Exploring Connections Between a Very Large Volume Ignimbrite and an Intracaldera Pluton: Intrusions Related to the Oligocene Wah Wah Springs Tuff, Western US

Brigham Young University BYU ScholarsArchive

Theses and Dissertations

2013-05-31

Chloe Noelle Skidmore Brigham Young University - Provo

Follow this and additional works at: https://scholarsarchive.byu.edu/etd

Part of the Geology Commons

BYU ScholarsArchive Citation Skidmore, Chloe Noelle, "Exploring Connections Between a Very Large Volume Ignimbrite and an Intracaldera Pluton: Intrusions Related to the Oligocene Wah Wah Springs Tuff, Western US" (2013). Theses and Dissertations. 4041. https://scholarsarchive.byu.edu/etd/4041

This Thesis is brought to you for free and open access by BYU ScholarsArchive. It has been accepted for inclusion in Theses and Dissertations by an authorized administrator of BYU ScholarsArchive. For more information, please contact [email protected], [email protected]. Exploring Connections Between a Very Large Volume Ignimbrite

and an Intracaldera Pluton: Intrusions Related to the

Oligocene Wah Wah Springs Tuff, Western US

Chloe Skidmore

A thesis submitted to the faculty of Brigham Young University in partial fulfillment of the requirements for the degree of

Master of Science

Eric H Christiansen, Chair Michael J. Dorais Bart J. Kowallis

Department of Geological Sciences

Brigham Young University

June 2013

Exploring Connections Between a Very Large Volume Ignimbrite and an Intracaldera Pluton: Intrusion Related to the Oligocene Wah Wah Springs Tuff, Western US

Chloe Skidmore, Department of Geological Sciences, BYU Master of Science

The Wah Wah Springs Tuff and the Wah Wah Springs Intrusive Granodiorite Porphyry (Wah Wah Springs Intrusion) both originated from the Indian Peak caldera complex, which was a major focus of explosive silicic activity in the middle Cenozoic Great Basin ignimbrite flareup. This caldera formed 30.0 Ma when an estimated 5,900 km3 of crystal-rich dacitic magma erupted to create the Wah Wah Springs Tuff. The Wah Wah Springs Intrusion later intruded the tuff, causing resurgence of the caldera. Field, modal, and geochemical evidence suggest the tuff and intrusion are cogenetic. The mineral assemblages of the two rocks are similar: both include similar proportions of plagioclase, quartz, hornblende, biotite, clinopyroxene, and Fe-Ti oxides, with trace amounts of titanite, apatite, and zircon. Whole rock geochemistry also matches, and both rocks have distinctively high Cr concentrations. Plagioclase, hornblende, and clinopyroxene have similar compositions but biotite and Fe-Ti oxides have been hydrothermally altered in the intrusion. Both hornblende and quartz provide clues to the magmatic evolution of the Wah Wah Springs Intrusion. Hornblende grains are either euhedral, have reaction rims, or are completely replaced by anhydrous minerals. Deterioration of hornblende was caused by decompression as the magma ascended and then stalled and solidified at shallow depths. Two stages of quartz growth are shown in cathodoluminescence (CL) imagery. Quartz first grew then was resorbed during eruption, then grew again at lower pressures indicated by CL-bright quartz rims and groundmass grains. The geochemical and mineralogical similarities, together with the distinctive hornblende and quartz characteristics suggest that after the Wah Wah Springs Tuff erupted, the unerupted mush rose to a shallow level where it crystallized at low pressure to form the Wah Wah Springs Intrusion. This indicates that the both rocks formed in the same chamber, and that tuffs and associated intrusions can be intimately related.

Keywords: intracaldera pluton, magma chamber, ignimbrite, Wah Wah Springs Tuff ACKNOWLEDGEMENTS

I would like to express my appreciation for all those who assisted in the completion of this work. Much appreciation goes to Dr. Eric Christiansen for answering my endless questions and providing guidance and direction. I am also grateful for the insight and comments of my committee, Dr. Bart Kowallis and Dr. Mike Dorais, with additional thanks to Dr. Dorais for his help with the electron microprobe. I would also like to acknowledge Dr. Myron Best and Kurtus

Woolf for their previous research on the Wah Wah Springs Tuff, without which this project could not have been done.

A special thanks also goes to my fellow students who helped me out a great deal: namely,

Doug Johnson for his instruction on lab equipment use, and Kat Robertson and Audrey Warren for their help with field work.

Last but not least, I’d like to thank Tanner Mills for always keeping me in good spirits while spending long hours in the grad cubes. TABLE OF CONTENTS

ABSTRACT ...... ii ACKNOWLEDGEMENTS ...... iii TABLE OF CONTENTS ...... iv LIST OF FIGURES ...... v LIST OF TABLES ...... vi INTRODUCTION ...... 1 GEOLOGIC SETTING ...... 3 METHODS ...... 5 FIELD RELATIONS ...... 7 RESULTS ...... 9 Whole Rock Analysis ...... 9 Petrography ...... 10 Mineral Chemistry ...... 11 Feldspar ...... 11 Hornblende ...... 12 Pyroxene ...... 14 Biotite ...... 15 Fe-Ti Oxides ...... 16 Quartz ...... 16 Thermobarometry ...... 19 DISCUSSION ...... 22 Comparison of the Wah Wah Springs Intrusion to the Wah Wah Springs Tuff ...... 22 Magma Decompression ...... 25 Magma Chamber and Caldera Development ...... 29 CONCLUSIONS...... 31 REFERENCES ...... 34 FIGURES ...... 41 TABLES ...... 56 APPENDIX A (Thin Sample Descriptions) ...... 63 APPENDIX B (Quartz Cathodoluminescence Images) ...... 66

LIST OF FIGURES

Figure 1. Index Map ...... 41 Figure 2. Geologic Map of Indian Peak ...... 42 Figure 3. Field Photographs ...... 43 Figure 4. Whole Rock Geochemistry Diagrams ...... 44 Figure 5. Modal Percentages Chart ...... 45 Figure 6. Feldspar Ternary Diagrams ...... 46 Figure 7. Hornblende Microphotographs ...... 47 Figure 8. Hornblende Classification Diagram ...... 47 Figure 9. Hornblende Exchange Mechanisms ...... 48 Figure 10. Pyroxene Classification Diagram ...... 49 Figure 11. Quartz Microphotographs ...... 50 Figure 12. Quartz Cathodoluminescence Imagery ...... 51 Figure 13. Pressure-Temperature Diagrams ...... 52 Figure 14. Mineral Stability Diagrams ...... 53 Figure 15. Ti in Quartz vs. Pressure Diagram ...... 54 Figure 16. Diagram of Caldera Development ...... 55

LIST OF TABLES

Table 1. Whole Rock Geochemistry ...... 56 Table 2. Representative Plagioclase Analyses ...... 58 Table 3. Representative Hornblende Analyses ...... 59 Table 4. Representative Pyroxene Analyses ...... 60 Table 5. Quartz Cathodoluminescence Characteristics ...... 61 Table 6. Summary of Geothermometry ...... 62

INTRODUCTION

The connections between volcanic and plutonic rocks are important to understanding magmatic systems. Traditionally, granitic intrusions have been thought to be intimately related to associated volcanic rocks and to have formed from large magma chambers that cool rather rapidly (<100,000 years) (Petford et al., 2000). For example, John (1995) (see also John et al.,

2008; Henry and John, 2013) concluded that granitic intrusions into the floors of two different middle Cenozoic calderas in Nevada formed from magma remaining after the eruption of large volumes of rhyolitic ignimbrites. However, this relationship has been questioned by the work of

Tappa et al. (2011) who concluded that post-caldera intrusions in northern New Mexico were not composed of the same magma as the earlier ignimbrites (i.e., Amalia Tuff) that they intrude.

They suggest that all of the magma was erupted during a volcanic phase and that the intrusive magma formed by a separate partial melting event in the lower crust. Lipman (2007) cites evidence that Cordilleran plutons in the Southern Rocky Mountain assembled incrementally and posits that there is incremental magma generation whether there is a magma chamber or not. On a more global scale, Glazner et al. (2004) contend that “plutons may commonly form incrementally without ever existing as a large magma body”. According to this widely cited hypothesis (e.g. Lipman 2007, de Silva and Gosnold, 2007, Tappa et al., 2011) large plutons form slowly by small accumulations of magma that solidify before other small batches of magma are intruded over time scales of millions of years, rather than cooling from a large magma chamber (or “big tank”).

Evidence that there are large magma chambers which could potentially cool into plutons exists in the sheer volume of some very large ignimbrites such as the Wah Wah Springs Tuff

(~5900 km3) and other similar tuffs which exceed 3000 km3 (Best et al., 2013; Woolf, 2008;

Maughan et al., 2002). If plutons only grow by incremental additions that solidify before the next intrusion this leaves room to wonder where all the magma for super eruptions like those from the

Indian Peak Caldera Complex originates. Such voluminous eruptions demonstrate that large magma chambers do indeed exist (Bachmann and Bergantz, 2008). Shaw (1985) suggested that extrusion rates are a function of intrusion rates, which could imply that the high volcanic effusivity that happened during an ignimbrite flare-up is indicative of voluminous magma chambers underground. Furthermore, Lipman (2007) proposed that the volatile-rich roof of a magma chamber erupts and the leftover magma solidifies and Hildreth (2004) suggests that large granodiorite plutons are the result of a leftover, melt-depleted, crystal-rich mush. On the other hand, Verplanck et al. (1999) theorized that the Organ pluton of New Mexico cooled rapidly right after the Squaw Mountain Tuff erupted from the same chamber (see also Zimmerer and

McIntosh, 2013). These studies imply that extrusive and intrusive rocks originate in the same magma chamber, a chamber which does not empty completely during volcanism.

The Wah Wah Springs Tuff and the associated intruding intracaldera granodiorite of the

Indian Peak caldera provide a worthwhile case study to test the hypothesis of Glazner et al.

(2004) and Tappa et al. (2011) that plutons may form incrementally and not be related to volcanic rocks. The tuff was first defined as a volcanic unit by Mackin (1960). The Wah Wah

Springs Tuff is a result of the middle Cenozoic ignimbrite flare-up, which was an important time in the tectonic history of western North America (Best et al., 2009). Slab rollback following the

Sevier and Laramide orogenies caused an increase in volcanism in southwestern Utah and eastern Nevada, forming large caldera complexes like the 80 x 120 km Indian Peak complex of which the Indian Peak caldera is a part (Best et al., 1989, Figure 1). Extensive eruptions from

2 supervolcanoes in the Great Basin produced many voluminous tuffs including the 30.0 Ma Wah

Wah Springs Tuff (Best et al., 2013).

Studying the geochemistry and structural relationship of these rocks sheds light on the mechanism of pluton formation and emplacement and helps clarify the often-ambiguous ties between intrusive and extrusive rocks. Geochemical, modal, and field evidence suggest that the

Wah Wah Springs Tuff and associated intrusion are indeed genetically related and originated in the same magma chamber.

GEOLOGIC SETTING

The Indian Peak volcanic field was a major focus of explosive silicic activity during the mid-Cenozoic (36-18 Ma) ignimbrite flareup of the southern Great Basin (Best et al., 1973,

1989, 2009). There are twenty-four ignimbrites in the Indian Peak-Caliente caldera complex alone (Best et al., 2013). Notable tuffs that erupted out of this caldera complex include the Lund tuff (~4400 km3), which erupted 29.0 Ma (Maughan et al., 2002) and the Cottonwood Wash Tuff

(~2000 km3), which erupted 30.9 Ma (Best et al., 2013). However, the Wah Wah Springs Tuff, which erupted between these two, is the most extensive and the voluminous at 5900 km3 (Best et al., 2013; Figure 1).

The ignimbrite flareup represents a time when North America was tectonically similar to the central Andes today (Best et al., 2009). In North America, enormous amounts of volcanism were generated by subduction of the Farallon slab (Lipman et al., 1971). Humphreys et al. (2003) attributed this magmatic flareup to removal, or break-off, of the shallowly dipping slab with high melt production caused by the asthenosphere being placed in direct contact with a thin and hydrated lithosphere. Christiansen and McCurry (2008) also suggested these rocks were derived

3 from mafic parent magmas produced by dehydration of oceanic lithosphere and melting in the mantle wedge above the subduction zone. These wet, mafic magmas rose, then mixed and fractionated in continental crust that was unusually thick (Best and Christiansen, 1991). In general, ignimbrites from this time are rhyolites and dacites and have been found to be calc- alkaline, magnesian, oxidized, wet, and relatively cool (<800oC) (Christiansen and McCurry,

2008).

The Indian Peak caldera is large; the inner ring fault extends 40 km north-south and the area within it, corrected for post eruption extension, is about 1000 km2 (Best et al., 2013). Best and colleagues (Best and Grant, 1987, Best et al., 1987a; 1987b) have mapped the caldera at a scale of 1:50,000, which spans across three mountain ranges, including the area described here in the southern Needle Range of southwestern Utah. They also named the Wah Wah Springs

Intrusive Granodiorite Porphyry, hereafter called the Wah Wah Springs Intrusion in this report for brevity.

When the Wah Wah Springs Tuff began erupting, it was initially deposited on the preexisting Marsden and Cottonwood Wash tuffs, which erupted earlier from the same area. It is believed that the Indian Peak caldera collapsed after the eruption of the outflow Wah Wah

Springs Tuff, which is supported by the presence of thick lenses of volcanic breccia (Best and

Grant, 1987; Best et al., 2013). This breccia is mostly made up of the previously-erupted

Cottonwood Wash Tuff but also includes older andesite and rhyolite tuffs. After the initial collapse, magma continued erupting, creating the lithic-rich intracaldera Wah Wah Springs Tuff.

Sometime following that, granodiorite magma initiated resurgent uplift as evidenced by vertical foliation of the tuff and the intracaldera pluton. Much later, the caldera was cut and exposed by

4 erosion and Basin and Range faulting, creating the modern-day Indian Peak Range (Best et al.,

2013).

METHODS

Field work included gathering 32 samples using a steel hammer and chisel and examining structural relationships between the tuff, intrusion, and surrounding bedrock. Twenty-eight of these samples were chemically analyzed, and 21 samples were considered for whole rock analysis in this report. These were selected based on freshness and relevancy since much of the intrusion has been propylitically altered, making the collection of fresh samples challenging.

