plot of pre-caldera volcanic rocks of the JMVF. Tschicoma Formation lavas are subdivided to include Tschicoma enclaves and La Grulla plateau lavas. Cerros del Rio evolved compositions are from Duncker (1988) and Duncker et al. (1991). Type I and Type II mafic lava data are from Duncker et al. (1991), Gibson et al. (1993) and Wolff et al. (2005). Cerros del Rio compositions within the trachyandesite field are dominantly benmoreites (Na2O ^ 2Á0 ! K2O sensu LeBas et al., 1986). Both types are found in all major divisions of the pre-caldera JMVF, but Type I samples are dominantly from Paliza Canyon and Cerros del Rio, and Type II dominantly from the Lobato Formation and Cerros del Rio and El Alto tholeiites. Annotated legend organized by location of formations.

are petrographically diverse; Goff et al. (1990) recognized eight distinct varieties of andesite and dacite. On the basis of age (9Á6 Ma, Goff et al., 1989) and chemical affinity, we include the dacite of Los Cerritos in the northeastern JMVF (Fig. 1), originally mapped as belonging to the Tschicoma Formation, in the Paliza Canyon Formation. Bulk compositions are dominantly weakly alkaline, and thus most Paliza Canyon lavas in the range 57^63% SiO2 are classified as trachyandesite; the remaining samples are andesitic. Trachydacites and dacites are equally abundant in the range 63^69% SiO2 (Fig. 3). Na/K is variable and, whereas the mafic lavas are dominantly sodic (Wolff et al., 2005), the trachyandesites include both benmoreites and latites (sensu LeBas et al., 1986). Some minor adjustment of alkali contents may have occurred as a result of

post-eruptive processes. Compared with Tschicoma Formation rocks of equivalent silica content, intermediatecomposition Paliza Canyon rocks exhibit more scattered major element variations, higher TiO2 and K2O contents, and little systematic variation in alkali concentrations (Figs 3 and 4). Trachyandesites and andesites have the phenocryst assemblage plagioclase (An30^An60) þ augite þ hypersthene þ opaque oxides Æ hornblende Æ olivine, and are characterized by disequilibrium textures: olivines invariably have reaction rims consisting of fine-grained clinopyroxene and opaque oxides, and plagioclase phenocrysts are commonly resorbed and exhibit complex zoning. Plagioclase (An16^53) is also the dominant phenocryst among the dacites, benmoreites and rhyodacites,






Fig. 4. Variations in major element concentrations of pre-caldera JMVF lavas and tephras with SiO2 wt % (symbols and mafic lava fields are the same as in Fig. 3).

with hornblende, biotite, clinopyroxene, opaque oxides and sparse apatite; glomeroporphyritic textures and crystal clots of plagioclase, augite, hypersthene and oxides are common.

Canovas Canyon Rhyolite
This formation consists of flows, domes, pumice fallout deposits, non-welded to densely welded and rheomorphic ignimbrites interbedded within the Paliza Canyon Formation, and minor rhyolitic intrusions. One tuff, dated

at 12Á4 Ma, locally forms the base of the JMVF (Gardner et al., 1986; Goff et al., 1990). Some Canovas Canyon units are low-silica rhyolites (69^72% SiO2, Fig. 3) that are petrographically very similar to, and compositionally overlap, the most silicic Paliza Canyon units that were mapped as ‘biotite dacite’ by Goff et al. (1990). The remainder are sparsely porphyritic high-silica rhyolites (76^78% SiO2) with phenocrysts of sanidine þ quartz þ opaque oxides Æ plagioclase Æ pyroxene Æ hornblende Æ biotite.


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Bearhead Rhyolite
The high-silica rhyolite domes, flows, tuffs and minor intrusions of the Bearhead Rhyolite have been described in detail by Smith et al. (1991), Gay & Smith (1993), and Justet & Spell (2001). They are nearly aphyric and are similar to the Canovas Canyon high-silica rhyolites. The bulk of the Bearhead rhyolite post-dates the Paliza Canyon Formation at 7Á 06^6Á52 Ma, but two late flows have ages of 6Á1Ma (Justet & Spell, 2001). Some of the El Rechuelos rhyolite domes in the northern JMVF (see below) also fall into this age range.

Tschicoma Formation
The Tschicoma Formation, with a volume of nearly 500 km3 (Gardner et al., 1986), dominates the northern JMVF. Most of this volume is dacite, with lesser amounts of trachyandesite, andesite and occasional rhyolite (Fig. 3). Most of the exposed Tschicoma Formation lies in the sector from north to east of the Valles Caldera (Fig. 1). Excluding the La Grulla Plateau volcanics (see below), ages range from 6Á9 to 2Á7 Ma, with most of the dacites having been erupted between 5 and 2Á7 Ma (Dalrymple et al., 1967; Leudke & Smith, 1978; Gardner & Goff, 1984; Goff et al., 1989; Woldegabriel et al., 2001; Goff & Gardner, 2004). Although early eruptions of the Tschicoma Formation coincide with a period of relative tectonic inactivity along the Jemez lineament and Rio Grande rift (Gardner & Goff, 1984; Aldrich, 1986), the large dacite domes in the NE JMVF, including Cerro Rubio, post-date the onset of renewed extension at 4 Ma. The degree of preservation of the thick Tschicoma lavas allows use of topography as a mapping aid in identifying different volcanic flows and eruptive centers. Dacitic flows, up to $6Á5 km in length in the case of Mesa de la Gallina (Fig. 1), originate from as many as 70 domes or eruptive centers that dominate the northern Jemez Mountains. Although there are a few crystal-poor to aphyric lavas, Tschicoma Formation dacites are typically coarsely porphyritic, containing 15^20% phenocrysts of plagioclase Æ clinopyroxene þ resorbed orthopyroxene Æ hornblende Æ biotite þ opaque oxides, sometimes accompanied by rounded quartz xenocrysts with fine-grained haloes of clinopyroxene. Plagioclase is frequently resorbed and hornblende phenocrysts have dehydration rims. Mafic enclaves, up to 25 cm in diameter, of basaltic andesite, andesite and trachyandesite compositionally similar to Cerros del Rio lavas are common in the coarsely porphyritic dacites. We include the two Cerro Rubio quartz latites, located in the T oledo Embayment and dated by K^Ar at 2Á18 Æ 0Á09 Ma and 3Á59 Æ 0Á36 Ma (Heiken et al., 1986), in the Tschicoma Formation on the basis of similar age, lithology and chemistry. Cerro Grande (Fig. 1) is a rhyolite

with abundant dacitic enclaves that are chemically identical to the bulk of the Tschicoma dacites. The La Grulla Plateau extends northward from the Valles Caldera (Fig. 1). It is geographically and tectonically separated from the rest of the Tschicoma Formation by a Bandelier Tuff-filled paleo-valley (Mesa del Medio) coinciding with a splay of the Canada de Cochiti fault zone (Fig. 1). The fault zone is a major down-to-the-east rift-bounding structure that has been active since at least 10 Ma. The plateau is formed of sparsely phyric andesite lavas capped by domes and flows of dacite and trachydacite (Singer & Kudo, 1986). The La Grulla Plateau suite has been dated at 7Á9^7Á4 Ma (Singer & Kudo, 1986), up to 1Myr older than the oldest ages reported for the rest of the Tschicoma Formation and overlapping with late Paliza Canyon activity (Fig. 2). The La Grulla Plateau rocks share major element, trace element and isotopic characteristics with both the Paliza Canyon Formation and the main pile of Tschicoma rocks to the east of the fault zone (Figs 3 and 5^7), but also have some distinct isotopic features (Fig. 7c). At least one dacite dome contains mafic enclaves, chemically similar to those in the more easterly Tschicoma dacites.

Puye Formation
The Puye Formation is an alluvial fan derived from Tschicoma Formation lavas on the eastern flank of the JMVF (Fig. 1) and covers a region of $200 km2 with a total thickness of at least 110 m and volume of 415 km3 (Waresback & Turbeville, 1990). Conglomeratic beds of the Puye Formation contain clasts that represent Tschicoma lavas that may no longer be preserved, having been removed by erosion and/or obliterated by later caldera-forming eruptions. Primary pyroclastic fallout units equivalent to Tschicoma dacite lavas are interbedded throughout the sequence, and the 2Á53 Æ 0Á1Ma Puye ignimbrite (Waresback & Turbeville, 1990), chemically identical to the Cerro Rubio dacites (Table 1), is prominently exposed in the middle of the Puye Formation section in medial exposures. The dominant cobble type in conglomeratic beds above the ignimbrite also closely resembles Cerro Rubio dacite. High-silica rhyolite pyroclastic fallout deposits occur near the top of the Puye Formation; one is correlated with the 1Á85 Ma San Diego Canyon ignimbrites (Turbeville & Self, 1988) that represent the first eruptions from the Bandelier magma system.

El Rechuelos Rhyolites
The El Rechuelos Rhyolite consists of six domes in the northern JMVF (Smith et al., 1970). Loeffler et al. (1988) dated these centers and found that they represent three periods of rhyolitic volcanism at 7Á5, 5^6, and 2 Ma (Fig. 2). The two older groups are similar in age to Bearhead Rhyolite (7Á1^6Á1Ma) but are chemically more






T 1: Major element, trace element, and isotopic analyses of selected intermediate and silicicJMVF lavas and tephras able
Sample: Formation: Latitude (8N): Longitude (8W): MR00-8 TAD 36 7Á016 106 24Á503 MR00-12 TAD 36 1Á943 106 28Á324 MR00-15 TAD 36 2Á773 106 19Á658 MR00-28 TAD 36 1Á818 106 19Á536 MR00-33 TAD 36 00Á750 106 23Á596 MR00-72 TAD 35 51Á851 106 26Á691 MR00-75 TAD 36 5Á72 106 23Á68 MR00-106 TAD 36 2Á29 106 23Á99 MR00-115 TAD 35 53Á823 106 24Á112 MR00-96 TAD 35 56Á394 106 23Á946 MR00-79 P 35 54Á973 106 13Á755

XRF major element (wt %) SiO2 66Á94 TiO2 0Á59 Al2O3 15Á36 FeOÃ 3Á44 MnO 0Á07 MgO 2Á08 CaO 3Á95 Na2O 3Á92 3Á38 K 2O P 2O5 0Á29 XRF trace element (ppm) Ni 34 Cr 37 Sc 7 V 66 Ba 1292 Rb 53 Sr 601 Zr 161 Y 14 Nb 19Á9 Cu 14 Zn 50 Pb 16 La 36 Ce 73 ICP-MS trace element (ppm) La 41Á4 Ce 68Á8 Pr 6Á97 Nd 25Á29 Sm 4Á79 Eu 1Á34 Gd 3Á77 Tb 0Á54 Dy 3Á04 Ho 0Á59 Er 1Á49 Tm 0Á22 Yb 1Á37 Lu 0Á22 Ba 1305 Th 8Á0 Nb 19Á6 Y 15Á9 Hf 4Á4 Ta 1Á4 U 2Á2 Pb 17Á5 Rb 52Á4 Cs 1Á6 Sr 608 Sc 8Á3 Zr 161 Ã Eu/Eu 0Á93 Isotope ratios (87Sr/86Sr)i 143 Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb 0Á70452 (2) 0Á51251 17Á454 15Á471 37Á323

63Á45 0Á80 16Á30 4Á74 0Á09 2Á41 4Á69 4Á05 3Á04 0Á43 16 18 4 78 1216 55 796 194 22 28Á5 11 64 17 62 85 61Á3 93Á0 10Á14 36Á81 6Á74 1Á78 5Á31 0Á78 4Á30 0Á82 2Á07 0Á30 1Á89 0Á30 1249 11Á9 28Á8 22Á4 5Á0 2Á0 3Á2 17Á4 57Á1 1Á9 727 10Á5 196 0Á87 0Á70442 (2) 0Á51257 17Á644 15Á474 37Á461

67Á89 0Á51 15Á50 3Á17 0Á07 1Á76 3Á59 4Á10 3Á21 0Á20 20 25 6 54 1104 68 495 155 15 20Á5 14 45 16 25 57 39Á3 60Á2 6Á24 22Á45 4Á31 1Á16 3Á58 0Á52 2Á99 0Á60 1Á57 0Á23 1Á46 0Á24 1107 10Á1 19Á1 16Á0 4Á1 1Á6 3Á0 16Á6 66Á4 2Á0 489 8Á2 147 0Á88 0Á70452 (1) 0Á51251 17Á517 15Á470 37Á334

