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In presenting this thesis in partial fulfillment of the requirements for an advanced degree from Idaho State University, I agree that the Library shall make it freely available for inspection. I further state that permission for extensive copying of my thesis for scholarly purposes may be granted by the Dean of Graduate Studies, Dean of my academic division, or by the University Librarian. It is understood that any copying or publication of this thesis for financial gain shall not be allowed without my written permission. Signature ____________________________ Date ____________________________
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THE PETROGENESIS OF QUATERNARY RHYOLITE DOMES IN THE BIMODAL BLACKFOOT VOLCANIC FIELD, SOUTHEASTERN
IDAHO
By Mark T. Ford
A thesis submitted in partial fulfillment
of the requirements for the degree of Masters of Science in the Department of Geosciences
Idaho State University July, 2005
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To the Graduate Faculty: The members of the committee appointed to examine the thesis of Mark T. Ford find it satisfactory and recommend that it be accepted. _________________________________ Michael McCurry Major advisor _________________________________ Scott S. Hughes Committee member _________________________________ Lisa Goss Graduate Faculty Representative
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ACKNOWLEDGEMENTS First, I would like to thank my advisors Dr. Mike McCurry and Dr. Scott Hughes. Their assistance has greatly improved both the scientific quality and syntax of this work and they helped me navigate the world of igneous petrology. Bobbie, Melissa and Connie along with many others in the main office have smoothed out the rough spots and without them, I might never have gotten paid. Special thanks to Diana Boyack as she has provided invaluable assistance to me, far above and beyond what she needed to do. Many others have impacted my stay here at Idaho State over the past few years. Arron Pope fed me too much coffee and the occasional undercooked meal but let me split some of his wood pile and has been a good friend, conversationalist, and pool shooting adversary. Jamie Blair provided hunting and fishing companionship and introduced me to many new fishing holes and some cool cats here in Pocatello. Duane DeVechio is a great friend and helped to widen my musical horizons plus he always had some Olympia in the fridge. Kate Pickett and Dan Narsavage are also great pals and I wish them continued happiness. Others here at ISU warrant special mention and I’m sure to forget to list a few names: Kaleb Scarberry - disk golf and high adventure, Butch and Diane Wheeler -Snake River fishing and some good eats in Blackfoot, Brian Hennings – Worms and rabbit stew, Mary Hodges – championship canoe rowing and Bannock Pass tufa, Rudy Ganske – petrology discussions and general banter, Songqiao Chen – expert in the lab (but please ware some safety glasses), and Rodger Rapp and Nagendra Singh – office mates over the years. I would also like to thank Dr. Jim Lia and Dr, Jagoda Urban-Klaehn for their support over the years. Thanks to a few other friends from my “living room” including Greg Bary, Dan Dawson, Paul Grayson, Don Lopez, Jim Olsen, Kevin Pitkin and Dennis Pullman. I would be remiss if I didn’t mention some of the folks from New York that have shown continued support over the years including Jason and Walt Strawser and the rest of the “Coon Hollow Gang”, Richard and Ellen Luce and the Clearing Corriedale Sheep and Wool Co., the entire Lynch family and all the old friends from Lamoka Lake and my taxidermist, Dave Fish. Finally, I would like to thank my parents. My father taught me, among other things, that with a good education, hard work and perseverance, one can achieve great things. He was a great fishing partner and my best friend. My mother was one of the happiest people I’ve ever met and, unless you were a snake, was kind and gracious to everyone. I wish I could grow plants as well as she did. I am saddened that they were not able to see me finish this segment of my educational odyssey and I miss them.
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TABLE OF CONTENTS Page # List of Plates vi List of Figures vii List of Tables ix Abstract x Chapter 1: Introduction 1 Section 1.1: Problem Statement and Objective 1 Section 1.2: Location of the Study Area 7 Section 1.3: Regional Geologic Setting 8 Section 1.4: Previous Work in the Blackfoot Volcanic Field 16 Chapter 2: Methods 19 Section 2.1: Mapping and Sampling Field Work 19 Section 2.2: Petrographic Methods 20 Section 2.3: Sample Geochemistry Preparation (ICP, INAA, EMP, and isotope) 20 Section 2.4: ICP-AES Methods and Uncertainties 22 Section 2.5: INAA Methods and Uncertainties 24 Section 2.6: Radiogenic Isotope Methods 25 Section 2.7: Heavy Mineral Separates 26 Section 2.8: Electron Microprobe Methods 29 Chapter 3: Results 30 Section 3.1: Field Relationships 30 Section 3.2: Petrography 43 3.2.1: Petrography of the CDF 44 3.2.2: Petrography of the NDF 49 3.2.3: Petrography of the SIR 51 3.2.4: Mafic Enclaves of the BVF 51 Section 3.3: Geochemistry of the BVF 53 3.3.1: Major-Element Geochemistry 54 3.3.2: Comparison of Major-Element Geochemistry of the BVF to Other Rhyolites 62 3.3.3: Trace-Element Geochemistry 67 3.3.4: Comparison of Trace-Element Geochemistry of the BVF to Other Rhyolites 70 Section 3.4: Electron Microprobe Results on Phenocrysts of the Central Dome Field 77 Section 3.5: Isotope Results 81
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Chapter 4: Discussion 82 Section 4.1: Physical Volcanology 82 Section 4.2: Petrography Interpretation 87 Section 4.3: The State of the Pre-Eruptive System 89 Section 4.4: Current Models for Nearby Physiographic Provinces 91 Section 4.5: Patterns of Rock Chemistry in the BVF 95 Section 4.6: MELTS Modeling to Determine Crystallization Sequences and Timing 98 Section 4.7: Isotope Modeling With Bulk Assimilation and EC-AFC Models 99 Section 4.8: NDF Petrogenesis and Regional Implications 115 Chapter 5: Conclusions 118 References: 122
APPENDICES Appendix 1: Summary of Electron Microprobe Standards, Crystals and Integration Times 132
PLATES Plate 1: Map, cross-section and tephra pit stratigraphy of the Central Dome Field back pocket
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FIGURES
Page # Figure 1: Map of southeastern Idaho including the Eastern Snake River Plain (ESRP) and Blackfoot Volcanic Field (BVF) 3 Figure 2: Hillshade DEM of the Blackfoot Volcanic Field 5 Figure 3: Geophysical based cross section of the crust under the ESRP 9 Figure 4: Physiographic province map of and location of topaz rhyolites of the western United States 10 Figure 5: Regional Archean craton map of the Wyoming terrane 11 Figure 6: Regional structure map of southeastern Idaho 12 Figure 7: Photo of lithophysae cavities in the Northern Dome Field 32 Figure 8: Photo of flow banding in the Northern Dome Field 32 Figure 9: Photo of Sheep Island Rhyolite outcrop 33 Figure 10: Photo of Hole in the Rock Lake graben 35 Figure 11: Photo of the graben south of China Cap 35 Figure 12: Photographic overview of the China Hat tephra deposit 37 Figure 13: Photo of palagonitized basalt and rhyolitic tephra in the 38 China Hat gravel pit Figure 14: Photo of block-sized rhyolite in the China Hat gravel pit 38 Figure 15: Photo of the coulee encircling China Hat 41 Figure 16: Overview photo of the Central Dome Field 41 Figure 17: Photo of spines on China Cap 42 Figure 18: Photo of crater between China Cap and North Cone 42 Figure 19: Photomicrograph of strongly embayed quartz 45 Figure 20: Photomicrograph of skeletal sanidine armoring plagioclase 45