In order to perform whole rock chemical analysis, samples were crushed and powdered using a tungsten carbide shatterbox. For major element analysis, the powder was mixed with lithium borates and heated to 1050° C in a Katanax electric furnace in order to make glass disks.

Loss-on-ignition (LOI) data was gathered by recording before and after weights on samples heated to 1000° C in a muffle furnace. For trace element analysis, pressed powder pellets were created using a binding agent and a hydraulic press. Both pressed powder pellets and glass disks were analyzed in a Rigaku Primus X-ray fluorescence spectrometer at Brigham Young

University. Major oxides analyzed included SiO2, TiO2, Al2O3, Fe2O3, MnO, MgO, Na2O, K2O, and P2O5. A Wah Wah Springs laboratory standard was also analyzed to ensure results were consistent with older analyses. Trace element concentrations (Ba, Rb, Th, Nb, La, Ce, Sr, Nd,

Sm, Zr, Cr, Ni, Sc, Zn, V, Cu, Ga, Y) were analyzed using the pressed pellets. Both major and trace elements were calibrated with international reference materials. The error for major elements is usually within 1% and trace element error is usually within 5% based on these reference materials.

Thin sections were examined petrographically and then used for mineral analysis in ten samples. To determine modal percentages, point counting was performed on a petrographic microscope. Between 500-1000 points were counted, depending on the grain size in the sample.

Major element compositions of hornblende, plagioclase, biotite, clinopyroxene, and Fe-Ti oxides were determined using a Cameca SX-50 electron microprobe at Brigham Young University.

Both traverses and point analyses were made, depending on the size and condition of the target grain. Hornblende, plagioclase, biotite, and clinopyroxene were analyzed with a 15 kV accelerating voltage, a 10 micron beam, and a 10 nA beam current. Magnetite and ilmenite were also analyzed under the same conditions but with a 2 micron beam.

Quartz cathodoluminescence (CL) was examined using a detector on a XL30 ESEM FEG scanning electron microscope (SEM) with a Gatan MiniCL detector at Brigham Young

University. Quartz grains were examined in thin sections then imaged using a 15 kV accelerating voltage, a 12.5° tilt, and scan time of 116 ms/line at 968 x 1268 pixels.

Reconnaissance analyses of trace elements in quartz were attempted, looking specifically at Ti,

Al, and Fe, with the electron microprobe at BYU. Titanium was calibrated on titanite, Al on an yttrium-aluminum garnet, and Fe on magnetite. Analyses were done with a 200 nA, 5 micron beam with 300 seconds on the peak and 150 seconds on both backgrounds. The theoretical detection limits for Ti, Al, and Fe, were 17 ppm, 24 ppm, and 72 ppm, respectively. Titanium and aluminum concentrations were above these detection limits but iron was sometimes below it.

However, since the concentrations of all three elements were so low, they could have been easily affected by carbon coating or other small changes with analysis environment, so these results must be regarded with caution.

FIELD RELATIONS

Field relations between the tuff and intrusion, as well as surrounding wall rock, were important in determining the sequence of events and whether or not the tuff and intrusion show any physical of evidence of forming from the same magma chamber.

The total area of exposed intrusion is 19.2 km2 with two large main bodies of outcrop and several smaller patches (Figure 2). These outcrops are either within or on the caldera ring fracture. Internal contacts within the intrusion were seen locally between small blobs of dark fine-grained rock and the “normal” coarse-grained, porphyritic rock that was found throughout most of the intrusion. From chemical and textural characteristics, these dark masses appear to be mafic enclaves of magmatic origin. In addition to the pervasive propylitic alteration, large quartz veins (to 1 m thick) are also present. Many are localized near fluorite mines related to the F-rich rhyolites of the Miocene Blawn Formation, so it is possible that the quartz veins are also related to this more recent stage of igneous activity. Trending dikes of the Miocene rhyolite cut the intrusion as well (Best et al., 1987a; Best et al., 1987b).

Outcrops of the granodiorite intrusion are texturally varied. Rocks range from equigranular, medium-grained, and phaneritic to aphanitic-porphyritic with finely crystalline groundmass that was similar to the tuff. Textural differences were somewhat regionally constrained; aphanitic-porphyritic rocks with finely crystalline matrixes are more common in the southern portion of the intrusion and there is a small patch of very coarse-grained rocks intruding

Paleozoic limestone near the ring fracture. The mineral assemblage of the intrusion includes plagioclase, quartz, biotite, hornblende, pyroxene, Fe-Ti oxides, and trace amounts of titanite,

7 zircon, and apatite. Plagioclase is the most common and usually largest phenocryst and was between 1-3 mm in size.

No regular compositional zoning of the pluton is apparent and the intrusion is granodioritic throughout which corresponds to the fact that Wah Wah Springs Tuff is a monotonous intermediate without significant zoning (Christiansen, 2005; Woolf, 2008).

However, a pegmatitic dike cuts through the body of the intrusion. This dike is about 2 m wide and poorly exposed, making its orientation indiscernible. Chemically, the dike is unlike the

Miocene rhyolites; thus, it could be a late differentiate of the granodiorite or related to other stages of Oligocene magmatism.

Although exposures of tuff-intrusion contacts are rare elsewhere in the world (John,

1995), Basin and Range faulting has fortuitously exposed contacts between the Wah Wah

Springs Tuff and the Wah Wah Springs Intrusion. The northern edge of the intrusion is marked by sub-horizontal contacts with the overlying tuff and also with Paleozoic carbonate wall rock

(Figure 3a). The carbonate rocks have been metamorphosed to coarse-grained white-gray marble. Contacts with tuff along the northern patches of the intrusion are sharp and there are distinct differences in phenocryst and matrix grain size between the tuff and granodiorite.

Vertical contacts are more common near the southern contact of the main mass where the intrusive rocks crop out in steep cliffs (Figure 3b). Contacts with tuff in the south are more gradational and the intrusion has an increasingly fine-grained groundmass towards the contact.

Both the tuff and intrusion are propylitically altered. There are also small, previously unmapped patches, of the intrusion along what is considered to be the northern ring fracture or structural margin of the caldera.

RESULTS

Whole rock analysis

Twenty-eight samples from the intrusion were originally chosen for whole rock chemical analysis, representing different geographic locations and textural variety (Table 1; for all data see supplemental material). Samples with high loss-on-ignition and perturbed alkali, iron, aluminum, and silica concentrations were determined to be hydrothermally altered and are not considered further in this report. Chemical analyses of the dark, fine-grained masses revealed that these are more mafic (57-61% SiO2) than the other intrusive rocks. Their distribution and chemistry suggests that they are genetically unrelated mafic enclaves. These samples as well as the hydrothermally altered samples are not considered further in this report but their whole rock geochemistry can be found in the supplementary tables. A total of 21 samples (out of the 28 analyzed) were used in the following discussion.

The intrusive samples range in SiO2 from 63.1 to 68.8 wt.% (normalized to 100% on a volatile free basis; excluding samples from the pegmatite dike) making most of them dacites or granodiorites (Figure 4a). These fall closely in line with the tuff samples, which vary from 62.5 to 70.0 wt. % (Woolf, 2008). The tuff also has one sample that is a rhyolite pumice and close in composition to the pegmatitic dike of the intrusion and could also be a sign of late differentiation

(Figure 4). On the modified alkali lime index (MALI), the intrusive rocks range from calcic to the upper boundary of calc-alkalic, which is also similar to the tuff (Figure 4b). The intrusion and tuff both range from medium-K to shoshonitic with most falling in the high-K region (Figure

4c). Normalized trace element patterns are also similar, showing negative Nb, Sr, P, and Ti anomalies and decreasing enrichment in the more compatible elements (Figure 4d). The Wah

Wah Springs Tuff has unique geochemical characteristics that are not observed in other contemporaneous monotonous intermediate tuffs (e.g., Lund and Cottonwood Tuff). Sr trends shows that there are difference in tuffs among the same caldera complex and that the Wah Wah

Springs Tuff and Intrusion are distinctly different from the Cottonwood Wash Tuff, although more similar to the Lund Tuff (Figure 4e). Most important, however, is that the Wah Wah

Springs Tuff has significantly higher Cr than both of the other tuffs from the Indian Peak Caldera

Complex and so does the intrusion (Figure 4f).

Petrography

Petrographic thin section analysis reveals that, modally, the intrusive rocks are also similar to the tuff. All contain plagioclase, quartz, biotite, and Fe-Ti oxides (Figure 5).

Accessory minerals in the intrusion include apatite, zircon, and titanite (found in mineral separates). Eight samples (out of ten thin sections examined) contain clinopyroxene, one sample

(MIN-0412) contains titanite, and two samples have alkali feldspar in their matrixes (MIN-0512 and MIN-8-60-1; for complete sample descriptions, see appendix). Titanite is rare in the tuff and intrusion and has been found only in mineral separates; it is so sparse it has not been reported in thin sections before. Only four samples contain hornblende that is unaltered or not replaced by anhydrous minerals. The condition of hornblende varies most among samples, while other minerals (feldspar, biotite, Fe-Ti oxides, pyroxene) are uniformly altered or unaltered throughout all samples. Orthopyroxene was not found in any of the samples. Fine-grained mafic xenolithic inclusions with defined borders are also present throughout the granodiorite and distinct from the dark masses (or “blobs”), which were also determined to be of mafic origin. These inclusions ranged from just a few centimeters to around 20 cm. In thin section, it is apparent that the

10 xenoliths are altered and composed of abundant Fe-Ti oxides and possibly quartz and plagioclase.

Texturally, the intrusive rocks are porphyritic-aphanitic to porphyritic-phaneritic. In thin section, it could be seen that matrixes are mostly composed of quartz, plagioclase, and, in two samples, alkali feldspar and vary from very fine-grained to coarse-grained (matrix grains up to 1 mm). Visible phenocrysts included plagioclase, hornblende, quartz, clinopyroxene, and biotite.

In samples with coarse-grained matrixes, quartz phenocrysts were sparse, as the matrix is made up of mostly late stage quartz.

Mineral Chemistry

Feldspar

All samples contained phenocrystic plagioclase with varying degrees of freshness. It varies in size, with the largest grains being up to 6 mm long. Hydrothermal alteration is common and many crystals show evidence of sericitization. Sieve texture is not prevalent and only minor resorption is observed. There are, however, instances of euhedral plagioclase grown on resorbed cores that contain abundant Fe-Ti oxide and pyroxene inclusions. Grains like these are also found in the Wah Wah Springs Tuff (Woolf, 2008). Broken plagioclase phenocrysts are present in one sample, but are not very common.

Plagioclase analyses were taken in traverses ranging from 230 to 2400 microns with steps from 24 to 240 microns. The plagioclase in the intrusion is chemically similar to that in the tuff

(Figure 6). The An content of fresh grains in the intrusion falls between An38 and An62 and plagioclase in the tuff ranged from An40 to An70, though the tuff has a higher proportion of grains at the An-rich end (Woolf, 2008). Sample MIN-0512 is an exception and has plagioclase with

11 notably lower An (An25 to An49) content than all other samples, which suggests it equilibrated at cooler temperatures (representative analyses can be seen in Table 2; for complete analyses see appendix).

Oscillatory zoning of plagioclase is visible in thin section and An varies by as much as

15% within a single grain. Rim to core traverses indicate that normal zoning was present but not very common. Only 8 out of 36 traverses (~22%) show substantial decreases in An toward the rims, while the rest either show small increases of An toward the rim or no pattern at all. Most decreases related to normal zoning are within 10% An; in contrast, the plagioclase in tuff shows the general trend of decreasing An towards the rim on the order of 20% (Woolf, 2008). Even though feldspar grains in sample MIN-0512 are compositionally varied (ranging from An17 to

An30) individual grains do not show significant zoning, which corroborates the conclusion that these crystals did indeed equilibrate completely instead of crystallizing new growth at cooler temperatures. These plagioclase also have slightly lower Or compared to the rest of the samples, which is also suggestive of lower temperatures.

Alkali feldspar is present in two samples: MIN-0512 and MIN-8-60-1. Both of these samples are coarse grained and the alkali feldspar is in the groundmass. These grains are anhedral to subhedral, around 0.5 mm in size, and locally display simple twinning. Only MIN-

0512 was available for microprobe analysis (MIN-8-60-1 was not polished) and alkali feldspar compositions in this sample range from Or75 to Or94, with two distinct clusters (Or75-Or80 and

Or88-Or92; Figure 6a).

Hornblende

In the Wah Wah Springs Intrusion, hornblende is one of the most texturally varied minerals. Hornblende is not found in every sample and is found as euhedral grains in three samples (out of ten thin sections analyzed) and as anhedral, interstitial grains in one sample.

Other samples had no hornblende at all or have hornblende remnants that are filled in with clots of Fe-Ti oxides, clinopyroxene, plagioclase and residual hornblende (Figure 7). In sample

PINTO-1412, hornblende grains have reactions rims that are composed of the same anhydrous minerals found in the hornblende remnants. These rims range in thickness from about 0.05 mm to almost 0.5 mm on grains up to 2.5 across (rim included). No other zoning was observed.

Electron microprobe traverses were made either from rim to rim or rim to core, depending on the size and quality of the grain (Table 3; for complete analyses see appendix).

Most hornblende grains are about 2 mm across. The majority of hornblende compositions are magnesiohornblende with the hornblende in one sample trending towards actinolite (Figure 8).

The hornblende in this sample (MIN-0512) is also texturally different from the others (Figure 7).

These hornblende grains are interstitial and have low Al (~5 wt. % vs. ~8 wt. %) and Ti (~0.9 wt.

% vs. ~1.9 wt. %) compared to the rest of the samples, which is consistent with low-temperature equilibration. Aside from this sample, elemental compositions of hornblende in the tuff and intrusion are very similar, except the intrusion had some grains with higher Al (~1.9 apfu vs. 1.5 apfu) and higher Na (~0.6 apfu vs. ~0.4 apfu). Hornblende grains with rims (sample PINTO-

1412) are chemically similar to euhedral rim-less hornblendes and show no anomalous chemical characteristics, thus suggesting they have the same magmatic origin.

No visible oscillatory zoning was observed but there is slight chemical variation across some of the grains (~2-3 wt. % Al2O3, ~0.5-1% TiO2). Zoning is only distinctive in one sample,

MIN-0312. This sample has hornblende with lower Al and Ti towards the rims, indicating

13 normal zonation and cooler temperatures as the crystal grew (Spear, 1981). This hornblende in this sample is also the most chemically similar to hornblende in the tuff, which contains both unzoned and normally zoned amphiboles (Woolf, 2008). Hornblende in both suites of rock have

TiO2 concentrations around 1.5-2.5 wt. % and Al2O3 around 7-9 wt. %. Hornblende grains with reaction rims lack low Ti and Al rims since their rims have been converted into other minerals, but the core compositions are similar to the cores of other amphiboles.

