Shallow storage of the explosive Earthquake Flat Pyroclastics magma body, Okataina Volcanic Center, Taupo Volcanic Zone, New Zealand: evidence from phase-equilibria experiments

Grant, Elizabeth R. G.; Blatter, Dawnika L.; Sisson, Thomas W.; Cooper, Kari M.

doi:10.1007/s00410-024-02151-y

Shallow storage of the explosive Earthquake Flat Pyroclastics magma body, Okataina Volcanic Center, Taupo Volcanic Zone, New Zealand: evidence from phase-equilibria experiments

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Published: 25 July 2024

Volume 179, article number 81, (2024)
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Shallow storage of the explosive Earthquake Flat Pyroclastics magma body, Okataina Volcanic Center, Taupo Volcanic Zone, New Zealand: evidence from phase-equilibria experiments

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Abstract

Rhyolitic tuffs range widely in their crystal contents from nearly aphyric to crystal-rich, and their crystal cargoes inform concepts of upper crustal magma reservoirs. The Earthquake Flat pyroclastics (Okataina Volcanic Center, Taupo Volcanic Zone, New Zealand) are 10 km³ of rhyolitic tuffs with abundant (~ 40 vol.%) plagioclase and quartz, minor biotite, hornblende, and orthopyroxene, and accessory Fe-Ti oxides, apatite, and zircon, set in high-silica rhyolitic glass. Major minerals form large, euhedral phenocrysts and abundant glomerocrysts with few disequilibrium textures excepting some faintly resorbed quartz. Plagioclase phenocrysts have thick rims of nearly constant composition near An₃₀, and hornblende is weakly zoned or unzoned. The abundant and texturally complex mineral assemblage contrasts with the nearby (~ 25 km), nearly synchronous, but more voluminous and crystal-moderate rhyolite tuffs from Rotoiti caldera. New H₂O-saturated phase-equilibria results on the erupted Earthquake Flat melt (glass) determine its co-saturation with the partial phenocryst assemblage of plagioclase, quartz, biotite, and Fe-Ti oxides at: 140 MPa, 755 ºC. These closely approximate the conditions of the pre-eruptive magma body assuming it was saturated with nearly pure H₂O and at an fO₂ of ~ Ni–NiO. Absence of hornblende and orthopyroxene from the synthesized assemblages may result from those minerals being in a peritectic reaction relation with melt to produce biotite, so they would not grow from the liquid used as starting material. Experimental results on Rotoiti rhyolite (Nicholls et al. 1992) show that the two bodies resided at similar pressures, temperatures, and fO₂s. Lower crystal abundance of the Rotoiti tuffs may result from slight compositional differences. We interpret that the Earthquake Flat pyroclastics were sourced from the crystal-rich periphery of a mushy reservoir system with the Rotoiti occupying a more melt-rich central location. Uncertain is whether this was a single intrusion zoned continuously in crystallinity, or discrete adjacent intrusions, but our results illustrate and quantify complexities of magma storage across relatively short distances.

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Introduction

The magmas that feed large rhyolitic eruptions are complex mixtures of melt, crystals, and gases with minerals that routinely have compositions, zoning, and ages inconsistent with having formed by the simple, progressive crystallization of a closed body of melt (Bachmann and Bergantz 2004; Befus and Gardner 2016; Caricchi and Blundy 2015; Cashman et al. 2017; Chamberlain et al. 2015; Ellis et al. 2014; Hildreth 2004; Hildreth and Mahood. 1986; Pichavant et al. 2007; Swallow et al. 2019). Instead, mineral interiors can have distinct compositional zones, some truncated by overgrown resorption surfaces, and even the rims of mineral grains need not be in equilibrium^{Footnote 1} with the melt that accompanied them upon eruption. These and other features point to varied and protracted pre-eruptive histories involving recharge, partial solidification — in some cases to advanced extents — remelting, and remobilization (Reid 2003, 2008), and the diverse crystal cargoes can impede inferring magmatic storage conditions (Ellis et al. 2014). One commonly endorsed scenario (the “mush model”) is that shallow silicic magma reservoirs reside over most of their spans of activity in highly crystalline states with small melt fractions, with eruptions being brought about by melt that segregates and accumulates as a transitory melt-rich zone or cap, or that, at times, the crystal mush is in some way partly remelted (or “defrosted,” Mahood 1990) locally or widely and thereby remobilized (Bachmann and Bergantz 2004, 2006; Cashman and Giordano 2014; Ellis et al. 2014; Hildreth 2004; Lipman et al. 1997). Melt segregation and defrosting are not mutually exclusive processes. Timespans over which shallow silicic magmas remain in eruptible melt-rich states are matters of active research, but estimates are 10³–10⁵ years (Bachmann and Bergantz 2004; Charlier et al. 2005; Cooper 2015, 2019; Reid and Vazquez 2017), considerably briefer than the overall lifespans of their respective magmatic systems.

Shallow silicic magmatic systems must be more complex and diverse than these simplified representations, with melt versus crystal abundances also differing with proximity to the system’s margins where heat is lost, and toward its sites of recharge where heat and replenishing melts are gained, as well as with the times elapsed since recharges, recharge volumes and compositions, reservoir depth and size, and other factors. Such complexities lead to multiple potential origins for the cognate crystals in silicic magmas including as true autocrysts formed by the progressive partial crystallization of a closed body of melt; as partially remelted relicts of a remobilized, formerly crystal-rich reservoir or domain; as grains entrained from highly crystalline regions by transiting melt; and as composite grains grown from multiple replenishing silicic melts. In all cases, if melt and crystals then remain together sufficiently long, mineral rims and melt can approach equilibrium. Discerning the processes that assembled a magma body and then brought it to eruption therefore requires examining its crystal assemblage and its state of equilibrium with coexisting melt.

Experimentally determined phase-equilibria assemblages can reveal important intensive parameters (P–T-H₂O-ƒO₂) of magma reservoirs (Rutherford et al. 1985; Scaillet and Evans 1999; Cadoux et al. 2014; Befus and Gardner 2016). We present new experimental phase-equilibria results for the Earthquake Flat (EQF) rhyolitic deposit (Okataina Volcanic Center, Taupo Volcanic Zone, New Zealand), which establish the storage conditions of the magma body immediately prior to eruption. The EQF crystal-rich ($\sim$ 40 vol.%) rhyolitic tuff provides an excellent example of a complicated assemblage of minerals that might have been entrained from different parts of the magma reservoir, including crystal-rich peripheral regions. Large ($\le 1 \text{mm})$, chemically evolved crystals (e.g., low An content across plagioclase rims), and abundant glomerocrysts suggest that the crystals had ample time to grow from evolved melt, to nucleate heterogeneously on one another forming glomerocrysts, or potentially to have been scavenged from highly crystalline regions.

These textural observations, alone, do not resolve whether the crystals in this eruption were scoured and entrained from the mushy regions or largely grew in equilibrium with the host melt. To address this problem, we performed phase-equilibria experiments using separated EQF glass as the starting material to establish the P–T conditions of the EQF magma immediately prior to eruption. The experimental results place the EQF magma storage conditions at 140 $\pm$ 10 MPa and 755 $\pm$ 5 ºC, if saturated with H₂O-rich vapor at $\Delta$ Ni–NiO = 0, based on the liquidus co-saturation of the natural melt with plagioclase, quartz, and biotite. These results and observations support a crystal-mush model, where the EQF crystals and melt were largely cogenetic, most crystals having grown from or equilibrated with the melt, rather than scoured from other portions of the system.

Geologic and petrologic background

The Taupo Volcanic Zone (TVZ), North Island of New Zealand, is the on-shore continuation of the Tonga-Kermadec arc and its spreading back-arc Havre Trough rift system. It is marked by 1.6 Myr of active rhyolitic volcanism and by high rifting rates (Wilson et al. 1995) (Fig. 1A). Back-arc rifting ranges from 15 mm/yr near the Bay of Plenty to 5 mm/yr near Lake Taupo and is partly controlled by the geometry of the adjacent subduction zone (Cole et al. 2014). The Pacific Plate subducts beneath the Australian Plate, but obliquity of the plate boundary increases southward as the Hikurangi Plateau subducts (Cole and Spinks 2009). The TVZ is defined by eight volcanic centers, the northernmost of which is the Okataina Volcanic Center (OVC). The most recent caldera-forming eruption from the OVC was the Rotoiti eruption at ~ 45 ka (Cole et al. 2014; Danišík et al. 2012; Gravley et al. 2016). The Rotoiti eruption produced > 100 km³ of moderately crystalline (15–25 vol.%) rhyolitic pyroclastic flow and fall deposits (Schmitz 1995; Davis 1985; Nairn 1972) (Fig. 1B). Shortly after the Rotoiti eruption, EQF erupted ~ 20–25 km to the southwest, producing 10 km³ of crystal-rich (25–45 vol.%), non-welded rhyolitic pyroclastic flow and fall deposits. The EQF tuffs consist overwhelmingly of juvenile pumiceous lapilli and ash; pumice are composed of filamentous and bubble-rich glass with dominant phenocrysts of plagioclase and quartz, minor-to-trace biotite, hornblende, and orthopyroxene, plus accessory apatite, zircon, and Fe-Ti oxides. Juvenile Rotoiti deposits differ from the EQF both by their lower abundance of crystals, by the lack of biotite and hornblende, and by the presence of cummingtonite, but the mineral assemblages are otherwise generally similar, although not strictly the same. Field relations suggest that the EQF erupted immediately following the Rotoiti as evinced by the EQF Rifle Range ash conformably overlying the reworked upper Rotoiti Unit 5 (Nairn and Kohn 1973). Furthermore, lack of paleosol development between the two units confirms a short duration between eruptions (Nairn and Kohn 1973). Subsequent eruptions from the OVC have obscured the Rotoiti vents; however, it has been inferred that the eruption originated from the NE–SW-trending Haroharo Linear Vent Zone in the OVC (Fig. 1B) based on pyroclastic distribution maps (Schmitz and Smith 2004). Lithic lag breccia exposed on the caldera rim suggests that vents were intra-caldera (Shane et al. 2005). EQF erupted from inferred NW–SE trending linear vents (Fig. 1B) along the SW corner of the extra-caldera OVC ring structure (Nairn 2002; Molloy et al. 2008). Despite the temporal and spatial proximity of these two eruptive bodies, and their similar compositions and crystal assemblages, subtle isotopic differences between the two deposits precludes their having derived from a homogenous magma that differed solely in extent of crystallization across ~ 20 km (Molloy et al. 2008; Schmitz and Smith 2004). Broad mineralogical similarities between the two eruptions suggest that the pre-eruptive storage conditions for the two bodies were likely similar; however, this has not previously been tested quantitatively.

