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Journal of Petrology Volume 42 Number 8 Pages 1491-1517 2001
© Oxford University Press 2001
Low-
18O Rhyolites from Yellowstone: Magmatic Evolution Based on Analyses of Zircons and Individual Phenocrysts
DEPARTMENT OF GEOLOGY AND GEOPHYSICS, UNIVERSITY OF WISCONSIN, 1215 W. DAYTON STREET, MADISON, WI 53706, USA
Received May 24, 2000; Revised typescript accepted February 8, 2001
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONSAMPLE COLLECTION, PREPARATION...GEOLOGY AND PETROLOGY OF...OXYGEN ISOTOPE RESULTSDISCUSSIONAPPENDIX: ZIRCON IN YELLOWSTONE...REFERENCES |
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The Yellowstone Plateau volcanic field is one of the largest centers of rhyolitic magmatism on Earth. Major caldera-forming eruptions are followed by unusual low-
18O rhyolites. New oxygen isotope, petrologic and geochemical data from rhyolites belonging to the 2·0 my eruptive history of Yellowstone are presented, with emphasis on the genesis of low-18O magmas erupted after the Huckleberry Ridge Tuff (2·0 Ma, 2500 km3) and Lava Creek Tuff (0·6 Ma,1000 km3). Analyses of individual quartz and sanidine phenocrysts, obsidian samples and bulk zircons from low-18O lavas reveal: (1) oxygen isotope variation of 1–2
between individual quartz phenocrysts; (2) correlation of zircon crystal size and 18O; (3) extreme (up to 5) zoning within single zircons; zircon cores have higher 18O; (4) 
18O disequilibria between quartz, zircon and homogeneous unaltered host glass where zircon cores and some quartz phenocrysts have higher 18O values. These features are present only in low-18O intra-caldera lavas that erupted shortly after caldera-forming eruptions. We propose that older, hydrothermally altered, 18O-depleted (18O 
0), but otherwise chemically similar, rhyolites in the down-dropped block were brought nearer the hot interior of the magma chamber. These rhyolites were remelted, promoting formation of almost totally molten pockets of low-18O melt that erupted in different parts of the caldera as separate low-18O lava flows. Alteration-resistant quartz and zircon in the roof rock survived early hydrothermal alteration and later melting to become normal 18O xenocrysts (retaining their pre-caldera 18O values) in the low-18O magma that formed by melting of hydrothermally 18O-depleted volcanic groundmass and feldspars. Zircon and quartz xenocrysts exchanged oxygen with newly formed melt through diffusion and overgrowth mechanisms leading to partial or complete isotopic re-equilibration. Modeling of the diffusive exchange of zircon and quartz during residence in low-18O magma explains 18O and (Qz–Zrc) disequilibria. The exchange time to form zoned zircons is between a few hundred and a few thousand years, which reflects the residence time of low-18O magmas after formation and before eruption. KEY WORDS: yellowstone; caldera; isotope disequilibria; zircon; 18O–16O
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONSAMPLE COLLECTION, PREPARATION...GEOLOGY AND PETROLOGY OF...OXYGEN ISOTOPE RESULTSDISCUSSIONAPPENDIX: ZIRCON IN YELLOWSTONE...REFERENCES |
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Low-
18O magmas and their abundanceMeteoric water is involved in sub-solidus exchange and in magma genesis of many shallow plutonic bodies and volcanic caldera complexes (Friedman et al., 1974
; Muehlenbachs et al., 1974; Hildreth et al., 1984; Criss & Taylor, 1986; Taylor, 1986; Gilliam & Valley, 1997). The more abundant hydrothermally altered rocks are characterized by irregular and large 18O fractionations between phenocrysts and glass. Low-18O magma are recognized by smaller fractionations, consistent with last equilibration at magmatic temperatures, suggesting their original crystallization from 18O-depleted magmas. Low-18O magmas with 18O(WR) <5·5 (where WR indicates whole rock) are depleted in 18O below the values attainable by normal mantle-derived magmas. Their genesis requires that tens of per cent of total oxygen must ultimately be derived from, or exchanged with, surface waters.
The origin of low-18O magmas has been proposed to require either direct, sometimes ‘catastrophic’ addition of meteoric water to magma (Friedman et al., 1974; Hildreth et al., 1984), or more gradual, multistage assimilation of hydrothermally altered rocks (Taylor, 1986; Hildreth et al., 1991), especially by hot, water-undersaturated magmas with large assimilation potential (Balsley & Gregory, 1998). Bacon et al. (1989) demonstrated that 18O depletion of normal-18O magma may proceed through bulk assimilation and be facilitated by partial melting of pre-existing 18O-depleted country rocks surrounding the magma chamber.
Permeable hydrogeological conditions in shallow extensional environments such as rifts, or calderas, coupled with proximity to magma, favor formation of low-18O magmas (Larson & Taylor, 1986). Calderas are important loci of large-scale fluid–rock isotopic exchange. The best and most extreme example of low-18O rhyolites occurs in the Yellowstone Plateau volcanic field, Wyoming, where dramatic (>5) 18O depletions occur in magmas erupted after caldera-forming eruptions. Such extreme depletions are not observed anywhere else in the world. Smaller decreases of 1–2 in 18O are more typical for magmas in other caldera complexes and associated plutonic bodies (Friedman et al., 1974; Lipman & Friedman, 1975; Gilliam & Valley, 1997; Balsley & Gregory, 1998; Bindeman et al., 2001b; Monani & Valley, 2001). Although the Yellowstone-scale 18O depletions are yet to be found in other calderas, we believe that the presence of surface-derived water and hence surface-derived oxygen in magmas may be a general feature of calderas world-wide. Small (1–2) depletions would be difficult to recognize because many silicic magmas are originally enriched above 18O(WR) = 5·5. Identification of such rocks awaits comparative studies of 18O composition in pre-caldera, post-caldera and extra-caldera rocks.
Yellowstone Plateau volcanic field
The Yellowstone Plateau volcanic field is one of the largest centers of silicic magmatism on Earth (Fig. 1). Voluminous ignimbrite eruptions, each followed by caldera collapse, occurred at 2·053 ± 0·006, 1·304 ± 0·011 and 0·640 ± 0·002 Ma, producing the Huckleberry Ridge Tuff (HRT, 2500 km3), Mesa Falls Tuff (MFT, 300 km3) and Lava Creek Tuff (LCT, 1000 km3), respectively (Christiansen, 1982, 1989; quoted 40Ar/39Ar ages are from M. A. Lanphere et al., personal communication, 2001). Voluminous and variably 18O-depleted rhyolites, predominantly lava and domes, erupted between tuff eruptions; the volume of intra-caldera lavas following the eruption of LCT alone is at least 900 km3 (Christiansen, 1989). Basalts represent <5 vol. % of all volcanic rocks, are mostly extra-caldera, and have normal, mantle-like 18O(WR) of 6 (Hildreth et al., 1991). Therefore, the Yellowstone system provides a unique opportunity to study the genesis of low-18O magma.
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Although the last eruption at Yellowstone occurred 70 ky ago (Obradovich, 1992), vigorous hydrothermal activity still occurs inside and outside of the 0·6 Ma Yellowstone caldera. Drilling has revealed the existence of modern hydrothermal circulation and steep temperature gradients (70–200°C/km), in all parts of the caldera, with most of the heat flux resulting from hydrothermal flow (Fournier, 1989). Meteoric waters in equilibrium with secondary minerals in drill holes and hot springs are shifted upwards by up to 10 in 18O relative to the average present-day value on the surface of -17 (Sturchio et al., 1990). Hydrothermal waters are also enriched in many trace elements leached from rhyolite (Sturchio et al., 1986; Lewis et al., 1997).
Friedman et al. (1974) and Hildreth et al. (1984) recognized unaltered Yellowstone rhyolites with unusually low 18O, and suggested that the low-18O values resulted from the crystallization of low-18O magmas in which tens of per cent of total oxygen had exchanged directly or indirectly with meteoric water. The most depleted lavas show dramatic 3 and 5 drops in 18O, and erupted after the HRT and LCT caldera-forming eruptions (Hildreth et al., 1984). These magmas are referred to as ‘low-18O’, and are the main focus of this study.
Objectives of the present study: a high-resolution look at low-18O magmas
In this work we report oxygen isotope analyses of over 380 individual phenocrysts and bulk zircon separates from 25 lavas spanning Yellowstone’s 2 my eruptive sequence.