Variation in IVAl via the edenite and tschermak exchange in hornblende plays an important role in determining changes in pressure and temperature (Shane and Smith, 2013;

Anderson and Smith, 1995). Figure 9 depicts the relevancy of four common Al substitution mechanisms in hornblende. The temperature-sensitive edenite and Ti-Tschermak exchanges play the most significant roles in IVAl variation in these samples, as well as in the tuff (Figure 9a;

Figure 9c). Although the Al-Tschermak exchange (Figure 9b) does not appear very well correlated, this exchange is pressure-sensitive and plays and important role in Al-in-hornblende barometry discussed later in this report.

Pyroxene

Clinopyroxene is present in all but two samples (MIN-0312 and MIN-0412) in the intrusion. Clinopyroxene is mostly subhedral and usually exists as small grains (~0.5 mm) although large grains are not uncommon (~2 mm). No orthopyroxene was found. In the tuff, orthopyroxene only exists as cores in hornblende (Woolf, 2008) and since hornblende in this granodiorite is scarce to begin with, the lack of orthopyroxene seems reasonable. However, hornblende in these samples has often been converted into clinopyroxene: reaction rims on hornblende are composed of clinopyroxene and clinopyroxene has also replaced hornblende in remnant clots (Figure 7).

All clinopyroxene compositions range from augite to diopside, and are similar in composition to clinopyroxene in the tuff, except that more clinopyroxene grains in the intrusion are diopside, especially sample MIN-0512 (Figure 10; Table 4, for complete analyses see appendix). High Ca grains in MIN-0512 are further evidence that this sample has a unique cooling history at lower temperatures. Clinopyroxene in both the intrusion and tuff have similar concentrations of Na (0.01-0.04 apfu), FeT (0.2-0.35 apfu), Al (0-0.1 apfu), Mg (0.7-0.8 apfu), and Mn (0.005-0.025 apfu). Two analyses of pyroxene from reaction rims around hornblende and in remnants in one sample have lower Ca content (~0.9 apfu) but otherwise have somewhat similar, unperturbed compositions. Although hydrothermal alteration has affected many minerals in these rocks, hydrothermal pyroxenes are usually Ca-enriched compared to their igneous counterparts (Manning and Bird, 1986), which implies that these anomalous grains are magmatic in origin. Most other analyses of pyroxene formed by hornblende decomposition fall in line with the main body of analyses (Figure 10).

Biotite

Although biotite is present in all samples of the intrusion, unaltered biotite is scarce and almost all grains show evidence of hydrothermal alteration in the form of oxidization and chloritization. Some electron microprobe analyses were performed on what appeared to be euhedral biotite, but even these showed highly anomalous chemical compositions, and they are probably biotite pseudomorphs that are now completely altered to chlorite. These mica grains are commonly filled with Fe-Ti oxide inclusions along cleavage planes. Consequently they are notably low in Fe, since Fe was mobilized to form the oxides (biotite in the intrusion ~12 wt.%

FeOT vs. tuff ~15.5 wt.%FeOT). Biotite is also increasingly Mg-rich in the intrusion (~16.8 wt.%

MgO vs. tuff~14.2 wt.% MgO); however, both of these changes represent low temperature alteration and are not magmatic (Ferry, 1979; Moore and Czamanske, 1973).

Fe-Ti Oxides

Magnetite is the predominant oxide phase in the intrusion; ilmenite grains are very sparse. Trellis style exsolution is common among the magnetite grains in these rocks but grains or regions of grains selected for analysis were homogenous. Both oxides are highly variable in composition – even within the same sample. The X’usp in magnetite ranges from 0.05 to 0.35 and the X’ilm for ilmenite ranges from 0.51 to 0.86. Compositions of oxides in the intrusion also differ dramatically from those in the tuff, especially in Mn (intrusion ~ 0.19 wt. % MnO vs. tuff

~0.60 wt. % MnO) and Mg (intrusion ~0.08 wt. % MgO vs. tuff ~1.19 wt. % MgO). These abnormal compositions are most likely the result of hydrothermal alteration of the Fe-Ti oxides

(e.g. Razjigaeva and Naumova, 1992).

Quartz

Quartz is present in all samples. Samples MIN-0512 and MIN-8-60-1 had very sparse phenocrystic quartz and most quartz grains in these samples were part of a coarse-grained groundmass. Some samples of the tuff are also quartz-poor (Figure 5). Quartz phenocrysts are up to 2 mm across and small (<0.5 mm) quartz grains comprised most of the matrixes of the intrusive rocks. Quartz grains exhibit textural variety from euhedral to resorbed to highly skeletal secondary grains probably grown during hydrothermal alteration (Figure 11). Phenoclasts of quartz appear in two (of ten) samples but are not very common.

Cathodoluminescence (CL) of quartz was used to interpret textures and derive crystallization histories. Quartz cathodoluminescence images depict zoning that is not seen

16 optically. Although the exact cause of CL variation is debated and could be the result of a number of factors, Ti concentration plays a significant role in the zoning observed with a CL detector (D’lemos et al., 1997; Muller et al., 2000; Muller et al., 2003; Rusk et al., 2008; Muller et al., 2009; Rusk et al., 2011; Vasyukova et al., 2013). Silicon can be replaced by Ti in quartz and areas of higher Ti concentrations are seen as brighter zones with a CL detector. These changes in Ti concentration are indicative of changing conditions within a magma; therefore, quartz zoning provides clues to its magmatic history. Higher Ti concentrations can be a result of several magmatic changes: increase in temperature, decrease in pressure, exsolution, or rapid growth (Wark, 2006; Allan and Yardley, 2007; Rusk et al., 2008; Thomas et al., 2010; Huang and Audetat, 2012;Vasyukova et al., 2013). Aluminum has also been noted to contribute to CL intensity in some cases (Watt et al., 1997; Rusk et al., 2008).

Previous studies on quartz cathodoluminescence were performed on grains from the tuff

(Lindsay, 2010). Of the grains studied from the tuff, most (73%) showed darker rims (Lindsay,

2010; Figure 12a-b). This was interpreted to be the result of crystal growth and cooling prior to eruption. Previously, Woolf (2008) concluded that unlike similar tuffs (e.g. Fish Canyon Tuff from Bachmann et al., 2002) thermal rejuvenation (melting of a nearly completely solidified magma) did not occur prior to eruption for the Wah Wah Springs Tuff since hornblende and plagioclase showed normal zonation (warm cores, cool edges). Quartz with Ti-poor rims corroborates this conclusion for the tuff.

Of all the intrusion grains imaged using the CL detector (n=29 from five samples), six showed more than two zones, thirteen showed two zones, and ten showed no zoning at all where a zone was defined as a notable difference between dark and light (Table 5). Zoning patterns were usually similar among grains in each sample. Wide, transitional zones such as those in

Figure 12c have been ascribed to gradual changes in the melt’s composition or temperature with time (D’Lemos et al., 1997). Sharp changes, however, are a result of more dramatic changes in magmatic conditions or growth in different stages. D’Lemos et al. (1997) suggests these distinct zones are caused by disequilbrium with the magma, either by xenocrystic transport or abrupt changes in magmatic temperature (Figure 12d), but pressure changes might also be the cause

(Huang and Audetat, 2012).

Fourteen (48%) quartz grains have euhedral faces (though some euhedral grains are broken), while the fifteen (52%) are resorbed and seven of the resorbed grains have deep embayments. This number is slightly biased however, since only two samples contained euhedral quartz and both were imaged. In whole, resorbed quartz was much more common on quartz grains in the intrusion. However, zoning clear enough to depict previous quartz grain shapes shows mild resorption and rounding followed by more growth. Deep embayments are only present as a last stage event in the quartz in the intrusion and no evidence of previous deep embayments was observed. This is different than quartz in the tuff, since some of those grains were deeply resorbed and then resumed growth (Figure 12a).

The trend among the majority of samples is for brighter rims (Figure 12d-g). Zoning was observed in 19 phenocrysts (65%). Of these zoned grains, twelve (63%) showed bright rims and five (26%) had dark rims (Figure 12c); the remaining two phenocrysts had sector zoning. It is also important to note that many groundmass quartz grains were also visible in CL images and most were brighter than any phenocrystic quartz (Figure 12b-c). Some of the quartz phenocrysts had thin bright rims that matched the brightness of the groundmass, suggesting the rim and groundmass grew under the same conditions and had higher Ti concentrations (Figure 12d).

Sector zoning is also present in a few quartz phenocrysts (Figure 12h). According to

Vasyukova et al., 2013; Rusk et al., 2008; and Wiebe et al., 2007, this is a result of hydrothermal growth of quartz. Since other signs of hydrothermal alteration are apparent (altered minerals, propylitic assemblage), this is a plausible reason for sector zoning.

Electron microprobe reconnaissance was performed in order to quantify changes in the concentration of titanium and aluminum in quartz phenocrysts. On average, aluminum is on the order of ten times more abundant than titanium; which has been observed in quartz from other rocks (Rusk et al., 2008; Rusk et al., 2011). Aluminum and titanium both roughly correlate to zoning. All analyses that showed high Ti and Al correspond to bright spots in CL imagery.

Analysis beyond this is limited since the concentrations of the elements were very low and hard to reproduce.

Thermobarometry

Thermobarometry proved challenging due to the degree of hydrothermal alteration in most minerals. The altered state of biotite and Fe-Ti oxides made these minerals unusable in thermobarometry calculations, so the choice of thermometers and barometers was limited.

The plagioclase-amphibole thermometer derived by Holland and Blundy (1994) was used by pairing hornblende analyses with low-An plagioclase on the assumption that the hornblende equilibrated with the last crystallizing plagioclase. Since only one sample has significantly zoned hornblende grains, both rim and core compositions were used to calculate overall temperature and pressure averages. Sample MIN-0512 was also not included in overall averages since its anomalously low temperature range (Table 6) provides more evidence it re-equilibrated at lower temperatures (Table 3). Holland and Blundy (1994) provide two thermometers: one that does

19 not require the presence of quartz (A) and one that does (B) (Figure 13). The thermometer that does not require quartz gives temperatures that match the experiments of Johnson and Rutherford

(1989a) and are more reasonable for a granodioritic pluton (791-797° C, avg. 793° C ± 20° C), while the thermometer that requires quartz gives higher temperatures (802-847° C, avg.

816° C±31° C). Similar temperatures ranges are calculated with this for the tuff (A: 773-797 ° C, avg. 791± 20° C; B: 812-834° C, avg. 818± 29° C) (Woolf, 2008).

The pressure-sensitive Al-Tschermak exchange is important in determining the accuracy and effectiveness of the Al-in-hornblende thermometer. Shane and Smith (2013) found the Al-in- amphibole barometers produce erroneous results when this exchange does not play a large role in the variability of hornblende in a sample. The Al-Tschermak exchange does not play a noteworthy role in any hornblende in these samples and is especially insignificant for sample

MIN-0512, which had low Al (Figure 9). However, it is possible that only one pressure prevailed and that the perceived variation is just scatter around one point representing a single pressure. If this is the case, then no strong correlation between VIAl and IVAl would be expected. Also, by subdividing the total aluminum into VIAl and IVAl, much of the analytical error associated with the other elements is compounded into the calculated VIAl which could make this scatter more exaggerated than it really is. There is another drawback of the Al-in-hornblende barometer, however, and that is that high temperature and low oxygen fugacity can lead to inflated pressure estimates (Blundy and Holland, 1990; Anderson and Smith, 1995). However, a mineral assemblage of quartz, alkali feldspar, plagioclase, hornblende, biotite, ilmenite, magnetite and titanite help negate these effects and moderate the barometer. The barometer of Johnson and

Rutherford (1989b) gives some unreasonably low pressure estimates, although the overall average is acceptable (2.3-2.6 kb, avg. 2.4±0.6 kb). The thermometer of Anderson and Smith

(1995) also gives very low and sometimes unreasonable (negative) pressure estimates (1.0-1.9 kb, avg. 1.7±0.8 kb). Pressure estimates for the tuff using hornblende compositions and the formulation of Johnson and Rutherford (1989b) were more similar (2.0-2.5 kb, avg. 2.2 kb), and slightly higher than Anderson and Smith (1995) (1.6-2.1 kb, avg. 1.8 kb) (Woolf, 2008).

Ridolfi et al. (2010) derived a method of calculating temperature, pressure, fO2, and H2O melt using solely hornblende compositions. This thermometer is Si-sensitive and was calculated using a Si index (Ridolfi et al., 2010). It gives significantly higher temperatures than the Holland and Blundy (1994) thermometer for the intrusion as well as the tuff (tuff: 788-914° C, avg.

861±28 ° C; intrusion: 753-909° C, avg. 849±25° C; Table 6). Ridolfi et al.’s (2010) barometer depends on Al like other barometers but its uncertainties increase with pressure and decrease with temperature. It produced pressures much lower than Johnson and Rutherford (1989b) (tuff:

1.5±0.3 kb; intrusion: 1.5±0.5). The hygrometer is also Al-sensitive and is most accurate for magnesiohornblendes (which these hornblendes are); the average estimate for the intrusion was

4.1±0.4 wt. % H2O which accords with the phase equilibrium experiments of Johnson and

Rutherford, 1989a) on the compositionally similar Fish Canyon Tuff.

Oxygen fugacity calculations were largely based on Mg proportions and gave an average estimate for log fO2 of -11.5±0.3 for the intrusion which is about 2 log units above the QFM buffer at the temperature given by their thermometer. Estimates for fO2 in the tuff were similar using the Ridolfi et al. (2010) method (-11.6±0.1 log units).