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Page # Figure 21: Photomicrograph of a blocky magnetite crystal 45 Figure 22: Photomicrograph of a silicic glomerocryst 48 Figure 23: Cross polar photomicrograph of Figure 22 48 Figure 24: Photomicrograph of granophyric textures in a silicic xenolith from the Northern Dome Field 50 Figure 25: Photomicrograph of oscillatory zoned plagioclase within a basaltic magmatic enclave 50 Figure 26: Photomicrograph of boxy cellular plagioclase within a basaltic magmatic enclave 50 Figure 27: Photomicrograph of a zone of hybridization of magmatic rhyolite and basalt 52 Figure 28: Photomicrograph of a pyroxene mantled quartz crystal 52 Figure 29: Photomicrograph of spongy cellular (fingerprint) texture in a sanidine crystal near a basaltic magmatic enclave 52 Figure 30: Major element variation diagram for the Blackfoot Volcanic Field rhyolites 61 Figure 31: Major element variation diagrams incorporating rhyolites from various physiographic regions of the western US 65-66 Figure 32: Spider diagram for rhyolites of the Blackfoot Volcanic Field 69 Figure 33: Trace element variation diagrams incorporating rhyolites from various physiographic regions of the western US 72-74 Figure 34: Chondrite normalized REE diagram comparing igneous rocks from the ESRP and BVF 76 Figure 35: Ternary plots of sanidine, plagioclase, magnetite and Ilmenite compositions 78 Figure 36: Hypothesized cross-section of the timing of the tephra and dome emplacement 85 Figure 37: Hypothesized cross section of crater formation 87
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Page # Figure 38: Atomic Mn/Mg concentration diagram for ilmenite
and magnetite equilibrium 90 Figure 39: Spider diagram normalized by primitive BVF basalt 95 Figure 40: Single stage bulk assimilation model 102 Figure 41: EC-AFC model using standard conditions from Table 10 107 Figure 42: EC-AFC model with elevated initial Sr and Nd concentrations 108
TABLES Table 1: Age data and references for the Blackfoot Volcanic Field 6 Table 2: Sample location, major and trace element geochemistry for the Blackfoot Volcanic Field (BVF) rhyolites 55-60 Table 3: Reference list for data given in Figures 31 and 33 64 Table 4: Summary of relative trace element abundances in the BVF rhyolites 68 Table 5: Average concentrations of trace elements from various Physiographic regions of the western US 71 Table 6: Electron microprobe phenocryst chemistry data for the Central Dome Field 79-80 Table 7: Nd and Sr isotopic ratios for the basalts and rhyolites of the BVF 81 Table 8: Pre-eruptive magma chamber properties of the CDF 91 Table 9: Sr and Nd isotopic and elemental concentrations for
xenoliths of the Snake River Plain 100 Table 10: General parameters used in EC-AFC models 105
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ABSTRACT
Seven Quaternary topaz-style rhyolite domes are present in the bimodal basalt-
rhyolite Blackfoot Volcanic Field (BVF). The high silica, slightly metaluminous
rhyolites average ~ 10% crystals in their glassy carapace (quartz > sanidine > plagioclase
> hornblende ≈ biotite > magnetite > ilmenite) and contain trace phases of zircon >>
apatite, +/- allanite, +/- thorite. As compared to other physio-tectonic rhyolites, the BVF
rhyolites are high in Cs, Ta, Yb, Rb, U, and Th and low in FeO(t), Sc, Sr, LREE, and Eu.
The three, aligned domes of the Central Dome Field (CDF) are especially high in Rb and
low in Sr resulting in Rb/Sr ratios approaching 250. Another distinction of BVF rhyolites
from other rhyolites are very low La/Yb ratios, with the CDF ratio near unity.
The young (~50 ka) CDF was emplaced after localized basalt flows and normal
faulting. It began with hydrovolcanic tephra production and subsequent tephra and
endogenous dome growth overlap. Other features in the CDF related to the rhyolite
emplacement include local tumescence and collapse crater formation and partial dome
collapse resulting in hummocky terrain.
All of the phenocryst phases, except quartz, are in textural equilibrium with the
melt. The CDF magma equilibrated at ~3.5 kbars (13 km deep), at a temperature of
~760oC with a log fO2 of -14.5. The presence of granophyric texture in the three, aligned
domes of the Northern Dome Field (NDF) may indicate a shallower equilibration depth
coupled with a depressurization and volatile loss event.
Fractional melting of either upper or lower crust or a combination of the two has
been ruled out as a possible genesis scenario for the BVF rhyolites based on both isotopic
and trace element constraints. Limited upper crustal assimilation coupled with fractional
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crystallization (AFC) models, similar to those models used to generate geochemically
similar Quaternary rhyolite domes in the adjacent eastern Snake River Plain (ESRP)
province (McCurry et al., 1999), can explain both the trace element and isotopic
signatures in the rhyolites using local basalts of the BVF as a parent material. Results
indicate that 9 to 18 percent upper crustal assimilation into the basalt followed by
extensive fractional crystallization can form the CDF rhyolites. Basaltic recharge to the
system that occurred just prior to eruption was very minor and did not affect the trace
element concentrations or isotope ratios in the rhyolitic magma.
The ability to melt the upper crust greatly controls the amount of assimilation in
these AFC systems. The ~1.4 Ma NDF contains ~5 percent more crustal assimilant than
the younger CDF because the crust under the CDF has become more refractory due
additional basalt injections over the past 1.4 Ma. The Quaternary ESRP rhyolites contain
significantly less crustal assimilant as compared to the BVF rhyolites because the crust
under the ESRP is even more refractory. This is a result of voluminous Tertiary rhyolite
eruptions and significantly greater amounts of basalt flooding the crust for the past 2 Ma
under the ESRP.
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Chapter 1 INTRODUCTION
1.1 Problem Statement and Objective:
This study focuses on the petrogenesis of the rhyolites in the bimodal basalt-
rhyolite Blackfoot Volcanic Field (BVF), located in southeastern Idaho (Figure 1). It
examines their possible genetic association with adjacent and coeval rhyolites of the
ESRP, Basin and Range or a hybrid combination of both provinces. The overlying,
larger issue is: how are rhyolites created in bimodal volcanic systems? A starting point to
answer this question is to determine a hypothetical reservoir for the basalt. The second
step is to determine if there is a petrogenetic link between the basalt and rhyolite of the
volcanic field. If so, the final step is to develop a model as to how this precursor or
parent magma must evolve via combinations of one or more of the following processes:
magma recharge, crustal assimilation, fractional crystallization or fractional melting to
produce the rhyolite. The primary focus of this work will be on the Central Dome Field
(CDF) (Figure 1, Figure 2).
In this study, I evaluate the following hypotheses:
1. Each set of spatially separated domes within the BFV (Figure 2) is chemically
homogeneous. Presuppositions to support this hypothesis include the fact that both sets
of domes, the Central Dome Field (CDF) and Northern Dome Field (NDF), are aligned
(Figure 2) and each dome within a set is within one kilometer of its neighbor, thus
indicating a similar magma source for each set. Radiometric ages (Table 1) imply that
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domes within each dome field are contemporaneous. Detailed major- and trace-element
geochemistry (Chapter 3) shows homogeneity within each dome field.