68Á09 0Á52 15Á39 3Á28 0Á06 1Á70 3Á43 3Á93 3Á40 0Á20 22 25 7 49 1121 70 467 159 16 19Á7 6 44 19 37 59 37Á5 60Á3 5Á91 20Á86 4Á09 1Á10 3Á29 0Á49 2Á81 0Á55 1Á47 0Á22 1Á38 0Á23 1105 10Á3 19Á4 15Á2 4Á2 1Á6 3Á0 16Á6 68Á8 2Á0 467 8Á1 147 0Á88 0Á70453 (2) 0Á51249 17Á498 15Á458 37Á286

67Á51 0Á53 15Á62 3Á43 0Á07 1Á84 3Á79 3Á80 3Á19 0Á21 24 24 11 63 1198 62 474 160 15 16Á1 14 53 19 31 67 36Á5 59Á2 6Á20 22Á44 4Á40 1Á17 3Á62 0Á53 2Á93 0Á58 1Á50 0Á22 1Á35 0Á21 1197 6Á8 15Á0 15Á6 4Á2 1Á2 1Á9 17Á1 58Á4 1Á4 492 9Á3 150 0Á87 0Á70478 (3) 0Á51244 17Á320 15Á460 37Á198

64Á42 0Á67 17Á11 3Á93 0Á08 2Á19 4Á70 4Á06 2Á57 0Á28 34 36 9 72 1282 40 635 174 17 16Á0 13 62 12 28 51 39Á1 62Á5 6Á63 24Á77 4Á90 1Á42 4Á03 0Á60 3Á45 0Á66 1Á77 0Á25 1Á55 0Á24 1319 6Á3 17Á6 17Á7 4Á6 1Á2 1Á6 17Á3 40Á6 1Á1 632 10Á9 175 0Á95 0Á70441 (2) 0Á51249 17Á528 15Á491 37Á409

64Á73 0Á67 15Á79 4Á38 0Á08 2Á56 4Á63 4Á00 2Á84 0Á31 35 35 7 89 1124 55 603 168 18 20Á2 15 58 19 30 75 45Á3 73Á7 7Á84 28Á89 5Á56 1Á50 4Á54 0Á66 3Á71 0Á70 1Á86 0Á26 1Á63 0Á25 1132 9Á4 21Á6 19Á5 4Á5 1Á6 2Á6 16Á7 57Á4 1Á7 590 11Á2 173 0Á89 0Á70444 (2) 0Á51251 17Á533 15Á464 37Á431

66Á06 0Á59 15Á68 4Á41 0Á09 2Á30 4Á24 3Á60 2Á70 0Á31 29 42 9 78 1251 43 709 169 15 17Á1 12 70 16 37 73 36Á4 57Á7 6Á05 22Á22 4Á24 1Á27 3Á54 0Á51 2Á78 0Á54 1Á35 0Á20 1Á25 0Á19 1236 6Á6 15Á7 14Á5 4Á0 1Á1 1Á7 16Á2 41Á2 1Á2 638 8Á5 151 0Á98 0Á70424 (2) 0Á51249 17Á502 15Á489 37Á378

68Á59 0Á45 15Á49 3Á31 0Á06 1Á52 3Á23 3Á97 3Á21 0Á16 25 34 8 64 1350 64 428 175 16 13Á2 15 53 18 34 74 35Á4 61Á9 6Á35 23Á31 4Á48 1Á19 3Á62 0Á54 2Á97 0Á56 1Á47 0Á21 1Á24 0Á20 1343 5Á1 11Á3 15Á6 4Á6 0Á8 1Á1 19Á2 60Á6 1Á0 399 7Á7 169 0Á88 0Á70490 (3) 0Á51235 17Á119 15Á418 36Á904

68Á43 0Á47 15Á84 2Á98 0Á06 1Á42 3Á42 3Á86 3Á37 0Á16 22 25 3 48 1491 50 469 190 15 11Á6 14 54 15 33 65 35Á2 59Á7 6Á21 23Á26 4Á49 1Á28 3Á74 0Á55 3Á04 0Á57 1Á49 0Á21 1Á31 0Á20 1499 4Á8 11Á2 15Á4 5Á0 0Á7 1Á1 17Á8 47Á4 1Á0 444 6Á8 181 0Á93 0Á70451 (1) 0Á51234 17Á181 15Á427 36Á981

68Á57 0Á45 15Á59 2Á87 0Á06 1Á51 3Á51 3Á62 3Á65 0Á17 30 31 7 40 1289 66 444 166 15 10Á9 16 54 18 42 45 34Á3 59Á6 6Á18 22Á73 4Á47 1Á12 3Á52 0Á52 2Á87 0Á56 1Á45 0Á21 1Á31 0Á21 1248 4Á8 11Á4 15Á8 4Á4 0Á8 1Á1 17Á4 64Á4 0Á9 430 7Á4 159 0Á83



ROWE et al.


T 1: Continued able
Sample: Formation: Latitude (8N): Longitude (8W): MR00-81 P 35 54Á973 106 13Á755 MR00-86B P 35 54Á973 106 13Á755 MR00-4 TE 35 51Á247 106 24Á776 MR00-13 TE 36 1Á943 106 28Á324 MR00-76 TE 36 5Á72 106 23Á68 MR00-1 TR 35 51Á247 106 24Á776 MR00-14 ER 36 1Á988 106 29Á122 MR00-56 ER 36 4Á603 106 25Á168 MR00-47 LGP 36 1Á122 106 30Á941 MR00-67 LGP 36 3Á374 106 33Á418 MR00-68 LGP 36 02Á440 106 32Á144

XRF major element (wt %) 68Á63 SiO2 TiO2 0Á45 Al2O3 15Á44 à FeO 2Á92 MnO 0Á06 MgO 1Á60 CaO 3Á49 Na2O 3Á68 K 2O 3Á56 P2O5 0Á17 XRF trace element (ppm) Ni 27 Cr 31 Sc 5 V 50 Ba 1297 Rb 65 Sr 435 Zr 168 Y 16 Nb 10Á4 Cu 10 Zn 53 Pb 19 La 33 Ce 57 ICP-MS trace element (ppm) La 35Á0 Ce 60Á7 Pr 6Á31 Nd 23Á10 Sm 4Á41 Eu 1Á17 Gd 3Á53 Tb 0Á53 Dy 3Á01 Ho 0Á56 Er 1Á48 Tm 0Á21 Yb 1Á28 Lu 0Á20 Ba 1274 Th 4Á7 Nb 11Á2 Y 15Á8 Hf 4Á4 Ta 0Á8 U 1Á0 Pb 17Á5 Rb 63Á3 Cs 0Á9 Sr 442 Sc 7Á6 Zr 162 Eu/Euà 0Á88 Isotope ratios (87Sr/86Sr)i 143 Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb

67Á47 0Á49 15Á77 3Á12 0Á06 1Á76 3Á84 4Á16 3Á13 0Á20 28 26 10 58 1298 50 597 146 14 13Á1 19 51 12 23 51 34Á2 54Á8 5Á53 20Á09 3Á87 1Á19 3Á07 0Á45 2Á47 0Á49 1Á26 0Á18 1Á11 0Á18 1272 6Á8 13Á1 13Á6 3Á9 0Á9 1Á9 15Á9 48Á7 0Á9 599 7Á6 140 1Á02

68Á31 0Á52 15Á61 3Á57 0Á07 1Á54 3Á07 4Á05 3Á09 0Á18 26 35 10 62 1199 55 426 197 23 19Á4 8 94 28 34 72

56Á17 1Á20 18Á12 6Á92 0Á12 3Á51 6Á80 4Á15 2Á28 0Á73 13 7 15 145 1297 32 1252 228 27 35Á9 14 87 14 73 136 81Á3 125Á9 13Á59 50Á11 9Á18 2Á65 7Á37 1Á04 5Á68 1Á07 2Á69 0Á38 2Á28 0Á35 1324 11Á9 36Á1 29Á8 5Á3 2Á1 2Á8 15Á2 33Á9 1Á1 1231 15Á1 222 0Á95 0Á70440 (1) 0Á51259 17Á876 15Á492 37Á689

57Á43 1Á50 17Á63 7Á08 0Á15 2Á98 5Á79 4Á36 2Á44 0Á64 5 6 15 106 1156 48 641 222 32 35Á7 27 89 14 52 76 48Á7 84Á4 9Á25 35Á50 7Á46 2Á21 6Á59 1Á05 6Á17 1Á20 3Á16 0Á44 2Á75 0Á43 1252 7Á5 37Á2 32Á1 4Á9 2Á3 2Á1 12Á2 47Á7 1Á4 673 13Á7 207 0Á94 0Á70413 (2) 0Á51267 17Á604 15Á477 37Á451

75Á15 0Á18 13Á21 1Á59 0Á06 0Á09 0Á32 4Á53 4Á86 0Á03 4 0 9 0 186 102 36 314 39 49Á3 0 59 18 72 142 76Á5 126Á0 13Á74 47Á01 9Á25 0Á42 7Á61 1Á26 7Á61 1Á50 4Á11 0Á61 3Á87 0Á61 203 14Á1 50Á4 39Á5 9Á1 3Á3 3Á4 17Á7 99Á8 1Á9 33 2Á6 322 0Á15 0Á7051 (33) 0Á51256 17Á938 15Á507 37Á728

73Á22 0Á27 14Á72 1Á47 0Á10 0Á34 1Á41 4Á34 4Á04 0Á09 12 6 5 16 1455 92 228 178 23 25Á2 5 42 19 45 72 44Á1 74Á0 7Á56 26Á53 5Á04 1Á13 4Á25 0Á68 4Á15 0Á85 2Á40 0Á36 2Á39 0Á39 1457 10Á2 26Á2 24Á4 4Á9 2Á0 2Á6 22Á2 87Á3 2Á5 225 3Á8 170 0Á73 0Á70520 (2) 0Á51246 17Á551 15Á467 37Á529

77Á15 0Á08 12Á87 0Á48 0Á06 0Á05 0Á50 3Á91 4Á88 0Á02 11 1 7 2 26 149 10 57 21 42Á3 1 29 24 20 36 18Á7 35Á8 3Á82 12Á99 3Á51 0Á20 3Á21 0Á58 3Á59 0Á72 2Á10 0Á32 2Á06 0Á31 36 18Á0 43Á4 21Á4 3Á1 4Á5 7Á8 28Á3 143Á8 4Á7 9 2Á9 59 0Á18

53Á94 1Á21 18Á20 7Á76 0Á13 4Á68 8Á18 3Á48 1Á81 0Á62 29 39 19 192 1079 36 1215 169 27 31Á5 83 85 10 68 88 61Á0 95Á6 10Á73 40Á71 7Á86 2Á31 6Á46 0Á93 5Á19 1Á01 2Á66 0Á37 2Á24 0Á35 1045 8Á2 30Á2 26Á1 4Á0 1Á7 2Á1 13Á3 37Á6 1Á5 1154 17Á8 145 0Á96

62Á16 0Á89 16Á92 5Á23 0Á15 2Á81 4Á90 3Á84 2Á80 0Á29 40 35 18 104 1204 59 626 220 22 24Á9 20 63 20 53 74 45Á4 72Á3 8Á22 30Á09 5Á95 1Á62 4Á86 0Á78 4Á49 0Á86 2Á32 0Á33 2Á07 0Á32 1223 11Á0 25Á7 22Á9 5Á8 1Á8 2Á9 21Á1 60Á8 2Á2 628 13Á6 216 0Á89 0Á70496 (1) 0Á51249 17Á806 15Á519 37Á804

68Á25 0Á50 15Á80 2Á67 0Á08 1Á37 2Á77 4Á51 3Á89 0Á16 24 14 10 45 3450y 80 310 317 26 21Á4 10 54 17 104 144 115Á9 169Á8 16Á02 54Á77 8Á35 2Á01 6Á39 0Á90 5Á09 1Á02 2Á80 0Á42 2Á64 0Á42 3496 15Á0 23Á5 28Á1 8Á1 1Á5 3Á0 20Á7 82Á3 3Á2 319 10Á4 324 0Á81







T 1: Continued able
Sample: Formation: Latitude (8N): Longitude (8W): MR00-111 LGP 36 4Á15 106 32Á45 JM93144 BH 35 41Á03 106 24Á38 JM93251 BH 35 42Á58 106 26Á83 JM9373-2 BH 35 44Á15 106 26Á55 JM93180 BH 35 42Á42 106 32Á42 JM9384 PC 35 39Á77 106 23Á00 JM9307 PC 35 46Á83 106 25Á03 JM93192 PC 35 40Á37 106 35Á97 MR0023 PC 36 1Á63 106 17Á13 MR0024 PC 36 1Á42 106 17Á27