With conflicting pressure and temperature estimates, it is difficult to decide which estimates are best (Figure 13). Based on a series of experiments, Johnson and Rutherford (1989a) put the probable temperature and pressure range for the Fish Canyon Tuff between 775-800° C

21 and 2.0-2.5 kb. Since the Wah Wah Springs Tuff and intrusion are similar to this tuff both chemically and genetically, it is useful to see which thermometers gave temperatures that correlate with these ranges. Of the three thermometers, Holland and Blundy (1994) thermometer

A (without quartz) gives the most plausible temperatures and is most similar to the Fish Canyon

Tuff experiments. These temperatures also agree most with the temperatures calculated from Fe-

Ti oxides in the tuff (Woolf, 2008). At the higher T’s calculated using the equation of Ridolfi et al. (2010) quartz would not be stable according to the experiments of Johnson and Rutherford

(1989a). Perhaps because several geobarometers rely on the Al-Tschermak exchange and it is poorly correlated here (and in other magmatic rocks), the barometer of Johnson and Rutherford

(1989b), which relies on total Al in hornblende, is more reliable. The barometers of Ridolfi et al.

(2010) and Anderson and Smith (1995) give lower estimates than expected based on the Fish

Canyon Tuff experiments (Johnson and Rutherford, 1989a. A pressure around 2.4 kb (intrusion average) would put the depth of crystallization between 7-9 km, which is common for an intermediate composition subduction-related magma (Rutherford and Hill, 1993; Johnson and

Rutherford 1989a).

DISCUSSION

Comparison of the Wah Wah Springs Intrusion to the Wah Wah Springs Tuff

Chemical and textural similarities between the tuff and intrusion are important in determining their relationship and crystallization history. Chemically, the tuff and intrusion are quite comparable. Both can be classified as dacites and granodiorites and in general their whole rock chemistry is very similar, especially in comparison to other tuffs from the same caldera complex. The Wah Wah Springs Tuff has been noted for its high Cr content in comparison to a

22 dozen other ignimbrites from the Indian Peak and Caliente calderas (Best et al., 2013). The intrusion also has the same distinctively high Cr concentration, suggesting that these rocks are from the same magma.

Modally, the tuff and intrusion are also similar and most differences can be explained by complete crystallization of the pluton or by pressure changes (explained below). For the most part, both rocks contain the same major minerals. One exception is that the intrusion lacks orthopyroxene (only found as cores of hornblende in the tuff.) Another notable difference is the presence of visible grains of alkali feldspar in the groundmass of two samples in the intrusion

(MIN-0512 and MIN-8-60-1). These two samples were collected from the same isolated patch of granodiorite and are both relatively coarse-grained (groundmass grains up to 0.5 mm) where alkali feldspar was the last mineral to crystallize. Alkali feldspars analyzed in MIN-0512 show two populations (Figure 6), one of which may represent a normal magmatic composition (Or75-

Or80) and one that represents a hydrothermally altered composition (Or88-Or92). This suggests that at least of some the K-feldspar is igneous in origin and later it was altered. Using the magmatic composition of K-feldspar, another temperature for this sample can be calculated using a two-feldspar thermometer (Nekvasil and Burnham, 1987): this method yields a temperature of 602° C, which again suggests low-temperature re-equilibration. Although samples with K-feldspar were sparse, it is likely that the tuff, as well as other samples of the intrusion, contain groundmass K-feldspar that was too small to analyze or formed by devitrification of the tuff. However, neither the intrusion nor the tuff contains phenocrystic alkali feldspar. Accessory minerals in the two deposits are also similar. It was previously cited that the Wah Wah Springs

Tuff did not contain titanite (Woolf, 2008) but heavy mineral separates confirm the presence of trace amounts of titanite in some samples of the tuff. One sample of the intrusion also contains

23 trace amounts of this mineral and was also found during the separation of heavy minerals.

Accessory zircon and apatite are present in both the tuff and intrusion.

Aside from the hydrothermally altered biotite and Fe-Ti oxides, mineral compositions in the two suites of rocks are similar. Plagioclase compositions are all between An38 and An70.

Most hornblende grains are magnesiohornblendes. Clinopyroxene compositions are also similar between the rocks except that the intrusion has a higher proportion of diopside grains than the tuff (Figure 10). Diopside is more Ca-rich and is thought to reflect crystallization at lower temperatures (Lindsley, 1983). It is likely that pyroxene continued crystallizing after the tuff erupted. Since pyroxene is stable at lower pressures than hornblende, it replaced amphibole as the stable mafic phase and crystallized at lower pressures and thus lower temperatures causing its more diopsidic composition (Figure 10).

Sample MIN-0512 is different from both the tuff and intrusion and shows chemical and textural evidence of re-equilibration or further differentiation. The whole rock geochemistry of this rock shows that it is generally more enriched in incompatible elements (e.g. Rb, Th, K, Na) and depleted in compatible elements (e.g. Fe, Mg, Zr) than other samples from the intrusion

(Table 1), Apparently, this part of the intrusion is slightly more evolved than the rest. The plagioclase in this sample also has lower anorthite content. Since anorthitic plagioclase is more likely to crystallize at higher temperatures, low An suggests this sample re-equilibrated at lower temperatures. This plagioclase also has slightly lower Or than the others, which is also consistent with the other indicators of lower temperature crystallization or re-equilibration. Hornblende in

MIN-0512 is interstitial and has significantly lower TiO2 and Al2O3 (~0.4-1.2 wt. % TiO2; ~4-6 wt.% Al2O3; Table 3). A study by Chambefort et al. (2013) on volcanic rocks from Yanacocha,

Peru shows that low-Al hornblendes crystallized at cooler temperatures and lower pressures so it

24 is likely that this portion of the Wah Wah Springs Intrusion did as well. A significant portion of the diopsidic clinopyroxene is from sample MIN-0512 as well and provides further evidence of this sample’s re-equilibration to cooler temperatures or crystallization at cooler temperatures

(Figure 10). Apparently, minerals in MIN-0512 re-equilibrated at cooler temperatures because this sample crystallized near the roof, against cold Paleozoic limestone while most of the granodiorite intruded the Wah Wah Springs Tuff, which was possibly still warm. This sample also has more calcic hornblende and pyroxene; small calcite crystals in thin section are evidence that it was contaminated with calcite.

Magma Decompression

The evidence suggests that the hornblende reactions rims found in PINTO-1412 were caused by decompression and magma re-equilibration outside the hornblende stability field

(Spear, 1981; Rutherford and Hill, 1993; Devine et al., 1998; Browne and Gardner, 2006).

Apparently, as the magma erupted, decompressed, and devolatilized, hornblende began to change into anhydrous minerals, namely, clinopyroxene, plagioclase, and Fe-Ti oxides. Reaction rims can be used to determine magma ascent rates since the thickness of a rim corresponds to this rate, or more generally, the time a hornblende crystal spent above its stability field before being quenched. In other words, thicker reaction rims indicate a slower magma ascent rate or more time spent at lower pressures beyond the hornblende stability field.

Rutherford and Hill (1993) used reaction rims on amphiboles to determine magma ascent rates for Mount St. Helens eruptions. Using decompression experiments that modeled an ascent of hornblende in a dacite from a depth of 8 km to the surface along a 900° C adiabat, they found that that no reaction rims developed on amphiboles within 4 days, a 10 micron reaction rim

25 developed in 10 days, and a 35 micron reaction rim developed during a 20 day decompression.

Browne and Gardner (2006) performed similar experiments on hornblende from the Redoubt volcano in Alaska. They formed reaction rims on hornblende in both single step (sudden pressure change) experiments and multiple step (gradual pressure change) experiments on an 840° C adiabat. Reaction rims in single step experiments formed between 0.6-0.7 kb (~1.2-2.3 km depth). Reaction rims in multiple step experiments formed between 0.1-0.4 kb (~0.2-1.3 km depth). All of these experiments used dacitic magma similar in composition to the Wah Wah

Springs Tuff.

The reaction rims in sample PINTO-1412 are much thicker than the rims in the experiments of Rutherford and Hill (1993) or Browne and Gardner (2006). While those rims are

<35 microns thick, reaction rims of PINTO-1412 range from 50 to 500 microns; other samples have entire millimeter-sized grains that are completely destroyed. Assuming similar conditions, this would imply slow ascent rates (year-scale) or that the magma stalled at a shallow depth beyond hornblende stability limits before complete crystallization. Johnson and Rutherford

(1989a) determined hornblende stability relations for the Fish Canyon Tuff of Colorado. It is chemically and genetically very similar to the Wah Wah Springs Tuff, so using these stability estimates a P-T path for the intrusion-forming Wah Wah Springs magma can be hypothesized

(Figure 14). The sample with reactions rims (as opposed to complete hornblende breakdown) is a near a contact and was probably quenched before the grains had to time to completely convert into other minerals. It is also important to note that the hornblende in the tuff has no reaction rims, indicating the eruptive magma rose from depths of 8 km or so in less than a few days

(Devine et al., 1998). Such rise times are consistent with the rapid eruption rates estimated by

Hildreth and Wilson (2007) for the Bishop Tuff of the Long Valley caldera in California. They

26 estimate that roughly 600 km3 of magma erupted from a depth between 5-11 km in a matter of six days.

Quartz resorption also provides clues as to what happened to this magma body. Resorbed quartz is common in samples of the intrusion and is suggestive of environmental changes (e.g.

Wiebe et al., 2007). Resorbed quartz is also very common in the tuff, so it is likely that the same factor caused this feature in both the tuff and intrusion. In several recent papers (Bachmann et al., 2002; Wark et al., 2007; Wiebe et al., 2007), quartz resorption has been linked to thermal rejuvenation, but since previous studies have found no other evidence of thermal rejuvenation of the Wah Wah Springs magma body (Woolf, 2008), this is most likely not the case. According to the experiments of Johnson and Rutherford (1989a) however, quartz saturation is very sensitive to water saturation and is unstable in water-saturated magma at the pressure and temperature estimates for the Wah Wah Springs magma; therefore, quartz resorption may be due to water- saturation in this magma as it became water-saturated during decompression and eruption

(Figure 14). Others have concluded that decompression alone can cause quartz resorption (e.g.

Nekvasil, 1991; Muller et al., 2009) and it is likely that water saturation and decompression were happening simultaneously.

Bright quartz rims and groundmass grains in CL imagery are important in determining the rest of the crystallization history of the intrusion. These bright features indicate a second stage of quartz growth when quartz was again stable after its previous period of instability and partial resorption. Semi-quantitative electron microprobe analyses showed that the bright zones in quartz are Ti-rich and also Al-rich. Al-enrichment could be due to hydrothermal growth (Rusk et al., 2008; Rusk et al., 2011) but this phenomenon is relatively understudied and in some cases the Ti and Al in quartz do not correlate (Vasyukova, 2013). Ti-enrichment of quartz phenocryst

27 rims and groundmass crystals, however, could be the result of several magmatic changes (see

Thomas et al., 2010; Allan and Yardley, 2007; Wark and Watson, 2006; Vasyukova et al., 2013;

Rusk et al., 2008). Ti-rich quartz is often cited as evidence for thermal rejuvenation of near solidus magma, but given the lack of evidence for thermal rejuvenation found by Woolf (2008) this seems unlikely. There is, however, evidence for decompression of this magma, which could cause Ti-enrichment of late crystallized quartz (Thomas et al., 2010, Huang and Audetat, 2012).

Therefore the most reasonable explanation is that the quartz rims and groundmass grains grew in a lower-pressure environment than the earlier-formed quartz phenocrysts. Thomas et al. (2010) and Huang and Audetat (2012) have experimentally derived Ti-in-quartz barometers that illustrate this. Using a fixed TiO2 activity of 1 and a temperature of 793° C (the preferred temperature), Figure 15 shows the relationship between pressure and Ti concentration (ppm) in quartz if it crystallized at depths ranging from ~7-8 km depth to a shallow depth. There is a clear negative correlation with both barometers: decreasing pressure results in significant increases of

Ti in quartz.

This evidence suggests that the bright rims indicate a second stage of quartz growth at lower pressure that accompanied the shallow emplacement of the Wah Wah Springs magma into the caldera filling tuff. Fluctuating water saturation in the magma or decompression may be responsible for both quartz resorption and for this second stage of growth at lower temperatures as the magma completely solidified. Previous studies have found that, at least for high- temperature quartz, low water is necessary for growth (Wark et al., 2007, Wiebe et al., 2007).

The Wah Wah Springs magma was not originally water-saturated, allowing the first stage of growth. Johnson and Rutherford (1989a) estimated that a “water-saturated” magma at 2 kb and

800° C would contain about 5.5 wt.% H2O. Using the hygrometer of Ridolfi et al., (2010), the

H2O melt wt. % of the Wah Wah Springs magma was calculated to be between 3.6-4.3 wt. %

(Table 6), thus making it water undersaturated at these conditions. As the lid to the magma chamber broke, the magma decompressed and became water-saturated which drove the voluminous eruption (Christiansen, 2005; Blake, 1984). This caused the quartz to become unstable and resorb (Figure 14). As eruption continued, the residual magma became water-poor.

Ongoing eruption eventually depleted the magma of its volatiles, leaving behind a rising body of non-explosive mush. The magma body, ascending in the caldera fill at depths less than 5 km, eventually stalled at shallow depth. This stalled crystal mush cooled allowing quartz to once again crystallize in a water-poor environment at lower pressures. These conditions were conducive to Ti-enrichment of quartz rims and late groundmass quartz (and eventually the growth of alkali feldspar). However, these conditions were not conducive to the growth of hornblende, and it is here that it decomposed to rims or masses of plagioclase, clinopyroxene, and Fe-Ti oxides in this low pressure, water-poor environment. The ascent of the magma also caused resurgence of the Indian Peak Caldera according to Best et al., (2013).

Magma Chamber and Caldera Development

The size of the magma chamber is an important factor in determining the relationship between the tuff and intrusion as well as whether or not the intrusion grew incrementally. It has been hypothesized that calderas reflect the size of the magma chamber, since the chamber roof falls into the area previously supported by the magma body (Radebaugh, 1999). Caldera size would indicate that the Wah Wah Springs magma chamber was indeed large, since the caldera has an area of approximately 1000 km2 and a diameter of over 40 km The granodiorite intrusion probably represents the final stages of what was once a large magma chamber since it is

29 geochemically complementary to the tuff which had an estimated volume of 5900 km3 (Best et al., 2013).