Recent advances in the microanalysis of stable isotope ratios (see Rumble & Sharp, 1998; Valley et al., 1998a; and therein) have made it possible to analyze mineral grains individually, including refractory phases such as zircon (e.g. Valley et al., 1994). Earlier conventional techniques required multiple grains of a mineral, thus averaging any intercrystalline or intracrystalline 18O differences. Laser fluorination and ion microprobes provide a new high-resolution look at the classic problem of the genesis of low-18O rhyolites at Yellowstone, and allow us to reinterpret many aspects of low-18O magma petrogenesis (Bindeman & Valley, 2000).
Zircon 18O analyses proved to be crucial for reinterpreting the petrogenesis of Yellowstone lavas (Bindeman & Valley, 2000). The ubiquitous presence of zircon as an early crystallizing, near-liquidus mineral in all magmas at Yellowstone makes its isotopic composition particularly diagnostic in inferring the near-liquidus relations in Yellowstone high-silica rhyolites. Zircon is also characterized by slow oxygen diffusion (Valley et al., 1994; Watson & Cherniak, 1997; Peck et al., 1999), sluggish growth relative to common minerals (Watson, 1996), and extremely sluggish cation diffusion (Cherniak et al., 1997a, 1997b). In addition, non-metamict zircon is particularly resistant to secondary hydrothermal alteration of 18O (Gilliam & Valley, 1997) and well known to survive melting episodes (Watson, 1996). Therefore, we use zircon as a ‘recorder’ of transient changes in 18O of a melt and the source rock, as well as age. The 18O zoning patterns in zircon (and quartz) when combined with knowledge of exchange rates, provide information about the rate and duration of low-18O magma genesis. We also present results of a comparative petrologic and mineralogic study of Yellowstone low-18O and normal-18O rhyolites which include: whole-rock chemistry; liquidus and zircon saturation temperature estimates; back-scattered electron and cathodoluminescence imaging, and trace element profiling of quartz and zircon grains; and zircon crystal size distribution measurements.
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SAMPLE COLLECTION, PREPARATION AND ANALYSIS |
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Lava and tuff samples were collected using published maps and field guides (Christiansen & Blank, 1972a
, 1972b, 1974; Christiansen, 1982, 1989), and were located using maps and the global positioning system (Table 1). We collected samples from all known post-caldera, low-18O lavas, all units of major tuff eruptions, and other samples covering the 2·0 my eruptive sequence (Table 1).
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Individual phenocrysts of quartz and sanidine, and obsidian spheres, were separated from hand specimens or from crushed rock by hand. Quartz was treated with cold fluoroboric acid to remove adhered feldspars and glass. All analyzed phenocrysts are primary and not the result of secondary precipitation from a hydrothermal fluid. The presence of glass inclusions and igneous morphology were taken as criteria of magmatic origin. Zircons were separated from 20 kg rock samples using standard techniques of crushing and density separation, and then purified with cold HF and HNO3. Separates of 50–250 mg of zircons were further subdivided by sieving into different sizes and by Frantz magnetic separator. No differences in 18O beyond 0·3 between different magnetic splits were found. Large-diameter zircons (>105 µm, or >149 µm) from 15 lavas and tuffs were air abraded in a corundum mortar for 0·5–5 days. Optical measurements, grain counting and mass of the starting vs remaining zircons provide an estimate of the decrease in the average zircon radius. After abrasion, zircons became ellipsoidal to nearly spherical with preferential rounding of pyramids and corners, i.e. along the c-axis. This procedure removes mass from zircon in exactly the reverse fashion to the crystal growth, i.e. shortening pyramidal faces grow several times faster than the prismatic faces (Vavra, 1994). Therefore, air abrasion retrieves cores and successive air abrasions provide a retrospective view of 18O during zircon crystallization.
The University of Wisconsin CO2-laser fluorination–mass-spectrometer system (Valley et al., 1995) provides accurate and precise determination of 18O for individual grains of quartz, sanidine, obsidian spheres and zircon splits. Samples were typically 1–2 mg, yielding 10–30 µmol of CO2. BrF5 was used as a reagent. Quartz phenocrysts, often single grains, were analyzed by rapid heating with defocused laser beam, yielding precise values (better than ±0·1, 1 SD) (Spicuzza et al., 1998a). For analyses of more reactive sanidine and obsidian, an airlock sample chamber was used to prevent partial pre-reaction (Spicuzza et al., 1998b).
Four to seven aliquots of UWG-2 garnet standard were measured at the beginning and end of each analytical session, and sample data were adjusted to the average value of the standards, typically by no more than ±0·2. The average reproducibility of 65 UWG-2 analyses is 0·10 (1 SD). Nine NBS-28 quartz analyses yielded an average value of 9·55 (Vienna Standard Mean Ocean Water; VSMOW).
Whole-rock chemical analyses for major and trace elements (Rb, Ba, Sr, Y, Nb, Zr) were carried out at XRAL Laboratories (Ontario, Canada) by X-ray fluorescence. The uncertainty of trace element determinations is ±2 ppm. Mineral and glass major element analyses, and cathodoluminescence imaging were performed on a Cameca SX-50 electron microprobe at the University of Wisconsin. For major element analyses, 15 kV accelerating voltage and 25 nA sample current was used, with synthetic crystals and minerals as standards. For trace element analyses of zircon and quartz, higher voltage (20–25 kV) and sample currents (100–300 nA), and 10–15 min counting times were used (Fournelle et al., 2000). For standards, we analyzed alloys of elements of interest at high concentrations, and Caltech–Corning synthetic glasses (at low currents) with trace element concentrations of 0·5–1 wt %.
Individual zircons in low-18O rocks were analyzed by Cameca ims-4f ion microprobe at the University of Edinburgh using 133Cs+ primary beam and energy-filtering (Valley et al., 1998a). A homogeneous KIM-5 zircon (18O = +5·04, Valley et al., 1998b) with similar Hf content to Yellowstone zircons (0·8–1·2 wt % HfO2) was used as a standard during the course of analyses. Ion microprobe analyses complement laser fluorination. Laser fluorination gives accuracy of better than ±0·10 (1 SD) for a milligram-size sample, but does not have the capability of analyzing a single zircon (100 µm x 200 µm). Ion microprobe analyses of 25 µm spots can measure zonation in a single zircon crystal, but are less precise (1, 1 SD in this study). Laser fluorination of air-abraded zircons was used to precisely measure the composition of zircon cores (in bulk), and the ion microprobe measured the composition of zircon rims and cores for individual crystals.
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GEOLOGY AND PETROLOGY OF YELLOWSTONE RHYOLITES |
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Newly defined low-
18O lava flowsAs a result of our fieldwork and oxygen isotope analyses, we propose a subdivision of the low-
18O Biscuit Basin Flow (BBF), which has previously been considered to result from a single eruption. BBF is exposed in three separate outcrops (hills), southern, middle and northern, along the Firehole River of the Upper Geyser Basin (see Fig. 1). These outcrops are separated by sinter and alluvial deposits. Drill hole Y-5 between the southern and middle outcrops does not show the presence of BBF (Christiansen & Blank, 1974). Previous studies analyzed only the middle outcrop across from the Midway Geyser (Hildreth et al., 1984, 1991; Obradovich, 1992; Gansecki et al., 1996). We found that although all three outcrops are equally glassy and rather similar macroscopically, 18O of quartz, zircon, feldspar and obsidian, and mineral–mineral fractionations are dramatically different (discussed below). Petrographically, the southern Biscuit Basin lava (SBB) is characterized by larger and more abundant phenocrysts (17–20 vol. %) than the middle (MBB, 10–12 vol. %), or the northern lava (NBB, 12–15 vol. %). Chemically, the three units are also distinct [Table 2; compare with analyses of Sturchio et al. (1986) for SBB and Hildreth et al. (1991) for MBB]. We suggest that each hill represents an independent lava flow or extrusive dome. We hence report analyses from all three lavas, calling them North, Middle and South Biscuit Basin Flows.