XRF major element (wt %) SiO2 69Á26 TiO2 0Á55 Al2O3 16Á25 FeOÃ 2Á76 MnO 0Á06 MgO 0Á63 CaO 1Á63 Na2O 4Á70 4Á05 K 2O P 2O5 0Á10 XRF trace element (ppm) Ni 23 Cr 7 Sc 9 V 44 Ba 3997y Rb 83 Sr 223 Zr 373 Y 25 Nb 23Á7 Cu 13 Zn 48 Pb 24 La 140 Ce 143 ICP-MS trace element (ppm) La 136Á7 Ce 167Á9 Pr 18Á76 Nd 62Á45 Sm 9Á41 Eu 2Á23 Gd 6Á94 Tb 0Á98 Dy 5Á37 Ho 1Á02 Er 2Á64 Tm 0Á39 Yb 2Á54 Lu 0Á39 Ba 3915 Th 17Á0 Nb 24Á0 Y 26Á5 Hf 9Á0 Ta 1Á5 U 3Á3 Pb 24Á1 Rb 80Á7 Cs 2Á6 Sr 208 Sc 8Á2 Zr 370 Ã Eu/Eu 0Á81 Isotope ratios (87Sr/86Sr)i 143 Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb 0Á70626 (3) 0Á51235 18Á142 15Á530 37Á796

77Á88 0Á10 12Á59 0Á61 0Á07 0Á05 0Á39 2Á29 5Á99 0Á02 5

79Á38 0Á11 11Á97 0Á54 0Á06 0Á00 0Á35 2Á90 4Á67 0Á02 4

77Á74 0Á12 12Á66 0Á63 0Á03 0Á08 0Á38 3Á69 4Á66 0Á02 5

77Á57 0Á13 12Á58 0Á65 0Á07 0Á06 0Á42 3Á84 4Á66 0Á03 4

60Á63 1Á40 17Á52 6Á36 0Á10 1Á09 5Á16 4Á54 2Á68 0Á52 16

63Á16 0Á90 16Á44 4Á89 0Á09 2Á42 4Á74 3Á92 3Á08 0Á37 21

64Á51 0Á77 16Á67 4Á18 0Á06 2Á25 5Á02 3Á52 2Á79 0Á21 20

65Á83 0Á56 16Á85 3Á38 0Á07 2Á11 4Á29 4Á49 2Á13 0Á29 29 24 9 70 1313 22 911 126 11 20Á2 11 55 18 38 59 55Á8 78Á5 9Á24 33Á56 6Á38 1Á66 4Á56 0Á67 3Á52 0Á63 1Á52 0Á21 1Á25 0Á19 1290 9Á1 21Á2 16Á8 3Á6 1Á2 2Á8 17Á6 28Á1 0Á9 947 8Á9 132 0Á89 0Á70390 (1) 0Á51263 17Á941 15Á510 37Á868

65Á89 0Á57 16Á90 3Á71 0Á07 1Á90 3Á91 4Á56 2Á21 0Á28 29 24 5 60 1273 28 880 136 13 20Á8 16 54 16 48 44 39Á0 52Á7 5Á45 19Á39 3Á61 1Á19 2Á92 0Á42 2Á36 0Á45 1Á14 0Á16 0Á98 0Á15 1320 8Á2 20Á1 12Á0 3Á3 1Á1 2Á5 17Á5 23Á3 1Á0 910 8Á9 122 1Á09 0Á70390 (2) 0Á51258 17Á979 15Á505 37Á854

142 16 89 26 26 1 32 22

113 8 91 27 27 31 24

120 26 90 25 26 2 25 24

124 25 87 24 25 21 25

42 764 302 27 37 11 49 13

54 778 253 26 29 21 58 19

44 746 141 14 15 42 45 18

23Á1 46Á7 4Á93 17Á01 4Á10 0Á46 3Á84 0Á69 4Á54 0Á93 2Á56 0Á38 2Á51 0Á38 419 11Á9 27Á1 25Á9 3Á4 2Á0 3Á8 22Á5 142Á2 4Á1 23 3Á6 81 0Á35 0Á7079 (20) 0Á51237 17Á858 37Á678

20Á4 42Á8 4Á55 16Á17 4Á16 0Á44 3Á93 0Á73 4Á55 0Á94 2Á65 0Á40 2Á58 0Á40 220 11Á5 29Á6 28Á5 3Á6 2Á1 4Á0 22Á3 118Á6 4Á3 12 3Á8 90 0Á32 0Á7078 (40) 0Á51236 17Á842 37Á607

25Á1 47Á7 4Á97 17Á28 3Á79 0Á51 3Á41 0Á63 3Á83 0Á79 2Á23 0Á34 2Á18 0Á35 675 12Á7 26Á6 23Á6 3Á4 1Á9 3Á9 23Á4 115Á7 4Á0 28 3Á3 85 0Á42 0Á7080 (20) 0Á51235 17Á853 37Á530

28Á2 52Á4 5Á21 17Á76 4Á04 0Á48 3Á72 0Á66 4Á10 0Á84 2Á36 0Á36 2Á33 0Á36 620 13Á5 27Á6 24Á6 3Á4 2Á1 4Á1 25Á6 126Á1 4Á4 30 3Á6 86 0Á37

50Á7 84Á1 8Á91 33Á97 6Á73 2Á04 5Á68 0Á88 5Á12 1Á00 2Á66 0Á39 2Á46 0Á39 895 9Á9 39Á6 27Á3 6Á4 2Á6 2Á8 10Á8 44Á3 1Á0 746 14Á5 281 0Á98

50Á8 83Á3 9Á03 33Á26 6Á38 1Á67 5Á27 0Á80 4Á79 0Á92 2Á44 0Á36 2Á22 0Á35 1185 9Á9 29Á6 24Á8 5Á8 1Á8 2Á9 17Á9 56Á8 1Á7 759 11Á0 233 0Á86 0Á70467 (2) 0Á51251 17Á814 37Á643

27Á8 42Á2 4Á66 17Á84 3Á36 1Á06 3Á16 0Á44 2Á76 0Á51 1Á33 0Á19 1Á18 0Á19 1338 5Á0 15Á4 13Á8 3Á5 1Á1 1Á6 18Á0 46Á9 1Á5 735 10Á6 133 0Á98 0Á70416 (2) 0Á51259 17Á535 37Á408

0Á7068 (20) 0Á70392 (2) 0Á51238 0Á51272 17Á497 18Á080 37Á160 37Á928



ROWE et al.


T 1: Continued able
Sample: JM9369 Formation: PC Latitude (8N): 35 47Á77 Longitude (8W): 106 24Á98 XRF major element (wt %) SiO2 66Á08 TiO2 0Á67 Al2O3 16Á57 Ã FeO 3Á77 MnO 0Á06 MgO 1Á49 CaO 4Á26 Na2O 3Á68 K2O 3Á20 0Á22 P2O5 XRF trace element (ppm) Ni 14 Cr Sc V Ba Rb 49 Sr 609 Zr 170 Y 16 Nb 14 Cu 23 Zn 58 Pb 16 La Ce JM93107 PC 35 43Á17 106 24Á83 JM9336 PC 35 44Á33 106 22Á08 JM93124 PC 35 44Á83 106 23Á27 JM93214 PC 35 47Á47 106 21Á40 JM9338 PC 35 44Á50 106 22Á07 JM93239 CC 35 46Á95 106 33Á40 JM93123 CC 35 44Á73 106 23Á32 JM93139 JM125 JM126 JM145 JM146 CC SDC A SDC A SDC B SDC B 35 44Á97 106 20Á83

66Á86 0Á70 15Á57 4Á11 0Á05 1Á31 3Á45 4Á15 3Á50 0Á28 36

67Á09 0Á80 16Á19 3Á70 0Á12 0Á65 2Á87 4Á83 3Á44 0Á31 5

68Á83 0Á61 16Á10 2Á42 0Á07 0Á41 1Á37 5Á19 4Á86 0Á14 4

70Á09 0Á64 14Á99 2Á52 0Á05 0Á49 1Á58 4Á89 4Á58 0Á17 4

70Á94 0Á76 13Á73 3Á55 0Á09 0Á80 2Á69 4Á08 3Á05 0Á31 8

61Á07 1Á02 16Á97 5Á52 0Á09 2Á65 5Á46 3Á89 2Á94 0Á39 17

77Á75 0Á10 12Á16 1Á29 0Á03 0Á00 0Á16 3Á91 4Á58 0Á02 4

76Á81 0Á12 12Á92 0Á71 0Á05 0Á05 0Á48 4Á13 4Á72 0Á02 4

75Á50 0Á17 13Á38 1Á56 0Á06 0Á33 0Á61 2Á97 5Á38 0Á03 6 0 5 11 152 151 31 241 40 61Á3 4 61 27 73 110

74Á62 0Á27 13Á22 1Á85 0Á07 0Á32 0Á87 3Á78 4Á93 0Á07 7 0 8 14 264 147 84 222 41 63Á2 5 68 24 61 95

76Á73 0Á10 12Á83 1Á27 0Á06 0Á12 0Á39 3Á23 5Á26 0Á01 1 0 5 7 45 169 9 229 44 68Á1 2 63 27 56 115

76Á89 0Á09 12Á60 1Á19 0Á06 0Á06 0Á34 3Á65 5Á10 0Á01 6 0 3 3 35 163 7 228 44 67Á5 0 64 28 53 98

67 677 227 21 26 18 48 19

63 608 331 24 38 9 48 17

107 264 477 35 53 8 40 28

103 277 405 34 49 2 35 19

54 485 291 25 35 10 31 17

49 801 242 25 28 21 58 15

148 17 215 46 65 5 49 16

113 47 107 21 48 3 23 28

ICP-MS trace element (ppm) La 30Á8 51Á7 56Á6 72Á3 66Á4 52Á5 45Á8 55Á6 33Á6 62Á6 46Á7 55Á3 55Á5 Ce 51Á8 72Á5 96Á0 113Á4 109Á4 85Á8 74Á8 77Á6 53Á6 109Á0 82Á0 97Á7 97Á2 Pr 5Á67 8Á81 9Á49 11Á83 11Á41 9Á11 8Á20 10Á58 5Á24 10Á98 8Á45 9Á95 9Á92 Nd 21Á45 32Á95 33Á43 41Á51 40Á79 33Á65 31Á32 35Á91 16Á84 38Á55 30Á33 35Á32 34Á95 Sm 4Á25 5Á71 6Á18 8Á01 7Á63 5Á75 5Á82 8Á26 3Á37 8Á02 6Á81 7Á82 7Á80 Eu 1Á25 1Á53 1Á73 1Á66 1Á69 1Á54 1Á63 0Á15 0Á46 0Á37 0Á69 0Á16 0Á17 Gd 3Á57 4Á84 4Á79 6Á50 6Á25 4Á90 5Á04 7Á05 2Á82 6Á85 6Á11 7Á22 7Á07 Tb 0Á54 0Á72 0Á78 1Á01 1Á00 0Á75 0Á80 1Á32 0Á51 1Á19 1Á09 1Á31 1Á29 Dy 3Á04 4Á17 4Á57 6Á41 5Á99 4Á30 4Á42 8Á28 3Á29 7Á23 6Á71 8Á01 7Á96 Ho 0Á58 0Á79 0Á90 1Á24 1Á17 0Á83 0Á87 1Á68 0Á68 1Á45 1Á36 1Á62 1Á60 Er 1Á53 2Á06 2Á45 3Á32 3Á19 2Á31 2Á37 4Á49 1Á95 4Á02 3Á76 4Á47 4Á46 Tm 0Á22 0Á31 0Á37 0Á50 0Á49 0Á34 0Á33 0Á66 0Á31 0Á60 0Á56 0Á68 0Á67 Yb 1Á28 1Á84 2Á32 3Á28 3Á12 2Á13 2Á00 4Á23 2Á11 3Á89 3Á62 4Á35 4Á30 Lu 0Á21 0Á29 0Á38 0Á50 0Á50 0Á34 0Á33 0Á62 0Á33 0Á60 0Á55 0Á67 0Á66 Ba 1172 1264 1300 1345 1178 1154 1055 83 389 128 307 42 42 Th 4Á4 11Á0 12Á6 21Á2 18Á7 10Á5 9Á1 18Á0 24Á1 22Á2 17Á5 21Á7 21Á3 Nb 13Á1 27Á8 40Á9 54Á4 54Á2 37Á1 26Á8 66Á2 50Á3 63Á6 55Á0 70Á8 68Á5 Y 16Á0 22Á6 23Á8 32Á9 33Á5 25Á1 24Á0 43Á0 19Á9 40Á1 37Á4 45Á1 45Á1 Hf 4Á3 5Á3 7Á3 10Á3 9Á6 6Á2 5Á4 7Á4 3Á6 8Á0 6Á8 8Á2 8Á1 Ta 0Á8 1Á8 2Á5 3Á6 3Á3 2Á3 1Á6 4Á5 4Á3 5Á0 4Á5 5Á5 5Á3 U 1Á2 3Á5 3Á9 6Á6 6Á4 1Á7 2Á8 4Á4 6Á7 6Á0 5Á4 6Á4 6Á5 Pb 15Á8 19Á5 15Á3 27Á4 18Á3 15Á3 14Á9 15Á0 28Á0 24Á0 21Á8 26Á9 26Á6 Rb 51Á9 67Á9 64Á1 107Á5 104Á8 54Á9 46Á6 141Á3 111Á4 145Á4 128Á0 166Á2 159Á7 Cs 1Á0 0Á9 1Á9 2Á6 2Á8 2Á1 1Á3 2Á9 2Á8 3Á7 3Á5 4Á1 4Á2 Sr 609 649 597 261 279 468 752 21 47 20 91 5 3 Sc 9Á4 10Á6 6Á6 6Á0 6Á6 7Á9 12Á3 1Á2 1Á9 3Á6 4Á5 3Á0 3Á0 Zr 161 204 301 404 400 259 215 196 98 238 197 230 229 Eu/Euà 0Á95 0Á86 0Á94 0Á68 0Á73 0Á86 0Á90 0Á06 0Á44 0Á15 0Á32 0Á06 0Á07 Isotope ratios 87 86 ( Sr/ Sr)i 0Á70477 (2) 0Á70470 (2) 0Á70464 (2) 0Á70463 (6) 0Á70488 (2) 0Á70468 (2) 0Á70462 (9) 143 Nd/144Nd 0Á51240 0Á51253 0Á51257 0Á51263 0Á51254 0Á51258 0Á51254 0Á51251 0Á51257 206 204 Pb/ Pb 17Á313 17Á815 17Á980 17Á987 18Á175 17Á986 17Á243 17Á833 18Á072
207 208