Figure 16 shows the step-by-step process of how this magma chamber affected caldera development and resurgence, specifically along the northern ring fault (Figure 2). The Wah Wah

Springs Tuff began erupting before the caldera formed. As the chamber began to empty, subsidence began and a potential void along the ring fault created a passageway for more magma. Erupting tephra was then trapped in the growing caldera and created a thick deposit of lithic-rich intracaldera tuff. Walls continued to collapse along reverse faults and formed collapse breccias, creating more intracaldera ejecta through which the tuff had to erupt. Trap-door style subsidence caused a maximum caldera depth of about 4.5 km and a large ring vent through which magma could continue to erupt. As the magma continued erupting though, it became volatile-poor and lost the driving force needed for eruption. What magma was left behind became the Wah Wah Springs Intrusion which rose and caused resurgence by reversing displacement on the preexisting ring faults. Vertical foliations of some portions of the tuff outcrop provide evidence that the previously horizontal tuff was domed. Sometime during this process the granodioritic magma stalled at shallow depths beyond the hornblende limits, probably only a few kilometers deep and began to crystallize Ti-rich quartz. Seismic studies done on more modern volcanoes (i.e. the Valles Caldera in New Mexico) show evidence that magma can get left behind in the magma chamber (Wilcock et al., 2013). These studies indicated that a cylindrical body was underlying the Valles Caldera, which was interpreted as residual magma left in the chamber. This supports our conclusion that magma was also been left behind in the Wah Wah Springs chamber. Furthermore, a study by Lipman (1984) on various tuff and intrusions in western North America, found that silicic intrusions are often a cause of resurgence,

30 especially in calderas with a > 10 km diameter, so this is a likely story for the Indian Peak

Caldera as well.

CONCLUSIONS

Chemical, modal, textural, and field evidence show that the Wah Wah Springs Tuff and the intracaldera Wah Wah Springs Intrusion are cogenetic. Whole rock geochemistry shows that both suites of rocks are dacites ranging from 62.5 to 70.0 wt. % SiO2 and both have distinctive concentrations of Cr and Sr that link them, especially in comparison to other tuffs from the same volcanic field. Compositions of magmatic plagioclase, hornblende, and pyroxene are indistinguishable. One sample (MIN-0512) experienced low-temperature hydrothermal alteration to form distinctive hornblende, feldspar, and pyroxene; this sample intruded cold Paleozoic carbonates instead of intracaldera tuff.

Hornblende in the intrusion is texturally varied, with some parts containing euhedral hornblende and other parts containing hornblende with reaction rims. But, most of the pluton has remnants of hornblende that have been converted into clinopyroxene, plagioclase, and Fe-Ti oxides. This hornblende decomposition is due to rapid ascent of the magma during the eruptive phase followed by stalling at a shallow depth beyond hornblende stability limits as the granodiorite finished its solidification.

Resorbed quartz phenocrysts are also an indicator of dynamic magmatic conditions.

Resorption occurred either during decompression or when the magma became water saturated during eruption. A second stage of quartz crystallization at low pressure is revealed by CL-bright phenocryst rims and groundmass grains and is suggestive of late Ti-enrichment. Ti concentrations in quartz increase with decreasing pressure so it appears that these quartz grains

31 crystallized as a late-stage event after the magma had already risen to a shallower depth and lost most of its water.

Overall, textural, structural, chemical, and mineralogical evidence suggest the Wah Wah

Springs Tuff and intrusion came from the same magma. After the tuff erupted some magma was left in the chamber. This magma did not erupt but rose to a shallow level and became the Wah

Wah Springs Intrusive Granodiorite Porphyry. The general sequence of events is as follows:

1. About 30 Ma, a large magma chamber filled with more than 6000 km3 of dacitic magma

at a depth of about 8 km.

2. The roof of the chamber broke and the magma became water-saturated and erupted

explosively forming the outflow tuff member of the Wah Wah Springs Formation.

3. The Indian Peak caldera collapsed as tuff continued erupting but was trapped in the

caldera.

4. Ongoing eruption caused the deeper magma to lose volatiles, leaving behind a dry mush.

5. This rising mush became the Wah Wah Springs Intrusion. As the eruption ceased, it

stalled at low pressure and caused resurgence of the Indian Peak Caldera. Hornblende

converted into anhydrous minerals and a second generation of high Ti quartz crystallized

at lower pressure during final cooling and solidification of the mush.

6. Some portions of magma re-equilibrated at lower temperatures, especially those that had

intruded colder limestone. Propylitic alteration of much of the intrusion destroyed

magmatic biotite and Fe-Ti oxides.

7. Much later, Basin and Range normal faulting and erosion exposed the intrusion.

The genesis of the Wah Wah Springs Intrusion adds more insight to the controversy of the relationship between plutons and volcanic rocks and about pluton amalgamation. Large volumes

(thousands of cubic kilometers) of magma were produced and present at least long enough to feed the eruption of the Wah Wah Springs Tuff. This suggests that even if the chamber filled incrementally it did not cool completely between insertions of new magma. Also, there was very little differentiation between the tuff and intrusion, or within the shallow intrusion itself. Judging from its discontinuous exposures, the shallow intracaldera pluton could extend across the entire floor of the caldera—which is more than 40 km across. The evidence points to the conclusion that this shallow intracaldera pluton was emplaced rapidly and did not assemble incrementally.

Thus, the Wah Wah Springs Intrusion represents the final stages of a large magma chamber that did not empty completely during volcanism.

REFERENCES

Allan, M.M., and Yardley, B.W.D., 2007, Tracking meteoric infiltration into a magmatic- hydrothermal system: A cathodoluminescence, oxygen isotope and trace element study of quartz from Mt. Leyshon, Australia. Chemical Geology, v. 240, p. 343-360.

Anderson, J.L., and Smith, D. R., 1995, The effects of temperature and fO2 on the Al-in- hornblende barometer: American Mineralogist, v. 80, p. 549-559.

Bachmann, O., Dungan, M.A. & Lipman, P.W., 2002, The Fish Canyon magma body, San Juan volcanic field, Colorado: rejuvenation and eruption of an upper crustal batholith: Journal of Petrology, v. 43, p. 1469-1503.

Bachmann, O., and Bergantz, G., 2008, The magma reservoirs that feed supereruptions: Elements, v. 4, p. 17-21.

Best, M.G., Shuey, R.T., Caskey, C.F., Grant, S.K., 1973, Stratigraphic relations of members of the Needles Range Formation at type localities in southwestern Utah: Geological Society of America Bulletin, v. 84, p. 3269-3278.

Best, M.G., Christiansen, E.H., and Blank, R.H., 1989, Oligocene caldera complex and calc- alkaline tuffs and lavas of the Indian Peak volcanic field, Nevada and Utah: Geological Society of America Bulletin, v. 101, p. 1076-1090.

Best, M.G., Christiansen, E.H., Deino, A.L., Gromme, S., Hart, G.L., and Tingey, D.G., 2013, The 36-18 Ma Indian Peak-Caliente ignimbrite field and calderas, southeastern Great Basin, USA: Multicyclic super-eruptions: Geosphere, in press.

Best, MG., and Grant, S.K., 1987, Stratigraphy of the volcanic Oligocene Needles Range group in southwestern Utah and eastern Nevada: U.S. Geological Survey Professional Paper 1433-A, p.3-28

Best, M.G., Grant, S.K., Hintze, L.F., Cleary, J.G., Hutsinpiller, A., and Saunders, D.M., 1987a, Geologic Map of the Indian Peak (Southern Needle) Range, Beaver and Iron Counties, Utah. U.S. Geological Survey Miscellaneous Investigation Series Map I-1795, scale: 1:50,000, 1 sheet.

Best, M.G., Hintze, L.F., Holmes, R. D., 1987b, Geologic map of the southern Mountain Home

and northern Indian Peak ranges (Central Needle Range), Beaver County, Utah: U.S. Geological Survey, Miscellaneous Investigations Series Map I-1796, scale: 1:50,000, 1 sheet

Best, M.G., and Christiansen, E.H., 1991, Limited extension during peak Tertiary volcanism, Great Basin of Nevada and Utah: Journal of Geophysical Research, v. 96, p. 13509- 13528.

Best, M.G., Barr, D.L., Christiansen, E.H., Gromme, S., Deino, A.L., Tingey, D.G., 2009, The Great Basin Altiplano during the middle Cenozoic ignimbrites flareup: insights from volcanic rocks: International Geology Review, v. 51, p. 589-633.

Blake, S., 1984, Volatile oversaturation during the evolution of silicic magma chambers as an eruption trigger: Journal of Geophysical Research, v. 89, p. 8237-8244.

Blundy, J.D., and Holland, T.J.B., 1990, Calcic amphibole equilibria and a new amphibole- plagioclase geothermometer: Contributions to Mineralogy and Petrology, v. 104, p. 208- 224.

Browne, B.L., and Gardner, J.E., 2006, The influence of magma ascent path on the texture, mineralogy, and formation of hornblende reaction rims: Earth and Planetary Science Letters, v. 246, p.161-176.

Chambefort, I., Dilles, J.H., and Longo, A.A., 2013, Amphibole geochemistry of the Yanacocha Volcanics, Peru: Evidence for diverse sources of magmatic volatiles related to gold ores: Journal of Petrology, v.54, 30 p.

Christiansen, E.H., 2005, Contrasting processes in silicic magma chambers: evidence from very large volume ignimbrites: Geological Magazine, v. 6, p. 669-681.

Christiansen, E.H., and McCurry, M., 2008, Contrasting origins of Cenozoic silicic volcanic rocks from the western Cordillera of the United States: Bull Volcanology, v. 70, p. 251- 267.

De Silva, S.L., and Gosnold, W.D., 2007, Episodic construction of batholiths: Insights from the spatiotemporal development of an ignimbrite flare-up: Journal of Volcanology and Geothermal Research, v. 167, p. 320-335.

Devine, J.D., Rutherford, M.J., and Gardner, J.E., 1998, Petrologic determination of ascent rates

for the 1995-1997 Soufriere Hills Volcano andesitic magma: Geophysical Research Letters, v.25, p.3673-3673.

D’Lemos, R.S., Kearlsey, A.T., Pembroke, J.W., Watt, G.R., and Wright, P., 1997, Complex quartz growth histories in granite revealed by scanning cathodoluminescence techniques: Geological Magazine, v. 134, p.549-552.

Ferry, J.M., 1979, Reaction mechanisms, physical conditions, and mass transfer during hydrothermal alteration of mica and feldspar in granitic rocks from south-central Maine, USA: Contributions to Mineralogy and Petrology, v. 68, p. 125-139.

Frost, B.R., Barnes, C.G., Collins, W.J., Arculus, R.J., Ellis, D.J., and Frost, C.D., 2001, A geochemical classification for granitic rocks: Journal of Petrology, v. 42, p. 2033-2048.

Glazner, A.F., Bartley, J.M., Coleman, D.S., Gray, W., Taylor, R.Z., 2004, Are plutons assembled over millions of years by amalgamation from small magma chambers?: GSA Today, v. 14, p. 4-11.

Henry, C.D., and John, D.A., 2013, Magmatism, ash-flow tuffs, and calderas of the ignimbrite flareup in the western Nevada Volcanic Field, Great Basin, USA: Geosphere, in press.

Hildreth, W., 2004, Volcanological perspectives on Long Valley, Mammoth Mountain, and Mono Craters: several contiguous but discrete systems: Journal of Volcanology and Geothermal Research, v. 136, p. 169-198.

Hildreth, W., and Wilson, C.J.N., 2007, Compositional zoning of the Bishop Tuff: Journal of Petrology, v. 48, p. 951-999.

Holland, T., and Blundy, J., 1994, Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry: Contributions to Mineralogy and Petrology, v. 116, p. 433-447.

Huang, R., and Audetat, A., 2012, The titanium-in-quartz (TitaniQ) themobarometer: A critical examination and re-calibration: Geochimica et Cosmochimica Acta, v. 84, p. 75-89.

Humphreys, E.D., Hessler, E., Dueker, K., Farmer, G.L., Erslev, E., Atwater, T., 2003, How Laramide-age hydration of North American lithosphere by the Farallon slab controlled subsequent activity in the western United States: International Geology Review, v. 45, 21 p.

John, D.A., 1995, Tilted middle Tertiary ash-flow calderas and subjacent granitic plutons, southern Stillwater Range, Nevada: Cross-sections of an Oligocene igneous center: Geological Society of America Bulletin, v. 107, p. 180-200.

John, D.A., Henry, C.D., Colgan, J.P., 2008, Magmatic and tectonic evolution of the Caetano caldera, north-central Nevada: A tilted, mid-Tertiary eruptive center and source of the Caetano Tuff. Geosphere, v. 4, p. 75-106.

Johnson, M.C., and Rutherford, M.J., 1989a, Experimentally determined conditions in the Fish Canyon Tuff, Colorado magma chamber: Journal of Petrology, v. 30, p. 711-737.

Johnson, M.C., and Rutherford, M.J., 1989b, Experimental calibration of the aluminum in hornblende geobarometer with application to Long Valley Caldera (California): Geology, v. 17, p. 837-841.

Leake, B.E., Woolley, A.R., Arps, C.E.S., Birch, W.D., Gilbert, M.C., Grice, J.D., Hawthorne, F.C., Kato, A., Kisch, H.J., Krivovichev, V.G., Linthout, K., Laird, J., Mandarino, J.A., Maresch, W.V., Nickel, E.H., Rock, N.M.S., Shumacher, J.C., Smith, D.C., Stephenson, N.C.N., Ungaretti, L., Whittaker, E.J.W., and Youzhi, G., 1997, Nomenclature of amphiboles: Report of the subcommittee on amphiboles of the International Mineralogical Association, commission on new minerals and mineral names. The Canadian Mineralogist, v. 35, p. 219-246.

Le Bas, M.J., LeMaitre, R.W., Streckeiesen, A., and Zanettin, B., 1986, A chemical classification of volcanic rocks based on the total alkali-silica diagram: Journal of Petrology, v. 27, p. 745-750.

Lindsay, A., 2010, Thermal characteristics of the Wah Wah Springs ignimbrite quartz grains: utilizing cathodoluminescence imagery [senior thesis]: Brigham Young University, 18 p.

Lindsley, D.H., 1983, Pyroxene thermometry: American Mineralogist, v. 68, p. 47-493.

Lipman, P.W., Prostka, H.J., Christiansen, R.L., 1971, Evolving subduction zones in the western United States, as interpreted from igneous rocks: Science, v. 174, p. 821-825.

Lipman, P.W., 1984, The roots of ash flow calderas in western North American: windows into the tops of granitic batholiths: Journal of Geophysical Research, v. 89, p. 8801-8841.

Lipman, P.W., 2007, Incremental assembly and prolonged consolidation of Cordilleran magma

chambers: Evidence from the Southern Rocky Mountain volcanic field: Geosphere, v.3, p. 42-70.