Petrologic summary of the Rotoiti and Earthquake Flat rhyolites

The bulk of the Rotoiti tuffs are chemically homogenous, with the exception of the small proportion (< 10 vol.% of total erupted materials) of biotite-bearing deposits that top the Rotoiti stratigraphy (Davis 1985; Schmitz and Smith 2004; Shane et al. 2005). Schmitz (1995) and Shane et al. (2005) both reported a lack of compositional diversity throughout the Rotoiti stratigraphic section in terms of mineralogy and glass composition (glass SiO₂ = 77.8 $\pm$ 0.3 wt.%, K₂O = 3.34 $\pm$ 0.10 wt.%). The small-volume upper units differ in mineralogy by the appearance of biotite, increase in hornblende content, disappearance of cummingtonite, and by subtle changes in glass composition (glass SiO₂ = 77.4 $\pm$ 0.3 wt.%, K₂O = 4.07 $\pm$ 0.17 wt.%). Single-clast pumice chemical compositions for the entire Rotoiti section range from 71.8–75.9 wt.% SiO₂, and average 74.5 wt.% SiO₂ (Schmitz 1995). Charlier et al. (2003) identified weak zoning based on mineral contents throughout the stratigraphy. Some studies argued that the late, biotite-bearing, small-volume magma was sourced from an injection of the EQF magma into the Rotoiti reservoir (Schmitz 1995), while others interpreted that the biotite-bearing magma was sourced from some other evolved rhyolitic magma adjacent to the Rotoiti reservoir and not from the Earthquake Flat magma (Shane et al. 2005; Schmitz and Smith 2004).

The Earthquake Flat tuffs consist of 14 intercalated flow and fall deposits that, where well exposed, range from 0.5 to 7 m in thickness, but with no systematic stratigraphic changes in mineralogy, crystallinity, or bulk composition as seen in other large silicic tuff deposits such as the Bishop Tuff (Davis 1985). Davis (1985) argued that EQF shows weak chemical zoning, with MgO, total Fe as Fe₂O₃, and TiO₂ slightly enriched in the upper EQF deposits relative to the lower deposits, but with no obvious trend in CaO concentrations throughout the entire sequence. Conversely, Molloy et al. (2008) found no evidence for chemical zoning in the EQF deposits. Reports of crystallinity range from 23–34 vol.% (Davis 1985; Schmitz 1995), but Molloy et al. (2008) expanded this range to 25–45 vol.% crystallinity, which is consistent with our observations. Davis (1985) interpreted the presence of two distinguishable magma types: “type 1” magmas are represented by single-pumice SiO₂ = 73.1–75.1 wt.% and K₂O = 2.80–3.15 wt.%, and “type 2” pumices that are defined by SiO₂ = 71.0–72.3 wt.% and K₂O = 2.36–2.47 wt.%. By Davis’ classification, our samples are of “type 1” magmas (single-pumice SiO₂ = 73.9–74.4 wt.%, K₂O = 3.02–3.30 wt.%). Schmitz (1995) confirmed that “type 1” samples dominate the deposits, but that the relationship between the two magma types is both discontinuous and unclear. From these reports, and from the compositional information we present below, we conclude that our samples are representative of the majority of the EQF tuff deposits.

Methods

Starting Material

The EQF pumice sample employed herein was collected from non-welded tuff exposed along New Zealand State Highway 38 (− 38.31112º, 176.3791º, NZGD 1949). Ten to fifteen large (6–25 cm) irregularly shaped pumice clasts were collected from a poorly sorted, matrix-supported, massive flow deposit. The single largest (~ 25 cm) pumice fragment without discoloration or visible vapor-phase alteration was selected for all analyses presented here. Analyzed materials were prepared by scrubbing the pumice clast clean of organic debris, drying, then coarse-crushing with a stainless-steel mortar and pestle. An aliquot of the cleaned, crushed single pumice was retained for whole-rock geochemical analyses and the remainder of the crushed sample was sieved into > 2, 1–2, 0.5–1, 0.125–0.5, 0.063–0.125, and < 0.063 mm size fractions. The 0.063–0.125 and 0.125–0.5 mm samples were passed through a Frantz magnetic separator at increasing amperages to separate minerals from glass. Some quartz and plagioclase grains in EQF pumice contain glass, Fe-Ti oxides, zircon, and apatite inclusions, which have a significant net effect on the magnetic susceptibility of such grains. As a result, the glass-rich split from the Frantz magnetic separator retained significant quartz and plagioclase. To remove the quartz and plagioclase from the glass (0.125–0.5 mm size fraction), we twice passed the concentrated glass split through a lithium heteropolytungstate (LST) solution of known density, resulting in a glass concentrate with < 5 vol.% mineral grain fragments (plagioclase > quartz). The resulting glass concentrate was washed and sonicated repeatedly, first in deionized water, then in acetone, and then again in deionized water to remove residual LST. This concentrate of natural glass, rather than the bulk pumice, was employed for experimental syntheses. Sparse remaining phenocryst fragments were easily identified in the experimental syntheses by their irregular, angular shapes, overgrowths, and their compositions. Splits of the purified glass concentrate and of the bulk crushed pumice were analyzed for major- and trace-element concentrations by X-ray fluorescence (XRF) at Hamilton Analytical Laboratory, Hamilton College, Clinton, New York (Table 1). Volcanic glass is prone to hydration and alteration, which can be accompanied by alkali exchange. Typical TVZ fresh glass alkali concentrations are > 3 wt.% Na₂O and < 4.5 wt.% K₂O (Shane 2000). Vesicular pumice glasses from EQF and from Rotoiti have Na₂O and K₂O concentrations indistinguishable from those of those deposits’ respective melt inclusions (Fig. 2) indicating little-to-no alkali losses. Our pumice glass and glass concentrate materials (Table 1) are well within the thresholds of what is considered hydrated but otherwise fresh glass.

Table 1. Anhydrous-normalized chemical compositions of starting materials

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Use of separated natural glass, versus bulk pumice, for experimental syntheses

The primary objective of this work was to identify the temperature and pressure (depth) representative of the bulk of the EQF magma body immediately prior to its eruption, and this was accomplished by finding the conditions at which the natural melt (the erupted glass) is saturated on its liquidus with the dominant phenocrysts. The textures and mineral chemistry of the EQF pumices are consistent with the phenocrysts in a growth regime: quartz, plagioclase, and biotite grains are large and sub-to-euhedral, and their rim compositions are homogenous or nearly so across relatively large distances (> 200 μm). Rim compositions are also similar from grain to grain, and the subtle oscillatory zoning of the thick, sodic plagioclase rims lack prominent resorption intervals. These features are consistent with the minerals having grown approaching the time of eruption. Some quartz grains are faintly resorbed, but this could have resulted from decompression rather than from heating (Luth et al. 1964; Tuttle and Bowen 1958). Practical advantages of using the separated glass, rather than bulk pumice, include that the natural glass’s liquidus is identified readily, its composition along the liquidus is guaranteed to be that of the natural melt at the time of eruption, and the run products are uncomplicated by the presence of abundant fragments of phenocrysts, antecrysts, and glomerocrysts that can remain unreacted or thinly armored in syntheses on crushed bulk pumices at the relatively low temperatures of rhyolites. An alternative approach is to prepare a glass with the composition of the bulk pumice, but nucleation and undercooled growth of crystals can be problematic in attaining the approximately 40 vol.% crystallinity, representative of the EQF magma at its time of eruption. A drawback of employing separated natural glass is that any peritectic reactant minerals will not grow from the melt despite being stable, and so will be absent from synthesized assemblages.

Capsule configuration and ƒO₂ buffering

Experiments were performed in H₂O-pressurized cold-seal vessels employing the classic double-capsule configuration developed by Eugster (1957) where the sample plus H₂O were sealed in an inner capsule (herein, Au) and this plus fO₂ buffer (herein, Ni–NiO) and H₂O were sealed in an outer Au capsule. The configuration was modified, owing to the long durations of the syntheses and the propensity of Ni to alloy with Au at the investigated temperatures. Modifications included containing the Ni–NiO buffers in two crimped but unsealed Pt tubes within the sealed outer Au capsule, thereby preventing Ni from touching Au, and by accompanying each synthesis with an additional Pt buffer capsule packed with Ni, NiO, and H₂O to prolong fO₂ buffered conditions (Electronic Supplementary Materials (ESM) 1).

Internal buffer capsules were 2 mm OD (outer diameter) Pt tubes loaded with a Ni–NiO (1:1 by weight) powder mixture (Ni metal and NiO powders) and a 1 mm diameter, ¼ inch (~ 7 mm) length of nickel wire, such that experiments were buffered at Ni–NiO, chosen based on calculated ƒO₂ from Fe-Ti oxide (titanomagnetite-ilmenite) touching pairs in the EQF pumice as discussed subsequently. The internal buffer capsules were welded at one end, loaded with buffer, and the second capsule end was squeezed shut and folded over with pliers so that buffer was contained but exposed to the H₂O environment.

The inner capsule containing the sample consisted of a 3/8-inch (9.5 mm) length of 3 mm OD annealed Au tubing into which 15–20 mg of the powdered natural glass concentrate and 7 wt.% distilled H₂O were loaded, and this was then triple crimped and welded closed using a tungsten-inert-gas welder.

The internal buffer capsules and the inner sealed Au capsule containing the wet sample were inserted in the outer Au capsule constructed of 4 mm OD annealed Au, along with ~ 60 $\mu$L H₂O, and this was finally triple crimped and welded shut. The external buffer capsules were constructed of 5 mm OD annealed Pt, loaded with ~ 50 mg of Ni–NiO (1:1 by weight) powder and ~ 50 $\mu$L distilled H₂O. The weigh, heat (several minutes at 112 °C), reweigh technique was employed at each stage to confirm welds were complete, and capsules that failed this test were rebuilt.

All experiments were H₂O-saturated, which yields ~ 2.7–7 wt.% H₂O dissolved in the melt, varying with pressure (Ghiorso and Gualda 2015; Liu et al. 2005; Moore et al. 1998; Iacovino et al. 2021; Table 2). These H₂O concentrations are consistent with those dissolved in Rotoiti melt inclusions (CO₂ < 10 ppm (~ limit of detection), and H₂O $\le$ 5.1 wt.%, Johnson et al. 2011; and CO₂ = 210–960 ppm, and H₂O $\le$ 7.5 wt.%, Smith et al. 2010), and available EQF melt inclusions (H₂O = 3.4–7.5 wt.%, Smith et al. 2010) and discussed in more detail below.