2. Rhyolites of the BVF form as a result of extreme fractional crystallization of a parent
magma with limited amounts of crustal assimilation, similar to the Quaternary rhyolites
of the ESRP (McCurry et al., 1999), not as a result of lower crustal melts. The BVF lies
off the southeastern flank of the volcanically active bimodal basalt-rhyolite Eastern
Snake River Plain (ESRP) in the Basin and Range physiographic province (Fiesinger et
al., 1982) (Figure 1). The BVF rhyolites are Quaternary in age (B. Nash, personal
communication referenced in Luedke and Smith, 1983; Leeman and Gettings, 1977;
Armstrong et al., 1975) (Table 1) and overlap in age with Quaternary ESRP (QESRP)
rhyolite domes (e.g. Big Southern Butte, East Butte, Unnamed Butte) (Armstrong et al,
1975; Kuntz et al., 1979; Spear, 1979; McCurry, 1999). Bulk major- and trace-element
geochemistry (Chapter 3) for the QESRP and BVF rhyolites are similar with extremely
low concentrations of Sr and Eu, indicating large amounts of feldspar fractionation. Nd
and Sr isotope ratios of the BVF rhyolites limit the amount of crustal assimilation based
on models using regionally exposed Archean rocks and crustal xenoliths
Figure 1 (overleaf) A hillshade composite DEM (digital elevation model) map of southeastern Idaho including Quaternary rhyolites of the ESRP (e.g. Big Southern Butte, East Butte, Unnamed Butte), presumed cryptodomes along the southern margin of the ESRP (e.g. Ferry Butte and Buckskin Dome) and the Central Dome Field and Northern Dome Field of the BVF. A set of trend lines that run from the Quaternary ESRP rhyolites into the BVF rhyolites and pass through the hypabyssal intrusions roughly approximates the same orientation of a geophysical anomalous zone (Peng and Humphreys, 1998). Dashed lines represent rift zones on the SRP and are roughly parallel to Basin and Range structures. Image modified from McCurry and Morse, (2003).
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(Leeman et al., 1985) present SRP volcanic rocks (Chapter 4). The parental (basaltic)
magmas for both the QESRP and BVF rhyolites have similar geochemical and isotopic
signatures (Pickett, 2004). The lower crust underneath the ESRP and BVF is linked by a
northwest-southeast geophysical anomalous zone, indicative of lower crustal melt (Peng
and Humphreys, 1998), and is roughly represented by the trend lines on Figure 1 and
displayed in cross-section in Figure 3. Lower crustal melting models can not model the
isotopic signatures of the BVF rhyolite and result in multiple subsequent small volume
partial melts of large areas of the lower crust.
3. The BVF rhyolites can be linked to the olivine tholeiite basalts of the BVF by a
petrogenetic model and the BVF basalts are the parent magma. Presuppositions are not
as strong for this hypothesis but isotopic data plotted on mixing hyperbolas indicates less
than 25% crustal component in the CDF rhyolites. Modeling such a large range in
chemistries is difficult due to changes in precipitating phases and the lack of data on
intermediate rocks in the bimodal field, but petrogenetic models like EC-AFC (Spera and
Bohrson, 2001; Bohrson and Spera, 2001) and MELTS (Ghiorso and Sack, 1995) show
that the basalts and rhyolites can be petrogenetically linked (Chapter 4). Partial melts of
Figure 2 (overleaf)
Hillshade DEM of the BVF showing the Quaternary rhyolites of this study (labeled, in black), basalts of the BVF (bright white) and Quaternary Basin and Range style normal faults exposed within the basalts (grey lines within the white basalt lava fields). Note that the dome fields form linear features that do not correspond with localized north-northwest Basin and Range faulting activity. Figure adapted from Oriel and Platt, 1980; Mitchell and Bennett, 1979; Pickett, 2004.
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lower crust, a model suggested by other workers for similar rhyolites (e.g. Christiansen et
al., 1986; Grunder, 1992; Orozco-Esquivel et al., 2002) including topaz rhyolites in the
Basin and Range (Figure 4) is ruled out by a combination of high Rb/Sr ratios and the
isotope systematics of the Archean aged lower crust. Hanna Nekvasil (personal
communication, 2004) has produced rhyolites from experimental melts starting with
Snake River Plain olivine tholeiites. Collectively, these three postulates strongly support
the given supposition.
Table 1. Radiometric ages and associated references for rhyolite domes and outcrops indicated in Figure 2. Note that the samples from the CDF (China Hat, China Cap and North Cone) are significantly younger than samples from NDF or Sheep Island. The nature of the listed uncertainties is unclear.
Sample Age (Ma) Method Source Northern Dome West 1.59 +/- 0.06 K-Ar Luedke and Smith, 1983* Northern Dome Center Not Determined -- Northern Dome East 1.41 +/- 0.15 K-Ar Luedke and Smith, 1983* Sheep Island 1.4 +/- 0.2 K-Ar Luedke and Smith, 1983* China Hat
~0.05 0.04 +/- 0.02 0.08 +/- 0.04
0.061 +/- 0.006
C-14 K-Ar K-Ar K-Ar
Leeman and Gettings, 1977 Armstrong et al., 1975 Armstrong et al., 1975 Pierce et al., 1982
China Cap
0.058+/- 0.007 ~0.05
0.1 +/- 0.1
Ar-Ar C-14 K-Ar
Heumann, 2004 Leeman and Gettings, 1977 Armstrong et al., 1975
North Cone ~0.05 C-14 Leeman and Gettings, 1977 * Referenced as B. Nash, personal communication in Luedke and Smith, 1983
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1.2 Location of the Study Area:
The study area is located in the Basin and Range province, near the juncture of the
Basin and Range and ESRP physiographic provinces (Figures 1, 2, 4) in southeastern
Idaho. Rhyolites of the BVF are located between 15 and 45 km north of Soda Springs,
Idaho and are contained on three adjacent north to south 7.5’ quadrangle maps: Little
Valley Hills, Henry, and China Hat (Figure 6). Three discrete groups of rhyolite domes
(Figure 2) and their field relationships with local geologic features were studied.
The Northern Dome Field (NDF) consists of three unnamed, locally erosionally
incised dome-like rhyolite outcrops located in the southwest corner of Bonneville
County, to the east of Crooked Creek Flat, 40 to 45 km north of Soda Springs. The three
domes are aligned approximately S60oE and I have informally named them Northern
Dome West, Northern Dome Center and Northern Dome East (labeled on Figure 2).
Sheep Island is a rhyolite outcrop located in the Blackfoot reservoir in Caribou
County and is accessible only by boat.
The Central Dome Field (CDF) contains three very conspicuous rhyolite domes
known locally as China Hat, China Cap (occasionally called Middle Cone in older
literature) and North Cone, located in north-central Caribou County 15 to 20 km north of
Soda Springs, at the southern end of the Blackfoot Reservoir (Figure 2). They are
aligned along a N33oE liniment.
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1.3 Regional Geologic Setting:
The crust in southeastern Idaho is composed of Archean crystalline basement
(Figure 5) topped by a thick succession of Proterozoic to Mesozoic sedimentary rocks
(Armstrong and Oriel, 1965; Mueller et al., 2002; O’brien et al., 1995). From the
Paleozoic to the early Mesozoic, the study area was in a transitional zone between a
miogeosyncline to the west and the corresponding platform to the east. The thickness of
units deposited in this setting is directly related to the positioning of the transition zone
and thus varies throughout the Paleozoic and early Mesozoic with the total thickness
between 4.5 and 10 km (Armstrong and Oriel, 1965). The thicknesses of the basal
Cambrian unit and Proterozoic rocks are unknown.