Mackin, J.H., 1960, Structural significance of Tertiary volcanic rocks in southwestern Utah: American Journal of Science, v. 258, p. 81-131.

Manning, C.E., and Bird, D.K., 1986, Hydrothermal clinopyroxenes of the Skaergaard intrusion: Contributions to Mineralogy and Petrology, v. 92, p. 437-477.

Maughan, L.L., Christiansen E.H., Best, M.G., Gromme, C.S., Deino, A.L., Tingey, D.G., 2002, The Oligocene Lund Tuff, Great Basin, USA: a very large volume monotonous intermediate: Journal of Volcanology and Geothermal Research, v. 113, p. 129-157.

Moore, W.J., and Czamanske, G.K., 1973, Compositions of biotites from unaltered and altered monzonitic rocks in the Bingham Mining District, Utah: Economic Geology, v. 68, p. 269-280.

Muller, A., Seltmann, R., Behr, H.J., 2000, Application of cathodoluminescence to magmatic quartz in a tin granite – cast study from the Schellerhau Granite Complex, Eastern Erzgebirge, Germany: Mineralium Deposita, v. 35, p. 169-189.

Muller, A., Wiedenbeck, M., Van den Kerkhof, A.M., Kronz, A., Simon, K., 2003, Trace elements in quartz-a combined electron microprobe, secondary ion mass spectrometry, laser-ablation ICP-MS, and cathodoluminescence study: European Journal of Mineralogy, v. 15, p. 747-763.

Muller, A., van den Kerkhof, A.M., Behr, H.J., Kronz, A., Koch-Muller, M., 2009, The evolution of late-Hercynian granites and rhyolites documented by quartz – a review: Earth and Environmental Science Transactions of the Royal Society of Edinburgh, v. 100, p. 185- 204.

Nekvasil, H., and Burnham, C.W., 1987, The calculated individual effects of pressure and water content on phase equilibria in the granite system: Magmatic processes: physicochemical principles: Special Publication- Geochemical Society, v. 1, p. 433-445.

Nekvasil, H., 1991, Ascent of felsic magmas and formation of rapakivi: American Mineralogist, v. 76, p. 1279-1290.

Petford, N., Cruden, A.R., McCaffrey, K.J.W., and Vigneresse, J.-L., 2000, Granite magma formation, transport and emplacement in the Earth’s crust: Nature, v. 408, p. 669–673.

Radebaugh, J., 1999, Terrestrial pluton and planetary caldera sizes: implications for the origin of Calderas: M.S. Thesis, Brigham Young University, Provo, UT.

Razjigaeva, N.G., and Naumova, V.V., 1992, Trace element composition of detrital magnetite from coastal sediments of northwestern Japan sea for provenance study: Journal of Sedimentary Petrology, v. 62, p. 802-809.

Ridolfi, F., Renzulli, A., Puerini, M., 2010, Stability and chemical equilibrium of amphibole in calc-alkaline magmas: an overview, new thermobarometric formulations and application to subduction-related volcanoes: Contributions to Mineralogy and Petrology, v.160, p. 45-66.

Rusk, B.G., Lowers, H.A., Reed, M.H., 2008, Trace elements in hydrothermal quartz: Relationships to cathodoluminescent textures and insights into vein formation: Geology, v. 36, p. 547-550.

Rusk, B., Koenig, A., Lowers, H., 2011, Visualizing trace element distribution in quartz using cathodoluminescence, electron microprobe, and laser ablation-inductively coupled plasma-mass spectrometry: American Mineralogist, v. 96, p. 703-708.

Rutherford, M.J., and Hill, P.M., 1993, Magma ascent rates from amphibole breakdown: and experimental study applied to the 1980-1986 Mount St. Helens eruptions: Journal of Geophysical Research, v. 98, p. 19667-19685.

Shane, P., and Smith, V.C., 2013, Using amphibole crystals to reconstruct magma storage temperatures and pressures for the post-caldera collapse volcanism at Okataina volcano: Lithos, p. 159-170.

Shaw, H.R., 1985, Links between magma-tectonic rate balances, plutonism, and volcanism: Journal of Geophysical Research, v. 90, p. 11275-11288.

Spear, F.S., 1981, An experimental study of hornblende stability and compositional variability in amphibolite: American Journal of Science, v. 281, p. 697-734.

Tappa, M.J., Coleman, D.S., Mills, R.D., Samperton, K.M., 2011, The plutonic record of a silicic ignimbrite from the Latir volcanic field, New Mexico: Geochemistry, Geophysics, and Geosystems, v. 12, 16 p.

Thomas, J.B., Watson, E.B., Spear, F.S., Shemella, P.T., Nayak, S.K., Lanzirotti, A., 2010,

TitaniQ under pressure: the effect of pressure and temperature on the solubility of Ti in quartz: Contrib Mineral Petrol, 17 p.

Vasyukova, O.V., Goemann, K., Kamanetsky, V.S., MacRae, C.M., and Wilson, N.C., 2013 Cathodoluminescence properties of quartz eyes from porphyry-type deposits: Implications for the origin of quartz: American Mineralogist, v. 98, p. 98-109.

Verplanck, P.L., Farmer, G.L., and McCurry, M., and Mertzman, S.A., 1999, the chemical and isotopic differentiation of an epizonal magma body: Organ Needle Pluton, New Mexico: Journal of Petrology, v. 40, p. 653-678.

Wark, D.A., and Watson, E.B., 2006, TitaniQ: a titanium-in-quartz geothermometer: Contributions to Mineralogy and Petrology, v. 152, p. 743-754,

Wark, D.A., Hildreth, W., Spear, F.S., Cherniak, D.J., Watson, E.B., 2007, Pre-eruption recharge of the Bishop magma system: Geology, v. 35, p. 235-238.

Watt, G.R., Wright, P., Galloway, S., McLean, C., 1997, Cathodoluminescene and trace element zoning in quartz phenocrysts and xenocrysts: Geochimicica et Cosmochimica Acta, v. 61, p. 4337-4348.

Wiebe, R.A., Wark, D.A., Hawkins, D.P., 2007, Insights from quartz cathodoluminescence zoning into crystallization of the Vinalhaven granite, coastal Maine: Contributions to Mineralogy and Petrology, v. 154, p. 439-453.

Wilcock, J., Goff, F., Minarik, W.G., Stix J., 2013, Magmatic recharge during the formation and resurgence of the Valles caldera, New Mexico, USA: Evidence from quartz compositional zoning and geothermometry: Journal of Petrology, v. 54, p. 635-664.

Woolf, K.S., 2008, Pre-eruptive conditions of the Oligocene Wah Wah Springs Tuff, southeastern Great Basin ignimbrite province: M.S. Thesis, Brigham Young University, Provo, Utah, 77 p.

Zimmerer, M.J., and McIntosh, W.C., 2013, Geochronologic evidence of upper-crustal in situ differentiation: Silicic magmatism at the Organ caldera complex, New Mexico: Geosphere, v. 9, p. 155-169.

ȋ͵ͲȌ

͵ͻι

400 4000 300 1000

͵ͺι 200 100

0 20 40 60 80 100 Km 0 20 40 60 Miles

ͳͳ͸ι ͳͳͷι ͳͳͶι ͳͳ͵ι ͳͳʹι

ͳǤ ϐ ȋȌǤ ͳͳ͵ͷ͹ǯ ͳͳ͵ͷͷǯ ͳͳ͵ͷ͵ǯ ͳͳ͵ͷͳǯ ͳͳ͵Ͷͻǯ ͳͳ͵Ͷ͹ǯ

ǦͲ͵ͳʹ ǦͲͳͳʹ ǦͲͷͳʹ ǦͺǦ͸ͲǦͳ ͵ͺͳ͹ǯ ǦͲʹͳʹ ǦͲͶͳʹ ǦͲ͸ͳʹ ǦͺǦ͸ͲǦʹ

͵ͺ ͳͷǯ ǦͲͺͳʹ ǦͲͷͳʹ

ǦͲ͸ͳʹ ǦͲ͹ͳʹ Twit

ǦͲ͵ͳʹ ǦͲʹͳʹ ǦͲͳͳʹ ͵ͺͳ͵ǯ ǦͲͶͳʹ ǦͲʹͳʹ ǦͲͳͳʹ ǦͲ͵ͳʹ ǦͲͻͳʹ ǦͳͲͳʹ ǦͲͻͳʹ Ǧͳͳͳʹ ǦͲ͹ͳʹ ǦͳͲͳʹ Ǧͳ͵ͳʹ ǦͲͺͳʹ Ǧͳʹͳʹ Ǧͳͳͳʹ ǦͳͶͳʹ ͵ͺͳͳǯ ǦͲ͸ͳʹ ǦͲͷͳʹ

͵ͺͻǯ

͵ͺ͹ǯ

Ͳͳʹ͵ͶͲǤͷ

ȋ Ȍ ϐ ȋǦȌ

ʹǤ Ǥǡȋͳͻͺ͹ȌǤǡȋͳͻͺ͹ȌǤ

͵Ǥ ǦȋȌ ȋȌǤǦ Ǥ Ǥ ǡ ϐǦ Ǥ

ͳʹ ͳͲ ͳͳ ͺ

ͳͲ Ȍ Ψ ͸ ͻ ȋǤΨȌ ʹ ͺ Ͷ

ǦȋǤΨȌ Ǧ ʹ

Ϊ ʹ ͹

ȋ Ȍ Ϊ ʹ ʹ Ǧ ͸ ȋ Ȍ Ͳ ͷ ȋȌ Ͷ Ǧʹ ͷͻ ͸ͳ ͸͵ ͸ͷ ͸͹ ͸ͻ ͹ͳ ͹͵ ͹ͷ ͹͹ ͹ͻ ͷͻ ͸ͳ ͸͵ ͸ͷ ͸͹ ͸ͻ ͹ͳ ͹͵ ͹ͷ ͹͹ ͹ͻ

ʹȋǤΨȌ ʹȋǤΨȌ

ͺ ͳͲͲͲ ͹

͸ ͳͲͲ

ͷ Ǧ Ͷ ͳͲ ȋǤΨȌ ʹ

͵ Ȁ ʹ Ǧ ͳ

ͳ Ǧ

Ͳ Ͳ ͷͲ ͷͷ ͸Ͳ ͸ͷ ͹Ͳ ͹ͷ ͺͲ

ʹȋǤΨȌ

͸Ͳ ͳͲͲͲ ͻͲͲ ͷͲ ͺͲͲ

͹ͲͲ ͶͲ ͸ͲͲ

ͷͲͲ ͵Ͳ ȋȌ ȋȌ ͶͲͲ ʹͲ ͵ͲͲ

ʹͲͲ ͳͲ ͳͲͲ

Ͳ Ͳ ͷͷ ͸Ͳ ͸ͷ ͹Ͳ ͹ͷ ͺͲ ͸Ͳ ͸ʹ ͸Ͷ ͸͸ ͸ͺ ͹Ͳ ͹ʹ ͹Ͷ ͹͸ ͹ͺ ͺͲ ȋǤΨȌ ʹ ʹȋǤΨȌ ͶǤ ǤȋȌ ϐ ȋǤǡͳͻͺ͸ȌǤȋȌϐ ȋ Ȍ

ȋ ǤǡʹͲͲͳȌǤȋȌʹ ϐ ǤȋȌ

ȋ ȌǤȋȌʹǤ ȋǦϐǡǦϐǢ

ǤǡʹͲͲʹǡǤ Ǥȏ ȐȌǤȋ ȌʹǤǡ Ǥ ͳͲͲ

ͻͲ

ͺͲ

͹Ͳ

͸Ͳ

ͷͲ

ͶͲ

͵Ͳ

ʹͲ Ǧ

ͳͲ

ͷǤ Ǥ ǤǦ ͲǤͷ ǤȋʹͲͲͺȌǤ Ǥȏ ȐǤ

ǦͲͶͳʹ ǦͲͷͳʹ

ǦͲ͵ͳʹATCH-0412 ǦͲͶͳʹMIN-0312 MIN-0412 ǦͲʹͳʹ ǦͲ͸ͳʹ ǦͲ͹ͳʹPINTO-0612 PINTO-0212PINTO-0712 ǦͳͶͳʹ ǦͲͷͳʹ

PINTO-1412MIN-0512

Bytownite

Anorthoclase Sanidine

Labradorite Anorthoclase Sanidine

Anorthoclase Sanidine Or

Anorthoclase Sanidine Andesine Anorthoclase Sanidine

Anorthoclase Sanidine

Oligoclase Anorthoclase Sanidine

Anorthoclase Albite Sanidine

Bytownite

Labradorite

Andesine

Oligoclase

Anorthoclase Albite Sanidine

͸ǤȋȌ Ǥ ǦͲͷͳʹ Ǧ ȋȌǤȋȌ ȋ ǦͲͷͳʹȌ ȋϐȌǤ ȋʹͲͲͺȌǤ

ǦͲ͵ͳʹ ǦͳͶͳʹ ǦͲ͸ͳʹ ǦͲͷͳʹ

Figure 7. Textural varieties of hornblende grains from the Wah Wah Springs Intrusion in both plane-polarized and cross-polarized light. (A) MIN-0312 contains euhedral hornblende grains with no reaction rims, which is only prevalent in one other sample. (B) PINTO-1412 has hornblende with reaction rims composed of clinopyroxene, plagioclase, Fe-Ti oxides, and residual hornblende. (C) PINTO-0612 contains hornblende grains that are almost completely destroyed and replaced with other minerals. (D) MIN-0512 has interstitial hornblende that has grown between plagioclase grains. Hornblende in this sample has chemical evidence suggesting it equilibrated to lower temperatures.