Table 2. Experimental conditions

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Experimental Setup

All experiments employed H₂O-pressurized, externally heated cold-seal pressure vessels without filler rods, other than the supplementary buffer capsule, with the vessels oriented vertically (synthesis conditions listed in Table 2). Temperatures were measured and controlled with a type-K thermocouple inserted into a 2.3 cm long well in the base of each vessel that places the thermocouple junction adjacent to the base of the vessel’s bore. Temperatures were controlled accounting for small differences between the control thermocouple’s temperature and the internal temperature at the sample’s position as determined by prior calibrations at run temperatures employing an internal type-K thermocouple. Calibrations included profiling the vessel’s internal temperature with the vessel inserted at different depths into the furnace to locate the optimal conditions of minimum thermal gradient across the sample position and a minimum difference between internal versus external temperature. Four vessels were generally run simultaneously in adjacent furnaces, and all the hotspots were located within the bottom 1–2 cm of the vessel’s bore, with offsets between external control thermocouple and internal sample position no greater than $\Delta$ 3 °C (e.g., 723 vs. 726 °C in bottom 15 mm of furnace #4). The sample capsule was placed in the bottom of the vessel’s bore with the sample positioned at the base to ensure it was in the hotspot and adjacent to the control thermocouple. The supplementary buffer capsule was placed above the sample capsule, and the bore was filled with water prior to closing of the vessel, connecting it to the H₂O pressure system, and inserting it the appropriate distance into the cold furnace. The vessels were water-pressurized using an air-driven pump, and pressure was monitored with a Heise Bourdon tube gauge. The vessels were brought to pressure before heating to prevent the sample capsule from expanding and bursting or (and) sticking in the vessel. Once the vessels were securely pressurized, the furnaces were turned on and adjusted to run temperature. During heating, the pressure was monitored and adjusted to not exceed the intended run pressure by more than about 50 bars. Once the intended temperature was reached, the pressure was adjusted to the intended value. Both pressure and temperature were maintained over the experimental durations to within ± 1 °C and ± 10 bars, therefore, the uncertainties on pressure and temperature measurements are small: temperature uncertainties are ≤ 5 °C based on replicate calibrations of temperature offsets and hotspot gradients, and pressure uncertainties are on the order of several bars. Syntheses brought directly to pressure and temperature were held at intended conditions for 8–22 days, generally increasing with proximity to the solidus. Reversal experiments, designed to approach equilibrium from high temperature conditions, were brought to silicate superliquidus conditions for 6.5–6.75 days, then temperature and pressure were adjusted to the conditions of prior direct syntheses and held for 14 days (Table 2). After the intended synthesis durations, vessels were extracted from the furnaces and cooled by a jet of compressed air. No evidence of quench crystallization was observed (no microlites in glass, no graphic textures, no quench overgrowths on stable crystals discernable at optical and electron-beam resolutions). A magnet was used to check that the internal and supplementary buffer capsules retained Ni metal, and internal buffer capsules that occasionally tore open on capsule disassembly consistently showed yellow-green NiO, collectively showing that fO₂s were buffered for the synthesis durations. The inner capsules containing the samples were then ground to expose the glassy slug, epoxy impregnated, mounted, and polished for analytical work.

Electron Microprobe (EMP)

The synthesized products and natural EQF pumice phases (polished thin sections) were analyzed on a five-spectrometer JEOL JXA-8530 F + electron microprobe at the U.S. Geological Survey, Menlo Park, California, using wavelength-dispersive methods. Instrument conditions were 15 kV accelerating potential, a 10 nanoamp, 2–5 µm diameter beam for minerals, or a 2 nanoamp, 5–20 µm diameter beam for glass. For minerals, elements were measured for 10–40 s, whereas in glass, Na and K were measured first for 10 s, and the other elements for 20–30 s. Backgrounds were determined every 6th point, and background-corrected count rates were reduced to concentrations with the JEOL-proprietary version of the ZAF method, standardized with natural and synthetic materials. Primary and secondary standards were analyzed as unknowns periodically during analysis sessions to monitor and post-process for drift. Analytical results, uncertainties, and instrument conditions are reported in Tables 3 and 4, analytical results of unknown and standards are both available in ESM 2 and 4.

Table 3 Average chemical compositions of minerals and glass in EQF pumice (oxide wt.%) measured by EMP

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Table 4 Average chemical composition (oxide wt.%) of synthesized glass, plagioclase (Plg), and alkali feldspar (Kfsp) measured by EMP

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ATR-FTIR

Dissolved H₂O concentrations of the experimental glasses, EQF pumice glass, EQF quartz-hosted embayment glasses, and EQF melt inclusions were determined by the attenuated total reflectance (ATR) method (Lowenstern and Pitcher 2013) using a germanium crystal on singly-polished surfaces with the Nicolet iN10 Fourier transform infrared imaging microscope (FTIR) at the U.S. Geological Survey in Menlo Park, California, with liquid-N₂-cooled MCT-B detector. This method limits light penetration to ~ 1 µm below the surface, allowing avoidance of crystals. The aperture size (30–150 µm) was adjusted according to the size of the glass pools in each sample and spectra were collected from 4,000 to 1,200 cm⁻¹ with 256 scans for each spot and spectral resolution of 8–16 cm⁻¹. A reference spectrum was collected in air prior to each sample measurement and a straight baseline was used to determine the height of the 3,450 cm⁻¹ band. An internal dacite glass standard with known H₂O concentration was analyzed periodically throughout each analytical session to verify that the ATR crystal and sample were oriented for optimal contact. Several spots were analyzed in distributed melt pools across each experimental or EQF pumice glass sample, and maps were collected for EQF melt inclusions (results not discussed here but available in ESM 3). Experimental products with high crystallinities prohibit accurate and precise H₂O measurements as there is not sufficient area for the germanium crystal to make complete contact without interfering with synthesized grains. Highly crystalline experiments yield H₂O concentrations that differ by up to 1.5 wt.% H₂O and are therefore not reported. For low crystallinity experiments, complete contact between glassy pools and germanium crystal yield H₂O concentration that vary no more than 0.6 wt.% H₂O within a single experiment (Table 2, ESM 3); H₂O concentrations of EQF melt inclusions are not reported in Table 2 as they presented analytical challenges beyond the scope of this study; however, these results are included in ESM 1 and 3.

Results

Characteristics of EQF pumices

The EQF pyroclastics are dominantly pumice lapilli and pumiceous ash with abundant crystals set in microlite-free highly vesicular glass. Compositions are presented for EQF pumice glass, glass concentrate starting material, and the average of seven EQF single pumices in Table 1, compositions of EQF pumice minerals are presented in Table 3, and those of experimentally synthesized minerals are presented in Table 4. The natural EQF mineral assemblage is characterized (all in vol.%) by plagioclase (16–23%) > quartz (4–11%) > biotite (1–3%) > hornblende (1–2%) > oxides (< 1.5%) > orthopyroxene (< 0.5%) > trace phases (clinopyroxene, apatite, zircon, pyrrhotite) (Fig. 3A–H) (Molloy et al. 2008). We note that the EQF pumice phase proportions are not directly comparable to our experimental results as we used the glass concentrate as a starting material. As presented previously, total crystallinity spans 25–45 vol.%, with the majority of samples near the upper end of that range (Molloy et al. 2008; and this study). The pumice glass and glass concentrate starting material share similar compositions, both have > 77 wt.% SiO₂, and identical concentrations of 12.5 wt.% Al₂O₃, but EMP alkali values are ~ 5 relative % lower than XRF values, either due to minor alkali migration during EMP analyses or to differences in standardization of the two methods. Molloy et al. (2008) reported 205 individual point analyses of EQF glasses. Of these, six have distinctly high K₂O concentrations (> 5 wt.%, normalized anhydrous), one has aberrantly high MgO (1.8 wt.%), possibly reflecting a transcription error, and one has slightly low CaO (< 0.65 wt.%). The remaining 197 analyses are highly homogeneous high-SiO₂ rhyolitic compositions yielding low standard deviations on concentrations of oxides sensitive to rhyolitic differentiation (e.g., average CaO = 0.80 $\pm$ 0.06 wt.%, average total Fe as FeO = 0.88 $\pm$ 0.22 wt.%, average K₂O = 4.53 $\pm$ 0.10 wt.%) and are effectively identical to pumice glass compositions determined in the present study (Fig. 2). The samples analyzed by Molloy et al. (2008) were collected by those authors and by Davis (1985) from locations throughout the EQF stratigraphy; they present data from four samples from the upper stratigraphy, six samples from the middle stratigraphy, two samples from the lower stratigraphy, and six samples from the lowest/basal stratigraphy. Collectively, these results indicate that, excepting a trace K-rich component, EQF melt was uniform in composition across the ~ 10 km³ of erupted magma. Because of high phenocryst contents, the average single-pumice composition is slightly less enriched in SiO₂ and K₂O (74.2 $\pm$ 0.2 wt.% SiO₂ and 3.15 $\pm$ 0.11 wt.% K₂O) than the glasses and has significantly higher concentrations of total Fe as FeO, MgO, and CaO. H₂O concentrations measured by ATR-FTIR in pumice glasses range from 1.4–1.9 wt.% H₂O (average 1.7 $\pm$ 0.1 wt.% H₂O) some of which could be from post-eruptive hydration but this was not assessed.

Quartz

Quartz grains range in size from 0.25 to > 2 mm, although many smaller quartz grains in thin section and in mineral separates appear to be shattered pieces of larger grains. Euhedral, bipyramidal quartz grains are commonly 1–2 mm across, although some rounded grains are present (Fig. 3A). The highest proportion of quartz grains are > 2 mm and tend to be anhedral shattered pieces of larger grains, suggesting that even larger grains existed prior to eruption. Almost all quartz grains have abundant inclusions of glass, apatite, zircon, or (and) opaque oxides; the exteriors of the grains have a pockmarked appearance (ESM 1).

Plagioclase

Most plagioclase grains are < 1 mm and occur as poly- and mono-mineralic glomerocrysts (Fig. 3B–D), although there is a less abundant population of individual plagioclase grains that are as large 2 mm. Most plagioclase crystals are subhedral and are commonly twinned. Many plagioclase crystals (Fig. 4) have narrow, resorbed, calcic cores, to An₆₇, overgrown over short distances by plagioclase zoned normally to ~ An₂₅, within thick sodic rims continuing that nearly constant composition to their margins (mole percent anorthite, An, is 100Ca/[Ca + Na + K]). Grain rims generally show subtle oscillatory zoning, some with rare, thin, calcic intervals, while most cores show oscillatory zoning and patchy zoning owing to resorption (Fig. 4).