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Petrography
Yellowstone rhyolites are glassy and contain 0–25 vol. % phenocrysts of sanidine, quartz, plagioclase, clinopyroxene, magnetite and apatite, with minor ilmenite, fayalite and orthopyroxene. Zircon is present in all rhyolites. Hydrous minerals are conspicuously absent in all volcanic rocks except member A of the Lava Creek Tuff. Low-18O and normal-18O rhyolites contain the same assemblage of phenocrysts, except for post-LCT low-18O lavas, which are poorer in quartz and richer in plagioclase (oligoclase) at the expense of sanidine. Plagioclase is characterized by a sieved texture and optically shows complex zoning.
Quartz in low-18O lavas often has a corroded morphology and smaller grain sizes than in normal-18O lavas (Fig. 2), especially in the glassy Canyon and Dunraven Road flows, where the quartz grain size is <0·5 mm and the concentration is <0·5 vol. %. Quartz is more abundant (1–2 vol. %, 1 mm) in the Biscuit Basin Flows (post-LCT) and Blue Creek Flow (post-HRT).
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Cathodoluminescence (CL) imaging of quartz phenocrysts by electron microprobe from low-18O and normal-18O rhyolites revealed morphological features indicative of multiple episodes of magmatic dissolution–reprecipitation. Elongated melt inclusions penetrate the interior of quartz grains, truncating growth zones (Fig. 2a). Concentric growth is recorded in other quartz phenocrysts. Many quartz phenocrysts exhibit a bright core surrounded by a dark rim by CL. Trace element profiling by electron microprobe through one quartz crystal in South Biscuit Basin Flow (Fig. 2b) indicates that brighter zones and layers are richer in Al and K (up to 0·08 wt %). Al is proposed to cause bright CL in quartz (Marshall, 1988). No definite correlation between 18O of quartz and its CL pattern was found. Secondary, hydrothermally precipitated quartz sporadically occurs in altered lavas. Precipitated quartz is distinguished from phenocrysts by perfectly faceted, elongated, singly terminated prismatic crystals, and the absence of melt inclusions. No healed cracks, observed elsewhere in hydrothermally altered quartz (Valley & Graham, 1996; King et al., 1997), were found at Yellowstone.
Zircon is ubiquitous in all rhyolites of Yellowstone and provides the most compelling constraints in interpreting petrogenesis (see Appendix). Careful examination of mineral separates revealed that some zircons included in sanidine, quartz and magnetite phenocrysts. HF dissolution of both magnetite and sanidine + quartz separates yielded many tens of grains of zircon, indicating that zircons were located inside these minerals and pre-date crystallization of sanidine, quartz and magnetite. In particular, up to a third of all extracted zircons resided in magnetite in units CF and DR. This conclusion is in accordance with high zircon saturation temperatures (800–910°C; see Table 2), confirming that zircon is a near-liquidus phase, and that almost all Yellowstone magmas were always saturated with respect to zircon.
SEM examination of zircon grain mounts revealed grains of chevkinite (units SBB, DR) and baddeleyite (CF). Baddeleyite forms euhedral prismatic crystals of up to 100 µm long and contains 1–2 wt % of HfO2 and TiO2, and 3 wt % of Nb2O5. The presence of baddeleyite in high-silica rhyolite is remarkable and has not been reported previously. Typically occurring in silica-undersaturated rocks, baddeleyite can be present in direct contact with quartz in late-stage silica-rich differentiates of mafic plutons (Heaman & LeCheminant, 1993), although it is not thermodynamically stable with quartz below 1676°C.
Whole-rock chemistry
Low-18O rhyolites of each eruptive cycle are chemically similar to normal-18O lavas (Table 2). In addition, whole-rock compositions are similar in major element chemistry to melt inclusions in quartz, determined by electron microprobe analysis (see Table 2). This similarity indicates lack of secondary alteration of volcanic glass, such as alkali removal, in fresh-looking glassy samples.
The comparison of low-18O and normal-18O rhyolites was carried out using Sr/Rb and Ba/Rb ratios, which provide a measure of differentiation in similar high-silica rhyolites elsewhere (e.g. Anderson et al., 2000). For example, earlier erupted members of Huckleberry Ridge Tuff and Lava Creek Tuff are more differentiated based on lower Sr and Ba (down to a few ppm), higher Rb, and lower Sr/Rb ratios (Fig. 3a). We stress here that Sr/Rb and Ba/Rb ratios are also vulnerable to change by hydrothermal fluids, as glass-hosted Rb is more soluble than mostly feldspar-hosted Ba and Sr. Ratios of fluid-immobile elements to alkalis, such as Zr/Rb, are another measure of alkali loss during hydrothermal alteration (Fig. 3b). An example of this is the pervasively altered sample of Canyon Flow (CF-1, Table 2), collected at Artist Point of Yellowstone Canyon, which contains virtually no Na, K and Rb, whereas other major and trace elements are undisturbed. It should be noted that concentrations of fluid-insoluble Zr are nearly identical in this hydrothermally altered sample to those in low-18O rhyolites or normal rhyolites, demonstrating the immobility of Zr (and zircon) in the course of 18O depletion. Even hot, acidic fluids have been shown not to disturb Zr in volcanic rocks (Terakado & Fujitani, 1998).
We find neither clear chemical differences between low-18O and normal-18O rhyolites, nor evidence of preferential leaching of major or trace alkalis. On the contrary, low-18O rhyolites roughly follow differentiation trends defined by normal rhyolites with scatter that can be explained by local and regional source rock heterogeneities (Fig. 3a and b). Most 18O-depleted, post-LCT rhyolites contain slightly more Mg, Sr and Ba, and less SiO2 and Rb. Therefore, they are slightly more ‘primitive’ than normal-18O post-LCT rhyolites. The relations are opposite for post-HRT low-18O rhyolites, which are slightly more evolved than HRT and later lavas. More significantly, low-18O rhyolites of the last eruptive cycle (post-LCT) display a large, within-cycle diversity (Fig. 3), suggesting that they do not represent a single well-mixed melt.
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The similar major element composition of low-18O rhyolites vs normal-18O rhyolites (Table 2) may indicate that hydrothermal fluids were saturated with respect to alkalis and other soluble trace elements, and did not quantitatively remove these elements from the magma generation zones before melting. The preservation of trace element concentrations comes as no surprise, as these hydrothermal fluids consist of 88% oxygen and only parts per million of trace elements (e.g. Lewis et al., 1997). Thus, for moderate fluid/rock ratios of 0·5–1, oxygen in a rock is diluted by over 10%, but trace elements are not measurably affected. In such an environment, oxygen isotopes are decoupled from trace elements and meteoric water is capable of significantly changing 18O, but not cations.
Liquidus temperatures of magmas
Temperature estimates for zircon saturation (Watson & Harrison, 1983) vary from 830 to 910°C for the three major tuff eruptions and are similar to those for low-18O intra-caldera rhyolites (800–875°C; see Table 2). Magnetite–ilmenite thermometry yields 820–920°C, and oxygen fugacities close to quartz–fayalite–magnetite (QFM) buffer (Hildreth et al., 1984). Neither method indicates a significant temperature difference between low-18O and normal-18O magmas, with the exception that the low-18O rocks are slightly more oxidized and plot 0·5 log unit above the QFM buffer (Hildreth et al., 1984). The bulk chemical compositions, plotted on an Ab–Or–Qz diagram, point to a relatively water-poor composition (PH2O < 0·5 kbar), shallow crustal pressures (0·5–2·5 kbar), and high temperatures (850–950°C) for Yellowstone rhyolite magmas (e.g. Doe et al., 1982).
Liquidus and solidus temperatures were also determined using the MELTS program (Ghiorso & Sack, 1995) based on whole-rock analyses and the composition of melt inclusions (see Table 2). We found no difference in calculated liquidus or solidus temperatures between low-18O and normal-18O rhyolites resulting from similar major element chemistry. At 3 wt % H2O and 1 kbar, the calculated liquidus temperatures are similar to zircon saturation temperatures with the average at 850°C (see Table 2). For these conditions, the minimum solidus temperatures, or eutectic, are between 710 and 750°C. A rise from 710 to 750°C leads to a spontaneous increase in melt fraction from 0% to >50%. This increase plays an important role in the segregation of rhyolite magmas, discussed below. These liquidus–solidus relations are consistent with melting–crystallization experiments with similar high-silica rhyolitic compositions (Whitney, 1988).