Beginning at approximately the Jurassic/Cretaceous boundary, thrust faulting
associated with the Sevier Thrust Belt began in and around the study area. While no
thrust faults are shown to lie directly beneath the study area, the Paris Thrust is located
approximately 15 km south of the CDF, and the Meade Thrust is approximately 5 km to
the northeast of the NDF (Figure 6) (Armstrong and Oriel, 1965; Dorr et al., 1987; Link,
1982). At least five major tectonic movements are recorded along the Paris Thrust
beginning at the Jurassic/Cretaceous boundary and ending in the Coniacian (~88 Ma)
while movement along the Meade Thrust began later (in the Albian ~110 Ma) and ceased
at approximately the same time (DeCelles, 2004). Minimum horizontal movement on
each of these faults was 8 to 16 km (Armstrong and Oriel, 1965; Wiltschko and Dorr,
1983) and may have totaled over 200 km (DeCelles, 2004). Thrust faulting to the east, in
western Wyoming, continued into the Eocene (Armstrong and Oriel, 1965).
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Figure 4. The locations of Cenozoic topaz rhyolites (black dots) and general physiographic provinces (dotted lines). The Blackfoot Volcanic Field (BVF) is shaded and location of Big Southern Butte (open circle) on the ESRP is given for reference. The solid line is the 0.706 Sr-isotopic line (inferred to separate accreted terrains from continental terrains with Proterozoic or older basement). Adapted from Finton et al., 1991; Pickett, 2004; Christiansen, et al., 1986).
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Figure 5. Regional cratonic map. The BVF (labeled with a “B”) overlies Wyoming Archean basement. Other abbreviations are Belt = Proterozoic Belt Supergroup, GFTZ = Great Falls Tectonic Zone, CH = Challis Volcanics, AR = Absaroka Range, Y = Yellowstone, SRP = Snake River Plain, (modified from O’Brien et al., 1995).
South of the BVF, thick successions of miogeosynclinal Paleozoic rocks continue
into the Nevada and western Utah portions of the Basin and Range Province with similar
thrust faulting (Royse, et al., 1975 in Link, 1982). Extensional tectonism started forming
block faulted Basin and Range type mountains in the Oligocene that generally trend north
to northwest in the BVF. The basins within the BVF are filled with Miocene or younger
sediments (Fiesinger et al. 1982; Hughes et al., 1997). Pierce and Morgan (1992)
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Figure 6. Map showing the locations of major thrust faults and down-dropped graben normal faults in the area. The three quadrangle maps containing the rhyolite domes investigated in this study are shown by the boxes (from north to south: Little Valley Hills, Henry, and China Hat quadrangles). For larger scale normal faults, see Figure 2. (Modified from Link, 1982, Royse, et al., 1975, and Blackstone, 1977)
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classified the normal faults in the study area as Belt III lesser late Pleistocene faults,
indicating a date of last offset greater than 15 ka with total escarpments of less than 200
m. Seismic activity currently exists along some of these faults. Some of these faults that
the cut the basaltic flows in the BVF are shown on Figure 2.
The Basin and Range province contains bimodal basalt-rhyolite volcanic systems
in addition to predominately basaltic and differentiated alkalic systems (Christiansen and
Lipman, 1972). Fitton et al. (1991) noted that the bulk of more recent basaltic volcanism
occurs at the margins of the Basin and Range province. Basaltic volcanism that is less
than 5 Ma is more mafic than older basalts in the Basin and Range, and is
compositionally similar to the basalts of the BVF (Pickett, 2004). Rhyolites from the
Basin and Range are variable ranging from voluminous oxidized ignimbrites erupted
from calderas to low volume topaz rhyolites emplaced as domes and flows (Christiansen
et al., 1986; Gans et al., 1989). A close spatial relation exists between the location of
topaz rhyolites (Figure 4) and late Cenozoic extensional faulting (Basin and Range, Rio
Grande Rift and the Lewis and Clark line in Montana), which would include the CDF
within the BVF (Christiansen, et al., 1983; Christiansen, et al., 1986). The central and
eastern parts of the Basin and Range Province are underlain by Proterozoic basement
(Gans et al., 1989), in contrast to the Archean rocks under the BVF (Figure 5).
Adjacent to the north of the study area, the Yellowstone - Snake River Plain (Y-
SRP) province is a 600 km long time-transgressive series of rhyolitic caldera forming
eruptions starting in southwestern Idaho at ~15 Ma and proceeding to the Yellowstone
plateau with eruptions from 2 to 0.6 Ma (Armstrong et al., 1975; Pierce and Morgan,
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1992). These rhyolites, emplaced by ignimbrites and lava flows are exposed along the
flanks of the SRP, in the Owyhee area of southwestern Idaho and on the Yellowstone
plateau (Hughes and McCurry, 2002; Pierce and Morgan, 1992) and may be up to 8 km
thick in places (Sparlin et al., 1982). Petrologic and geochemical data indicate a change
in the genesis of the rhyolites at ~114o west longitude (the Twin Falls eruptive center),
which may represent a change in lithospheric terranes (Hughes and McCurry, 2002).
Tertiary rhyolites in the eastern Snake River Plain (ESRP) are covered by
numerous individual flows of dominantly olivine-tholeiite basalt (OTB), each generally
less than 10 m thick, intercalated in some regions with sediments, that cumulatively form
a layer 0.5 to 1.5 km thick. Basalt flows are ubiquitous on the ESRP with the greatest
thickness found along the Axial Volcanic Zone (Hackett and Smith, 1992; Kuntz et al.,
1992; Whitehead, 1992). They range in age from subsurface basalts at ~4 Ma to the
Holocene such as the 2.1 ka Craters of the Moon (COM) evolved basalts (Figure 1,
Kuntz, et al., 1992) and Wapi-Kings Bowl OTB and, in contrast to the Tertiary rhyolites,
exhibit no time-transgressive characterizations in their spatial distributions. Basalts are
generally erupted from vents aligned along northwest-trending volcanic rift zones, nearly
parallel to the northwest-trending Basin and Range structures. While the basalts are
somewhat variable in chemistry, no correlation exists between location or timing of
eruptions and chemistry (Kuntz et al., 1992).
Quaternary rhyolitic volcanism is expressed only in isolated locations such as Big
Southern Butte, East Butte, Unnamed Butte, and the high silica components of Cedar
Butte (Figure 1) along the Axial Volcanic Zone which have been studied by Leeman
(1982a), Spear and King (1982), Bretches (1984), Hackett and Smith (1992), Kuntz et al.
15
(1992), Hughes et al. (1997), McCurry et al. (1999), and McCurry and Ganske (2005).
These rhyolites are petrologically and geochemically distinct from the voluminous SRP
Tertiary rhyolites and the Quaternary rhyolites associated with the current position of the
proposed hot spot plume located at Yellowstone National Park (McCurry et al., 1999;
Pierce and Morgan, 1992). Petrogenetic studies indicate that these rhyolites formed
primarily by extensive fractional crystallization of mafic parental magma with very
limited crustal assimilation (McCurry, et al., 1999, McCurry and Ganske, 2005).
McCurry et al. (1999) further suggest that there may also be a NW-SE trending rhyolite
association that includes the domes listed above and some cryptodomes along the
southern margin of the ESRP (e.g. Ferry Butte and Buckskin Dome – Figure 1). Peng
and Humphreys (1998) describe an anomalous geophysical zone beneath and southeast of
the ESRP that they interpret as partially molten lower crust (Figure 3). This trend
roughly parallels the trend mentioned above and an extension of this trend to the
southeast would include the rhyolite domes located in the study area (Figure 1). They
place the upper-crust, lower-crust boundary at 21 km under both the ESRP and BVF.