Glomerocrysts

We identify two populations of glomerocrysts: (1) “large” (mineral grains 0.5–1 mm) plagioclase-dominated clusters which also contain quartz and mafic minerals (hornblende, biotite, oxides, and rare orthopyroxene), and which are easily identifiable in thin section, but which are not well preserved in mineral separates (ESM 1); and (2) “small” (mineral grains < 0.25 mm) clusters dominated by plagioclase and hornblende (ESM 1). The “small” clusters are rarely observed in thin section but are relatively common in mineral separates.

Biotite

Biotite grains are large (0.5–2 mm) and euhedral, commonly forming thick books ($\le$ 1.5 mm) with characteristic hexagonal habits (Fig. 3F). Most biotite grains occur individually within vesicular glass, but some occur in the “large” glomerocrystic clusters. Biotite cores and rims have nearly homogenous compositions with Mg# of cores = 47–51 and those of rims = 46–51 (Mg# is 100 Mg/[Mg + Fe], molar, employing total Fe).

Hornblende

Amphibole grains are large (0.5–2 mm), euhedral, and commonly prismatic. Pleochroism ranges from light green to dark brown. They occur as sparse individual grains within vesicular glass and abundantly in both populations of glomerocrysts (Fig. 3B, 3C, 3E; ESM 1). The amphiboles are magnesio-hornblende (Hawthorne 1981) approximating their ferric iron concentrations as the average of all-ferrous and 13eCNK methods. Hornblende compositions are relatively homogenous with cores ranging from Mg # = 56–57 and rims ranging from Mg # = 55–58. Molloy et al. (2008) reported transect and spot analyses of EQF hornblendes and an overall tendency for Al₂O₃ concentrations to increase modestly from grain cores to rims, an attribute those authors interpret as evidence that temperatures increased as the hornblendes grew. Excluding one outlier low-Al₂O₃ analysis, the average Al₂O₃ concentrations of hornblende rims (within 100 μm of grain edges), and their standard deviations, are similar for that and the present study (6.49 $\pm$ 0.37 and 6.35 $\pm$ 0.48 wt.%, respectively). No cummingtonite has been observed in any EQF samples.

Fe-Ti Oxides

Titanomagnetite and ilmenite are abundant as individual grains in vesicular glass, in glomerocrysts, and as inclusions in other minerals (Fig. 3G, 3H). Two populations of euhedral titanomagnetite-ilmenite pairs were examined for temperature and fO₂ determinations: 1) touching pairs in contact with glass and 2) euhedral grains within glass separated by no more than 25 μm. Both touching pairs and pairs separated by glass (n = 8) yield temperatures of 694–722 °C (average 703 °C) by the Ghiorso and Evans (2008) calibration, 718–746 °C (average 733 °C) by Lepage (“ILMAT”, 2003), and 708–731 °C (average 714 °C) by Andersen et al. (1993) (“QUILF”) (Fig. 7). Average oxide compositions are X_ulsp mol = 0.275 and X_ilm mol = 0.926 (n = 8). All grains pass the Mg/Mn partitioning test of Bacon and Hirschmann (1988) within the 2 $\sigma$ uncertainty envelope. Titanomagnetite-ilmenite pairs yield ƒO₂ estimates between $\Delta$ Ni–NiO = 0 to − 1, which guided our choice of experimental ƒO₂ buffering (ESM 1). Trellis and lamellar intergrowths of titanomagnetite and ilmenite are also present and result from oxidation (Buddington and Lindsley 1964), probably during deposit cooling, and are excluded from consideration. Molloy et al. (2008) reported compositions of 29 titanomagnetite-ilmenite pairs that, employing the Ghiorso and Evans (2008) calibration, yield a wider range and higher average temperature: 702–805 °C (average 748 $\pm$ 24 °C), including 705–778 °C for pairs within a single pumice. The principal difference is that TiO₂ concentrations of the Molloy et al. (2008) Fe-Ti oxides average ~ 5 relative % lower than those in the present study; average Fe-Ti oxide compositions are otherwise effectively identical in the two studies. The disparity in TiO₂ concentrations, and therefore in derived temperatures, may result from a standardization difference, since mineral and glass compositions were determined by Molloy et al. (2008) with energy-dispersive (EDS) methods on a scanning electron microscope that are less precise and accurate than the wavelength-dispersive electron-microprobe methods employed herein.

Orthopyroxene

Rare orthopyroxene microphenocrysts are present as euhedral to anhedral grains and range from 0.1–0.25 mm with a range of textures (Fig. 3G). Common orthopyroxene textures include ragged edges, broken grains, and sieved cores. Compositionally, orthopyroxene grains are nearly identical across cores (En₄₈) and rims (En₅₀) (En = 100 Mg/[Mg + Fe + Ca], molar, employing total Fe).

Inclusions in ferromagnesian phases

All mafic mineral phases contain diverse inclusions. Hornblende and biotite contain abundant inclusions of glass, apatite, zircon, and titanomagnetite-ilmenite. Notably, we have observed a hornblende grain with a clinopyroxenitic core (Fig. 3E), and biotite grains with plagioclase inclusions (Fig. 3F). Titanomagnetite grains have inclusions of glass, apatite, zircon, and pyrrhotite (Fig. 3H). Orthopyroxene grains have the fewest inclusions but still contain devitrified glass, titanomagnetite, and apatite.

Melt inclusions and embayments

Melt inclusions are present in nearly all mineral phases. Those in quartz and plagioclase range from glassy to fully devitrified, whereas those in mafic minerals are almost all devitrified or contain significant post-entrapment crystals grown along host-mineral boundaries. In quartz, there are both melt inclusions and melt embayments. The melt inclusions are large ($\le$ 200 μm), rectangular in cross section, and are partially or entirely devitrified (ESM 1). We attempted reconnaissance analyses of H₂O concentrations in the quartz-hosted melt inclusions by ATR-FTIR and obtained diverse results from 0.8 to 6.0 wt.% H₂O, with an average of 3.9 $\pm$ 1.7 wt.% H₂O. The higher H₂O concentrations may have been influenced by crystallization of anhydrous minerals, thereby concentrating H₂O in the melt, whereas the low concentrations could result from inadvertent analysis of inclusions that were not fully sealed and so partly degassed upon eruption; post-eruptive hydration of incompletely sealed inclusions is also possible. A focused study of acid-treated quartz grains to isolate truly sealed inclusions, coupled with experimental re-homogenization of those so identified, would be necessary to better understand the dissolved H₂O concentrations at the time of entrapment. The quartz–hosted embayments are elongate and almost always glassy (ESM 1) with H₂O concentrations ranging from 0.6–1.9 wt.% H₂O (average 1.2 $\pm \, 0.3$ wt.% H₂O) and SiO₂ between 77.1 and 77.9 wt.% (average 77.7 wt.% SiO₂), both of which are similar to the host glass (Table 3). Since the matrix glass and the melt embayments were exposed to the surficial environment for ca. 45 kyrs, hydration is likely to have raised H₂O concentrations of both. Assessing hydration would require measurements by transmission FTIR to distinguish molecular H₂O from that dissolved as hydroxyl groups, or by measurements of stable isotopes of glass. Plagioclase-hosted melt inclusions are an order of magnitude smaller (10 s of μm in size) than the quartz-hosted melt inclusions and are restricted to the interior zones of the grains (ESM 1). We observe hundreds of small melt inclusions clustered into arrays, just outside of heavily resorbed zones. Resorbed surfaces are effective substrates for new melt inclusion development and entrapment (Lowenstern and Thompson 1995), which are then efficiently captured within the grain by subsequent plagioclase growth (ESM 1). We were unable to measure H₂O concentrations in plagioclase-hosted melt inclusions due to size limitations, but other studies report H₂O concentrations (3.4–7.5 wt.% H₂O, Smith et al. 2010), which are further supported by volatile solubility models.

Experimental results

The unaltered EQF pumice glass composition (Tables 1 and 3) was chosen and carefully refined as the starting material for this experimental study as it is representative of the melt composition present in the majority of EQF magma immediately prior to eruption. The goal of the approach is to identify the conditions at which that melt is multiply saturated on its liquidus with the major minerals present in the natural pumice, and with the natural compositions for those minerals capable of solid solution. By comparing our experimental results to the EQF pumice, we can resolve the pressure and temperature, assuming H₂O saturation and oxygen fugacity, of the magma body immediately prior to eruption. We present experimental results, including phase assemblages and proportions (where feasible) in Table 2 and phase compositions in Table 4. We could not adequately quantify the phase compositions or proportions for low pressure/temperature experiments #16 and #43 due to small crystals (< 1 μm) and melt pools ($<$ 10 μm). As our experiments were loaded with H₂O in amounts exceeding saturation, and the runs products contain vapor bubbles, we estimated H₂O concentrations dissolved in melt (glass) as: H₂O by difference from microprobe analysis (Table 2), H₂O measured directly by ATR-FTIR (Table 2 and ESM 3), and H₂O estimated from volatile solubility models assuming saturation with pure H₂O (Ghiorso and Gualda 2015; Liu et al. 2005; Moore et al. 1998; Iacovino et al. 2021) (Table 2).

Approach to equilibrium

The combination of long experimental durations, precise temperature and pressure calibrations and control, successful ƒO₂ buffering ($\Delta$ Ni–NiO = 0 with all buffer components present after a given experiment was completed), H₂O saturation, and consistent phase appearances and shifts in mineral composition point to a close approach to equilibrium conditions. Nineteen experiments were brought directly to experimental conditions, and to confirm the stability limits for quartz and plagioclase, four reversal experiments were held above the liquidus and then brought to the conditions of prior direct syntheses (Table 2). These reversals produced the same phase assemblages and abundances as the direct-to-temperature experiments, thereby confirming quartz and plagioclase stability limits. Phase compositions and abundances change systematically with temperature and pressure as shown by plagioclase anorthite content, with isopleths subparallel to the plagioclase stability curve (Fig. 5). All of our synthesis conditions were above the H₂O-saturated haplogranitic solidus, although two experiments (#16 at 100 MPa/720 °C, and #43 at 50 MPa/750 °C) were highly crystalline with low melt abundances, suggesting that these conditions were very near the solidus (> 95 vol.% crystallinity).