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OXYGEN ISOTOPE RESULTS |
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Oxygen isotope ratios of minerals and glass from 25 eruptive units of Yellowstone are listed in Table 3. In Fig. 4,
18O is plotted against the age of each eruptive unit (see Table 1). Each sample is represented by the vertical line on which 18O of coexisting minerals is plotted allowing visual estimation of mineral variability and mineral–mineral isotopic fractionation.
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General observations
Individual quartz phenocrysts
Individual quartz phenocrysts (Fig. 4a) show an evolution pattern with extreme
18O depletions after LCT and HRT eruptions. This pattern follows the trend of bulk phenocrysts shown by Hildreth et al. (1984) with one important difference. Analyses of individual quartz phenocrysts reveal strong variability from grain to grain in low-18O rhyolites which erupted immediately after the LCT, with 1–2 variations among different crystals within a single hand specimen. This variability exceeds 2 for Canyon Flow and South Biscuit Basin Flow (Fig. 4a and b). Post-HRT low-18O rhyolites also show 18O variability among quartz phenocrysts. In contrast, analyses of individual quartz phenocrysts from other rhyolites are homogeneous to within ±0·2. Thus, many quartz phenocrysts in the low-18O rocks are not in equilibrium with each other, the host obsidian or feldspar. Values of 18O(Qz–Obs) and 18O(Qz–Fsp) are larger than equilibrium values of 0·35 and 0·75 for magmatic temperatures (see below, Table 3) and many quartz phenocrysts have higher than equilibrium 18O values. There is no correlation between crystal size and 18O.
Zircon
Oxygen isotope ratios of zircon broadly follow the quartz trend (Fig. 4a and b). 18O(Qz–Zrc) are in high-T equilibrium for normal-18O pre-caldera and extra-caldera lavas and tuffs, and 18O of different size fractions of zircon are on average homogeneous within ±0·15 (1 SD). In contrast, low-18O lavas are characterized by zircon–quartz disequilibrium (Fig. 4c and d; Table 3). Values of 18O(Qz–Zrc) between individual quartz phenocrysts and bulk separates of zircon deviate by >2 below the equilibrium value of 2·19 ± 0·07 (Fig. 4d). Values of 18O(Obs–Zrc) and 18O(Fsp–Zrc) are often less than 1·76 and 1·45, respectively. Thus, zircon is 18O richer than expected for equilibrium with quartz at magmatic temperatures. In addition, differences of up to 0·7 in 18O exist between the smallest (<53 µm diameter) and largest (>149 µm) zircons from all low-18O lava flows erupted after LCT, and in Blue Creek Flow erupted after HRT.
Individual obsidian spheres and feldspar phenocrysts
Analyses of sanidine phenocrysts (and their parts) and obsidian spheres demonstrate that sanidine is a fraction of a per mil lower in 18O than obsidian, and obsidian is a fraction of a per mil lower than quartz. These fractionations are characteristic of equilibrium at magmatic temperatures for silicic rocks (e.g. Taylor & Sheppard, 1986). In low-18O lavas of the post-LCT cycle, 18O values of plagioclase (andesine) overlap with the 18O range of obsidian, confirming that their magmatic 18O values were not hydrothermally altered following emplacement. Furthermore, analyses of individual sanidine phenocrysts and obsidian spheres in low-18O lavas that show heterogeneity among individual quartz phenocrysts, demonstrate remarkable homogeneity (Table 3).
Mineral–mineral and mineral–obsidian fractionations
With the exception of low-18O lavas appearing after LCT and HRT, all other rocks exhibit similar, small and regular isotopic mineral–mineral, and mineral–obsidian fractionations, consistent with magmatic temperatures. All lavas in this study are high-silica rhyolites and have similar calculated temperatures of 800–900°C (Table 2). Thus, isotopic fractionations and calculated isotopic temperatures are also expected to be similar, despite significant variations in whole-rock 18O. Figure 5 plots 18O(Qz) vs 18O(Zrc). Equilibrium quartz–zircon fractionation (Fig. 5, open squares) is determined to be 2·19 ± 0·07 (1
, n = 20 samples). This value is between that predicted by empirical (2·0 ± 0·2, Valley et al., 1994) and theoretical (2·57, Zheng, 1993) calibrations at 850°C. Other fractionations for Yellowstone high-silica rhyolites (excluding low-18O samples) are: 18O(Qz–San) = 0·75 ± 0·08 (1, n = 15 samples); 18O(Qz–Obs) = 0·35 ± 0·04 (1, n = 6 samples); 18O(San–Zrc) = 1·45 ± 0·11 (1, n = 15 samples); and 18O(Obs–Zrc) = 1·76 ± 0·19 (1, n = 6 samples). The predicted quartz–sanidine equilibration temperature (Clayton et al., 1989), based on our measured fractionation, is 850 ± 60°C, which is in excellent agreement with other temperature estimates (see Table 2).
Oxygen isotope zoning in zircon
The disequilibrium relations for zircons and quartz in low-18O rhyolites (see Figs 4d and 5) suggest that variability or zoning is present among single zircons. Several tests were designed to evaluate this possibility. Zircons were separated by size and large zircons were air abraded. Larger zircons, as well as cores of air-abraded zircons, are 1–3 higher in 18O than zircon in equilibrium with obsidian (see Fig. 4b). Smaller zircons are inferred to have grown later, or exchanged oxygen by diffusion more completely because of large surface/volume ratio. Smaller zircons thus represent the same timing as rims of large zircons. This is the first time oxygen isotope zoning in zircon has been demonstrated and it is significant for tracing magma evolution. Future studies will pursue this using ion microprobe analysis for stable isotope ratios as well as U–Pb age. The 18O differences between larger and smaller zircons prompted the search for isotopic differences between cores and rims of zircons.
Zircon air abrasion
Analysis of groups of air-abraded zircon cores from low-18O rocks closely approach the normal 18O value for zircon in pre-caldera magmas in some cases. Successive air abrasions were made on abundant zircons in the Blue Creek Flow and Middle Biscuit Basin Flow to measure 18O after each step of rim removal. Values of 18O are plotted as a function of average radii of residual zircon cores (Fig. 6). The upper mark of each bracket shows measured 18O of residual cores and the lower mark is the calculated 18O of the rims. The value in per mil on each curve is the bulk 18O for the whole zircon crystal. The resulting calculated profile shows an average gradient of 1 per 2–3 µm at the rim, and core-to-rim zoning approaches 5. Rims are in overall equilibrium with the host glass.
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Ion microprobe analyses
The
18O values of the zircon rims were determined by ion microprobe analysis on prismatic crystal faces of zircon grains pressed into indium metal, whereas the 18O of cores of individual crystals were determined in polished grain mounts (Table 4). The average zircon rim 18O is consistent with that predicted from mass balance calculations based on zircon abrasions and for equilibrium with obsidian. Zircons in Lava Creek tuff (sample LCT-3a) do not show 18O zoning. Air abrasion and ion microprobe-measured 18O values of rims in this sample are similar to those of cores, as predicted. In post-LCT and post-HRT lavas, rims of zircon on average are several per mil lower in 18O than cores (samples YL96-4, unit DR, Fig. 1; YL96-20, unit MBB; and BC-1, unit BC, see Fig. 6). Repeated analysis by ion microprobe on the same spot on the face allowed monitoring of 18O vs depth. A typical 20 min analysis penetrated 2 µm into the crystal after a preliminary 1 µm burn-in, thus making the average depths of the first and second analysis at 2 and 4 µm, respectively. In seven out of nine analyses of zircon faces the second analysis yielded higher 18O with a typical gradient of 0·5 per 2 µm (Table 4), consistent with the average gradient predicted from air abrasion (Fig. 6).
The 18O values of zircon cores in two samples (YL96-4 and YL96-2) show large crystal-to-crystal variations. Likewise, rims in YL96-2 (unit SBB) and YL96-18 (unit CF) show larger variability from crystal to crystal, up to the magnitude of core–rim variability (Table 4). Rims on zircons in CF are on average similar to cores and are 5 higher in 18O than zircon expected to be in equilibrium with the host rock (Table 3).
The ion microprobe analyses provide the only data for 18O within single zircon crystals, but the spot analyses do not perfectly match the zoning predicted by bulk analyses. Individual zircons represent a mixture of grains of different age and source (Bindeman et al., 2001a), and thus may preserve initial 18O differences between zircon crystals. Although 18O of bulk zircon does not vary significantly in pre-LCT Yellowstone rocks, the range in 18O is from 4 to 6. Air abrasion tends to remove the fastest grown pyramids, thus preferentially removing any overgrowths. Variability of rim compositions may also be related to a sampling bias; the compositions of rims were measured only on the largest crystals (100 µm x 200 µm) and on the slowest-growing prismatic crystal faces.