Pb/204Pb Pb/204Pb










XRF and ICP-MS analyses performed at Washington State University. All major element analyses normalized to 100% volatile-free. [See text and Wolff et al. (2005) for isotope laboratory details.] TAD, Tschicoma andesite and dacite; P, Puye; TE, Tschicoma Enclaves; TR, Tschicoma Rhyolites; ER, El Recheulos; LGP, La Grulla Plateau; BH, Bearhead; PC, Paliza Canyon; CC, Canovas Canyon; SDC A, San Diego Canyon A; SDC B, San Diego Canyon B. Error in calculated (87Sr/86Sr)i reported to fifth decimal place. Ã Total Fe given as FeO. y Denotes values 4120% of highest standards.






similar to Tschicoma rhyolites than to the Bearhead main group lavas of Justet & Spell (2001). Three domes, located along the east flank of Polvadera Peak, are dated at 2 Ma (Loeffler et al., 1988), consistent with the original stratigraphic definition (Bailey et al., 1969). They are high-silica rhyolites with sparse microphenocrysts of sanidine, sodic plagioclase and quartz.

Cerros del Rio volcanic field
This dominantly mafic volcanic field, with bothType I and II mafic compositions, is exposed in the axis of the Espanola basin, to the east of the JMVF (Fig. 1). ‹ Domes and flows of benmoreite (trachyandesite with Na2O ^ 2Á0 ! K2O) with up to 63% SiO2 occur in the northern part of the field (Duncker et al., 1991). These lavas temporally overlap and post-date late Tschicoma activity and form a compositional continuum with Tschicoma dacites, their enclaves, and the dominant mafic lavas of the Cerros del Rio (Wolff et al., 2005). Duncker et al. (1991) found that 143Nd/144Nd systematically decreases, and d18O increases, with increasing silica content among the Cerros del Rio lavas, and concluded that the more silicic (4 57% SiO2) flows are related to the mafic magmas of the field through mixing with silicic melts derived from regional crust, or assimilation^fractional crystallization of Type I mafic magma with continental crust as the assimilant, or some combination of the two processes.

San Diego Canyon ignimbrites
These two high-silica (75^78% SiO2) rhyolitic ignimbrites were erupted at 1Á85 Ma and are now found beneath the Bandelier Tuff in upper Canon San Diego (Fig. 1), and in ‹ wells drilled inside the caldera (Hulen et al., 1991). A correlative pumice fallout deposit occurs interbedded with sediments in the uppermost part of the Puye Formation (Turbeville & Self, 1988). All three units show strong geochemical affinities with the Bandelier Tuff and are regarded as early products of the Bandelier magma system (Spell et al., 1990). They are much smaller than the climactic Bandelier ignimbrites, with probable volumes 510 km3 (Turbeville & Self, 1988).

(ICP-MS), and Sr, Nd and Pb isotope ratios by thermal ionization mass spectrometry (TIMS) and multicollector (MC)-ICP-MS, plus analyses of Type I and II mafic lavas and regional Precambrian crustal rocks referred to in subsequent sections of the paper have been reported by Wolff et al. (2005). It should be noted that, as a result of difficulties experienced during analysis, we do not report 207 Pb/204Pb ratios for the JM93- series of samples. 87 Sr/86Sr values are age-corrected using ICP-MS Rb and Sr data. Most samples have low Rb/Sr, thus most corrections are trivial (50Á00005), but there are errors associated with Rb and Sr precision and with uncertain sample ages in many cases; a potential age range of 8^12 Ma has been assumed for each Paliza Canyon sample. Similarly, where ages for Tschicoma and El Rechuelos Formation lavas were unknown a potential age range of 6Á9^2Á7 Ma is assumed (7Á9^7Á4 Ma for La Grulla Plateau lavas). Calculated (87Sr/86Sr)i values, with errors, are reported in Table 1. Initial 87Sr/86Sr values for the two Canovas Canyon high-silica rhyolites were rejected because they have experienced post-eruptive alteration, and were thus open to Rb and Sr exchange since eruption. Bearhead Rhyolite samples are fresh, and consequently (87Sr/86Sr)i values are reported in Table 1, although with large errors; the ages of these samples are accurately known (Justet & Spell, 2001), but significant imprecision arises from uncertainties associated with Rb and, especially, Sr contents in the trace element analyses of these high-Rb/Sr samples.

Major elements
Variations among major elements (Figs 3 and 4) show some systematic differences between the various JMVF formations. Paliza Canyon Formation dacites have a significantly greater range in total alkali content than the Tschicoma Formation dacites from the NE and east JMVF (6Á3^8Á9% and 6Á3^7Á1% respectively at 65% SiO2, Fig. 3) and, with some exceptions, are enriched in K2O relative to Tschicoma Formation lavas between 60 and 70 wt% SiO2 (Fig. 4). However, Tschicoma dacites from the La Grulla Plateau show a similar range to that of the Paliza Canyon Formation dacites. For most major oxides, the compositions of Paliza Canyon lavas typically overlap with those of the Tschicoma dacites at constant SiO2. Paliza Canyon Formation lavas are nonetheless skewed to slightly higher TiO2 and lower MgO over most of the compositional range, and to lower CaO among the dacites, without any clear systematic contrasts between the La Grulla and the NE to east groups of Tshicoma lavas (Fig. 4). With the exception of alkalis, these differences are not apparent at the mafic end of the compositional range. Trachyandesites from the Cerros del Rio tend to have slightly elevated Na2O and lower TiO2 contents, but are otherwise chemically similar to mafic enclaves recovered from Tschicoma Formation dacites (Table 1). In general, major element

Representative chemical compositions, including all samples that were analyzed for Sr, Nd and Pb isotopes, are given in Table 1. The complete set of chemical analyses used in this study (n ¼196) can be found in Electronic Appendix 1 (available for downloading from http://www. and the data are plotted in Figs 3^5. Analytical procedures for major and trace elements by X-ray fluorescence (XRF), trace elements by quadrupole inductively coupled plasma mass spectrometry


ROWE et al.


variations among the NE to east Tschicoma lavas tend be more linear than among Paliza Canyon flows. For example, in the MgO vs SiO2 plot (Fig. 4c), the Paliza Canyon lavas have an overall larger range in MgO despite generally lower concentrations relative to Tschicoma Formation lavas such that the maximum MgO at a given SiO2 concentration in Paliza Canyon Formation lavas is equivalent to that in Tschicoma Formation lavas. Rhyolites of all ages (Canovas Canyon, Bearhead, Tschicoma, El Rechuelos, and San Diego Canyon

Formations) range from 70 to 79 wt% SiO2 and formations are essentially indistinguishable based solely on major element contents (Figs 3 and 4).

Trace elements Strontium
Strontium concentrations vary by over two orders of magnitude, from 8 to 1380 ppm. In all formations, Sr concentrations decrease with increasing SiO2 (Fig. 5a). As with major elements, Paliza Canyon Formation lavas

Fig. 5. Selected trace element variations vs SiO2 wt % for pre-caldera JMVF lavas and tephras (mafic lava fields as Fig. 2). Otowi lithic fragments reported by Wolff et al. (2005). Simple mixing curves are plotted for mixtures between Cerros del Rio benmoreite E6-8B and Otowi lithic fragments 18L-3 and CCL-1 (dashed lines). Fractional crystallization of Paliza Canyon trachyandesite JM93190 is represented by continuous line with crosses (þ) at 10% increments (see text for fractionating assemblage). Model model compositions and parameters are given in Tables 2 and 3.






Fig. 5. (Continued)

exhibit larger variability in Sr concentration at constant SiO2 (e.g. Sr 960^1286 ppm at 58% SiO2; Sr 114^506 ppm at 69% SiO2) than do NE to east Tschicoma rocks, which have a range in Sr content of less than 150 ppm at constant SiO2 (Fig. 5a). Mafic enclaves recovered from the Tschicoma Formation exhibit a large range in Sr concentration from 640 to 1380 ppm. Cerros del Rio trachyandesites have relatively restricted Sr abundances, varying from 775 to 921ppm. Low-silica rhyolites have varying concentrations of Sr that are comparable with those for dacites of the Paliza Canyon and Tschicoma Formations. High-silica rhyolites are extremely depleted in Sr, with concentrations 84 ppm.

The overall range in Ba concentrations for the Paliza Canyon and NE to east Tschicoma lavas is less than a factor of two (913^1562 ppm). Most flows from these two formations have Ba concentrations between $1100 and 1300 ppm with only very slight enrichment between 60 and 70% SiO2, and high-silica rhyolites are Ba-depleted (Fig. 5b). Ba concentrations in La Grulla Plateau lavas are generally equivalent to those for the rest of the Tschicoma Formation lavas, although a few dacites have extremely high Ba concentrations (3450^4000 ppm; Fig. 5b). Excluding one anomalously low-Ba sample, Ba concentrations in Cerros del Rio


ROWE et al.


trachyandesites increase over a restricted range of SiO2, a pattern distinctly different from trends observed in either the Paliza Canyon or Tschicoma Formation.

Pb concentrations in JMVF lavas on average increase from $12 ppm to $24 ppm with increasing SiO2 over the entire compositional range (Fig. 5f). Little systematic variation is evident within each formation. Tschicoma Formation lavas from the La Grulla Plateau have the greatest overall variation in Pb, ranging from 11 to 31ppm at $67% SiO2; however, the average concentration of $17 ppm is similar to that of both the Paliza Canyon and Tschicoma Formation lavas.

Below $65 wt% SiO2, Rb concentrations show little systematic covariation with SiO2; concentrations range from 22 to 85 ppm (Fig. 5c). As with Ba, there is little difference between Paliza Canyon Formation and Tschicoma Formation lavas. Above $65 wt% SiO2, Rb concentrations increase as silica increases, with the high-silica rhyolites containing the greatest concentration of Rb (up to 173 ppm). Cerros del Rio trachyandesites show little variation in Rb, with concentrations varying from 26 to 40 ppm.

Thorium and uranium
U, and to a greater extent, Th appears to behave similarly to Nb and Zr. In general, the variability in U and Th concentrations increases with increasing SiO2 in Paliza Canyon Formation lavas (Fig. 5g and h). U and Th concentrations broadly decrease with increasing SiO2 in Tschicoma and Puye formation samples. Overall Tschicoma and Puye Formation samples have a more restricted range in U and Th concentrations (1^3Á4 ppm and 4Á5^13Á3 ppm respectively, excluding La Grulla Plateau lavas) than Paliza Canyon Formation lavas (1Á2^6Á6 ppm U and 4Á4^23Á7 ppm Th). Bearhead Formation rhyolites are generally depleted in both U and Th relative to other high-silica rhyolites and have concentrations similar to those in Paliza Canyon Formation dacites.