The ESRP also contains intermediate compositions (e.g. COM, Cedar Butte,
Unnamed Butte) that represent varying degrees of magma mixing coupled with fractional
crystallization. The combination of the intermediate composition COM and Cedar Butte
rocks produce smooth, curved trends that link the least evolved COM magmas and
Quaternary ESRP rhyolites (Big Southern Butte type). These trends show the complete
evolution via fractional crystallization between the evolved basalts and rhyolites with
very limited crustal assimilation (McCurry et al., 1999). Unnamed butte and correlated
16
rocks from core hole CH-1 indicate mixing and hybridization of ESRP basalts and
rhyolite formed by fractional crystallization (McCurry et al., 2003, McCurry and Ganske,
2005).
1.4 Previous Work:
While significant work including extensive mapping, petrology, geochemistry,
and radiometric dating has been accomplished on the basaltic rocks in and around the
study area (Armstrong et al., 1975; Fiesinger et al., 1982; Pickett, 2004), rhyolites of the
area have received less attention. Locals and westward travelers along the Oregon Trail
used the conspicuous China Hat as a landmark in the mid to late 1800’s (Carney, 1998).
Mansfield (1927) was the first to document the rhyolites in the study area during
reconnaissance surveys between 1906 and 1912. He noted and roughly mapped both the
size and position of eight rhyolite exposures including the three domes of the CDF,
(China Cap was called Middle Cone by Mansfield), three domes of the NDF (Figure 1,
Figure 2), and two islands in the Blackfoot Reservoir, presumed to be Sheep Island and
Gull Island. He also published a brief petrological characterization of the rhyolites and
hypothesized that basalt flows both pre- and post-date the emplacement of the CDF
rhyolite. Mansfield’s emplacement history involves three rhyolitic eruptive episodes
separated by three basaltic episodes and he attributes many of the faults in the area (Oriel
and Platt, 1980) to collapsed lava tubes.
Since Mansfield’s study, little work outside of some radiometric dating has been
published concerning rhyolites in the study area. Mabey and Oriel (1970) note an
anomalous gravity low near China Hat and hypothesize that one cause could be a granitic
17
intrusion at depth. Dayvault and others (1984) evaluated the CDF as a potential source
for uranium, identified post-emplacement topaz crystals in lithophysal cavities, and
classified the domes as “nonproductive topaz-bearing rhyolites”. Christiansen et al.
(1986) agreed with this classification the CDF rhyolites and noted that the chemistry of
the CDF rhyolite is similar to that of other topaz rhyolites in the western United States
that are enriched in the traditionally incompatible elements. Radiometric dating by
previous workers for the three domes of the CDF, Sheep Island, and two of the NDF
outcrops are given in Table 1. Nd- and Sr-isotope ratios from the China Cap dome were
obtained by Dr. Michael McCurry prior to the initiation of this study and indicate that
there may be more crustal component present in the BVF rhyolites than in the Quaternary
ESRP rhyolites (M. McCurry, personal communication, 2001). Most recently, Heumann
(2004) conducted U-series isotope studies on the CDF to investigate magma residence
times. He concludes that open-system magma evolution has disturbed the Sr isochron
and cannot be used to evaluate magma residence times. He determined that a U-Th
fractionation event occurred shortly before the eruption but that further study of the U-
series disequilibrium and glass–mineral ages would be needed to evaluate the residence
time of the magma.
Additional studies have characterized the basalts in the bimodal BVF. In addition
to Mansfield’s reconnaissance mapping, other workers have refined the extent of the
basalt flows (e.g. Oriel, 1968; Armstrong, 1969; Mitchell and Bennett, 1979; Oriel and
Platt, 1980). Fiesinger et al. (1982) suggested that the basalts in the study area share
many characteristics with the olivine tholeiites of the ESRP whereas Christiansen et al.
(1986) “…prefer to place the development of the Blackfoot (Volcanic) Field in the
18
context of Basin and Range extension”. Pickett (2004) demonstrated that the BVF
basalts share eruptive styles with Basin and Range basalts but have major, trace, and
isotopic signatures that are more similar to ESRP olivine tholeiites. She concluded that
the basalts of the BVF are derived originally from old mantle lithosphere (as are ESRP
basalts – Hughes et al., 2002A) and that they have undergone extensive fractional
crystallization of olivine and plagioclase (and perhaps clinopyroxene). Along with the
fractional crystallization, between 10 and 20 percent bulk assimilation of upper
continental crust component, depending on crustal xenolith isotopic compositions, is
required (Ford and McCurry, 2003) to produce the isotopic signatures of the most
chemically evolved basalts from the most primitive.
19
Chapter 2 METHODS
2.1 Mapping and Sampling Field Work:
Over 55 rhyolite samples were collected for this work. Possible sample collection
locations were determined by finding high points (summits), spines or ridges on the
domes as these areas provided the greatest chance of finding in situ rhyolite as opposed to
colluvium. In the NDF, incised valleys also contain in situ rocks. Only non-weathered,
unaltered, dense and glassy samples were collected for geochemistry. Dense rocks that
contained less pumice-like texture were sampled preferentially over light and frothy
samples. The dense rocks have less surface area available for weathering. Obsidian
without lithophysae or spherulites was considered to be the highest quality sample for
geochemistry. Sample locations were identified using topographic maps and a GPS
receiver, and noted in a field book along with a preliminary hand sample descriptions.
Mapping, including the overall extent of the rhyolite, sample positions and
volcanic features, was done in the CDF using topographic base maps with Mylar overlays
and used to create a geologic map and cross section at a 1:12,000 scale (Plate 1). In
addition to the domes, areas adjacent to the domes including fault scarps, crater-like
depressions and locally exposed tephra were examined to constrain the timing of the
rhyolite emplacement with respect to local basalts and faulting events. Animal burrows
in loess near the domes were investigated to see if they contained any rhyolitic ash or
colluvium. There was no systematic measurement of flow banding attitudes within the
domes. Two areas that received particular attention were the northwest side of China
20
Cap and the south side of China Hat. Both were mentioned by Mansfield (1927) as
places where basaltic activity might have occurred after the dome formation.
2.2 Petrographic Methods:
Twenty-six standard 27 X 46 mm thin sections were produced by two separate
outside labs for petrographic analysis (Spectrum Petrographics, Vancouver, Wash and
University of Utah Thin Section Laboratory, Salt Lake City, Utah). Thin sections from
all three dome sets were examined using a standard petrographic microscope for the
purpose of determining 1. phenocryst populations, 2. textural relationships, 3. basaltic
enclave properties and their relationship to the surrounding rhyolite (where present), 4.
paragenesis , and 5. which phases were in textural equilibrium with each other. Specific
samples that represented the textures and phenocryst populations of the collected rocks,
but without alteration products, were selected for more detailed analysis including point
counts and electron microprobe analysis (see section 2.8).
2.3 Sample Geochemistry Preparation:
A total of 39 dense, non-weathered samples or obsidian, twenty-one from the
CDF, thirteen from the NDF, one from Sheep Island and four “grab” samples from Big
Southern Butte (on the ESRP), were prepared for major- and trace-element geochemical
analysis using standard operating procedures at the Idaho State University Laboratory for
Environmental Geochemistry (ISU-LEG). First, dirt and weathering rind, or other
surface alteration was removed via compressed air, a stiff brush and tap water, or with a
trim saw. Samples were then either air-dried or dried in an oven at ~40oC. A four cm-
21
thick steel plate, used as a base for crushing the sample, was scrubbed with a wire brush,
cleaned with compressed air and wiped with alcohol. The steel plate was then pre-
contaminated by reducing a small (50 to 100 gram) piece of the cleaned sample to pea-
sized gravel with a four pound crack hammer. The resultant rock pieces and dust were
removed using a dedicated synthetic hair brush and discarded, thus increasing the
likelihood that any remaining fragments of rock lodged in the irregular surface of the
steel plate were from the sample of interest.