In contrast to crystal growth, dissolution into mell would act more homogeneously and be less face-selective (e.g. Watson & Harrison, 1983). In particular, if the dissolution of the prismatic faces is faster than the oxygen isotope exchange by diffusion, it will result in retaining relatively higher 18O of the prismatic faces, as seen in some of the ion microprobe analyses. Dissolution and/or reprecipitation, therefore, may compete with diffusive exchange, and are most likely to expose higher-18O cores on prismatic faces. Additionally, not all zircons float freely in the melt and experience isotropic diffusive exchange with the melt. Many crystals are attached to, or included in, other phenocrysts such as magnetite, quartz and feldspar (see above). The crystal faces of these zircons may not have been in equilibrium with low-18O melt and each zircon may have experienced different degrees and rates of diffusive exchange with its host melt or phenocrysts. We hypothesize that individual zircon grains may differ in their initial 18O, entrapment, overgrowth vs dissolution, and zircon–mineral exchange histories.
Products of major tuff eruptions
Our individual quartz phenocryst 18O analyses in the three major tuff eruptions, Huckleberry Ridge Tuff (HRT, three members: A, B and C), Mesa Falls Tuff (MFT, one member) and Lava Creek Tuff (LCT, two members: A and B), agree well with bulk analyses of quartz phenocrysts from Hildreth et al. (1984) (Fig. 4a). Even the 18O(Qz) of the earliest airfall deposits of each eruption is similar to the quartz in bulk magma (Hildreth et al., 1984). Zircon from these voluminous normal-18O tuffs is in equilibrium with quartz, and the 18O of smaller zircons is similar to 18O of large zircons in the same rock. This implies that these tuffs originated as magmas that were homogeneous and equilibrated with respect to oxygen isotope ratios. The only exception is the Member C of HRT. Despite being more radiogenic in terms of Pb and Sr isotopes (Hildreth et al., 1991), we find that 18O of quartz and zircon in Member C are 0·5 higher than for members A and B (Table 3). Member C erupted last in the sequence and is likely to represent deeper layers of an isotopically zoned, batholithic-scale magma chamber, in which there is a subtle downward increase in 18O. Values of 18O of individual quartz phenocrysts in members A and B of LCT from different parts of the caldera overlap and are similar to analyses of bulk phenocrysts in Hildreth et al. (1984) from 22 locations, confirming that there are no significant lateral or spatial variations. Furthermore, among individual quartz phenocrysts (14 in LCT, six in MFT and 20 in HRT), we found no low-18O xenocrysts.
Post-caldera lavas
Post-HRT (first) cycle
Two low-18O lava flows followed the Huckleberry Ridge Tuff eruption: the Blue Creek Flow and Headquarters Flow. The Blue Creek Flow is similar with respect to trace elements, and in Pb and Sr isotope ratios to member C of HRT (Hildreth et al., 1991), whereas the Headquarters Flow is distinct. We found differences between the two flows in 18O of zircon and quartz. Zircons in the Blue Creek Flow are zoned by at least by 3, as revealed by air abrasion and ion microprobe analysis (Fig. 6). Individual quartz phenocrysts also show larger variability (0·7) than is typical for equilibrated rocks. The Headquarters Flow, in contrast, contains fully equilibrated quartz, and abraded zircon cores are only 0·18 higher in 18O than the bulk (see Table 3). Bishop Mountain Flow, which postdates the low-18O flows, contains fully equilibrated quartz and zircon with more normal 18O values.
Post-MFT (second) cycle
No significant 18O depletion in or around Henry Forks Caldera has been found after eruption of the Mesa Falls Tuff. Quartz, zircon, sanidine, and obsidian are in overall equilibrium. However, abraded zircon cores in Warm River Butte Flow (WRB), which erupted shortly after MFT, and in Harlequin Flow, erupted later than WRB, are 0·4–0·6 higher in 18O (by >2 SD) than the value for bulk zircons or smaller zircons. This suggests a modest core–rim zoning, 1·5. Post Warm River Butte lavas have nearly constant 18O values for both quartz and zircon, and are similar to those of LCT. Given the distribution of post-MFT rocks around the Yellowstone caldera (Table 1, Fig. 1), it can be assumed that post-MFT lavas (and their plutonic counterparts) form much of the basement and may serve as country (source?) rocks beneath the LCT and rhyolites of the last eruptive cycle.
Post-LCT (third, or last) cycle
A dramatic 18O depletion is seen in lavas and tuffs erupted inside Yellowstone Caldera after the Lava Creek Tuff eruption. Post-LCT lavas erupted outside of the caldera have higher 18O values than LCT (Fig. 4a; also Hildreth et al., 1984). Both quartz and zircon in low-18O lavas show significant 18O heterogeneity (see Fig. 4c), whereas obsidian and feldspar are homogeneous within analytical precision. Values of 18O in many quartz phenocrysts in South Biscuit Basin Flow approach +6, similar to 18O in quartz from LCT. Cores of abraded zircons from South Biscuit Basin and Canyon Flows are +3·5, approaching the pre-caldera value of +4. Successive air abrasion of larger zircons (>105 µm) from Middle Biscuit Basin Flow, and ion microprobe analyses of zircon rims (Table 4) allow reconstruction of a zoning profile (Fig. 6). Given the steep gradient of the zoning profile, it is likely that innermost cores (unattainable by abrasion) in MBB (and likewise DR flows) also approach pre-caldera values of +4. It is significant that most smaller zircons (<53 µm) have 1 higher 18O than the calculated rim composition of larger zircons based on zircon abrasion. This suggests that even the smallest zircons are isotopically zoned, with higher 18O cores and lower 18O rims. In this case, the high-18O zircon cores suggest that most of the zircons are xenocrysts and that isotopic exchange with low 18O magma shifted the bulk 18O values of smaller zircons more than that of larger zircons because of their higher surface area/volume ratio. The zircon zoning and disequilibria are confined to 18O-depleted intra-caldera rhyolites erupted <150 ky after LCT. This zoning and disequilibria are not seen in rhyolites erupted >150 ky after LCT (Fig. 4d). In particular, the voluminous Scaup Lake, West Yellowstone, and Aster Creek Flows, and the less voluminous North Biscuit Basin Flow, show no significant heterogeneity in 18O of quartz phenocrysts, no 18O zoning in zircon and no quartz–zircon disequilibria.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONSAMPLE COLLECTION, PREPARATION...GEOLOGY AND PETROLOGY OF...OXYGEN ISOTOPE RESULTSDISCUSSIONAPPENDIX: ZIRCON IN YELLOWSTONE...REFERENCES |
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The 2–3
variability among individual phenocrysts of quartz in low-18O lavas has not been observed in other calderas, in part because individual quartz phenocryst analyses are not routinely performed. Quartz phenocrysts in single pumice clasts of the Bishop Tuff, and intra-caldera lavas and voluminous low-18O tuffs of Timber Mountain–Oasis Valley caldera complex (I. N. Bindeman & J. W. Valley, unpublished data, 2000) do not show such variable values. Analyses of individual quartz phenocrysts are nearly identical to the bulk, and the variability is uniformly low (<0·4). Thus, the variability of phenocrysts is specific to Yellowstone’s low-18O volcanic rocks and is indicative of their genesis.
We emphasize that because zoned and normal-18O zircons and quartz phenocrysts occur in homogeneous low-18O glass, exchange of oxygen is likely, and probably took place through diffusion and solution–reprecipitation. Even if only diffusion is considered, any initial 18O variations in melt disappear several orders of magnitude faster (Zhang et al., 1991) than variations in quartz and zircon (Farver & Yund, 1991; Watson & Cherniak, 1997).