Niobium and zirconium
In Paliza Canyon Formation lavas, Nb and Zr concentrations show large variations at constant SiO2 (Fig. 5d and e). At $58 wt% SiO2, Nb and Zr concentrations are clustered at $35 ppm Nb and $300 ppm Zr. As SiO2 increases, both Nb and Zr fan out to higher and lower concentrations (13^64 ppm and 123^539 ppm, respectively). In distinct contrast, Nb in Tschicoma Formation lavas with 570% SiO2 (including those of the La Grulla Plateau), mafic enclaves, and Cerros del Rio trachyandesites consistently decreases with increasing SiO2. Tschicoma Formation dacites further form distinct highNb and low-Nb groups ($19 and 11ppm Nb respectively at 68Á5% SiO2; Fig. 5d). These Nb groupings are geographically restricted; the low-Nb group lavas are found on the eastern side of the JMVF, exemplified by Pajarito Mt. and Cerro Rubio dome samples (Fig. 1), and conglomerate clasts and pumice from the Puye ignimbrite of Turbeville et al. (1989). The high-Nb group dominates the NE Tschicoma dacites, exemplified by Tschicoma Mt. and Polvadera Peak lavas. La Grulla Plateau lavas cover the range in Nb observed in both the high- and low-Nb groups. Zr concentrations in the NE to east Tschicoma lavas, like Nb, exhibit an overall decrease with increasing SiO2 but unlike Nb do not fall into two distinct groups among the dacites; instead, there is little systematic variation in Zr between 65 and 69% SiO2. Zr in the La Grulla lavas is highly variable, and therefore more similar to the Paliza Canyon than to the NE to east Tschicoma Formation. Nb and Zr are highly variable in the rhyolites, with some tendency for different formations to fall into distinct groups; for example, the Bearhead rhyolites are notably low in Zr and Nb compared with the rest of the JMVF.

Rare earth elements
Rare earth elements (REE) are highly variable in all JMVF formations (see Ce and Yb vs SiO2; Fig. 5i and j). As with Nb, Zr, U and Th, the Paliza Canyon Formation has the greatest overall variability in REE concentrations. REE inTschicoma Formation lavas and mafic enclaves and Puye clasts define a broad negative correlation with increasing SiO2, although the high-Ba La Grulla Plateau lavas (Fig. 5b) are also light REE (LREE) enriched with Ce concentrations up to 150 ppm (Fig. 5i). Cerros del Rio trachyandesites fall within the broad range of compositions observed from Tschicoma lavas. In high-silica rhyolites, LREE concentrations span the entire range exhibited by the other JMVF formations. However, most high-silica rhyolites have much lower LREE/HREE (heavy REE) ratios than less evolved compositions (Table 1; Electronic Appendix 1). JMVF lavas with less than 67 wt% SiO2 lack a significant Eu anomaly, with a Eu/Euà ratio between 1Á0 and 0Á8, identical to the range observed in the ‘Type I’ mafic lavas (Fig. 5k; Wolff et al., 2005). Eu is depleted with respect to Sm and Gd in several of the low-silica rhyolites and, like Sr and Ba, shows extreme depletion among the high-silica rhyolites.






Radiogenic isotopes Lead isotopes
Pb isotope ratios in Paliza Canyon lavas do not correlate with SiO2 content (Fig. 6a and b); 206Pb/204Pb ¼17Á24 ^ 18Á18 and 208Pb/204Pb ¼ 36Á95 ^ 37Á87, within the ranges reported by Wolff et al. (2005) for the JMVF mafic lavas. In contrast, 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb in Tschicoma lavas, excluding the La Grulla Plateau, are

negatively correlated with SiO2 at 570% SiO2. La Grulla Plateau lavas have more radiogenic Pb signatures than other Tschicoma rocks. Mafic enclaves from Tschicoma lavas have Pb isotope ratios (206Pb/204Pb ¼17Á60 ^ 17Á88, 208 Pb/204Pb ¼ 37Á45 ^ 37Á69) that encompass the small range in the least evolved ($63% SiO2) Tschicoma lavas and Cerros del Rio trachyandesites (Fig. 6a and b). Rhyolitic lavas and tuffs overall have a large range in Pb

Fig. 6. (a) 206Pb/204Pb, (b) 208Pb/204Pb, (c) 87Sr/86Sr, and (d) 143Nd/144Nd vs SiO2 for pre-caldera volcanics. Dashed lines are simple mixing curves between a Cerros del Rio basalt (H I-5) and Otowi lithic fragment 18L-3, a Bearhead rhyolite (JM9373-2), and a Tschicoma Formation rhyolite (MR00-1). Continuous lines are simple mixing curves between Cerros del Rio benmoreite (E6-19a) and Otowi lithic fragments CCL-1 and 18L-3 [extending to 0Á72484 and 0Á74349 in (c)] and Bearhead rhyolite JM9373-2, with crosses (þ) at 10% increments. Bold box is the isotopic range observed in the Bandelier and San Diego Canyon ignimbrites, excluding very low-Sr rhyolites that have been affected by minor late-stage contamination (Wolff & Ramos, 2003, and unpublished data). Sr isotope data for the La Grulla Plateau (SK86; Singer & Kudo, 1986), and Sr and Nd isotopic data for the El Rechuelos rhyolites (L88; Loeffler et al., 1988) also are plotted for comparison. Both Type I and Type II mafic lavas are included in ‘Mafic lavas’ field.


ROWE et al.


isotope ratios, with 206Pb/204Pb varying from 17Á50 to 18Á07 and 208Pb/204Pb from 36Á95 to 37Á73, similar to the overall variability observed in the Paliza Canyon lavas (Fig. 6a and b). The San Diego Canyon ignimbrites and calderaforming Bandelier Tuff (206Pb/204Pb ¼17Á79 ^ 18Á02, 208 Pb/204Pb ¼ 37Á52 ^ 37Á87; Wolff & Ramos, 2003, and unpublished data) overlap the radiogenic ends of the ranges seen in the Paliza Canyon lavas and earlier rhyolites, but are more radiogenic than most Tschicoma dacites.

P E T RO G E N E S I S Identification of components
Intermediate and silicic lavas and tuffs make up 490% of the volume of the JMVF, and the diversity of their major element, trace element and isotopic compositions attests to continual magma^crust interaction throughout the lifetime of the volcanic field. Several petrogenic models, including crustal melting, simple mixing, fractional crystallization, and energy constrained assimilation^fractional crystallization (EC-AFC; Bohrson & Spera, 2001, 2003; Spera & Bohrson, 2001) may be invoked to explain this diversity. T test different models against the geochemical o dataset, some assumptions regarding the choice of starting compositions must be made. Below, we outline a rationale for selection of basaltic and crustal end-member compositions used in petrogenetic modeling. In addition to the basaltic and crustal end-members, some rhyolites and trachyandesites have been used in mixing calculations: a Tschicoma rhyolite (MR00-1), Bearhead rhyolite (JM9373-2), Cerros del Rio benmoreites (E6-19a and E6-8a) and Paliza Canyon Formation trachyandesites (JM9384 and JM93190). End-member compositions used in various models are reported in Table 2 and model parameters for EC-AFC and fractional crystallization simulations are provided in Table 3.

Strontium isotopes
Paliza Canyon lavas have a relatively restricted range in initial 87Sr/86Sr from 0Á70392 to 0Á70490 (Fig. 6c). The bulk of the Tschicoma lavas and mafic enclaves also have a narrow compositional range (0Á70413^0Á70490). However, a Tschicoma Formation rhyolitic lava is slightly enriched in 87Sr/86Sri at 0Á7051, and lavas from the La Grulla Plateau [87Sr/86Sri ¼ 0Á70496 ^ 0Á70626, similar to the range reported by Singer & Kudo (1986) of 0Á7051^0Á7069] are significantly more radiogenic than the rest of the Tschicoma and Paliza Canyon flows (Fig. 6c). An El Rechuelos rhyolite has 87Sr/86Sri ¼ 0Á70520, similar to that of the Tschicoma Formation rhyolite and consistent with previously reported values (87Sr/86Sri ¼ 0Á7050 ^ 0Á70566; Loeffler et al., 1988). Bearhead high-silica rhyolites have the highest 87Sr/86Sr of the pre-caldera JMVF volcanics, varying from 0Á7068 to 0Á7080. The exceptionally Sr-depleted San Diego Canyon and Bandelier Tuff highsilica rhyolites range to still higher values of (87Sr/86Sr)i, attributed by Wolff et al. (1999) and Wolff & Ramos (2003) to trivial amounts of wall-rock contamination following rhyolite genesis. The least radiogenic Bandelier (87Sr/86Sr)i ratios of 0Á7041^0Á7056 (Skuba and Wolff, 1990; Wolff et al., 1999; Fig. 6c) correspond to the bulk of the Paliza Canyon and Tschicoma lavas.

T I basalts ype
The abundances of, and ratios between, many major and trace elements in intermediate magmas form continua with those of the Type I mafic lavas but are dissimilar to Type II tholeiites (Figs 3, 4f, g, 5 and 7). High concentrations of P, Sr, Ce, Nb, Zr, U, and Th in Type I basaltic lavas coincide with those in the evolved trachyandesites (i.e. $57^60 wt% SiO2) and are distinctly enriched relative to Type II basaltic lavas (Figs 4f and 5). Type I mafic lavas are therefore considered to represent the mafic parent or end-member involved in petrogenesis of JMVF intermediate and silicic magmas. A representative Type I Cerros del Rio hawaiite (H I-5; Table 2) is used for modeling purposes.

Neodymium isotopes
As with Pb and Sr isotopes, 143Nd/144Nd ratios in Paliza Canyon Formation lavas show poor correlation with silica content, with a range from 0Á51240 to 0Á51272 (Fig. 6d). The 143Nd/144Nd of Tschicoma Formation lavas (including La Grulla Plateau) and mafic enclaves decreases systematically with increasing SiO2 from 0Á51267 (57Á4 wt% SiO2) to 0Á51231 (68Á6 wt% SiO2). Bearhead rhyolites are isotopically distinct from Canovas Canyon rhyolites, with a narrow 143Nd/144Nd range of 0Á51238^0Á51236. El Rechuelos, Canovas Canyon, and Tschicoma Formation rhyolites have a wider range in 143 Nd/144Nd from 0Á51245 to 0Á51257. San Diego Canyon and Bandelier Tuff high-silica rhyolites (143Nd/144Nd ¼ 0Á51249 ^ 0Á51266; Skuba, 1990) show a strong similarity to the bulk of the Paliza Canyon rocks (Fig. 6d).

Precambrian crust
Wolff et al. (2005) evaluated regional basement lithologies as representatives of the crustal component(s) in JMVF mafic lavas and found that, perhaps not surprisingly, Precambrian crystalline rocks from beneath the volcanic field itself provide the best approximation of silicic compositions required for generation of the isotopic and trace element characteristics of both types of mafic lavas. The Precambrian rocks are found as rare amphibolite and granitoid lithic fragments in the Otowi Member of the Bandelier Tuff (Eichelberger & Koch, 1979). Wolff et al. (2005) showed that, with the exception of Sr isotopes, isotopic and trace element variations among JMVF mafic






T 2: Compositional end-members used in petrogenic modeling able
Formation: Sample: Reference: CdR H I-5 2, 3 CdR E 6-19a 3 CdR E 6-8B 3 PC JM9384 1, 4 PC JM93190 1, 4 Otowi JM 18L-3 2 Otowi JM CCL-1 2 TR MR00-1 4 BH JM9373-2 1, 4

SiO2 (wt %) K2O (wt %) Ce Ba Nb Pb Rb Sr Zr U Th Yb (87Sr/86Sr)i
143 206 207 208

52Á29 2Á06 122Á5 1203 49Á7 9Á8 29Á1 1300 240 — — — 0Á70413 0Á51267 18Á305 15Á529 38Á129

57Á88 1Á80 79 1084 28 11Á5 26 790 176 — — — 0Á70421 0Á51258 17Á825 15Á477 37Á632

58Á70 2Á12 95 1247 32 8 32 892 194 — — — — — — — —

60Á63 2Á68 84Á1 895 39Á6 10Á8 44Á3 746 281 — — — 0Á70392 0Á51272 18Á08 15Á501 37Á928

60Á85 2Á99 89Á2 1192 33 18 57 917 284 3Á6 11Á4 2Á3 — — — — —

71Á75 4Á42 50Á4 1279 16Á4 26Á2 177Á4 188 245 — — — 0Á74349 0Á51205 17Á163 15Á433 36Á998

73Á41 4Á86 40Á1 1800 4Á0 18Á5 178Á1 283 52 — — — 0Á72484 0Á51201 17Á568 15Á472 37Á235

75Á15 4Á86 126Á0 203 50Á4 17Á7 99Á8 33 322 — — — 0Á7051 0Á51256 17Á939 15Á504 37Á756

77Á74 4Á66 47Á7 675 26Á6 23Á4 115Á7 28 85 — — — 0Á7080 0Á51235 17Á853 15Á468 37Á53

Nd/144Nd Pb/204Pb Pb/204Pb Pb/204Pb

Formation abbreviations are as in Table 1 (CdR, Cerros del Rio). Tschicoma Pb isotopes measured by MC-ICP-MS at Washington State University. References: 1, Ellisor et al. (1996); 2, Wolff et al. (2005); 3, Duncker et al. (1991); 4, this study.