Following pre-contamination, samples ranging in mass from 550 to 950 g were
first reduced to three centimeter sized pieces or smaller and visually inspected for any
“foreign material”. Any pieces containing xenoliths, basaltic enclaves or remaining
weathering rind were discarded and the sample was further reduced to pea-sized gravel.
The powdered and graveled sample was then removed from the plate with the synthetic
hair brush and passed through a riffle splitter until two separate aliquots of approximately
30 mL were produced from the original sample. For major-element geochemistry (i.e.
ICP-AES), a tungsten-carbide puck mill or shatterbox reduced the sample to a fine
powder (<200 mesh or about 75 microns) in less than one minute. Trace-element
geochemistry (conducted with INAA) required the use of an alumina puck mill for two to
five minutes to obtain a fine powder. Powdered samples were stored in new poly vials to
hold for future processing. Cross contamination between samples was kept to a
minimum by wearing vinyl gloves and by thoroughly cleaning all items used with
compressed air and alcohol wipes between each sample processed.
Samples analyzed at the University of Florida Department of Geosciences (by Dr.
John Chadwick) for radiogenic isotopic analysis, followed a similar process to that
22
outlined for INAA above, that is, using the alumina puck mill. Additional samples sent
to the University of Arizona Department of Geosciences (Dr. Paul Wetmore) were
processed differently. Because they were analyzed for lead isotopes, the samples could
not be broken with steel instruments, cut with a trim saw or otherwise come in contact
with metals. Samples were closely examined to identify old hammer marks or previous
trim saw surfaces and these areas were marked with permanent marker. Samples were
then put into a brown paper bag and placed in a heavy cardboard box and then struck
with heavy wooden dowel rod and broken into one to two cm pieces. These pieces were
then examined and any piece with a hammer marked surface was discarded. A portion of
the broken sample (50 to 100 g) that had never been struck with metal instruments was
forwarded to UA for further preparatory work. At this time, isotopic results from UA
have not been completed but may be obtained from Dr. Michael McCurry at Idaho State
University.
2.4 ICP-AES Methods and Uncertainties:
Inductively Coupled Plasma-Atomic Emission Spectroscopy (ICP-AES) was
performed at the ISU-LEG following standard operating procedures, briefly outlined
below, on a Jobin-Yvon 70 Plus ICP-AES with both a sequential monochronometer with
a one meter focal length and simultaneous polychronometer with a ½ meter focal length.
This instrumentation and methodology was used to measure concentrations of the
following oxides (wt. percent) and elements (ppm): SiO2, TiO2, Al2O3, FeO, MgO, CaO,
Na2O, K2O, P2O5, Ba, Zr, and Sr.
23
A 0.1000 gram (+/- 0.0003 grams) aliquot of rock powder (prepared with the
tungsten carbide puck mill) was combined with 0.500 grams of Lithium Metaborate
(LiBO2) flux in a graphite crucible and fused in a muffle furnace at 1025oC for 25
minutes. The resultant molten glass was then poured into a beaker containing ~80 mL of
7 percent nitric acid and stirred using a magnetic stir bar until the sample was completely
dissolved. This solution was then transferred to a 100 +/- 0.8 mL class A volumetric
flask, and the beaker and stir bar rinsed repeatedly with 7 percent nitric acid with the
effluent of this rinsing process added to the flask. The total volume of the solution was
then brought to 100 mL resulting in a 0.1 percent solution (0.1 g of rock in 100 mL)
which is used to determine the concentration of Ba, Zr, and Sr. A 10 +/- 0.02 mL aliquot
of this concentrated solution, measured using a class A calibrated pipette, was then
diluted to 100 mL (in a class A volumetric flask) to yield a 0.01 percent solution which
was used to determine the oxide concentrations.
Samples were analyzed using the ICP-AES on several different dates resulting in
slightly variable uncertainties from run to run. The following are the largest analytical
uncertainties for standards, expressed as relative standard deviations (RSD’s) at the 95
percent confidence interval: SiO2, Al2O3, CaO, and Na2O within 2 percent, FeO within 4
percent, and K2O, Ba, and Zr within 10 percent while TiO2, MnO, MgO, and P2O5 were
large due to the very small (<0.1 wt. percent oxide) quantities in the samples. The
uncertainty for Sr was also large as it was under 20 ppm for all samples and under 10
ppm for many samples.
24
2.5 INAA Methods and Uncertainties:
Instrumental Neutron Activation Analysis (INAA) was used to determine the
following oxides (wt. percent) and elements (ppm): FeO, Na2O, Sc, Rb, Cs, La Ce Nd,
Sm, Eu, Tb, Yb, Lu, Hf, Ta, Th, and U. Approximately 0.7 grams (measured to the
nearest 0.0001 gram) of the previously prepared powder (using the alumina puck mill)
was heat sealed in a 2/5-dram reactor safe polyethylene vial which in turn was heat sealed
in a 2-dram polyethylene vial (double encapsulation). Encapsulated samples, along with
appropriate NIST and in-house standards, were sent for activation in the 1 MW Mark II
TRIGA research reactor at Oregon State University where they underwent irradiation for
2 hours at a neutron flux of 3 X 1012 n cm-2 s-1. Samples were then sent back to the ISU-
LEG where gamma emissions were sequentially counted with a high purity Ge (HPGe)
detector over three time intervals. Short-lived nuclides were counted in an initial count
sequence, nuclides with half-lives between ~100 and ~1000 hours were counted in an
intermediate sequence that uses a longer counting time, and long-lived (half lives >1000
hours) nuclides were counted after the activities of the short-lived and some intermediate
activities had decreased, generally 4 to 8 weeks after irradiation. In some cases, more
than one counting sequence yielded acceptable results for a particular decay series (e.g.
Fe). Data was then reduced using in-house software written by Dr. S. S. Hughes of Idaho
State University.
The following were the largest analytical uncertainties for standards, expressed as
RSD’s at the 95 percent confidence interval: FeO, Sc, La Ce, Sm, and Th within 2
percent, Na2O, Rb, Cs, Nd, Tb, Yb, Lu, and Hf between 2 percent and 4 percent, Ta and
Eu between 5 percent and 10 percent and U up to 15 percent. Although INAA can detect
25
Sr and Ba, results were poor and thus the ICP data was used for these elements. For the
oxides FeO and Na2O, results from both techniques gave similar results in most cases and
provided cross-checks on the two techniques.
2.6 Radiogenic Isotope Methods:
Five samples (four from the CDF and one from the NDF) were submitted to Dr.