For quartz, solution–reprecipitation is potentially a faster process of exchange than diffusion, as silica (unlike zirconia) is a major constituent of rhyolites. Furthermore, the diffusion of silica in water-bearing rhyolite at 750–850°C is 3 x 10-10 to 9 x 10-11 cm2/s (Baker, 1991). Diffusion of silica in melt is 1–2 orders of magnitude faster than even ‘wet’ diffusion of oxygen in quartz at 750–850°C [1·4 x 10-12 to 1·1 x 10-13 cm2/s, Farver & Yund (1991) and therein]. Solution–reprecipitation features, such as truncation of growth zones by penetrating melt inclusions, are seen in CL images of some quartz phenocrysts (e.g. Fig. 2), but mostly quartz exhibits concentric CL zoning indicative of steady growth. Therefore, diffusive exchange following crystallization appears to be the dominant mechanism of isotope exchange in quartz, although diffusion distances may be significantly reduced by melt inclusions.
For zircon, diffusion may dominate over new growth or solution–reprecipitation. Oxygen diffusion coefficients at PH2O >70 bars in zircon residing in 850°C water-bearing magmas are determined experimentally to be of the order of 10-17 cm2/s (Watson & Cherniak, 1997). At this diffusion rate, a typical zircon grain of 50 µm radius would exchange and nearly equilibrate after a few tens of thousands of years. The rate of zircon growth in magmas is controlled by sluggish Zr diffusion in melt. Dissolution experiments suggest that it would take as long or longer to crystallize a new 50 µm zircon grain as to exchange it by diffusion. For instance, a zircon of 50 µm radius would grow in 150 ky at a linear growth rate of 10-15 cm/s (Watson, 1996). Zircon crystal size distributions (Appendix) would be consistent with 1·56–2·27 µm/ky growth at this rate. We conclude that both diffusive exchange and new overgrowth played a role in the isotopic exchange of zircon, quartz and magma.
Modeling of diffusive isotope exchange and re-equilibration
The core-to-rim 18O zoning observed in zircons from low-18O rhyolites (Fig. 6) indicates that zircons (and quartz) with normal 18O values were immersed into low-18O melt. Such xenocrysts of zircon and quartz will exchange at different rates creating disequilibrium values of 18O(Qz–Zrc) as shown in Figs 4d, 5 and A2 (Appendix). If given enough time, exchange will become complete and xenocrysts will re-equilibrate at lower 18O.
Diffusive exchange of oxygen isotopes in quartz and zircon was modeled [equation (6.18), Crank, 1975] for diffusion in a sphere surrounded by an infinite reservoir (magma) at 750–850°C. Below, we describe results of modeling at 850°C, the near-liquidus conditions for Yellowstone magmas (Table 2); rates of diffusive exchange at near-solidus temperatures of 750°C are 10 times slower. Most exchange occurs at higher temperatures and the model below constrains the minimum time necessary for diffusive exchange.
The zoning profiles and the bulk composition of whole zircons (Fig. 6) can be compared with those of the diffusion model. The measured zoning profile of 18O in zircons from Blue Creek Flow (Fig. 6) fits between that predicted for 5–10 ky of residence of normal-18O zircon (+5·6) in low-18O melt (+2·8, in equilibrium with +1·0 zircon). Normal (+4) zircons in Middle Biscuit Basin Flow would require 500–5000 years to develop the observed zoning.
Figure 7 puts the observed variations in 18O of both quartz and zircon, and (Qz–Zrc) for lava flows, into a single model of isotopic exchange (Bindeman & Valley, 2000). Crystals of different size were chosen (similar to those observed) and the 18O (core) in each crystal was calculated and compared with the average 18O for the whole crystal. For zircons, we also consider the case in which a new overgrowth forms at a constant rate. The curves of 18O exchange and equilibration result from different rates of diffusive exchange. The key observation in Fig. 7 is that intervals of the maximum 18O zoning within quartz and within zircon do not generally overlap because diffusion in zircon is much slower. The maximum deviations from equilibrium for core or bulk compositions occur at transient times (Fig. 7b). During these times, (Qz–Zrc) shows wide variations similar to that observed in low-18O lavas (Fig. 4d). Quartz–zircon disequilibrium, therefore, serves as a gauge for estimating the time of diffusive exchange in magmas.
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In Fig. 7a, we also plot the measured range of 18O in zircon, and the observed variability for individual quartz phenocrysts in each unit between the calculated core and bulk curves. Variability of quartz in SBB and CF suggests shorter times of exchange (200–500 years), which can be inferred from the observation that quartz in these units has the largest variability, whereas slower-exchanging zircon has barely exchanged. DR erupted on top and shortly after CF (Gansecki et al., 1996; Bindeman et al., 2001a). DR may represent the same magma as CF, in which quartz had more time to exchange. Quartz phenocrysts in MBB and DR show only limited remaining heterogeneity, whereas zircons are most zoned (Fig. 4c). This suggests that the time scale for MBB and DR magmas was long enough for quartz equilibration, but not long enough for zircons to fully equilibrate [equivalent to 102–103 years for (Qz–Zrc) bulk in Fig. 7b]. The diffusion calculations indicate that the time between the beginning of exchange and eruption was of the order of hundreds to thousands of years. For other low-18O lavas, which show no zoning in zircon and erupted significantly after LCT (e.g. WY, AC, ML), or after HRT (BM), the residence time could have been significantly longer.
Origin of low-18O magmas of Yellowstone
The equilibration calculations (see Figs 6 and 7) suggest that at least a few hundred years are necessary to develop 18O zoning in quartz and zircon. These calculations also bracket the maximum time limit needed to preserve this heterogeneity, 25 ky. This time interval is an important constraint for any model.
Previous models
The original and attractive model of ‘catastrophic’ meteoric water–magma interaction during caldera collapse (e.g. Hildreth et al., 1984) is not supported by our data on 18O zoning in zircon and quartz from low-18O lavas. The length of time required to create zoned xenocrysts is too great. Assimilation (bulk or partial) and magma mixing can create low-18O magmas (e.g. Taylor, 1986; Bacon et al., 1989; Balsley & Gregory, 1998; Bacon et al., 2000), but we present three arguments against these possibilities.
First, the observed low-18O magmas of Yellowstone have 18O from 0 to +3 and such values are restricted by mass balance. In the extreme case of mixing equal proportions of normal (+6) magma and 18O-depleted magma or rock, the 18O (bulk assimilant) would have to be <-6 to create magma at 0. Mixing with an ultra-low-18O magma (<-5) is unlikely because no ultra-low-18O individual phenocrysts of quartz and zircon (Tables 3 and 4) have been found. Furthermore, there is no known source for the genesis of such very low-18O magmas nor are any wall rocks at Yellowstone observed with sufficiently low 18O.
Second, it is difficult to mix tens to hundreds of cubic kilometers of two viscous rhyolite magmas and achieve the textural, chemical and isotopic homogeneity observed in volcanic glass of Yellowstone lavas. The uniformity of 18O (within ±0·5, Table 3) in individual obsidian spheres in low-18O rocks is not consistent with the magma mixing model.
Third, in assimilation of low-18O rocks models, the heat balance controls the maximum amount of assimilation. If all the heat comes from the crystallization of the parent magma, then, of the maximum possible, assimilation is 50% before the magma becomes 100% crystalline, given the similar chemistry (and heat capacities) of both the magma and the assimilant. The observed low-18O rhyolites are, on the contrary, crystal poor (<15%), showing that the percentage of any assimilation was small.
The bulk melting–caldera collapse (BMCC) model
We propose that low-18O magmas originate as a result of the nearly total, bulk melting of hydrothermally altered, low-18O rocks after caldera collapse. The source rock must be strongly depleted in 18O as is expected in the roof zone of large, shallow magma chambers. Depletion results from circulation of meteoric water, and hydrothermal alteration of volcanic glass and possibly all other phenocrysts, except alteration-resistant quartz and zircon (Criss & Taylor, 1986; Gilliam & Valley, 1997). One example of such alteration is exposed in outcrop: the quartz and zircon in pervasively altered Canyon Flow (sample CF-1, see Tables 1–3) preserve 18O values similar to those of quartz and zircon in unaltered samples.
We have demonstrated above that quartz and zircons in low-18O magmas are higher in 18O than would be in equilibrium with their host low-18O magmas, and zircons contain cores similar to pre-caldera rhyolites. Thus, these zircons did not crystallize from the magma in which they now reside. The most feasible way to disperse zircon xenocrysts in a viscous low-18O magma would be by near-total fusion of low-18O source rock(s) containing these zircons. The deficiency of smaller zircons with larger surface to volume ratio (Fig. A1, Appendix) may support the evidence of preferential dissolution of smaller grains, and survival of larger cores in the event of melting. The same argument applies to quartz, which shows lesser abundance, smaller sizes and resorbed morphology in low-18O lavas.