Phase stability

Mineral stability curves are plotted relative to phase appearance at specific synthesis conditions (Fig. 5). Five of the experimental charges were sufficiently above the liquidus such that they did not saturate with any of the major silicate minerals despite the long duration run times. These syntheses (12, 14, 26, 27, and 31) contain only melt (glass), rare relict apatite and zircon, and H₂O vapor (vesicles). Plagioclase is the highest temperature silicate mineral at low pressure but is replaced by quartz and biotite at pressures $\ge$ 140 MPa. K-feldspar is the lowest temperature silicate mineral and the last silicate mineral to join the crystallizing mineral assemblage upon cooling. At high pressure ($\ge$ 200 MPa) and low temperature (< 710 ºC), plagioclase overtakes K-feldspar as the lowest temperature silicate mineral. For the investigated range in pressure and temperature conditions, the mineral appearance curves have negative slopes. Plagioclase has the shallowest negative slope, approximately − 1.25 °C/MPa. K-feldspar has a steep negative slope, approximately − 5.8 °C/MPa, between 250 and 75 MPa, but shallows to a slope of − 0.5 °C/MPa between 75 and 50 MPa. Quartz has a slope of − 3 °C/MPa between 225 and 75 MPa, and shallows to a slope of − 2.5 °C/MPa. Biotite has a slope of − 8.75 °C/MPa. Below 75 MPa the biotite curve is dashed to approximate the well-established behavior of biotite at low pressures in which the saturation curve transitions to a positive slope (Piwinskii and Wyllie 1968; Gardner et al. 2014; Waters and Andrews 2016; First et al. 2021). The quartz, biotite, and plagioclase saturation curves intersect at ~ 140 MPa and ~ 755 °C. Given the spacing in our synthesis conditions (50 MPa and 20 °C steps), there is a small amount of imprecision in the exact conditions of the quartz–plagioclase–biotite saturation curve intersection point, but the steep slope defined by the biotite-in curve anchors the intersection to 140 $\pm$ 10 MPa and 755 $\pm$ 5 °C.

Experimental phase descriptions

Characteristic images of experimental phases are presented in Fig. 6. Table 4 presents the compositions of the experimentally synthesized glass and feldspars. Synthesized quartz grains have euhedral, sharp rims, equant habits, and range from 1–10 μm. Quartz grains are nearly indistinguishable from the glass in back-scattered electron (BSE) images except for their relatively bright rims (an edge-effect artifact in high-contrast BSE imaging). Some relict plagioclase grain fragments from the glass concentrate starting material are present in the experimental products and are distinguished by their angular shapes and sizes atypically larger than newly synthesized plagioclase grains. Newly synthesized plagioclase grains are euhedral to subhedral with sharp boundaries, where grown as individual grains, and new grains commonly nucleated on and/or overgrew relict plagioclase grains (< 100 μm) (Fig. 6B); minor numbers of plagioclase grains are skeletal (Fig. 6B). Traces of relict plagioclase grains are present above the depicted plagioclase saturation curve in experiments brought directly to pressure and temperature and are distinguished as unstable by their rounded edges, dissolved portions of grains, and lack of overgrowths or coexisting homogeneously nucleated plagioclase grains. Such relict grains are absent from reversal experiments initially brought to superliquidus conditions, confirming the location of the plagioclase saturation curve. K-feldspar grains are euhedral with sharp grain edges and equant to tabular habits ranging from 1 to 20 μm in length. In several runs, K-feldspar overgrew the rims of relict plagioclase grains (Fig. 6A) and, in some instances, encase an entire relict grain (Fig. 6B). Individual K-feldspar grains also commonly encase other co-precipitating phases, and in higher pressure experiments ($\ge$ 200 MPa) they exhibit sector zoning indicating rapid growth (Fig. 6C). Compositions of synthesized plagioclase and K-feldspar range from An_13-29, and Or_49-73(Or is 100K/[Ca + Na + K + Ba]) respectively, with plagioclase grains decreasing in anorthite component approaching the solidus (Fig. 5). K-feldspar in melt-poor, near-solidus experiments (#43 at 50 MPa/740 °C and #16 at 100 MPa/720 °C) has compositions of Or_{55 & 49}, whereas K-feldpsar in somewhat cooler but higher pressure and more melt-rich experiments (#11 at 200 MPa/700 °C and #42 at 250 MPa/680 °C) are Or_{73 & 72}. Biotite grains form thin laths < 5 μm in length, and < 1 μm in width (Fig. 6A–C, E) and were too small to analyze quantitatively by EMP; however, biotite was confirmed by EDS, as well as by their characteristic tendency to erode preferentially during polishing. Mineralogy of the natural EQF pumices indicate that hornblende and orthopyroxene should also be members of the equilibrium assemblage, but we did not identify either mineral in our experiments despite extensive, high-magnification examination by EDS. Absence of hornblende and pyroxene are discussed subsequently. Fe-Ti oxide grains are < 1 μm, too small for quantitative analyses, but their identities as titanomagnetite and ilmenite were confirmed by EDS. Since magnetic methods were included in purifying the natural glass, these Fe-Ti oxides are unlikely to be relicts from the natural pumice. Traces of relict apatite and rare relict zircon are also observed in some of the run products, and micron-scale monazite was found in the four most crystalline experiments (#16 at 100 MPa/720 °C, #43 at 50 MPa/750 °C, #42 at 250 MPa/680 °C, #11 at 200 MPa/700 °C). As monazite is not part of the EQF pumice mineral assemblage, and it only appeared in run products much more crystallized than the natural pumice, the monazite grains were synthesized and are not relicts. Additionally, upon finding monazite and subsequently examining phases with high backscatter brightness at high magnification, we found a few exceptionally sparse, sub-micron specks of a tungsten-rich phase in a few experimental products. These lack a Ca peak on EDS spectra, so they are not scheelite. They are probably WO₃ derived from traces of the lithium heteropolytungstate employed in separating the glass. Their exceptionally low abundance, along with the saturation with H₂O that would absorb lithium, are evidence that the trace contaminant would not have affected phase stability within the resolution of our study.

Phase proportions

Phase proportions were calculated by linear least-squares multiple regression (Table 2) mass-balancing the composition of the average super-liquidus glass with those of coexisting minerals and melt all measured by electron microprobe. H₂O was included as a phase and component in the linear regressions since experimental capsules were loaded with more than sufficient H₂O to saturate the melt at run conditions, and the runs were pressurized with water throughout their durations. Four near-liquidus experiments (#s 28, 29, 40, and 41) returned inconsistent phase-proportion results but superior results were obtained using the XRF-determined glass as the bulk composition (ESM 1). Attempts to measure phase proportions by processing X-ray elemental maps failed owing to an inability to set thresholds that consistently differentiated between glass, quartz, and feldspars. Crystallinity is low (< 10 wt.%) for about 40–60 °C below the silicate liquidus and then increases sharply once K-feldspar joins the assemblage. For example, syntheses 18 and 16 differed only by 20 °C (740 vs. 720 °C at 100 MPa), yet the crystallinity of their products differ by 73 wt.% (6 and 79 wt.% crystals, respectively). The abrupt increase in silicate minerals (plagioclase, K-feldspar, quartz) approaching the solidus excludes biotite. For all syntheses, biotite was $\le$ 1 wt.% of total products.

Discussion

Experimental studies of high-silica rhyolitic compositions are relatively few in comparison to their more mafic counterparts. Slow diffusivity in silicic melts inhibits nucleation and crystal growth, requiring long experimental durations which can be difficult to sustain, especially with appropriate ƒO₂ buffering. Nucleation and crystal-growth issues in such experiments extend to the analytical realm where small crystal sizes hinder reliably analyzing phase compositions. Although challenging, such studies are important because petrologic tools such as thermometers, barometers, and thermodynamic models are calibrated on experimental results, and limited work in high-silica rhyolites could lead to inaccurate extrapolations. To improve our fundamental understanding of these potentially highly explosive systems and to test and calibrate models, it is important to produce quality high-silica rhyolitic experimental data. Our experimental results determine the co-saturation conditions of quartz, plagioclase, and biotite, all of which grew in equilibrium between 75–200 MPa and 710–760 ºC. Notably, the three stability curves intersect on the natural melt’s H₂O-saturated liquidus at 140 $\pm$ 10 MPa and 755 $\pm$ 5 ºC with minimal uncertainty since the three curves are well defined by experimental constraints and differing pressure–temperature slopes (Fig. 5). Despite the EQF tuff’s large erupted volume, its melt was nearly homogenous, matching that employed for the experiments, and all the EQF rhyolitic tuffs contain at minimum quartz, sodic plagioclase, biotite, and Fe-Ti oxides. Collectively, these observations and results indicate that the experimentally determined pressure–temperature conditions are representative of those that prevailed across much of the immediately pre-eruptive EQF magma reservoir. Based on these findings, we explore how these conditions compare to those estimated from other petrologic methods and models and consider the implications of these storage conditions for the assembly and eruption of the EQF magmatic system.

Thermometry

Fe-Ti oxide pairs in the EQF pumice samples provide estimates of pre-eruptive temperatures. Fe-Ti oxide compositions are sensitive to changes in temperature and oxygen fugacity and, due to rapid elemental diffusion at magmatic temperatures, can reset their recorded temperatures to conditions shortly prior to eruption; times to re-equilibrate are by some estimations on the order of days to months (Devine 2003; Hammond and Taylor 1982; Venezky and Rutherford 1999). Figure 7 illustrates the range in the temperatures calculated from five different geothermometers, which should caution against overinterpreting the accuracy of the results given the relatively small temperature range (694–761 ºC). While the precision is relatively good for the presented thermometry, the accuracy is less well understood as demonstrated by the dispersion in each set of calculated thermometry data, which itself is a function of the calibration of each thermometer.

The basic observation (Fig. 7) is that temperatures derived with various calibrations of the Fe-Ti oxidation-exchange reaction are on average lower than that determined experimentally by tens (Powell and Powell 1977; Andersen and Lindsley 1985) to a few tens of degrees (Spencer and Lindsley 1981; Andersen et al. 1993), to as much as 40–50 degrees (Ghiorso and Evans 2008). The spread in derived temperatures is also greater for the newer calibrations. Higher temperatures can be derived with the Fe–Mg exchange reaction, and although Ghiorso and Evans (2008) interpreted that such a divergence in temperatures may reflect disequilibrium within a magma body, Mg concentrations are low in the Fe-Ti oxides of silicic magmas, including those of EQF, hindering their accurate and precise analysis. Commencing with Buddington and Lindsley (1964), experimental studies have focused mainly on the Fe-Ti oxidation-exchange reaction owing to those elements’ higher concentrations and that reaction’s ability to provide information on oxidation state, so the Fe-Ti oxidation-exchange reaction is also more solidly calibrated. For these reasons, most published studies rely on the Fe-Ti results, as do we herein. As noted previously, titanomagnetite and ilmenite compositions presented by Molloy et al. (2008) for the EQF have TiO₂ concentrations about 5 relative % lower than those determined herein and yield temperatures closer to the experimental determination. This highlights the sensitivity of Fe-Ti oxide temperature determinations to the accuracy of oxide analyses, probably also in the calibrating experiments. Given these factors, we refrain from assigning geological significance to the temperature difference between our experimental determination for the EQF magma and those provided by the various Fe-Ti oxide models.