John Chadwick for 87Sr/86Sr and 143Nd/144Nd analyses at the University of Florida
Department of Geosciences. Approximately 0.1 g of rock powder, prepared following
the same procedure as detailed for INAA powders above, was dissolved in a 100oC HF-
HNO3 mixture in Teflon vials. The dry residuum, left after evaporating the acid mixture
in a class 1000 laminar flow hood, was converted to a chloride salt by dissolving it in 2
mL of HCl and re-drying. These salts were then re-dissolved in HCl and run through a
centrifuge to concentrate any remaining solids so that they could be removed. The
resultant, solid-free liquid was then passed through calibrated cation exchange resin
quartz glass columns to separate Sr and Nd. Re-dried solutions were then loaded onto
filaments where isotopic Sr measurements were collected using a Micromass Sector 54
Thermal Ionization Mass Spectrometer (TIMS) and isotopic Nd measurements were
collected on a Nu Plasma Multiple-Collector magnetic-sector Inductively Coupled
Plasma Mass Spectrometer (MC-ICP-MS). For Sr isotopes, NIST NBS-987 was
measured four times during the suite of analysis in which the BVF samples were
measured and had an average 87Sr/86Sr value of 0.710244 as compared to the accepted
value (the average of this standard over the past year on this instrument) of 0.710243 with
26
standard errors better than +/- 0.000002 (2-sigma). For Nd isotopes, USGS standard
reference material BCR-1 yield 2-sigma values of +/- 0.000011.
2.7 Heavy Mineral Separates:
Mineral separation was performed on three samples so that accessory phases
could be concentrated, made into a grain mount and analyzed on an electron microprobe.
The general process is outlined as follows:
(1) Samples were cleaned and trimmed of weathering rind and the steel plate used to
gravel the sample was pre-contaminated as described in section 2.3 above.
(2) Samples were then reduced using a crack hammer in a stepwise process that involved
lightly crushing the sample, passing it through a 0.81 mm brass sieve, and returning what
did not pass to the steel plate for further reduction. By repeating this process many times,
more material that is of the appropriate size is produced and the production of large
quantities of powder that is too small to use in later processing is avoided. The
percussion method (i.e. hammer and steel plate) for breaking the rock is preferred as it
tends to break the specimen along grain boundaries (Hutchison, 1974).
(3) The smaller than 0.81 mm fraction was then washed to remove the less than 50 µm
fine dust by placing the sample in a large plastic container and using a jet of tap water to
thoroughly mix the sample and separate grains that might adhere to each other. After a
short settling time (about 30 seconds), the murky water containing the small fraction
(dust) of the sample was decanted off the top of the larger, settled fraction. This process
was repeated a number of times until the washing water was nearly clear after mixing.
The removal of the fine fraction is important because if it is not removed, the heavy
27
liquids used later in this process can be hopelessly contaminated as the fine particles will
not settle out of the heavy liquids.
(4) The washed sample was then transferred to disposable aluminum bread pans and
placed in an over at ~50oC to dry.
(5) After drying, a 60 mesh (250 µm) disposable nylon sieve (from the Gilson Co.,
Lewis Center, Ohio) was used to obtain a sample that would range in size from less than
250 µm to greater than approximately 50 µm. This fraction was passed through the
heavy liquids, as detailed below, and the larger than 250 µm fraction was reserved in case
more sample was required at a later date.
(6) Approximately 100 mL of the above prepared sample was passed though a separatory
flask that contained bromoform. The sample was stirred many times and heavy portion
was allowed to settle to the bottom of the flask while the light portion (e.g. glass, quartz,
feldspars) floated to the top. The bromoform was then recovered and the light fraction of
the sample discarded. This process was repeated until a sufficient amount of heavy
minerals were collected.
(7) The heavy minerals portion was collected onto filter paper and rinsed multiple times
with acetone and allowed to dry in a fume hood. It is important to note that vinyl gloves
must not be used in this process as they are permeable to acetone. The author
recommends neoprene.
(8) The clean, dry heavy separate was then spread over a clean sheet of paper and a large
hand magnet, covered with a kimwipe, was passed over the sample to remove the
magnetic fraction which was collected and reserved in a glass vial. If there is sufficient
28
sample available, it can be passed through additional heavy liquids (i.e. methylene
iodide) to form additional separates.
(9) The non-magnetic heavy separate was then passed though a Frantz isodynamic
separator with a forward slope of 15o, a side tilt of 15o and a current of 0.5 amps. This
process resulted in two fractions, a magnetically susceptible fraction (e.g. hornblende,
ilminte, biotite, grains with magnetite inclusions) and a non-susceptible fraction (e.g.
zircon, apatite, thorite). It is important to thoroughly clean the Frantz as there are many
places where particle contamination can occur. Hutchinson (1974) gives a generalized
review of the above process and details for Frantz settings for other minerals.
(10) Mineral grains from all three separates (the magnetic, magnetically susceptible, and
non-susceptible) were placed in holes drilled into a labeled, hard, one inch round, 3/8
inch thick nylon spacer (generally available at hardware stores).
(11) These holes were then filled with Epo Tek 301, a low viscosity epoxy (from Epoxy
Technology Inc, Billerica, Mass). Poking air bubbles with a needle or other sharp device
was effective in removing bubbles from the epoxy.
(12) The grain mount was then placed in a ~50oC oven to cure overnight before getting
sent offsite for a microprobe polish.
The above method produced high quality one inch round mineral grain mounts for
microprobe analysis and there were no problems with the nylon substrate in polishing the
mounts at the University of Utah Thin Section Laboratory (Quintin Sahratian, personal
communication, 2004).
29
2.8 Electron Microprobe (EMP) Methods:
One thin section from the CDF (sample G1) and a heavy mineral separate from
the CDF were examined using a JEOL 8900 R Electron Probe Microanalyzer under the
direction of Karen Wright at the Idaho National Laboratory Research Center (IRC) in
Idaho Falls, Idaho. Beam current was set at 20 nA with a 15 kV accelerating voltage and
a beam diameter 10 microns (plagioclase, sanidine, biotite, hornblende, magnetite, matrix
glass and melt inclusions in quartz) or 5 microns (ilmenite and melt inclusions in
magnetite). Integration times were 20 seconds for major elements, 30 seconds for minor
elements and 10 seconds for background. Both rim and core measurements were
performed on plagioclase, sanidine, hornblende and biotite with single measurements on
magnetite, ilmenite, melt inclusions and matrix glass. Thirty second long energy
dispersive spectroscopy (EDS) profiles were made of selected mineral unknowns in order
to define their major-element composition and mineral identification. Specific EMP
methods are documented in Appendix 1.
30
Chapter 3 RESULTS
This section focuses on the results of field, petrographic and geochemical studies.
Field observations are described and photographs of representative areas and a map (Plate
1) of the CDF are included. Phenocryst assemblages and percentages along with textual
descriptions and representative photomicrographs from the three dome fields are
presented in the petrography section. Geochemical results include major- and trace-
element chemistry as well as isotope results and electron microprobe chemistry on
phenocrysts from the CDF. The geochemistry of the BVF rhyolites is then compared to
rhyolites from other physiographic provinces.
3.1 Field Relationships:
Rhyolite emplacement in the BVF occurs at two distinct times (~1.4 Ma and
~0.05 Ma; Table 1). The two dome sets (NDF and CDF) and Sheep Island Rhyolite
(SIR) are spatially separated (Figure 2). Stratigraphic relationships of the incised domes
of the NDF to surrounding geology are obscured both by loess and weathering products.
Likewise, contacts between the SIR and surrounding rock units are covered either by
water or deep mud. The NDF and SIR areas will be briefly discussed first followed by a
more detailed examination of field relations on the younger CDF.