We propose that remelting of hydrothermally altered roots of down-dropped blocks following caldera collapse creates low-18O magmas with normal-18O xenocrysts. Pre-caldera lavas and tuffs are the most likely source rock for remelting. They have less variable 18O(Qz) (5·5–7·5) and 18O(Zrc) (3·5–5·5) than low-18O lavas. Trace element similarity of low-18O rhyolites to normal-18O pre-caldera rhyolites (e.g. Fig. 3), and not to Archean basement or Eocene Absaroka volcanic rocks, makes them possible intra-caldera substrate for melting.
SHRIMP dating of 89 grains of zircon in LCT, HRT and post-LCT lavas (Bindeman et al., 2001a) demonstrated no pre-Quaternary ages, and few LCT-aged zircons in post-LCT lavas. These results suggest that pre-Quaternary rocks were not involved in melting, and also that LCT magma (or its crystal rinds, e.g. Mahood, 1990) was not a dominant source for post-LCT low-18O rhyolites. The latter conclusion is also supported by 40Ar/39Ar individual feldspar age data (Gansecki et al., 1996). Zircon age spectra in each post-LCT low-18O lava (CF, DR, MBB) consist of three distinct age groups: 2·1 ± 0·4 Ma (consistent with HRT), 1·0–0·7 Ma (pre-LCT volcanic rocks) and the eruption age of 0·5–0·6 Ma. These age spectra in each low-18O lava suggest that post-LCT low-18O melts originated from more than one source rock, notably HRT and pre-LCT volcanic rocks, residing in the down-dropped intra-caldera block.
Caldera collapse and low-18O rhyolites
We consider caldera collapse as essential for formation of low-18O rhyolites at Yellowstone. The absence of low-18O xenocrysts in HRT and LCT argues against the existence of a significant low-18O magma layer(s) in any of the dominant magma chambers, which could result from the melting of low-18O roof rock before caldera formation.
Figure 8 illustrates the main points of the bulk melting–caldera collapse (BMCC) model. Before caldera collapse, intensive hydrothermal alteration by heated surface waters lowers the 18O of country rocks (Stage I). The roof of the magma chamber (source rocks to low-18O rhyolites) becomes low in 18O(WR) through alteration of feldspars and glass, but not quartz and zircon. Studies of 18O in whole rocks, glass and quartz in drill holes of 2 km depth in Bishop tuff within Long Valley caldera show that quartz survives hydrothermal alteration, and that the degree of whole-rock 18O depletion increases with depth and temperature (Smith & Suemnicht, 1991; McConnell et al., 1997). Similar results were obtained at Yellowstone (Sturchio et al., 1990).
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Caldera collapse (Stage II, Fig. 8) brings the base of the hydrothermally altered, low-18O(WR) down-dropped block closer to the hot interior of the magma chamber. The failure of non-rhyolitic magma to erupt during major caldera-forming events suggests that most magma was not erupted and remained in batholith-scale magma chambers, parental to LCT and HRT (e.g. Smith, 1979). The conservative estimate of vertical drop as a result of caldera subsidence after the eruption of 1000 km3 of LCT is 500 m based on an eruptive volume–caldera area argument (Smith, 1979; Hildreth et al., 1984). R. L. Christiansen (personal communication, 2001) estimates the minimum down-drop of 600 m after LCT, and 700 m after HRT, based on the projection of thickness of LCT and HRT from caldera walls. It is also possible that the vertical drop is a few times larger. For example, the Long Valley caldera comparable in size with Yellowstone caldera formed by eruption of 650 km3 of Bishop Tuff, and subsided 2–3 km, as demonstrated by the thickness of intra-caldera tuff (McConnell et al., 1995). The vertical drop could be even larger following eruption of 2500 km3 of HRT (Smith & Christiansen, 1980), but still smaller than the vertical dimension of most batholiths (of 5–15 km; Ague & Brandon, 1996; Anderson et al., 1997).
It is possible that caldera collapse caused fresh basalt to underplate the floor of the magma chamber soon after climatic eruption, as the denser basalt would compensate for the isostatic deficiency after removal of up to thousands of cubic kilometers of rhyolite. Basalt would also provide heat for subsequent melting of low-18O rocks. However, no basalt erupted with, or immediately after, eruption of low-18O magmas. If basalt underplating occurred, the layer of remaining LCT magma between the down-dropped block and the underplating magma appears to have served as a density filter.
In Stage III (Fig. 8), remelting of the roots of the down-dropped block forms pockets of low-18O magma near the roof of the magma chamber. As most Yellowstone rhyolites are high-silica, near-cotectic melts, crystallization modeling (MELTS; Ghiorso & Sack, 1995) above shows that a temperature rise of from 710°C to over 800–850°C yields 100% melting. The melting could have proceeded at lower temperatures than the temperature in the interior of the magma chamber, possibly facilitated by the water-richer nature of hydrothermally altered source rocks. The pockets of newly formed rhyolitic melt may have been near, but not necessarily at, the contact with the main magma chamber itself. There could have been screens between low-18O magma and the main chamber composed of (1) pre-existing or newly formed crystal rinds of the magma chamber, and/or (2) older (pre-Quaternary) rocks with higher melting temperature. Pre-Quaternary rocks probably participated in exchange with magma following HRT eruption, as HRT member C and post-HRT Blue Creek low-18O lava flow are more radiogenic (Hildreth et al., 1991). However, by the LCT cycle, either pre-Quaternary rocks in the intra-caldera block were completely digested or lower-temperature melting components were exhausted.
The most 18O-depleted source rocks were originally concentrated close to the magma chamber, where Stage I hydrothermal alteration was most intense. They thus melted first, giving rise to low-18O magmas, which erupted independently near the eastern and western resurgent domes (ML and SCRD in Fig. 1) of the last Yellowstone caldera, preserving local chemical (e.g. Fig. 3, Table 2) and 18O variability of the parent source. The coeval relations between post-caldera resurgence and the appearance of low-18O lavas may serve as geological evidence of the melting in the down-dropped block, which caused local isostatic rebound. Each lava formed by this process and was homogenized within itself, but contained xenocrystic quartz and zircon inherited from the source (Fig. 8, Stage III). We favor the formation of isolated ‘pockets’ of melt, rather than a well-mixed layer of low-18O magma on the top of the magma chamber. As post-collapse, low-18O magmas are chemically similar to pre-caldera normal-18O lavas, they must have been derived from remelting of chemically similar, but 18O-depleted (hydrothermally altered) rhyolites, products of previous magmatic activity. This explains why the first magmas to appear after LCT (CF, MBB, DR), and after HRT (BC, HQ) caldera collapses are the most 18O depleted. The isotopic ‘recovery’ toward the less 18O-depleted compositions of younger intra-caldera lavas could result from mixing with normal magma from the interior, as was suggested previously (e.g. Hildreth et al., 1984, 1991). We suggest that this recovery is more probably the result of progressive remelting of less 18O-depleted rocks of the caldera block farther from the magma chamber.
Other calderas and the origin of intra-caldera volcanism
The BMCC model is not restricted to the genesis of low-18O magmas. This process of shallow remelting is likely to be widespread and may characterize many other calderas that do not have recognizable 18O depletions. Total remelting of geochemically similar rhyolites of pre-caldera rocks using only heat but not matter from the dominant magma is an important newly proposed process of intra-caldera volcanism. It is largely cryptic and can best be discovered using oxygen isotope analyses, but it may be important in calderas where distinctive, low-18O magmas do not occur. The BMCC model is significantly different from older models in which the dominant magma assimilates roof rocks. The BMCC model is helpful for explaining post-caldera volcanism and resurgence, and assessing chemical evolution, and the energy balance that is pertinent to volcanic hazards.
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APPENDIX: ZIRCON IN YELLOWSTONE MAGMAS |
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TOPABSTRACTINTRODUCTIONSAMPLE COLLECTION, PREPARATION...GEOLOGY AND PETROLOGY OF...OXYGEN ISOTOPE RESULTSDISCUSSIONAPPENDIX: ZIRCON IN YELLOWSTONE...REFERENCES |
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Zirconium distribution
To determine how much of the total whole-rock Zr resides in zircon, 657 g of a glassy rhyolite (sample YL96-4, unit DR, 393 ppm of Zr, Table 2) was digested in HF, and yielded 93 mg of zircon. This corresponds to 141·5 ppm of zircon in the rock, or 70·5 ppm of Zr. Therefore, only 18% of the total zirconium in this sample resides in zircon, whereas most Zr remains in the glass.