Molina et al. (2021) presented a new amphibole–plagioclase geothermometer, which updates previous iterations, such as the plagioclase–amphibole model of Holland and Blundy (1994). Using the pressure-independent Molina et al. (2021) expression, EQF samples yield temperatures averaging 703 and 696 °C for amphibole rim compositions of this study and Molloy et al. (2008), respectively (Fig. 8). Molina et al. (2021) stated the precision of the geothermometer as approximately ± 50 °C, so the calculated temperatures do not strictly differ from the experimental determination of 755 ± 5 °C. A further caution is that only 15 of the 203 experimental amphibole–plagioclase pairs used to calibrate and validate the geothermometer have plagioclase ≤ An₃₀, similar to plagioclase rims in the EQF and many other rhyolites, so application to rhyolites is at the compositional limit of the geothermometer’s calibration. Although we did not recognize amphibole in our synthesis products, hornblende grains in the EQF pumice samples are large, euhedral, and lack breakdown rims so they were probably stable up to the time of eruption as part of the equilibrium assemblage. Molloy et al. (2008) documented modest rim-ward increases in Al₂O₃ concentrations of some EQF amphibole phenocrysts that those authors interpret as evidence that magmatic temperatures increased as the amphiboles grew. Perhaps, but the lack of research on amphibole chemistry (especially under different oxygen fugacity conditions) urges circumspection. A robust study quantifying how tetrahedral alumina behaves at different P–T-ƒO₂ conditions is necessary in order to interpret such slight gradients in Al₂O₃ concentration. The Putirka (2016) amphibole–liquid geothermometer estimates the temperatures at which amphiboles may have been in equilibrium with host melt. Using the Putirka (2016) geothermometer, amphibole rim and glass compositions from this study and from Molloy et al. (2008) yield apparent equilibration temperatures averaging 755 ± 53 and 751 ± 53 °C, respectively. Again, within uncertainties, these overlap with amphibole–plagioclase and the directly determined experimental temperatures for the EQF magma body (Fig. 8). While the EQF magma body may have increased in temperature approaching its eruption (Molloy et al. 2008), the uniform composition of its erupted melt and the consistency of its phenocryst assemblage are evidence that temperatures were at or close to the experimentally determined value of 755 °C across the wide portions of the magma body that erupted.

Barometry

Melt inclusion barometry effectively provides the pressure that an inclusion of melt was entrapped within its crystal host, contingent on the melt having been saturated with an H₂O–CO₂ vapor; such pressures can be converted to approximate depths employing a mean density for overlying magma and roof rocks. These depths can be those of magma accumulation and storage before eruption, although complexities include the melt may not have been vapor saturated, some melt inclusions may have entrapped deeper than the site of storage, other inclusion-bearing crystals may have been entrained from shallower levels during eruption, as well as the prosaic issue that incompletely sealed inclusions can go unrecognized and if analyzed give aberrant low pressures owing to eruptive partial degassing. Nevertheless, inclusion-derived pressures are plausible proxies for those of magma storage if melt inclusions are present in late-crystallizing quartz and feldspar crystals, if compositions of included melts are similar to that of co-erupted host glass, and if the inclusions give consistent results. There are very few melt-inclusion-derived volatile concentration determinations for EQF, but abundant determinations are available for the nearby Rotoiti system, allowing assessment of whether magmas of the two systems resided at similar depths prior to eruption. Using the Liu et al. (2005) solubility model, a study by Johnson et al. (2013) derived Rotoiti entrapment pressures of 76–147 MPa. A similar study by Bégué et al. (2014) used the rhyoliteMELTS geobarometer of Gualda and Ghiorso (2014) to estimate pressures where Rotoiti inclusion glasses could be melts in equilibrium with quartz, plagioclase, and H₂O-vapor, obtaining 75–223 MPa. In this study, we use melt inclusion compositions presented in Smith et al. (2010) and the MagmaSat model of Ghiorso and Gualda (2015) to estimate Rotoiti entrapment pressures isothermally (740 ºC), yielding 81–261 MPa. Such pressures equate to shallow crustal depths, as is found in studies of the products of other explosive silicic eruptions and that are expected from the Rotoiti caldera environment. The derived pressure ranges are relatively large, however, reaching particularly low values, but the significance of those low values is unclear for the complexities previously presented and cannot be assessed without further focused study.

Mutch et al. (2016) presented an updated iteration of Al-in-amphibole barometry, which uses compositions of near-solidus amphiboles co-saturated with plagioclase, quartz, and melt to calculate an equilibration pressure; this expression has an estimated minimum uncertainty of ± 60 MPa based on the limiting precision of Al analyses in amphiboles. EQF amphibole rim compositions from this study and Molloy et al. (2008) yield pressures of 168–233 and 150–202 MPa, respectively (Fig. 9). We also derive pressures for the small volume, uppermost Rotoiti deposits, which are similar in mineralogy and composition to EQF (trace biotite and hornblende, absence of cummingtonite) (Schmitz and Smith 2004), yielding crystallization pressures of 105–552 MPa, which includes outliers. Excluding outliers, the barometer gives pressures of 179–282 MPa. The similarity of the Al-in-amphibole derived pressures supports that the hornblende in Rotoiti and EQF crystallized at similar shallow, sub-caldera depths, and as with temperatures, the pressures are indistinguishable within formal uncertainty with the experimental determination of 140 ± 10 MPa.

Absence of experimental hornblende and orthopyroxene

As mentioned, we did not observe synthetically grown hornblende or orthopyroxene in our experimental products, despite those minerals being relatively large, well formed, and compositionally little zoned in the natural pumice. The only hydrous Fe–Mg silicate that we produced was biotite, which was too small to analyze quantitatively by EMP, but which was confirmed by EDS spectra. Probable causes for the absence of hornblende and orthopyroxene are either that the natural melt employed as starting material is too evolved to grow those minerals appreciably upon cooling below the liquidus, despite their being equilibrium phases, or that those minerals are equilibrium phases but are in reaction relation with the natural EQF melt to produce biotite. Consistent with these hypotheses is that hornblende generally disappears in common arc granitoids as whole-rock SiO₂ concentrations exceed 70–72 wt.% (Dodge et al. 1982), and evidence for a reaction relation between hornblende and biotite is long known from petrographic observations of plutonic rocks (Speer 1987). First et al. (2021) presented BSE images and other experimental evidence for an orthopyroxene to biotite reaction in the presence of rhyolitic melt. It’s possible to write many amphibole to biotite reactions owing to the many amphibole endmembers, but two simple H₂O conservative ones are:

$${\text{KAlSi}}_{3} {\text{O}}_{{8 }(sanidine, \, melt)} \, + \,{\text{Ca}}_{2} \left( {{\text{Mg}},{\text{Fe}}^{{2} + } } \right)_{3} {\text{Al}}_{2} {\text{Si}}_{6} {\text{Al}}_{2} {\text{O}}_{{22}} \left( {{\text{OH}}} \right)_{{2 }(tschermakite)} \, = \,{\text{2 CaAl}}_{2} {\text{Si}}_{2} {\text{O}}_{{8 }(plagioclase)} \, + \,{\text{K}}\left( {{\text{Mg}},{\text{Fe}}^{{2} + } } \right)_{3} {\text{AlSi}}_{3} {\text{O}}_{{1}0} \left( {{\text{OH}}} \right)_{{2 }(biotite)} \, + \,{\text{2 SiO}}_{{2 }(quartz)}$$

and:

$${\text{KAlSi}}_{3} {\text{O}}_{{8 }(sanidine, \, melt)} \, + \,{\text{Ca}}_{2} {\text{Fe}}^{{2} + }_{5} {\text{Si}}_{8} {\text{O}}_{{22}} \left( {{\text{OH}}} \right)_{{2 }(ferro - actinolite)} \, + \,\raise.5ex\hbox{$ 1$}\kern-.1em/ \kern-.15em\lower.25ex\hbox{$ 2$} {\text{ O}}_{{2 }(vapor, \, melt)} \, + \,{\text{2 Al}}_{2} {\text{O}}_{{3 }(melt)} \, = \,{\text{KFe}}^{{2} + }_{3} {\text{AlSi}}_{3} {\text{O}}_{{1}0} \left( {{\text{OH}}} \right)_{{2 }(annite)} \, + \,{\text{2 CaAl}}_{2} {\text{Si}}_{2} {\text{O}}_{{8 }(plagioclase)} \, + \,{\text{4 SiO}}_{{2 }(quartz)} \, + \,{\text{Fe}}_{2} {\text{O}}_{{3 }(ilmeno - hematite)}$$

An orthopyroxene consuming reaction is:

$${\text{KAlSi}}_{3} {\text{O}}_{{8 }(sanidine, \, melt)} \, + \,{3 }\left( {{\text{Mg}},{\text{Fe}}^{{2} + } } \right){\text{SiO}}_{{3 }(orthopyroxene)} \, + \,{\text{H}}_{2} {\text{O}}_{(vapor, \, melt)} \, = \,{\text{K}}\left( {{\text{Mg}},{\text{Fe}}^{{2} + } } \right)_{3} {\text{AlSi}}_{3} {\text{O}}_{{1}0} \left( {{\text{OH}}} \right)_{{2 }(biotite)} \, + \,{\text{3 SiO}}_{{2 }(quartz)}$$

Quartz saturation fixes the activity of SiO₂, and subsequent fractionation increasing the activity of sanidine component and decreasing the activity of anorthite component drives the reactions to the right, forming biotite. Amphiboles are expected to have low concentrations of octahedral Al at the upper crustal pressures inferred for the EQF magma body, so the actinolite consuming reaction may dominate with a secondary effect of increasing the Al concentration of the residual amphibole. This could provide a potential alternative to heating to account for the weak rim-ward increases in Al reported by Molloy et al. (2008) in some EQF hornblendes. The amphibole consuming reactions also produce anorthite component, potentially contributing to the nearly constant compositions of the thick rims on EQF plagioclase grains. A straightforward experimental test of reaction relations would involve seeding the natural melt with grains of hornblende and orthopyroxene, similar to Blatter et al. (2017), who investigated pyroxene to amphibole reactions in dacitic melt, but such tests are beyond the resources available to us at the time of this report.