ͳǤͲ ͲǤͻ

ͲǤͺ

ʹΪ ͲǤ͹ ͲǤ͸

ȀȋΪ ͲǤͷ

ͲǤͶ ͲǤ͵

ͲǤʹ ȋα͵Ȍ ͲǤͳ ǦͲͷͳʹ

ͲǤͲ ͷǤͷ ͸ǤͲ ͸Ǥͷ ͹ǤͲ ͹ͷ͹Ǥͷ ͺͲͺǤͲ ȋȌ

Figure 8. Hornblende classification diagram (Leake et al., 1997) based on four samples of the intrusion that contained hornblende. Most hornblende in the intrusion is magnesiohornblende, which matches the Wah Wah Springs Tuff (red field). Sample MIN-0512 is shown separetely since these analyses are consistently anomalous and trend towards actinolite. The anomalous chemistry of MIN-0512 suggests it equilibrated to lower temperatures. ȋʹͲͲͺȌ Ǥ Ǥȏ ȐǤ ͲǤͻ ͲǤͶͲ Ǧ ͲǤͺ Ϊ྅αȋ ȌΪȋΪȌ ͲǤ͵ͷ ȋǡ ȌΪα Ϊ

ͲǤ͹ ͲǤ͵Ͳ

ͲǤ͸ ͲǤʹͷ αͲǤͳͳͷͻǦͲǤͲͺͶʹ ʹαͲǤͲͻͶͷ ͲǤͷ ͲǤʹͲ

ͲǤͶ ȋȌ

ȋΪȌȋȌ ͲǤͳͷ ͲǤ͵ ͲǤͳͲ ͲǤʹ αͲǤͶ͸ͶͺǦͲǤʹͷ͵ ͲǤͳ ʹαͲǤ͹Ͷͳ ͲǤͲͷ

ͲǤ͵ͷ ʹǤͳͲ Ǧ ȋα͵Ȍ ΪαΪ ʹǤͲͷ Ϊα Ϊ ͲǤ͵Ͳ ʹǤͲͲ ǦͲͷͳʹ

ͳǤͻͷ ͲǤʹͷ ͳǤͻͲ

ͲǤʹͲ ͳǤͺͷ

ȋȌ ȋȌ ͳǤͺͲ ͲǤͳͷ ͳǤ͹ͷ

ͳǤ͹Ͳ ͲǤͳͲ αͲǤͳ͹ͶͻǦͲǤͲ͵Ͷͷ αͲǤͲ͵ͲͶΪͳǤ͹ͷͳ͹ ʹαͲǤ͸ͳͷ ͳǤ͸ͷ ʹαͲǤͲͶʹʹ

ͲǤͲͷ ͳǤ͸Ͳ ͲǤ͸ ͲǤͺ ͳǤͲ ͳǤʹ ͳǤͶ ͳǤ͸ ͳǤͺ ʹǤͲ ͲǤ͸ ͲǤͺ ͳǤͲ ͳǤʹ ͳǤͶ ͳǤ͸ ͳǤͺ ʹǤͲ ȋȌ ȋȌ

ͻǤ ȋϐȌǤ ȋȌǡǦ ȋȌǡǦ ȋȌǡ ȋȌǤ ǦͲͷͳʹǡ Ǥ ǤȋʹͲͲͺȌ Ǥ Ǥȏ ȐǤ

ʹ͸ ʹ͸ ͷͲͲ ͸ͲͲ ͹ͲͲ ͺͲͲ ͻͲͲ ȋα͹Ȍ ǦͳͶͳʹ ǦͲ͸ͳʹ ǦͲͷͳʹ

ʹʹ͸ ʹʹ͸

ͳͲǤ Ǥ ϐǤ Ǥ ǦͲͷͳʹǡ Ǧ Ǥ ȋͳͻͺ͵ȌǤ ȋʹͲͲͺȌǤ Ǥȏ ȐǤ

PINTO-0712 ͲǤ͵͹ͷ ͲǤ͹ͷ ͲǤͳʹͷ

ǦͲ͸ͳʹ ǦͲͶͳʹ ǦͲ͹ͳʹ

ͳͳǤ ǤȋȌǡȋȌǡȋȌǤ

ǦͺǦͲȋȌ Ǧ͵ͳ͵ȋȌ

ǦͲͶͳʹ ǦͳͶͳʹ

ǦͲͷͳʹ ǦͲ͵ͳʹ

ǦͲ͵ͳʹ ǦͲ͸ͳʹ

ͳʹǤ ǤȋȌȋȌ ȋȌȋǡʹͲͳͲȌǤȋȌ Ǧ ǤȋȌ Ǥ Ǥ ȋȌǦȋ Ȍ Ǥȋ Ȍ Ǥ

ȋͳͻͻͶȌǦ ͵ ȋͳͻͺͻȌ

ȋͳͻͻͶȌ Ǧ ȋͳͻͺͻȌ

ȋȌ ȋͳͻͻͶȌ Ǧ ȋͳͻͻͷȌ ϐǤȋʹͲͳͲȌǦϐǤȋʹͲͳͲȌ

ȋͳͻͻͶȌ Ǧ ȋͳͻͻͷȌ

Ͳ ͹ͲͲ ͹ͷͲ ͺͲͲ ͺͷͲ ͻͲͲ ͻͷͲ ͳͲͲͲ

ȋιȌ ͳ͵ǤǦ ȋ ǡͳͻͻͶǢ ǡͳͻͺͻǢǡͳͻͻͷǢϐǤǡʹͲͳͲȌ ǤǤǦ ȋȌ ȋͳͻͻͶȌ ȋͳͻͺͻȌ ǡ Ǥ ǡ Ǥ

ͷǤͲ ϐ ϐ ʹ αͳ ʹ αͲǤͷ

ͶǤͲ ǦͲ͵ͳʹ ǦͲ͵ͳʹ ǦͲͶͳʹ ǦͲͶͳʹ ǦͳͶͳʹ ǦͳͶͳʹ ͵ǤͲ

ǦͲͷͳʹ ǦͲͷͳʹ )

ʹǤͲ ȋ ͳǤͲ

ͲǤͲ ͸ͷͲ ͹ͲͲ ͹ͷͲ ͺͲͲ ͺͷͲ ͻͲͲ ͸ͷͲ ͹ͲͲ ͹ͷͲͺͲͲ ͺͷͲ ͻͲͲ ȋȌ

ͳͶǤ Ǧ ȋȌǦȋȌǤ ȋ Ȍ ȋͳͻͺͻȌ ȋʹͲͲ͸ȌǤ ǦǦǤȋȌ Ǥ

ʹǤͷ ȋʹͲͳʹȌ ǤȋʹͲͳͲȌ

ʹǤͲ

ͳǤͷ ȋȌ ͳǤͲ

ͲǤͷ

ͲǤͲ Ͳ ʹͲͲ ͶͲͲ ͸ͲͲ ͺͲͲ ͳͲͲͲ ͳʹͲͲ

ȋȌ ͳͷǤ ǤȋʹͲͳͲȌ ȋʹͲͳʹȌǡ ͺȋ̱ͲǤ͸ȌǤ ȋ͹ͻ͵ȌǤ Ǥ

Andesitic NNWlava SSE NNW SSE w w ccw c m m m c Pz Pz Pz Pz Future Pz e Pz

r s ring fault t e l

Hypothetical normal v u

e a

r f Potential void

e g

r n i u r ring fault t

Magma Magma

ȋȌ Ǧ Ǥ ȋȌ Ǧ Ǥ Ǥ Ǧ Ǥ Ǥ Ǥ ̱ͺǡ Ǥ ̱ͺͲͲǡǦǤ

Future Indian SSE Peak NNW SSE NNW wl w m w w w w r l+r l+r Pz wl m Pz m Pz m Pz m m m Pz m Pz Pz Pz Pz Pz

c g m m c Magma m Pz

ȋȌ ͷ ȋȌǦ Ǧ Ǥ ϐǤ Ǥ ǡ ǡ Ǧ Ǥ Ǥ Ǥ Ǥ Ǥ Ǥ

ͳ͸Ǥ Ǧ r w c ȋϐǤǡʹͲͳ͵Ȍ

l m 2 wl Pz km g 0 0 2

Table 1. Whole rock major and trace element compositions of the Wah Wah Springs Intrusive Granodiorite Porphyry. ATCH- ATCH- ATCH- ATCH- ATCH- ATCH ATCH ATCH- ATCH- MIN- MIN- 0112 0212 0412 0512 0612 -0712 -0812 1012 1112 0212 0312 wt. %

SiO2 66.06 64.53 64.65 65.86 78.03 75.63 68.76 63.30 65.69 68.97 66.19 TiO2 0.59 0.62 0.63 0.61 0.12 0.18 0.50 0.78 0.63 0.75 0.56 Al2O3 15.02 15.30 15.29 15.59 12.15 13.14 15.18 15.81 15.74 14.95 15.04 Fe2O3 5.03 5.44 5.44 5.02 0.88 1.20 4.30 5.76 5.11 4.83 4.86 MnO 0.08 0.11 0.09 0.08 0.01 0.01 0.05 0.07 0.09 0.08 0.08 MgO 2.39 2.24 2.39 2.06 0.13 0.24 1.76 2.63 2.30 1.44 2.21 CaO 4.16 5.55 5.01 4.22 0.64 1.17 2.14 4.43 3.01 3.77 4.19

Na2O 2.92 2.66 2.78 2.95 2.76 2.27 2.94 3.27 3.16 2.46 2.91 K2O 3.59 3.36 3.53 3.44 5.25 6.10 4.22 3.72 4.07 2.57 3.80 P2O5 0.16 0.19 0.19 0.17 0.03 0 0.15 0.22 0.20 0.18 0.17 Total 100 100 100 100 100 100 100 100 100 100 100 LOI* 1.37 3.58 3.10 2.09 1.17 1.24 2.02 0.73 1.92 2.04 1.93 Anal Total 99.85 99.84 99.83 99.85 99.93 99.87 99.84 97.59 99.79 99.79 99.84 ppm Sc 15 19 16 15 1 3 11 16 12 12 15 V 106 116 108 107 13 14 89 120 102 83 104 Cr 28 31 28 29 0 0 22 31 24 10 29 Ni 16 17 16 15 7 4 15 19 16 6 15 Cu 22 23 17 26 1 3 18 17 23 4 22 Zn 62 66 59 53 12 10 53 53 60 100 57 Ga 17 18 17 18 13 12 16 18 18 19 17 Rb 125 115 119 122 142 159 146 135 151 71 140 Sr 447 515 489 504 212 246 453 574 550 453 446 Y 21 21 20 19 17 6 20 22 23 23 19 Zr 161 165 159 161 82 82 156 181 185 232 166 Nb 14 12 12 13 20 6 13 13 14 16 13 Ba 722 727 793 731 294 783 820 783 785 766 800 La 42 36 39 38 32 18 44 41 43 50 40 Ce 74 74 81 74 88 48 78 64 80 101 70 Nd 36 35 38 34 31 19 37 36 38 46 34 Sm 4 4 4 4 3 3 5 4 4 5 4 Pb 32 22 23 21 24 21 23 15 25 46 19 Th 23 20 20 22 32 12 26 19 23 19 23 U 4 3 3 3 4 3 5 3 5 3 4

*Loss on ignition at 1000°C.

Table 1 Continued. MIN- MIN- MIN- PINTO- PINTO- PINTO- PINTO- PINTO- PINTO- PINTO- 0412 0512 0612 0112 0512 0612 0712 1212 1312 1412 wt. %

SiO2 66.15 66.48 65.56 64.76 65.63 65.53 64.16 63.13 65.65 66.55 TiO2 0.62 0.58 0.59 0.66 0.64 0.58 0.69 0.73 0.60 0.57 Al2O3 14.95 15.04 15.06 14.84 15.44 15.03 16.08 15.62 15.03 14.84 Fe2O3 5.03 3.34 5.12 5.90 5.52 5.14 5.50 6.44 5.17 4.81 MnO 0.06 0.06 0.09 0.12 0.09 0.08 0.08 0.11 0.09 0.08 MgO 2.35 2.02 2.34 2.61 2.51 2.37 2.66 3.15 2.37 2.15 CaO 4.30 4.38 4.56 4.33 2.88 4.22 3.47 4.17 4.50 4.11

Na2O 2.74 3.55 2.85 2.92 3.20 2.78 3.35 3.25 2.78 2.76 K2O 3.64 4.39 3.67 3.66 3.90 4.08 3.80 3.16 3.64 3.97 P2O5 0.17 0.16 0.17 0.20 0.20 0.17 0.21 0.22 0.17 0.16 Total 100 100 100 100 100 100 100 100 100 100 LOI 3.01 0.97 1.21 3.79 2.17 1.23 2.80 2.36 1.15 0.75 Anal Total 99.53 99.78 99.57 99.83 99.83 99.84 99.82 99.63 99.84 99.84 ppm Sc 16 14 16 17 13 15 15 17 15 14 V 108 104 111 129 116 109 115 140 112 104 Cr 32 21 28 39 36 31 28 43 30 27 Ni 16 14 15 18 19 18 17 23 16 16 Cu 12 4 20 118 16 26 15 20 22 22 Zn 46 38 64 84 56 67 58 70 66 57 Ga 17 17 18 17 18 18 18 18 17 17 Rb 124 163 132 122 141 144 140 105 125 143 Sr 476 514 466 470 489 475 567 636 449 434 Y 20 22 22 23 22 21 22 23 21 21 Zr 164 149 170 188 169 168 189 183 169 167 Nb 14 16 14 14 13 14 14 13 14 14 Ba 685 626 718 802 781 716 870 772 736 747 La 41 35 46 40 42 42 43 40 39 43 Ce 76 75 85 73 73 81 72 73 80 77 Nd 36 33 41 38 38 37 38 39 37 36 Sm 4 4 5 4 5 4 5 5 4 4 Pb 23 25 26 16 15 32 20 19 45 25 Th 23 27 23 21 22 25 23 20 25 24 U 5 6 5 4 4 6 5 3 6 4

Table 2. Representative chemical analyses of feldspars from the Wah Wah Springs Intrusion. ATCH- ATCH- MIN- MIN- MIN- PINTO- PINTO- PINTO- PINTO- 0412 0512 0312 0412 0512 0212 0612 0712 1412 Plagioclase

SiO2 56.91 56.57 56.46 56.81 58.44 56.94 56.04 56.48 56.29

Al2O3 27.49 27.57 26.99 27.55 26.48 27.36 27.52 27.20 27.28

Fe2O3 0.44 0.22 0.50 0.28 0.17 0.24 0.35 0.32 0.41 CaO 9.51 9.43 9.99 9.58 8.20 9.28 9.67 9.63 9.44

Na2O 5.11 5.07 5.34 5.89 6.87 6.05 5.78 5.62 5.91

K2O 0.58 0.52 0.76 0.46 0.27 0.43 0.33 0.60 0.51 Total 100.05 99.38 100.04 100.57 100.43 100.31 99.69 99.84 99.84

Ab 0.48 0.48 0.4703 0.51 0.59 0.53 0.51 0.50 0.52 An 0.49 0.49 0.49 0.46 0.39 0.45 0.47 0.47 0.46 Or 0.04 0.03 0.04 0.03 0.02 0.02 0.02 0.03 0.03 Alkali Feldspar