Crystal size distribution and shape
Zircon morphology and crystal size distribution (CSD) were used to estimate the residence time of zircon populations at nominal growth rates (Fig. A1). The CSD was measured on bulk zircon separates using digital images and the Scion-NIH image program. The advantage of using individual whole grains for size measurements (as opposed to two-dimensional thin-section images) could be compromised if there was selective loss of smaller grains during zircon separation procedures. To confirm that conventional separation procedures did not bias the original size distribution, the CSD of HF-extracted zircons was compared with that for conventionally extracted zircons (sample YL96-4, unit DR). HF-extracted zircon size distribution overlapped that obtained using the conventional technique (Fig. A1), thus showing no introduced bias.
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Most zircons are 30–350 µm long, but rocks with smaller quantities of zircons are characterized by smaller zircons. The majority of zircons, regardless of size, are euhedral tetragonal dipyramids. The average length-to-breadth ratio (L/B) for zircon increases linearly from 1·5:1 for 50 µm length, to 3:1 for zircons of 250 µm length; few zircons have L/B as high as 20. No statistical difference was observed between the rate of L/B increase with L between different rocks, nor was any correlation found as a function of the amount of extracted zircons, rock crystallinity or 18O. Similar L/B increase with length was cited to be an intrinsic property of igneous zircons worldwide (e.g. Larsen & Poldervaart, 1957) and is due to faster growth of pyramids vs prismatic faces (Vavra, 1994).
The results of CSD measurements are plotted as the natural log of linear population density vs average crystal length (Fig. A1) with 10 µm sample bins. For each sample, typically >350 zircons were measured. According to CSD theory, slopes taken from Fig. A1 are proportional to -1/G
, where G is a linear growth rate, and is the residence time (Cashman & Marsh, 1988). Assuming a constant growth rate G, this relationship can be used to determine the population residence time . Zircon growth rates are likely to be limited by diffusion of zirconium in the melt to the growing crystal (Watson & Harrison, 1983; Watson, 1996). Zircon dissolution experiments suggest growth rates on the order of 10-15–10-17 cm/s for silicic rocks (Watson, 1996). Given a rate of 10-15 cm/s and the range of slopes in Fig. A1, the residence time of zircon populations is of the order of 100–200 ky. Zircon residence times determined elsewhere using U–Th disequilibria also are found to be >100 ky (Reid et al., 1997). Zircon CSD in lavas erupted in adjacent localities without a significant time gap are remarkably similar. In particular, Dunraven Road Flow erupted after Canyon Flow (as evidenced by field relations and 40Ar/39Ar age, Table 1). These two flows have similar size distribution and estimated residence time, which suggests that no new zircon growth occurred in the magma chamber between the CF and DR eruptions. Progressively decaying nucleation rates, Ostwald ripening or ‘annealing’ (Cashman & Ferry, 1988) leads to the preferential growth of large zircon grains at the expense of small ones and these processes may account for the unusual positive slopes of CSD for smaller crystals (Fig. A1). Size distributions of zircons in low-18O lavas are similar to those in other rocks, as are their average zircon saturation temperatures and corresponding whole-rock concentrations of zirconium (Table 2). However, the zircon CSD trends shown in Fig. A1 do not necessarily indicate a single episode of growth for a single zircon population from a cooling magma. CSD are a time-integrated picture. We argue below that many zircons are xenocrysts and thus the size distribution is partially inherited from the source rock, and may include stages of zircon dissolution and overgrowth.
Imaging and trace element profiling
Yellowstone zircons, like many igneous zircons world-wide, exhibit euhedral morphologies but complex internal structures caused by trace element zonation (Fig. A2). Examination of external shapes of zircons by secondary electron microscopy reveals that many zircons form aggregates and show features indicative of zircon growth on, or together with, magnetite and apatite. Back-scattered, secondary electron, and CL images of polished zircon grain mounts reveal that most zircons have oscillatory zonation, often with sector zoning (80% of all crystals, Fig. A2a and b). Zircons with resorbed cores, where zoning in the cores is truncated by the zoning in the rims, constitute 5–20% (Fig. A2c). Fewer than 5% of grains have dark and irregularly shaped cores (Fig. A2d).
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Trace element profiles of Hf, U, Th, P, Y and Ce concentrations were made by electron microprobe across the four zircon types described above (Fig. A2). Trace elements show a clear correlation with CL: darker CL zones are richer in Y, Th and U. Different sector zones in a single crystal exhibit a factor of two variation in concentrations of trace elements, especially Y (Fig. A2b). Dividing the measured Y concentrations in zircons by the appropriate zircon–melt partition coefficient (Hinton & Upton, 1991) yields magmatic concentrations. For most zircon grains with <2x variations within different sectors in one grain, the reconstructed equilibrium magmatic concentrations of Y are within that measured in Yellowstone rhyolites, 40–60 ppm (Table 2).
Quantitative maps of trace element concentrations were made on several grains (Fig. A2f) (Fournelle et al., 2000). The results show complicated patterns of trace element distribution within many zircon crystals from low-18O lavas, but more regular patterns in LCT and HRT. In particular, wide variations in concentrations (Y and Th vary by 2·5 orders of magnitude in a single crystal, whereas Th/U ratio varies by a factor of three) cannot be ascribed to a single episode of zircon growth from a magma with equilibrium partitioning, given the published ranges of zircon–melt partition coefficients (Hinton & Upton, 1991). Rather, these zircons are derived from related plutonic (subvolcanic) rocks (e.g. Bacon et al., 2000) and record nearly complete crystallization, or a series of crystallization–dissolution events.
We conclude that trace element abundance and distribution in zircon, unlike isotopic ratio and age information, is of limited importance for making unambiguous petrogenetic inferences on the composition of coexisting melt. Trace elements in zircon are more dependent on local variations in concentration in the source, and kinetics of zircon crystallization–overgrowth (e.g. Hoskin & Ireland, 2000).
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ACKNOWLEDGEMENTS |
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We are grateful to Mike Spicuzza (Madison) and John Craven (Edinburgh) for support during isotope analyses; John Fournelle for help with SEM and CL imaging of quartz and electron microprobe analysis of zircons; Carrie Gilliam, Curtis Manley and Lena Bogolubova for help with sample collection; Kristin Greene for help with zircon separation; and Brian Hess for polishing thin sections and grain mounts. Mike Spicuzza, Liz King and Jade Star Lackey provided informal comments on the manuscript. William Peck helped with ion microprobe analyses. Reviews by Wes Hildreth, William Nash and Steve Balsley, and editorial suggestions of George Bergantz are gratefully acknowledged. This research was supported by the DOE (grant FGO2-93ER14389) and NSF (grant EAR96-28142).
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FOOTNOTES |
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*Corresponding author. Telephone: 608 262 7118. Fax: 608 262 0693. E-mail: inbindem{at}geology.wisc.edu
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REFERENCES |
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TOPABSTRACTINTRODUCTIONSAMPLE COLLECTION, PREPARATION...GEOLOGY AND PETROLOGY OF...OXYGEN ISOTOPE RESULTSDISCUSSIONAPPENDIX: ZIRCON IN YELLOWSTONE...REFERENCES |
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I. N. Bindeman Crystal sizes in evolving silicic magma chambers Geology, April 1, 2003; 31(4): 367 - 370. [Abstract] [Full Text] [PDF]
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J. W. Valley and J. W. Valley Oxygen Isotopes in Zircon Reviews in Mineralogy and Geochemistry, January 1, 2003; 53(1): 343 - 385. [Full Text] [PDF]
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 R. Macdonald and H. E. Belkin Compositional variation in minerals of the chevkinite group Mineralogical Magazine, December 1, 2002; 66(6): 1075 - 1098. [Abstract] [Full Text] [PDF]
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C.-S. Wei, Y.-F. Zheng, Z.-F. Zhao, and J. W. Valley Oxygen and neodymium isotope evidence for recycling of juvenile crust in northeast China Geology, April 1, 2002; 30(4): 375 - 378. [Abstract] [Full Text] [PDF]
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