Alternative explanations for the absence of amphibole and orthopyroxene are kinetic barriers to their nucleation and growth, that the fO₂ of our experiments differed appreciably from that of the natural magma body, and that the EQF glass lost alkalis during hydration and chemical weathering. Hammer and Rutherford (2003) presented experimental evidence for sluggish re-equilibration between amphibole and melt at temperatures similar to those investigated herein, but barriers to nucleation and growth are expected to be greater for quartz than for mafic silicates owing to slow diffusion of Si in melts and the need for all other chemical components to diffuse away from growing quartz grains. However, quartz’s stability limits and growth proportions vary regularly with pressure and temperature, and its grains lack skeletal habits, consistent with a close approach to equilibrium. The essential chemical components in hornblende and orthopyroxene all diffuse faster through melt than does Si, and experiments on less evolved melts are notoriously prone to quench-crystallizing amphibole and pyroxene, so kinetic barriers (disequilibrium) are unlikely to account for the absence of hornblende and orthopyroxene in the EQF experiments. First et al. (2021) showed, with H₂O saturated experiments on a dacite, that distinctly low fO₂ conditions reduced the low-pressure stability limits of amphibole and biotite. At < 850 ºC and ΔNi–NiO = − 2, amphibole is stable at > 110 MPa, whereas at those temperatures and ΔNi–NiO = + 0.2 it’s stable to as low as 75 MPa (First et al. 2021). High fO₂ can also destabilize amphibole and biotite, as shown by isobaric studies on end-member compositions (ferropargasite: Gilbert 1966; ferrotremolite: Ernst 1966, annite: Eugster and Wones 1962, among others), and high fO₂ also can destabilize amphibole in melting experiments on igneous rocks (Spear 1981; Sisson et al. 2005). The influences of fO₂ on amphibole and biotite stability limits are not abrupt, however, and since our experiments were directly Ni–NiO buffered to closely match the fO₂ given by EQF Fe-Ti oxides, incorrect fO₂ is unlikely to account for the absence of amphibole and orthopyroxene from the synthesis products. Loss of alkalis by chemical weathering of pumice glass was previously assessed for both the Rotoiti and the EQF deposits by comparing their glass’s Na₂O and K₂O concentrations with those of their respective melt inclusions. Alkali concentrations are indistinguishable in the pumice glasses and their respective melt inclusions (Fig. 2), so any losses by chemical weathering are below the measurement resolution. For these reasons we prefer amphibole – melt and orthopyroxene – melt reactions to explain the absence of those minerals from the synthesis products.

RhyoliteMELTS model comparison

To investigate similarities and differences between our experimental results and thermodynamic predictions from RhyoliteMELTS (v 1.1.0) (Gualda et al. 2012), we input the composition of the EQF glass concentrate and modeled this at H₂O-saturated conditions (H₂O = 8.0 wt.%) and ƒO₂ = Ni–NiO to generate a phase diagram (Fig. 10). Neither amphibole nor biotite were predicted as stable in the RhyoliteMELTS simulations, although no significance is assigned to this since thermochemistries and activities of those minerals are not well established in that program. The resultant predicted mineral saturation curves generally align with our experimental results; however, there are some key differences. First, RhyoliteMELTS derives the crossover between liquidus quartz and plagioclase at 100 MPa and 795 ºC, whereas we found this crossover experimentally at 140 MPa and 755 ºC. Second, RhyoliteMELTS simulates the plagioclase and K-feldspar appearance curves as intersecting at 300 MPa and 732 ºC, whereas we find this experimentally around 210 MPa and 700 ºC. The latter difference reflects that in the RhyoliteMELTS model, plagioclase and K-feldspar stabilities respond nearly equally to changing H₂O-pressure, whereas in the experiments, plagioclase is the more sensitive of the two for the investigated bulk composition. For geological and volcanological applications the agreement is generally good in that both the experiments and the simulations indicate shallow crustal storage if the EQF magma was saturated with pure or nearly pure H₂O. The temperature difference appears relatively large but because output from RhyoliteMELTS, and the other members of the MELTS family of programs, do not propagate and report uncertainties on phase stability limits, the significance of the differences in temperature and pressure cannot be assessed quantitatively – perhaps the uncertainties on the MELTS-predicted crossover at 100 MPa and 795 ºC are sufficiently large as to overlap the direct experimental determination at 140 MPa and 755 ºC.

Mineralogical controls on magma storage

A puzzling aspect of the modern (~ 45 ka to present) Taupo Volcanic Zone has been the absence of K-rich alkali feldspar in all melt-bearing volcanic products, including highly evolved rhyolites (Danišík et al. 2012; Wilson et al. 1995); however, Burt et al. (1998) reported the presence of rare K-feldspar-bearing plutonic lithics in the Rotoiti lag breccia and EQF flow deposits. Sanidine in plutonic lithics is chiefly preserved as granophyric intergrowths, which Burt et al. (1998) interpreted as having formed by high-temperature, rapid ascent from mobilization of the Rotoiti magma body. Burt et al. (1998) presented two-feldspar temperatures between 700 and 800 °C, confirming that the lithics were at or close to solidus conditions, potentially as part of the outer, crystal-rich ‘rigid sponge’ or roof portions of the magma reservoir where cooler conditions dominate. Experimentally, we found that the erupted EQF melt composition saturates with K-feldspar only at temperatures < 760 °C at 50 MPa through < 700 °C at 200 MPa (Fig. 5), close to the probable H₂O-saturated granite (rhyolite) solidus. Thus, the EQF magma, stored at 140 MPa and 755 °C, would have had to cool a further ~ 60 °C and would have attained high crystallinity in order to produce K-feldspar.

Rhyolites in other parts of the world have erupted with alkali feldspar phenocrysts and at temperatures greater than the effective H₂O-saturated solidus, and the absence of K-feldspar from EQF and other modern Taupo rhyolites may simply be a product of their relatively K₂O-poor compositions. Among rhyolitic tuffs globally, those lacking K-feldspar typically have low K₂O concentrations in their glasses (< 4.5 wt.%). For example, Rotoiti, Novarupta, Usu, Krafla, and EQF tuffs all have K₂O < 4.5 wt.%, lack K-feldspar in the pumice samples, and correspondingly, did not grow K-feldspar in any experimental studies (Coombs and Gardner 2001; Nicholls et al. 1992; Rooyakers et al. 2021; Tomiya et al. 2010). Conversely, we stabilized K-feldspar in four runs for a similarly low-K₂O starting composition, thereby defining a K-feldspar saturation curve. Potentially this is due to our use of a highly evolved glass separate as our starting material, which allowed productive experiments to near-solidus conditions, whereas the other studies used whole-rocks whose less-evolved compositions and relict phenocrysts yielded more crystalline and difficult to interpret products at similar conditions that experimentalists tend to avoid. Figure 11 shows glass and whole-rock compositions of some well-known global rhyolites plotted in the haplogranitic ternary and the tonalite-trondhjemite-granite (TTG) discrimination diagram (Barker 1979). This comparison shows that the K-feldspar-free magmas plot closer to the quartz–albite join than do the K-feldspar-bearing magmas in the haplogranitic ternary. Similarly, the K-feldspar-free magmas plot closer to the trondhjemite region of the TTG ternary relative to the K-feldspar-bearing magmas which plot in the granite region. Notably, the EQF glass composition plots in between the K-feldspar-free and the K-feldspar-bearing groups in the haplogranitic plot and in the TTG diagram. The EQF whole-rock composition falls within the K-feldspar-free region on the haplogranitic ternary plot and on the granodiorite boundary region on the TTG plot. Since the EQF pumice lacks K-feldspar, the EQF glass composition must be at the threshold of K-feldspar saturation–a modest increase in normative orthoclase component or a moderate decrease in temperature and increase in crystallinity would push the EQF glass composition into the K-feldspar-bearing group.

The absence of K-feldspar provides a secondary limit on the storage conditions of the EQF magma body, which must have resided at higher temperature and greater H₂O-pressure than K-feldspar-in, or > 760 °C if close to 50 MPa and > 700 °C if close to 200 MPa. In our K-feldspar-bearing experiments, we saw a wide range in crystallinity between 17 and 95 wt.% crystals. The highly crystalline experimental products show that K-feldspar appears very near the system solidus. Therefore, we suggest that K-feldspar phenocrysts are absent in TVZ volcanic products because those magmas are poor in K₂O (trondhjemitic or low-K granitic), so they only saturate in K-rich alkali feldspar close to the solidus where the magmas are much too crystalline to erupt. Similarly, if such compositions are heated, K-feldspar melts out close to the solidus, precluding K-feldspar in erupted products that originate by remobilizing (defrosting) near-solidus and sub-solidus intrusions.

Magma storage in a broader context

The widely advocated “mush model” for large, subvolcanic silicic systems, illustrated in Fig. 12 (after Hildreth 2004), envisions a dominantly crystal-rich shallow intrusion with highly evolved melts percolating upward through a viscous crystalline framework (rigid sponge) and accumulating in a melt-rich central domain or domains. Although such conceptual diagrams often show gradual transitions from crystal-rich to melt-rich, little is actually known of such details. Recent work shows that multiple batches of crystal-poor melts may coexist within the larger crystal-rich body (Ellis et al. 2014; Storm et al. 2014; Rubin et al. 2016; Swallow et al. 2019), an interpretation presaged by discovery of the isotopic heterogeneity of some zoned plutons (Hill 1988; Kistler et al. 1986). The mineral grains within the crystal-rich region grew by various processes and in various settings, including as disaggregated cumulus crystals, autocrysts grown simply and directly from the extant melt, and as grains or zones of grains grown from prior replenishments as the shallow intrusion was constructed and sustained. In this conceptualization, an erupting rhyolite is sourced from some crystal-poor portion of the intrusion, and as that melt-rich portion evacuates, it may be followed by progressively more crystal-rich materials. This aspect of the model was developed to account for voluminous, stratigraphically zoned tuff sequences emplaced during single eruptions. Such zoned tuffs commonly increase in their crystal abundances and magmatic temperatures in successively deposited products as those products’ bulk-magma, melt, and mineral compositions and assemblages also become less evolved (e.g., Hildreth and Wilson 2007). In contrast, the Rotoiti tuff is largely chemically homogenous, and lacks pronounced stratigraphic zoning other than a slight variation in the proportion of trace phases (biotite, hornblende, and orthopyroxene) in the uppermost units and the presence of mingled pumice containing two distinct magma components (Shane et al. 2005; Schmitz and Smith 2004; Molloy et al. 2008). Based on the similarity between the EQF and Upper Rotoiti mineral assemblage (viz. trace biotite, and no cummingtonite), it has been suggested that the EQF magma is the second magma component identified in the mingled Rotoiti pumice, and that the EQF magma fed into the Rotoiti reservoir post-caldera collapse; however, this interpretation is unsupported by isotopic measurements (Schmitz and Smith 2004). And although mineralogical and chemical differences between Rotoiti and EQF preclude direct parental lineage, the pressure and temperature constraints support that magmas of the two volcanic centers were adjacent, coeval neighbors, residing at similar shallow depths and temperatures.