lavas could be accounted for by interaction of primitive nephelinite^basanite and tholeiitic magmas with these granitoids. For Nd and Pb isotopes and incompatible trace elements, either simple mixing of bulk granitoid or partial melting accompanied by mixing with primitive magmas (modeled as EC-AFC; Bohrson and Spera, 2001, 2003) reproduces the observed compositions of the mafic lavas. However, Wolff et al. (2005) identified a Sr^(Pb, Nd) paradox, characterized by low and uniform Sr isotope ratios relative to large variability in Pb and Nd isotopes. Wolff et al. (2005) showed that the Sr isotope characteristics could be explained by crustal melt production in the presence of residual feldspar (DSr ¼ 2), thus suppressing 87 Sr/86Sr variations in the resulting contaminated basalts, but also noted that this model is difficult to reconcile with the lack of a measurable Eu anomaly in the mafic lavas. A second possibility is that the actual assimilant is characterized by low 87Sr/86Sr, which has been previously invoked for contaminated basalts in the Rio Grande rift region (Dungan et al., 1986; McMillan & Dungan, 1988; Duncker et al., 1991), but is otherwise isotopically similar to its surface equivalents. Low-87Sr/86Sr crust is in fact found regionally (Wolff et al., 2005), but has inappropriate Pb isotope ratios for components in many JMVF lavas.

JMVF lavas with SiO2 between 57 and 67% also lack Eu anomalies (Fig. 5k), and we now consider Eu^Sr coupling in more detail. Eu partitioning between feldspar and silicate melt is strongly dependent upon oxygen fugacity (Drake, 1975; Wilke & Behrens, 1999), because Eu2þ is expected to behave similarly to Sr2þ, whereas Eu3þ should behave similarly to Sm and Gd. In oxidized magmas where Eu is mostly present as Eu3þ, the relative abundance of Eu is not sensitive to the presence of plagioclase. Nine iron^titanium oxide pairs from two representative Paliza Canyon trachyandesites, which fulfill textural and compositional (Bacon & Hirschmann, 1988) equilibrium criteria, yield T^fO2 results that cluster close to the nickel^nickel oxide (NNO) buffer. Regression of the plagioclase^liquid Eu partitioning experimental results of Drake (1975) for basaltic compositions and Wilke & Behrens (1999) for tonalitic compositions predicts D(Eu, plag/liq) values of $0Á4 and $1Á0 respectively at NNO. D(Sm, plag/liq) and D(Gd, plag/liq) values from the same experiments are 0Á1. Hence, a large role for plagioclase during magma genesis under oxidation conditions approximating the NNO buffer should be accompanied by development of a Eu anomaly; the Paliza Canyon magmas are not so oxidized that Sr and Eu are strongly decoupled. Therefore, we conclude that the 87Sr/86Sr


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T able 3: EC-AFC and fractional crystallization model parameters
EC-ACF thermal parameters

Magma Assimilants Magma liquidus temperature (8C) Magma initial temperature (8C) Assimilant liquidus temperature (8C) Assimilant initial temperature (8C) Solidus temperature (8C) Equilibration temperature (8C) Specific heat of magma (J/kg per K) Specific heat of assimilant (J/kg per K) Crystallization enthalpy (J/kg) Fusion enthalpy (J/kg)

JM9384 18L-3, JM9373-2 1250 1250 900 800 800 1090 1593 1373 396000 281000

Partition coefficients plagioclase clinopyroxene orthopyroxene oxide Bulk D

North American continent is 1Á7^1Á8 Ga (Bennett & DePaolo, 1987), which presumably represents the time of extraction from the mantle and hence a maximum possible crystallization age. Consistent with this model age, basement rocks in the vicinity of the JMVF have crystallization ages of 1Á62^1Á44 Ga (Brookins & Laughlin, 1983). Attempts to calculate initial 87Sr/86Sr ratios for the granitoid lithic fragments using measured Rb/Sr and a minimum age of 1Á4 Ga yield unreasonably low values (5 0Á69), hence these rocks have been open to Rb and/or Sr since crystallization at 4 Ga. A variety of meta1Á4 somatic and metamorphic processes could be responsible, but it is noteworthy that simple bulk mixing (see below) of the lithic compositions with Type I mafic magma satisfies the compositions of the 67^68% SiO2 dacites (representing 70^80% of the crustal end-member; Fig. 7) for all trace elements except Rb (Fig. 5). We conclude that the Bandelier granitoid lithic fragments are, in fact, good representatives of the Precambrian crust that is involved in Tschicoma Formation magmatism and mafic magma genesis for most geochemical tracers except Rb and 87Sr/86Sr.

Sr Rb Ce Pb K Ba Nb Zr U Th Yb Mode

3Á4(1) 0Á3(1) 0Á12(1) 0Á18(5) 0Á0837(6) 0Á36(3) 0Á008(5) 0Á2(11) 0Á01(13) 0Á01(1) 0Á02(3) 0Á8

0Á5(1) 0Á03(1) 0Á2(3) 0Á005(4) 0Á00344(6) 0Á05(3) 0Á0081(9) 0Á29(1) 0Á03(12) 0Á03(13) 0Á9(14) 0Á11

0Á01(1) 0Á01(1) 0Á02(3) 0Á29(5) 0Á016(8) 0Á23(3) 0Á27(5) 0Á11(1) 0Á023(5) 0Á14(1) 0Á29(3) 0Á06

0Á11(2) 0Á15(2) 0Á12(4) 2Á9(2) 0Á045(7) 0Á028(7) 0Á7(10) 0Á51(11) 0Á16(11) 0Á13(11) 0Á24(4) 0Á03

2Á78 0Á25 0Á12 0Á25 0Á07 0Á31 0Á04 0Á21 0Á02 0Á02 0Á14

Southern JMVF: Paliza Canyon Formation lavas with 57^69% SiO2
No single petrogenetic model can explain the compositional diversity, for example the increasing diversity of Nb, Zr and REE concentrations with increasing SiO2, among lavas of the Paliza Canyon Formation (Figs 5 and 6). Given the longevity of the Paliza Canyon magmatic system(s), this is perhaps not surprising. In the following discussion, we consider the roles of fractional crystallization, incorporation of continental crust (including mixing with crustal melts), and mixing between mafic and silicic magmas.

Fractional crystallization
Average SiO2

Approximate SiO2 of mineral phases






Mode based on sample JM93190. Numbers in parentheses are partitioning references: (1) Bacon & Druitt (1988); (2) Ewart & Griffin (1994); (3) Luhr & Carmichael (1980); (4) Beattie (1993); (5) Dunn & Sen (1994); (6) Philpotts & Schnetzler (1970); (7) Sisson (1991); (8) Okamoto (1979); (9) Hauri et al. (1994); (10) Haskin et al. (1966); (11) Mahood & Hildreth (1983); (12) Villemant (1988); (13) Dostal et al. (1983); (14) Nicholls & Harris (1980).

values of the lithic fragments do not represent those of the assimilant(s) involved in petrogenesis. The lithic fragments themselves yield some evidence that their 87Sr/86Sr values have been modified since crystallization. The Nd model age for this portion of the

The increase of Zr and Nb concentrations with increasing SiO2 among some Paliza Canyon Formation lavas is consistent with fractional crystallization (Fig. 5; Table 3). A 50^55% crystallization of a plagioclase^clinopyroxene^ orthopyroxene^oxide assemblage (0Á8:0Á11:0Á06:0Á03) from a trachyandesite bulk composition (JM93190) reproduces the highest Nb and Zr concentrations measured in some Paliza Canyon trachydacites (Fig. 5d and e), which have isotopic compositions that lie within the range defined by the Type I mafic lavas. In the same group of lavas, Ba/Nb and K/Nb overlap with Type I mafic lavas and vary little with increasing SiO2 and Nb concentrations (Figs 7 and 9). Ba/Nb and K/Nb are expected to remain approximately constant during fractional crystallization that does not involve potassium feldspar or biotite, and hence are good discriminators between crustal assimilation and fractional crystallization in JMVF trachyandesites and trachydacites.






Fig. 7. Ba/Nb vs Nb concentrations for intermediate and silicic volcanics. Rhyolites are excluded because Ba is a strongly compatible element in systems crystallizing alkali feldspar. Dashed lines are mixing curves between Cerros del Rio benmoreite (E6-8B) and Otowi lithic fragment CCL-1, with crosses (þ) at 10% increments. Fractional crystallization model same as in Fig. 5. Also shown is the composition of Bandelier lithic fragment 18L-3.

Although zircon fractionation will strongly influence Zr concentrations, most andesitic and dacitic liquids are expected to be zircon undersaturated (Watson & Harrison, 1983; Miller et al., 2003). Zircon saturation temperatures calculated after Miller et al. (2003), for a range of Paliza Canyon and Tschicoma Formation lavas with 568% SiO2, vary from $7508C to 8508C, lower than magmatic temperatures calculated from Fe^Ti oxides ($850^10508C). It is possible that zircon saturation could have been achieved in some high-Zr, high-SiO2 (68^71wt%) Paliza Canyon compositions. However, Nb and Zr show very similar behavior among lavas with 572% SiO2 (Fig. 5), and there is no Nb-phase analogous to zircon that could precipitate from these liquids and cause a decrease in Nb concentrations. The two elements are decoupled in high-silica rhyolites of all ages, with Zr showing relative depletion (Fig. 5), as expected when zircon joins the fractionating assemblage. We conclude that zircon fractionation did not occur during petrogenesis of the lavas with 57^68% SiO2. Therefore, the approximate coincidence of the fractional crystallization curve with the boundary of the data distribution for Zr, Nb and several other trace elements (Fig. 5) strongly suggests that fractional crystallization represents a limiting case for the Paliza Canyon intermediate rocks.

Type I Paliza Canyon mafic lavas have Pb/Ce, which in JMVF mafic lavas is a sensitive indicator of open-system processes, up to $0Á15 (Wolff et al., 2005). It is significant, therefore, that Pb/Ce is elevated above this value among some of the dacites and trachydacites that we have identified as dominantly the products of fractional crystallization (Fig. 9b). Among intermediate magmas, preferential removal of Ce during fractionation may occur as a result of separation of phases for which DCe 4DPb, such as hornblende and apatite. Pb/Ce may be increased by up to 50% because of fractionation of a hornblende- and apatitebearing assemblage over the interval 57^70% SiO2, and values of Pb/Ce up to 0Á25 are consistent with fractionational crystallization.

Open-system processes
Trace element abundances in the bulk of the Paliza Canyon intermediate lavas deviate from the predictions of closed-system fractional crystallization. In particular, Zr and Nb and other incompatible elements either remain approximately constant or show a sharp decrease in concentration with increasing SiO2 (Fig. 5), whereas K/Nb, Ba/Nb, and Pb/Ce all increase (Figs 7 and 9). Lavas with the lowest Nb and highest K/Nb, Ba/Nb and Pb/Ce also have low 206Pb/204Pb and can be modeled as


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Fig. 9. (a) K/Nb and (b) Pb/Ce vs SiO2 for intermediate and silicic volcanics.‘Mafic lavas’ field as in Fig. 6. Short dashed lines are simple mixing curves between Cerros del Rio benmoreite (E6-8B) and Otowi lithic fragments CCL-1 and 18L-3, with crosses (þ) at 10% increments. Long-dashed lines are mixing curves between Cerros del Rio basalt (H I-5) and Bearhead and Tschicoma Formation rhyolites (Table 2). Fractional crystallization model same as in Fig. 5.