SiO2 64.73

Al2O3 18.63

Fe2O3 0.09 CaO 0.19

Na2O 2.60

K2O 13.14 Total 99.37

Ab 0.23 An 0.01 Or 0.76

Table 3. Representative chemical analyses of hornblende from the Wah Wah Springs Intrusion. MIN-0312 MIN-0412 PINTO-1412 MIN-0512

SiO2 45.98 44.99 46.53 50.60

TiO2 1.81 2.17 1.81 0.47

Al2O3 8.03 8.25 7.90 4.07

FeOt 12.61 12.63 12.56 12.26 MnO 0.29 0.35 0.33 0.29 MgO 14.82 14.53 14.64 15.84 CaO 11.51 11.60 11.36 12.32

Na2O 1.44 1.62 1.72 1.01

K2O 0.75 0.81 0.70 0.46 F 0.40 0.30 0.70 0.79 Cl 0.06 0.05 0.07 0.10 Subtotal 97.53 97.15 98.00 97.83

H2O* 1.86 1.89 1.72 1.67 Total 99.39 99.04 99.71 99.50 *calculated

Table 4. Representative chemical analyses of pyroxene from the Wah Wah Springs Intrusion. ATCH- ATCH- MIN- PINTO- PINTO- PINTO- PINTO- 0412 1012 0512 0212 0612 0712 1412

SiO2 52.63 52.68 53.02 52.64 52.22 52.96 52.87

TiO2 0.13 0.15 0.08 0.11 0.14 0.15 0.12

Al2O3 1.01 0.95 0.56 1.15 1.51 0.92 0.90

FeOt 9.04 8.41 8.37 9.19 9.69 8.46 8.41 MnO 0.62 0.61 0.52 0.67 0.56 0.26 0.59 MgO 13.75 13.99 13.24 14.35 13.27 14.71 13.84 CaO 22.16 22.51 23.68 21.69 21.48 21.80 22.80

Na2O 0.35 0.23 0.31 0.39 0.42 0.28 0.32 Total 99.70 99.51 99.77 100.18 99.29 99.54 99.86

Table 5. Quartz cathodoluminescence in Wah Wah Springs Intrusive Granodiorite Porphyry

Zoning Rim color Groundmass

ATCH-0512-1 Present Bright Zoned ATCH-0512-2 Present Bright ATCH-0512-4a Present Bright ATCH-0512-4b Present Bright

MIN-0412-1 Absent N/A Bright MIN-0412-2 Absent N/A MIN-0412-3 Absent N/A MIN-0412-4 Absent N/A MIN-0412-5 Present Dark MIN-0412-7 Present Dark

PINTO-1412-1 Present (sector) N/A Bright PINTO-1412-2 Absent N/A PINTO-1412-3 Present Bright PINTO-1412-4 Present Bright PINTO-1412-5 Absent Bright

MIN-0312-1a Present Bright Bright MIN-0312-1b Present Bright MIN-0312-2 Present Bright MIN-0312-3 Present Bright MIN-0312-4a Present (sector) N/A MIN-0312-4b Absent N/A MIN-0312-5 Absent N/A

PINTO-0612-1 Present (sector) N/A N/A (altered) PINTO-0612-2 Present Bright PINTO-0612-3a Absent N/A PINTO-0612-3b Absent N/A PINTO-0612-5a Present Dark PINTO-0612-5b Present Dark PINTO-0612-5c Present Dark

Table 6. Summary of geothermobarometry for the Wah Wah Springs Intrusion and Tuff Overall Avg. Overall MIN-0312 MIN-0412 PINTO- MIN-0512 (Tuff*) Avg.† 1412 (Intrusion) T (oC) Holland and Blundy, 791±20o 793±20o 791±19o 793±29o 797±14o 656±28o without quartz (1994) A§ T (oC) Holland and Blundy, 818±29o 816±31o 802±27o 847±19o 830±29o 753±28o with quartz (1994) T (oC) Ridolfi et al. (2010) 861±28o 849±25o 841±23o 858±21o 862±26o 776±27o

P (kb) Johnson and 2.2±0.2 2.4±0.6 2.3±0.6 2.3±0.6 2.6±0.7 0.3±0.6 Rutherford (1989) P (kb) Anderson and Smith 1.8±0.2 1.7±0.8 1.9±0.5 1.6±0.7 1.0±0.9 0.0±0.5 (1995) P (kb) Ridolfi et al. (2010) 1.5±0.3 1.5±0.5 1.4±0.3 1.4±0.3 1.9±0.7 0.7±0.1 fO2 (log units) Ridolfi et al. (2010) -11.6±0.1 -11.5±0.3 -11.6±0.3 -11.3±0.3 -11.3±0.4 -12.6±0.5

H2O melt (wt. %) Ridolfi et al. (2010) 4.3±0.2 4.1±0.4 4.3±0.2 3.6±0.3 4.1±0.5 3.7±0.2 *Tuff P-T estimates from Woolf (2008). Other estimates from this study. † MIN-0512 was excluded from overall averages since it contains non-igneous hornblendes. § Holland and Blundy (1994) A does not require quartz in the mineral assemblage, B does require quartz.

APPENDIX A Thin section descriptions

ATCH-0412 Porphyritic granodiorite with 59.3% quartz rich matrix, 27.0% plagioclase, 4.5% quartz, 4.3% biotite, 2.0% clinopyroxene, 1.9% Fe-Ti oxides, with the remaining 1.0% composed of indistinguishable altered remnants. Plagioclase is euhedral, up to 4 mm across, and displays oscillatory zoning and polysynthetic twinning. Quartz is resorbed and up to 2 mm across. Some biotite remains in euhedral condition and is around 1mm across. This sample contains no euhedral hornblende but there are remnants of either biotite or hornblende that are filled with oxides, clinopyroxene, and plagioclase. Clinopyroxene phenocrysts are up to 1 mm in size and have resorbed edges.

ATCH-0512 Porphyritic granodiorite with 58% quartz-rich matrix, 25.6% plagioclase, 7.3% quartz, 3.8% biotite, 1.6% clinopyroxene, 1.1% Fe-Ti oxides with the remaining 2.6% composed of indistinguishable altered remnants. Feldspar grains display zoning, are up to 5 mm in size, and sometimes form composite grains. Quartz grains are highly resorbed and up to 2.5 mm across. Some biotite remains in euhedral condition and is around 1 mm. No hornblende is seen but there are ambiguous remains of either biotite or hornblende that are filled with oxides, clinopyroxene, and plagioclase. Clinopyroxene grains are up to 1.5 mm in size and have resorbed edges.

ATCH-1012 Heavily altered porphyritic granodiorite containing a fine-grained mafic inclusion surrounded by phenocrysts of 51.4% plagioclase feldspar, 17.8% quartz, 12.7% biotite, , 5.6% clinopyroxene, 2.3% Fe-Ti oxides, with the remaining 10.2% composed of indistinguishable altered remnants. Most plagioclase are euhedral and up to 2.5 mm across. Almost all quartz is highly resorbed, some is skeletal, and some is graphic. Quartz grains are up to 2 mm across, but most are smaller. Biotite is present but chloritized and up to 1 mm across. Euhedral pyroxene grains are up to 1.5 mm across, twinned, and include Fe-Ti oxide and apatite grains.

MIN-0312 Porphyritic granodiorite with 62.0% quartz-rich matrix, 22.4% plagioclase, 4.1% quartz, 6.6% hornblende, 3.3% biotite, and 1.2% Fe-Ti oxides. Euhedral plagioclase grains are oscillatory zoned, polysynthetically twinned, and up to 5 mm across; one plagioclase grain has a restitic, poikilitic core with euhedral plagioclase grown around it. Quartz grains are highly resorbed and some show skeletal growth around the edges. Euhedral hornblende up to 2 mm across exists in glomerocrysts, and some grains narrow reaction rims. Biotite is highly altered and full of oxide inclusions.

MIN-0412 Porphyritic granodiorite with 53.2% fine-grained aphanitic groundmass; phenocrysts include 22.2% plagioclase, 7.8% biotite, 8.4% quartz, 6.0% hornblende, and 2.4% Fe-Ti oxides. Euhedral plagioclase grains, up to 4 mm across, display oscillatory zoning and polysynthetic twinning. Quartz is highly resorbed, often deeply embayed, and up to 1.5 mm across. Euhedral hornblende grains up to 2.5 mm across display cleavage, have Fe-Ti oxide inclusions, and exist

63 in glomerocrysts. Many hornblende grains are chloritized on fractures. Chloritized remnants of biotite are also present.

MIN-0512 Coarse-grained porphyritic granodiorite with 53.7% matrix, 33.3% plagioclase, 10.2% hornblende, 1.8% quartz, 0.4% clinopyroxene, 0.3% Fe-Ti oxides, and 0.3% carbonate. The matrix is composed of quartz, plagioclase, and K-feldspar and some matrix grains are up to 1 mm across. These K-feldspar grains do not exhibit simple twinning. Highly altered, euhedral plagioclase grains are up to 6 mm in size and have polysynthetic twinning and oscillatory zoning. Slightly resorbed quartz grains are up to 1 mm in size, however most are smaller. Anhedral to euhedral hornblende is most commonly interstitial and grown between plagioclase grains; however, there is one large 6 mm phenocryst. Hornblende grains have Fe-Ti oxide and pyroxene inclusions. Clinopyroxene phenocrysts are up to 1.5 mm across.

MIN-8-60-1 Porphyritic granodiorite with 70.7% coarse-grained matrix, 26% plagioclase, 2% hornblende, 0.6 % quartz, 0.5% biotite, and 0.2% Fe-Ti oxides. The matrix consists of quartz, plagioclase, and K-feldspar and some grains are up to 1 mm across. These K-feldspar grains exhibit simple twinning. Plagioclase is euhedral and exhibits oscillatory zoning and polysynthetic twinning; there is one large, 12 mm, glomerocrystic grain of plagioclase. Quartz grains are most abundant in the groundmass and phenocrysts are sparse. Hornblende, up to 2 mm across, exists interstitially between matrix quartz grains, and Fe-Ti oxides are only present as inclusions in hornblende. Some biotite is intergrown with hornblende. This sample contains comparatively few mafic minerals compared to other samples. This thin section was not available electron microprobe analysis.

MIN-8-60-2A Porphyritic granodiorite with 42.9% fine-grained quartz-rich matrix, 23.8% plagioclase, 7.8% hornblende, 4.8 % biotite, 3.6% quartz, 1.6% Fe-Ti oxides, and 1.3% clinopyroxene. Plagioclase grains are strongly altered and up to 2 mm. Quartz grains are highly resorbed and up to 2.5 mm across. Hornblende exists in glomerocrysts and is highly altered but without reaction rims. Biotite is chloritized. Clinopyroxene grains are up to 2.5 mm across but sparse and significantly weathered. This thin section was not available electron microprobe analysis.

MIN-8-60-3 Porphyritic granodiorite with 69.6% fine-grained groundmass, 19% plagioclase, 3.4% quartz, 3.2% biotite, 3.6% hornblende, and 1.1% Fe-Ti oxides. Most plagioclase is euhedral but heavily altered and up to 4 mm across. Quartz is deeply resorbed and around 1 mm in size. Hornblende exists as euhedral phenocrysts and in glomerocrysts and some show reaction around the edges. Biotite is distinguishable by shape but strongly altered and poikilitic. This thin section was not available electron microprobe analysis.

PINTO-0212 Partially coarse-grained and partially fine-grained rock with 49.4% plagioclase feldspar, 25.6% quartz, 17.8% altered biotite, 4.4% Fe-Ti oxides, and 1.6% clinopyroxene. All grains are highly altered in both the coarse-grained and fine-grained zones of the rock. Plagioclase is

64 euhedral, up to 2.5 mm in size, and occasionally shows oscillatory zoning. Quartz is skeletal and appears to have “decayed” from quartz phenocrysts. Biotite remnants are up to 2.5 mm in size and often contain many Fe-Ti oxides and pyroxene inclusions. Clinopyroxene grains are < 1 mm and highly fractured.

PINTO-0612 Porphyritic granodiorite with a fine-grained, highly altered inclusion. This sample contains 73.3% of fine-grained and altered groundmass; phenocrysts include 6.7% quartz, 7.7% plagioclase, 5.2% hornblende, 4.4% biotite, 1.6% Fe-Ti oxides, and 1.1% clinopyroxene. Both hornblende and biotite only exist as altered remnants full of oxide and other inclusions, but are distinguishable by their shape. Plagioclase, up to 3 mm across, is also very altered and has oxide inclusions. Quartz grains can be up to 2 mm across but most are smaller and highly resorbed; some are broken and groundmass quartz is skeletal. Subhedral to euhedral pyroxene grains are up to 1 mm across, contain apatite inclusions, and are sometimes intergrown with biotite.

PINTO-0712 Porphyritic-phaneritic granodiorite with 39.7% plagioclase, 20.9% quartz, 10.6% biotite, 2.9% clinopyroxene, 2.4% Fe-Ti oxides, and 0.7% hornblende with 22.8% matrix of smaller quartz, Fe-Ti oxides, pyroxenes, and altered remnants. Heavily altered plagioclase grains are up to 4 mm across. Quartz grains are up to 2.5 mm across and are resorbed and embayed and groundmass grains are skeletal. Biotite and hornblende remnants exist interstitially between other small matrix grains. Pyroxene is up to 1.5 mm across and is euhedral to anhedral, sometimes exhibiting holes and altered edges.

PINTO-1412 Porphyritic granodiorite with 52.1% very fine-grained (grains <50 microns) quartz-rich matrix, 21.9% plagioclase, 9.9% biotite, 9% quartz, 2.7% hornblende, 2.3% Fe-Ti oxides, and 2.1% clinopyroxene. Plagioclase grains exhibit oscillatory zoning, polysynthetic twinning, and are up to 2 mm in size. Quartz is also up to 2 mm in size and some look euhedral while some have slightly resorbed edges. Biotite is euhedral, poikilitic, and up to 1.5 mm in size. Clinopyroxene is around 0.5 mm across; most have defined edges but some show beginning stages of resorption. Hornblende is sometimes intergrown with biotite, and contains 0.05 mm- 0.5 mm thick reaction rims of clinopyroxene, Fe-Ti oxides, plagioclase and residual hornblende.

APPENDIX B Quartz cathodoluminescence images

PINTO-0612

PINTO-1412

MIN-0312

MIN-0412

ATCH-0512