A previous experimental study by Nicholls et al. (1992) limits the Rotoiti magma’s storage conditions loosely at < 300 MPa and < 750 $^\circ$C, based on the presence and stability of cummingtonite, whereas our results indicate that the EQF magma was stored at $\sim$ 140 MPa and ~ 755$^\circ$C, based on liquidus co-saturation of the natural melt in quartz, plagioclase, and biotite. The pressure results are permissively the same, and based on the experiments, the Rotoiti magma may have been cooler, but this is not reflected in pronounced differences in the composition of EQF versus Rotoiti glasses, both of which are high-silica rhyolites with similar concentrations (wt.%) of TiO₂ (EQF: 0.13, Rotoiti: 0.19), FeO_T (EQF: 0.78, Rotoiti: 1.00), and CaO (EQF: 0.82, Rotoiti: 0.87), employing average Rotoiti glass compositions from Schmitz and Smith (2004).

Where the glasses do differ notably is in their alkali concentrations, Na₂O (EQF: 3.45, Rotoiti: 3.75) and K₂O (EQF: 4.49, Rotoiti: 3.44). The reaction:

$$0.{\text{5 K}}_{2} {\text{O}}_{(melt)} \, + \,0.{\text{5 Al}}_{2} {\text{O}}_{{3 }(melt)} \, + \,\left( {{\text{Mg}},\!\!\!{\text{ Fe}}^{{2} + } } \right)_{2} \left( {{\text{Mg}},\!\!\!{\text{ Fe}}^{{2} + } } \right)_{5} {\text{Si}}_{8} {\text{O}}_{{22}} \left( {{\text{OH}}} \right)_{{2 }(cummingtonite)} \, = \,{\text{K}}\left( {{\text{Mg}},\!\!\!{\text{ Fe}}^{{2} + } } \right)_{3} {\text{AlSi}}_{3} {\text{O}}_{{1}0} \left( {{\text{OH}}} \right)_{{2 }(biotite)} \, + \,{4 }\left( {{\text{Mg}},\!\!\!{\text{ Fe}}^{{2} + } } \right){\text{SiO}}_{{3 }(orthopyroxene)} \, + \,{\text{SiO}}_{{2 }(quartz)}$$

indicates that for melts saturated with quartz and orthopyroxene, and with similar Al₂O₃ concentrations (~ activities), lower K₂O concentration enhances cummingtonite at the expense of biotite. Smith et al. (2005) also showed that glasses in cummingtonite bearing OVC deposits have lower K₂O concentrations than for glasses that contain biotite. Cummingtonite is also known from other distinctly low-K₂O silicic magmas (Geschwind and Rutherford 1992; Nadeau et al. 2015). Potentially, the Rotoiti tuffs contains cummingtonite not because it's storage conditions were appreciably cooler than the EQF but because its melt was less potassic, a possibility best explored by narrowing the experimental limits on Rotoiti mineral–melt stabilities. The petrologic observations and experimental results, combined with the near contemporaneity of the eruptions and the adjacent proximity of their vents (~ 20–25 km) support that the Rotoiti and EQF were lateral members of the same overall intrusive complex (pluton), with the crystal-moderate Rotoiti magma in its central area, and the crystal-rich EQF magma occupying the mushy periphery (Fig. 12).

Conclusions

Our experimental results determine the pre-eruptive storage conditions for the Earthquake Flat (EQF) magma body by locating the liquidus co-saturation of its natural melt with quartz, plagioclase, biotite, and H₂O-rich fluid. Based on our experiments, this corresponds to 140 MPa ± 10 MPa and 755 ºC ± 5 ºC for a H₂O-saturated rhyolitic magma at ƒO₂ = Ni–NiO. By using a glass concentrate starting material, our experimentally defined conditions represent the conditions under which the magma body was staged immediately prior to eruption. Our experiments on the natural erupted glass did not produce hornblende and orthopyroxene that are sparse but widespread in the EQF pumices, a difference we provisionally ascribe to those minerals being in reaction relation with melt. Our experimental results are within error of results from Fe-Ti oxide thermometry, amphibole-plagioclase thermometry, amphibole barometry, and melt inclusion volatile barometry, and are similar to results from RhyoliteMELTS, which predicts the intersection of the quartz and plagioclase saturation curves moderately shallower and hotter than the experimental determination. If the temperature of the EQF magma body increased approaching its eruption (Molloy et al. 2008), then temperature had stabilized sufficiently such that there was negligible diversity in melt composition and mineral assemblage across nearly all of the ~ 10 km³ of erupted EQF magma. A general absence of K-feldspar in rhyolites of the larger Taupo Volcanic Zone (TVZ), including EQF, is largely a consequence of their low-K₂O concentrations more similar to trondhjemites than to granites. K-feldspar appears in our experiments approximately 60 ºC cooler than the melt’s H₂O-saturated liquidus at upper crustal pressures, close to the probable solidus. The low-K₂O concentration also accounts for the presence of cummingtonite in some TVZ rhyolites, which is controlled by a combination of the system’s temperature and the low K₂O concentration of the melt in the larger TVZ system. We show that the EQF magma system did not have sufficient K₂O component in the pre-eruptive melt and was stored slightly too hot for K-feldspar saturation. Our findings provide pressure, temperature, and compositional constraints for H₂O-saturated, high-fO₂ ($\sim \! \Delta$ Ni–NiO = 0) rhyolite systems globally.

Compositional differences between the Rotoiti and EQF glasses, and mineralogical differences between the two tuffs, require that the source magmas must have evolved along separate lines of descent, and therefore, that they were not stored in a well-mixed magma reservoir, although they could have been portions of the same intrusion assembled from slightly differing supplied magmas. Combined with our experimental results, which show that the two bodies were broadly stored at similar conditions, we propose a model of lateral variability in the shallow crustal intrusion, where multiple melt-rich bodies developed at similar storage conditions, but from different parental magmas, yielding different chemical signatures. Based on the observed textures in the EQF tuff, we interpret that it originated from a higher crystallinity marginal portion of the active intrusion (pluton) whose crystal-moderate center sourced the Rotoiti rhyolites only shortly earlier.

Data availability

Data are available in the Tables (1–4), and electronic supplemental material files (ESM 1 – 4). All data are also available in the U.S. Geological Survey data release from Grant et al. (2024; https://doi.org/10.5066/P13IM4QQ).

Notes

Growth of minerals from melt is a disequilibrium process, so minerals capable of solid-solution need not have compositions strictly in equilibrium with the melt from which they are growing, but for simplicity of discussion this growth disequilibrium is ignored.

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Acknowledgements

The project was funded through an NSF GRIP grant written by Elizabeth Grant, with input from Dawnika Blatter, Laura Waters, and Kari Cooper, and from an NSF INTERN grant, supplemental to Kari Cooper’s existing NSF grant. This research was supported by the USGS Volcano Hazards Program. We would like to thank our reviewers Emily First, Michelle Coombs, and an anonymous reviewer, whose input and guidance greatly improved this manuscript. We are grateful to Greg Baxter for sample preparation, Dawn Ruth for guidance on the SEM and EMPA, and Darren Gravley and Chad Deering for their field guidance. Any use of trade, firm, or product names is for descriptive purposes only and does not imply endorsement by the U.S. Government.

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National Science Foundation, EAR-1654506, Kari M. Cooper.

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University of California, Davis, CA, USA
Elizabeth R. G. Grant & Kari M. Cooper
U.S. Geological Survey, California Volcano Observatory, Moffett Field, CA, USA
Dawnika L. Blatter & Thomas W. Sisson

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Elizabeth R. G. Grant
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Thomas W. Sisson
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Kari M. Cooper
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This work was conceived of and designed by Elizabeth Grant, Dawnika Blatter, and Tom Sisson. Sample preparation and analyses were performed by Elizabeth Grant. The manuscript was written by Elizabeth Grant with significant input from Dawnika Blatter, Tom Sisson, and Kari Cooper.

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Correspondence to Elizabeth R. G. Grant.

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Grant, E.R.G., Blatter, D.L., Sisson, T.W. et al. Shallow storage of the explosive Earthquake Flat Pyroclastics magma body, Okataina Volcanic Center, Taupo Volcanic Zone, New Zealand: evidence from phase-equilibria experiments. Contrib Mineral Petrol 179, 81 (2024). https://doi.org/10.1007/s00410-024-02151-y

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Received: 17 October 2023
Accepted: 31 May 2024
Published: 25 July 2024
DOI: https://doi.org/10.1007/s00410-024-02151-y

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Shallow storage of the explosive Earthquake Flat Pyroclastics magma body, Okataina Volcanic Center, Taupo Volcanic Zone, New Zealand: evidence from phase-equilibria experiments

Abstract

Introduction

Geologic and petrologic background

Petrologic summary of the Rotoiti and Earthquake Flat rhyolites

Methods

Starting Material

Use of separated natural glass, versus bulk pumice, for experimental syntheses

Capsule configuration and ƒO2 buffering

Experimental Setup

Electron Microprobe (EMP)

ATR-FTIR

Results

Characteristics of EQF pumices

Quartz

Plagioclase

Glomerocrysts

Biotite

Hornblende

Fe-Ti Oxides

Orthopyroxene

Inclusions in ferromagnesian phases

Melt inclusions and embayments

Experimental results

Approach to equilibrium

Phase stability

Experimental phase descriptions

Phase proportions

Discussion

Thermometry

Barometry

Absence of experimental hornblende and orthopyroxene

RhyoliteMELTS model comparison

Mineralogical controls on magma storage

Magma storage in a broader context

Conclusions

Data availability

Notes

References

Acknowledgements

Funding

Author information

Authors and Affiliations

Contributions

Corresponding author

Ethics declarations

Conflict of interest

Additional information

Publisher's Note

Supplementary Information

Supplementary file1 (PDF 508 KB)

Supplementary file2 (XLSX 181 KB)

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Capsule configuration and ƒO₂ buffering