Fig. 8. (a) 206Pb/204Pb vs 208Pb/204Pb, (b) 206Pb/204Pb vs 143Nd/144Nd, and (c) 208Pb/204Pb vs 143Nd/144Nd in JMVF volcanics. Thin-dashed lines are simple mixing curves between Cerros del Rio basalt (H I-5) and Otowi lithic fragment 18L-3, Bearhead rhyolite JM9373-2, and Tschicoma Formation rhyolite MR00-1. Continuous lines are simple mixing curves between Cerros del Rio benmoreite (E6-19a) and Otowi lithic fragments CCL-1 and 18L-3 and a Bearhead rhyolite, with crosses (þ) at 10% increments. Short^longdashed lines are EC-AFC curves between Paliza Canyon trachyandesite and 18L-3 and JM9373-2. Modeling parameters and end-member compositions are given in Tables 2 and 3. Bold box represents the isotopic range of the Bandelier Tuff (San Diego Canyon isotopic data presented separately). Taos Range amphibolites, which may represent some of the crust present beneath the JMVF (Wolff et al., 2005), are shown for comparison.

simple mixtures of mafic magma with low-206Pb/204Pb, low-Nb crust as represented by the Otowi lithic fragments (Figs 5^9). Their petrogenesis is thus similar to that of Tschicoma Formation andesites, trachyandesites and dacites (Fig. 7; see next section). A significant subset of lavas exhibit increasing Pb/Ce with decreasing Nb content whereas 206Pb/204Pb remains within the range of the Type I basalts, and Nb and SiO2 show little correlation. The latter observation suggests a petrogenesis involving both magma^crust interaction

and fractional crystallization, such as assimilationfractional crystallization (AFC). EC-AFC simulations (Figs 8 and 10) are successful in modeling some of these compositions, but no single composition in the suite of Otowi xenoliths has the combination of high Pb/Ce, appropriate Pb isotope ratios, and low Ba/Nb to be a suitable contaminant for all cases. Of course, it is possible that an appropriate crustal composition for generating the Paliza Canyon intermediate rocks is not represented in our sample suite, especially given the highly heterogeneous nature of the Proterozoic basement (Magnani et al., 2004). Further possibilities are interaction between mafic magma and isotopically similar crust, perhaps hybridized crust produced earlier in the history of the volcanic field, or mixing between mafic and rhyolitic magma. Within the limitations of our dataset, the high-206Pb/204Pb, highPb/Ce subset of Paliza Canyon andesites and dacites may be most closely modeled through fractional crystallization or as mixtures between mafic lavas and broadly contemporaneous rhyolites, particularly Bearhead Formation and early Tschicoma Formation rhyolites (Figs 6 and 8^10), as well as Otowi granitoid lithic fragments. T extural






(at least 6 Myr; Fig. 2) the diversity is perhaps not surprising, although most of the geochemical variation we observe was developed in a much shorter period. The detailed geochronology of the Paliza Canyon Formation is imperfectly known, although many dates cluster at 8Á5 Æ1Ma (Goff et al., 1990; Justet, 2003). Most of our samples for which an eruptive age can be established either directly or by stratigraphic bracketing fall in this range, hence it appears that several magma chambers, or a large complex system allowing different crustal lithologies to be melted, were established in the crust beneath the volcanic field during this 2 Myr period.
Fig. 10. Pb/Ce vs 206Pb/204Pb. Dashed lines are mixing curves between Paliza Canyon trachyandesite JM9384 and Otowi lithic fragments CCL-1 and 18L-3 and Bearhead rhyolite JM9373-2. Continuous lines are EC-AFC model curves between JM9384 and 18L-3 and JM9373-2. Modeling parameters and end-member compositions are given in Tables 2 and 3.

North to east JMVF: Tschicoma and Puye Formations, and Cerros del Rio trachyandesites
Mafic enclaves in Tschicoma Formation dacites have the same geochemical traits as Cerros del Rio lavas in the range 53^60% SiO2, with the characteristic low K/Nb and high Th/Nb and La/Nb of Type I basalts (Wolff et al., 2005). The most silicic Cerros del Rio lavas likewise resemble Tschicoma lavas of the same silica content (63% SiO2). The contemporaneity of Cerros del Rio and later Tschicoma Formation activity (Fig. 2), the presence of mafic enclaves with trace element concentrations and ratios identical to Cerros del Rio benmoreites, and the compositional overlap between the Tschicoma Formation and Cerros del Rio lavas (Figs 3^6) strongly indicates that the Tschicoma Formation and Cerros del Rio lavas are genetically related. Tschicoma Formation andesites and dacites and Cerros del Rio benmoreites exhibit variations in incompatible trace elements (especially high field strength elements; HFSE) and radiogenic isotopes that are relatively well correlated with SiO2 (Figs 4^6). Systematic variation of isotopic ratios (especially Nd and Pb; Fig. 6) with increasing SiO2 and near-ubiquitous disequilibrium textures (see descriptions above) indicate a significant role for mixing and crustal assimilation in the generation of Tschicoma Formation lavas. In addition, the lack of divergence in trace element or isotopic compositions with increasing SiO2, in contrast to Paliza Canyon Formation lavas, indicates a lack of diversity in petrogenetic mechanisms and among the crustal end-member(s) involved. 207 Pb/204Pb vs 206Pb/204Pb of Tschicoma Formation lavas, as well as Otowi basement xenoliths, define a linear covariation that intercepts the Stacey^Kramers terrestrial Pb growth curve at 1Á8 Ga (Fig. 11), which is the Nd model age for this part of the North American craton (Bennett & DePaolo, 1987). This observation is consistent with the model that less evolved Tschicoma lavas are assimilating Proterozoic granitic crust, compositionally similar to the Otowi granitoid xenoliths.

evidence for disequilibrium among phenocrysts and groundmass, as described above, also supports magma mixing as a significant process in Paliza Canyon Formation lavas. It should be noted that Pb/Ce may be very high in rhyolitic magma as a result of fractionation of LREE-rich phases such as chevkinite and allanite. In contrast, the combination of low 206Pb/204Pb with low Pb/Ce (Fig. 10) in a few Paliza Canyon (and Tschicoma Formation) lavas may simply be the result of changing the Pb/Ce ratio during melting of granitic basement. Melting of all the Ce-bearing phases with Pb still retained in residual feldspars or simply having an assimilant with a higher proportion of allanite would be two possible scenarios. The petrological and geochemical complexity of the Paliza Canyon lavas calls for more detailed study to answer the questions raised above, but it is clear that there is a major role for interaction of mafic magma (itself carrying a significant crustal component; Wolff et al., 2005) with Proterozoic crust. At least three distinct crustal reservoirs must be involved to account for the respective isotopic uniqueness of the Bearhead, Canovas Canyon, and Tschicoma Formation^El Rechuelos rhyolites (Fig. 8; see also ‘Rhyolites’ section below) and the granitoid component, invoked by Wolff et al. (2005) as a contaminant of the basalts. We may recall that Bearhead and Canovas Canyon Formation rhyolites are erupted either contemporaneously with, or immediately following, the eruption of the Paliza Canyon Formation lavas, and that early El Rechuelos and Tschicoma Formation rhyolites temporally overlapped with the Paliza Canyon Formation (Loeffler et al., 1988). This requires consumption of at least four different types of crust beneath the volcanic field, of the order of hundreds of cubic kilometers, during magma genesis. Given the longevity of the Paliza Canyon Formation


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Fig. 11. 206Pb/204Pb vs 207Pb/204Pb for Tschicoma Formation, El Rechuelos Formation, and Cerros del Rio lavas, and Otowi lithic fragments. A regression through the isotopic compositions of Tschicoma Formation lavas intersects the terrestrial Pb growth curve of Stacey & Kramers (1975) at 1Á8 Ga.

Rio benmoreites have already incorporated up to 25% crust to generate the characteristic trace element patterns (Wolff et al., 2005) such that higher-silica dacites are dominantly crustal melts. That the dacites are dominated by the products of near-complete crustal melting represents renewed consumption of crust by magma at 54 Ma following the hiatus since the Paliza Canyon Formation volcanic maximum at about 8Á5 Ma. Taking into account the $25% crustal contribution to the Cerros del Rio benmoreites and assuming a conservative average of 50% crustal component and 50% benmoreite for the bulk of the Tshicoma Formation dacites, over $300 km3 of granitoid crust must be involved in the generation of the Tschicoma Formation lavas.

NW JMVF: Tschicoma Formation, La Grulla Plateau
The La Grulla Plateau lavas temporally overlap late Paliza Canyon Formation activity (Singer & Kudo, 1986) and share geochemical characteristics with both the north and east Tschicoma Formation and Paliza Canyon lavas. Singer & Kudo (1986) noted a positive correlation between 87 Sr/86Sr and SiO2 among La Grulla lavas (Fig. 6c), and also that La Grulla plateau lavas have higher 87 Sr/86Sr at a given SiO2 content relative to Tschicoma Formation lavas to the east of Mesa del Media and the Canada de Cochiti fault zone (Fig. 1). Our limited La ‹ Grulla isotopic data are consistent with this pattern and 143 Nd/144Nd is negatively correlated with 87Sr/86Sr and silica as expected. La Grulla lavas resemble the Tschicoma Formation in 143Nd/144Nd and Nb content, but Pb isotope ratios resemble the high-206Pb/204Pb subset of Paliza Canyon Formation lavas (Fig. 6), and one La Grulla plateau rhyodacite has some Nd isotopic affinities with Bearhead rhyolite (Figs 6 and 8). Singer & Kudo (1986) modeled 87Sr/86Sr variations in the La Grulla lavas by AFC using a Lobato basalt as the parent and an upper crustal assimilant with highly radiogenic 87Sr/86Sr (0Á74). Nothing in our data contradicts this model, although mixing between intermediate magma and silicic liquid resembling Bearhead Rhyolite (albeit with higher 206Pb/204Pb) is equally plausible. Clearly, more work is required on the La Grulla lavas. Nonetheless, the salient feature emerges that the consumed crust is probably distinct from both that in the main Tschicoma Formation east of the Canada de Cochiti fault zone and that in ‹ Paliza Canyon Formation lavas. A striking feature of some La Grulla Plateau lavas is their enrichment in Ba, which reaches concentrations up to 4000 ppm in lavas from Cerro Pavo (Fig. 1) yielding extremely high Ba/Nb ratios (4 160; Fig. 7). The same lavas are LREE-enriched with La up to 140 ppm (Table 1), roughly twice that of comparable Tschicoma Formation dacites, and higher La/Sm (13Á6^14Á5) than the rest of the La Grulla Plateau suite (7Á7^9Á9). Distinctive LREE and

Among the dacites, there appear to be two slightly different end-member compositions at 68% SiO2. The first has Nb %11ppm, 143Nd/144Nd % 0Á51235, and 206 Pb/204Pb %17Á1 (‘eastern dacite’, exemplified by Pajarito Mt. and Cerro Rubio; Figs 1 and 5). This composition is also seen in the Puye ignimbrite of Turbeville et al. (1989) and in conglomerate cobbles from the Puye fan above the level of the ignimbrite, and hence may be volumetrically more significant than is suggested by the present size of the dome remnants in the east (Fig. 1). The second dacite composition has Nb % 20 ppm, 143Nd/144Nd % 0Á51250, and 206Pb/204Pb %17Á5 (‘northern dacite’, exemplified by Tschicoma Mt., the Mesa Gallina flow, and Polvadera Peak; Figs 1 and 5). Simple mixing relations (Figs 5 and 6) suggest that the northern dacite composition is the more important end-member for most of the intermediate to silicic Tschicoma and Cerros del Rio magmas. The two distinct dacite types cannot themselves be related to each other by simple mixing with mafic magma and may instead represent melts derived from slightly different crustal assimilants (Figs 5d and 6). The eastern dacite type may require a crustal end-member slightly less radiogenic in Pb (206Pb/204Pb $ 17Á0) than is present among the Otowi basement xenoliths and may indicate a crustal component not present in the current suite of basement samples. Simple bulk-mixing of a Cerros del Rio benmoreite E6-8B (30^60%) and Otowi lithic fragments CCL-1 and 18L-3 (Table 2) can reproduce much of the trace element and isotopic variability of Tschicoma dacites (Figs 5^9). Eastern dacites, with low Nb and high Ba/Nb, culminating with Cerro Rubio lavas, require up to 60^70% crust mixed with Cerros del Rio benmoreite (Fig. 7). These high crustal proportions are supported by the Pb isotope data; Cerro Rubio Pb isotopic ratios are nearly identical to those of the inferred basement (Fig. 6). The evolved Cerros del






Ba enrichments in lavas from a single small center may point to the involvement of heterogeneous crust in the genesis of the La Grulla suite; Singer & Kudo (1986) also noted the ‘outlier’ character of this chemical type.

JMVF rhyolites record extensive separation of feldspar in their very low Sr, Ba and Eu/Euà (Fig. 4). The depletion in Ba requires that at least some of the fractionating assemblage is alkali feldspar. Despite such clear indications, high-silica rhyolites (4 75% SiO2) present a special problem for geochemically based petrogenetic interpretation. The high probability of multiple saturation with small amounts of accessory phases (e.g. allanite, chevkinite, monazite, zircon) complicates understanding the inheritance of large ion lithophile elements (LILE), HFSE and REE from a parental magma. For example, Zr is highly variable at near-constant silica content in JMVF rhyolites, probably as a result, in part, of fractionation of variable amounts of zircon consequent upon a range of zircon solubilities and the actual physical separation of the phase. A further problem is that the very low Sr concentrations put magmatic 87Sr/86Sr values at the mercy of contamination by very small quantities of wallrock that are otherwise difficult to detect (Wolff et al., 1999). In the case of older rhyolites, very high Rb/Sr ratios, themselves subject to change during weathering, lead to large errors in age corrections. Nd and Pb isotopes are more reliable guides to the sources of rhyolitic liquids because, relative to Sr, high concentrations of Nd and Pb render these systems insensitive to minor amounts of late-stage contamination (Wolff & Ramos, 2003). Uncertainty about effective partition coefficients, especially for Nd, plus the numerous possible process-dependent responses of magma to contamination (e.g. Bohrson & Spera, 2003), make it difficult to identify precise petrogenetic pathways. Some indications can nonetheless be gained from the positions of rocks and potential source materials on isotope^isotope plots. Nd and Pb isotope ratios of JMVF rhyolites, including the Bandelier Tuff, show no systematic variation with the ages of their formations. However, individual formations do exhibit restricted isotopic ranges (Fig. 6). Canovas Canyon high-silica rhyolites and one Paliza Canyon rhyodacite have unusually low 208Pb/204Pb (Fig. 8a) and cannot be simple fractionates of the Paliza Canyon andesites, trachyandesites and dacites with which they are interbedded. Several rocks, distinct from the Otowi lithic fragments, which may represent sub-JMVF basement (Wolff et al., 2005) also have low 208Pb/204Pb, especially Taos Range amphibolites (Fig. 8). If these represent one component in the Canovas Canyon rhyolites, Pb^Pb relations indicate that a second component lies among the low-206Pb/204Pb group of Paliza Canyon Formation mafic to intermediate magmas (Fig. 8a). Otowi lithic fragments,

with low 206Pb/204Pb ratios, are ruled out by 208 Pb/204Pb^143Nd/144Nd relations (Fig. 8c). Alternatively, the Canovas Canyon rhyolites may be derived by partial melting of a composition not represented among the basement suite. Bearhead Formation rhyolites and the La Grulla plateau rhyodacite form another group that may contain an additional crustal component, with relatively low 143Nd/144Nd at high 206Pb/204Pb and 208Pb/204Pb. This group is distinct from the broad trend defined by non-rhyolitic Paliza Canyon and Tschicoma Formation rocks, and the Otowi amphibolite xenolith CCL-2 is a candidate for the crustal component (Fig. 8b and c). For both Canovas Canyon and Bearhead rhyolites, the exact petrogenic processes by which high-silica rhyolite is derived from a mixture of basement partial melt and mafic or intermediate magma are not well constrained; however, it may be significant that each is accompanied by a rhyodacite with broadly similar isotopic composition that may represent a parent for the respective high-silica rhyolites. The Pb and Nd isotope compositions of the San Diego Canyon and Bandelier ignimbrites coincide with the middle of the ‘JMVF isotopic array’, although they more closely resemble the Paliza Canyon Formation lavas rather than the immediately preceding Tschicoma Formation (Figs 6 and 8). The caldera-forming Bandelier magmas therefore overlap with JMVF mafic lavas in Pb^Nd isotope space. Nonetheless, the suggestion that this is due to fractionation from basaltic magma (Perry et al., 1993) seems hard to sustain in light of the fact that no isotopically similar intermediate magmas, with the exception of Cerros del Rio benmoreites located 30 km away in the rift axis, were erupted from the JMVF during the previous 5 Myr. Thus, there appears to be no systematic chemical pattern among JMVF rhyolites as a whole. This, and the possible status of the Canovas Canyon and Bearhead Formation rhyolites as the crustal component in some of the intermediate magmas, is more consistent with a model of heterogeneous rhyolite production under the influence of mafic intrusion than it is with an origin by fractional crystallization. In the case of the Bandelier Tuff and associated rhyolites, Nd and Pb isotope data indicate that the likely protolith consists wholly of buried intrusions associated with Paliza Canyon volcanism. Because such intrusions crystallize from crustally contaminated, mantle-derived magma, they constitute hybridized crust (sensu Riciputi & Johnson, 1990; Riciputi et al., 1995). Based on mixing models and erupted volumes of the Paliza Canyon (1000 km3) and Tschicoma Formations ($500 km3), making the conservative assumption of no cognate cumulates and that the volume deficit created by eruptions was compensated by magma, the minimum volume of hybridized crust is of the order of 1500 km3, and is likely to be


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much greater. Furthermore, such young intrusions were probaby still warm, reducing the thermal penalty associated with their rejuvenation. The much greater volume of Bandelier Tuff than either Canovas Canyon or Bearhead rhyolite, both of which involved melting of previously untapped volumes of Precambrian crust, may thus be explained.

remelting young, warm intrusions may in part explain the larger volume of the caldera-related rhyolites.

Pre-caldera volcanic formations of the JMVF have distinctive isotopic compositions and compositional variations with each new major volcanic phase, consistent with consumption of new volumes of crust. This belies the apparent overall simple evolution of the three successive major formations: the andesite^trachyandesite-dominated Paliza Canyon Formation (1000 km3, 13^7 Ma), through the dacite-dominated Tschicoma Formation (500 km3, 7^2Á2 Ma), to the high-silica rhyolite caldera-related San Diego Canyon ignimbrites and Bandelier Tuff (4 600 km3, 52 Ma). The caldera rhyolites do not mark a new isotopic excursion, but are more similar to the Paliza Canyon Formation (4 Myr older) than to the 5 Tschicoma Formation (mostly 53 Myr older), and are probably derived wholly through partial remelting of intrusive rocks, consisting of crust^mantle hybrids, related to the Paliza Canyon Formation. The temporal geochemical record does not support any simple liquid-line-ofdescent model, with or without minor accompanying contamination, for the ultimate generation of large volumes of rhyolite. Throughout the history of the JMVF, an important petrogenetic mechanism for the production of magmas in the range 57^70% SiO2 appears to have been mixing between mafic liquids and crustal melts, particularly in the case of the Tschicoma Formation. Heating of crust above its solidus by mafic magma implies partial crystallization of the latter, and there is additionally a significant role for fractional crystallization in Paliza Canyon petrogenesis. It is very likely, therefore, that significant volumes of cumulates cognate with the Paliza Canyon Formation, and perhaps also the Tschicoma, exist at depth. A lack of evidence for caldera formation or subsidence on the volume scale of the two formations suggests volume compensation of erupted magma by new liquid. Complex plutons resulting from the assembly of cumulates, nonerupted, and recharged magma now make up much of the crust beneath the JMVF, which consists isotopically of a crust^mantle mixture. Melting of significant volumes of pristine Precambrian crust effectively ceased between 2 and 3 Ma, after the generation of the late Tschicoma dacites. The tempo of JMVF magmatism is controlled by Rio Grande rift tectonic activity (Gardner et al., 1986), which is fully consistent with a lithospheric origin for the mafic magmas (Wolff et al., 2005). Significant rift-related extension at 10^7 Ma (Gardner et al., 1986; Self et al., 1986) coincides with the Paliza Canyon volcanic maximum at

Summary of petrogenesis
Geochemical variations among pre-caldera JMVF rocks permit the following conclusions. (1) Mafic lavas produced throughout the history of the volcanic field carry a crustal component that has the overall composition of Precambrian granitoid rocks excavated by the Bandelier Tuff eruptions (Wolff et al., 2005). (2) Granitoids at depth may have lower Rb contents and 87 Sr/86Sr than extant lithic fragments which, at least for the JMVF, solves the Sr^(Nd, Pb) paradox. (3) The same granitoids experienced near-complete melting to produce the large Tschicoma dacite domes of the north and east JMVF, which carry only a minor component of mantle-derived basalt. Magma of the same composition mixed with Type I basalt to produce the intermediate lavas and tuffs of the main Tschicoma Formation and Cerros del Rio, and some intermediate lavas of the Paliza Canyon Formation. (4) Many Paliza Canyon, Canovas Canyon, and Bearhead Formation samples and La Grulla Plateau volcanic rocks in the range 63^73% SiO2 carry crustal component(s) that are isotopically distinct from the Bandelier lithic fragments; at least three additional crustal components are required. Petrogenetic processes were similarly diverse and involved fractional crystallization and AFC in addition to simple mixing of mafic magmas with crustal melts. These lavas therefore record earlier melting events in the crust that peaked at $8Á5 Ma. Mantle and crustal components were apparently blended in a complex magmatic system with more than one storage chamber. (5) JMVF rhyolites have isotopic compositions that are distinct from contemporaneous or immediately preceding intermediate volcanic rocks, and there is no direct line-of-descent (by FC or AFC) of rhyolites from more primitive compositions. Instead, rhyolites appear to be generated by melting of both preexisting and hybridized crust. The Bandelier Tuff and associated caldera-related rhyolites originated through rejuvenation of intrusions associated with Paliza Canyon Formation magmatism, whereas older JMVF rhyolites are isotopically distinct and in part represent melting of ‘pristine’ Proterozoic crust. The reduced thermal penalty associated with






8Á5 Æ1 Ma. It has been previously stated that Tschicoma dacitic volcanism was associated with a tectonic lull from 7 to 4 Ma (Gardner & Goff, 1984; Gardner et al., 1986; Self et al., 1986), but the major dacite domes such as Tschicoma Mountain and Polvadera Peak, representing significant crustal melting, erupted after the onset of renewed extension at about 4 Ma. The dacite domes contain mafic enclaves, and this period of activity continued after 2Á7 Ma in the adjacent rift-floor Cerros del Rio field until 2Á3 Ma (Woldegabriel et al., 1996), with mafic and intermediate lavas that chemically and isotopically overlap the Tschicoma dacites and their enclaves. Hence it is clear that the major Tschicoma dacites are ‘fundamentally basaltic’ and represent advective heat transfer from mantle to crust. Cerros del Rio vents migrated westward between 2Á7 and 2Á3 Ma (Woldegabriel et al., 1996; Dethier, 1997), and we therefore speculate that repeated intrusion of basalt beneath the central JMVF after 2Á3 Ma ultimately led to the formation of the Bandelier magmatic system, which first vented at 1Á85 Ma. A continuing role for mafic magma in powering the JMVF volcanic system is hinted at by an andesite lava between the two Bandelier tuffs (Smith et al., 1970) and a basaltic andesite component among the products of the most recent rhyolitic eruption from the caldera (Wolff & Gardner, 1995). Major transitions in JMVF activity therefore occur on timescales of 105^106 years, but petrological patterns and styles of activity, once established, seem to persist for a few million years. These patterns may then relate to sites of melt lodgement and maximum heat transport into the crust that act as traps for later rising magma. Major transitions in style may correspond to times when the system is starved of mantle input, hence the crust becomes sufficiently rigid to permit, when activity is renewed, mantle-derived magma to rise and lodge at a new location in the still-warm crustal column. A similar result may occur where a lateral shift in the focus of intrusion occurs; this may be the case for the La Grulla Plateau suite, which is tectonically, geographically and chemically distinct from the rest of the Tschicoma and Paliza Canyon Formations. If this model is generally valid, it then becomes difficult to predict at what stage in its development an individual continental intermediate volcanic field may enter a major rhyolitic, caldera-forming phase.

Winters, Pam Hartman, Keith Brunstad, and Geoff Cook, for discussion, data, and/or assistance in field and laboratory. Diane Johnson Cornelius, Rick Conrey and Charles Knaack of the WSU Geoanalytical Laboratory generated many of the data presented here. This manuscript has been substantially improved by reviews from Wendy Bohrson, Nancy McMillan, Calvin Barnes and Thomas Vogel. None of the aforementioned individuals carry any responsibility for the conclusions presented herein. Research in the JMVF has been supported by Associated Western Universities, Inc., the US Department of Energy, and the National Science Foundation, most recently under EAR-9909700 to J.A.W.

Supplementary data for this paper are available at Journal of Petrology online.

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We wish to thank the many who have contributed to this paper in tangible and intangible ways over several years, including Rob Creaser, Steve Self, Fraser Goff, Philip Kyle, T erry Spell, Scott Baldridge, Jon Davidson, Phil Leat, Bob Thompson, Mike Dungan, Stephen Moorbath, Anita Grunder, Jack Flannery, Katherine Romanak ¤ (nee Duncker), Bruce Turbeville, Damon Waresback, Steve Balsley, Dave Kuentz, Marty Horn, Wade Aubin, Lee


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