The NDF, located on the Little Valley Hills 1:24,000 quadrangle map, consists of
three domes, roughly aligned along a S60oE liniment. Samples from Northern Dome
West (NDW) and Northern Dome East (NDE) yield K/Ar dates of 1.59 +/- 0.06 and 1.41
+/- 0.15 Ma respectively (B. Nash, Personal Communication, referenced in Luedke and
31
Smith, 1983; Table 1). Luedke and Smith (1983) did not indicate whether the
uncertainties are at the 1-sigma or 2-sigma confidence interval and thus the domes of the
NDF and SIR could overlap in age. Domes of the NDF have a cumulative volume of
approximately 0.45 km3 of rhyolite with Northern Dome West containing 85 percent of
this volume. It is also the largest in aerial coverage, retaining much of its original glassy
carapace although it is deeply incised by tributaries to Crooked Creek on its southern
side. NDW does contain at least one lobe, expressed to the west and southwest, that
appears to have flowed out onto the Crooked Creek Flat. Between NDW and NDM,
there are large pieces of quartz pebble conglomerate float, some over one meter in size,
that contain slickensides. Reasonable access is available with a high clearance vehicle
only and permission to trespass should be secured from Gentile Valley Land and Cattle
Company, Grace, ID for all visits.
The NDF rhyolite ranges in color from massive white to pink and purple flow
banded sections (Figure 7). Vesicle content ranges from near zero in dense black and
brown obsidian to highly vesiculated pumiceous rhyolite. Lithophysae cavities are
common and some are up to eight cm in diameter (Figure 8). Visible phenocrysts include
quartz, plagioclase, sanidine and biotite in a dense to vesiculated glassy matrix. The
rhyolite also contains sparse basaltic enclaves, some with crenulated contacts with the
rhyolite and some with angular contacts.
The SIR, located in Blackfoot Reservoir on the Henry 1:24,000 quadrangle map,
is similar in age to rhyolites of the NDF (B. Nash, Pers. Com. referenced in Luedke and
Smith, 1983; Table 1). Sheep Island is entirely composed of rhyolite and covers 0.75
km2, reaching an elevation of approximately 57 m above the surface of the reservoir.
32
Figure 7. Typical flow banding in the NDF with alternating layers of pink and purple colored rhyolite. Located on Northern Dome West in an incised valley, representing the interior of the original dome. Four pound crack hammer for scale.
Figure 8. Lithophysae cavities, up to 8 cm, located on the southwest lobe of Northern Dome West near Crooked Creek Flat. These features are more common and larger in the NDF than in the CDF. Four pound crack hammer for scale.
33
Mansfield (1927) reports the existence of two rhyolite outcrops in the Blackfoot
Reservoir. Gull Island, located approximately 600 m to the south-southeast of Sheep
Island is composed of rhyolite and is likely the second outcrop mentioned by Mansfield.
It reaches a height of approximately 12 m above the reservoir covering only 0.03 km2 and
was not investigated. Total volume of both outcrops was not specifically calculated but
is less than 0.03 km3.
The rhyolite at Sheep Island is generally dense and glassy containing few vesicles
and is sparsely porphyritic containing plagioclase and quartz in hand sample. It contains
alternating purple and pink flow-banded layers (Figure 9) inclined at ~45o. While there is
no evidence of enclaves or lithophysae cavities, the island was not thoroughly searched
for them. Access is by boat only and Sheep Island is public land.
Figure 9. Typical dense, glassy flow banded rhyolite from Sheep Island in the Blackfoot Reservoir. There are few outcrops on Sheep Island. Four pound crack hammer for scale.
34
The CDF, the youngest of the three study locations at ~0.05 Ma (Table 1),
contains three domes aligned approximately N33oE located on the China Hat 1:24,000
quadrangle map. Total volume for the CDF is approximately 0.7 km3 with China Hat
containing 0.5 km3, China Cap with 0.12 km3 and the North Cone with 0.08 km3. The
total volume of rhyolite in the BVF, including the NDF and CDF is approximately 1.2
km3. The glassy carapaces are very much intact with essentially no erosion or dissection
of the domes. In addition to the dome formation, a number of other volcanic and local
tectonic events were examined in and around the CDF including graben formation due to
normal faulting, voluminous basaltic flows, rhyolitic tephra production, and the
formation of craters between two of the domes. These are discussed below.
Numerous north and northwest trending normal faults cut Quaternary basalts in
the BVF (Figure 2; Michell and Bennett, 1979; Oriel and Platt, 1980). These faults are
most commonly exposed in the lowest, central part of the large, down-dropped Basin and
Range valley, where the rhyolite domes are located, and often produce small grabens
(Figure 10). Four of these grabens are especially noteworthy including the grabens
containing Hole in the Rock Lake (Figure 10) and Dike Lake, a small graben along the
south-west flank of China Hat, likely an extension of the Hole in the Rock Lake graben,
and a subtle graben directly south of China Cap (Figure 11) and are labeled on Plate 1.
Hole in the Rock Lake partially fills a graben (Figure 10) that has an offset up to
35 m and has a nearly continuous surface expression in excess of eight kilometers. The
northern half of this distance was mapped as a fault by Oriel and Platt (1980) and they
continue the fault to the north-northwest at least 15 km further. The surface expression of
this fault passes the western side of China Cap and extends past the south-west flank of
35
Figure 10. Hole in the Rock Lake partially filling a graben with scarp offsets up to 35 m in the Blackfoot Volcanic Field basalt. This graben has a continuous surface expression over 8 km long. Picture taken from China Cap, looking northwest
Figure 11. Subtle graben south of China Cap (edge of China Hat on the right side of the photo). Thick loess, actively farmed, has muted the surface expression of the graben. Photo taken from China Cap looking south.
36
China Hat, though the apparent offset has decreased (Plate 1). The talus apron around
China Hat obscures any direct evidence of the fault cutting the dome but there is a small
lobe where the rhyolitic lavas appear to have flowed over the edge of a preexisting
precipice or over a fault scarp. Another large fault is mapped directly south of China Hat
(Oriel and Platt, 1980) with offsets up to 35 m does not appear to offset the dome.
Dike Lake also fills a graben (Plate 1), although the offset is not as great as the
Hole in the Rock graben and the linear continuity is hidden by the Blackfoot Reservoir.
The lake filled graben is exposed just to the northeast of China Cap and if the linear trend
of the graben was extended, the fault would pass either under or through the west side of
China Cap and directly into the north side of China Hat. There is no evidence that this
fault offsets either dome but again, the talus apron and dense vegetation on the north side
of China Hat obscure most intact, continuous units. Additionally, unlike the Hole in the
Rock graben, the Dike Lake graben, while visible to within 1000 m of China Cap, does
not have a continuous surface expression directly adjacent or leading into the domes and
therefore may not exist at these points. No evidence of offset tephra beds due to faulting
is evident in the exposed deposit on the north side of China Hat.
A small graben directly south of China Cap (Figure 11, Plate 1) was not mapped
by Oriel and Platt (1980) as a fault. The offset on this graben is approximately 15 m and
its topography is generally muted by an unknown thickness of loess that is actively tilled
and farmed. A laterally continuous band of dense obsidian glass that exists along the
southern edge of China Cap is not offset by this fault. All of the faults described above
are within four kilometers from each other from north to south and within two kilometers
from east to west and presumably cut the same laterally continuous basalt flow. The
37
basalt flows in the area do not show evidence of on lapping the rhyolite domes.
A tephra deposit at least 17 m thick is exposed on the northern flank of China Hat
and (Figure 12, Plate 1). The beds of the deposit have a 5o dip to the north (away from
the dome) and it contains rhyolite and basalt ash, lapilli and block sized constituents.
Some of the larger blocks of rhyolite are up to 0.5 m and contain rounded edges (Figures
12 and 14). Angular, palagonitized basalt fragments (Figure 13) are common, especially
in the bottom four meters of the section. The percentage of basalt clasts appears to
decrease towards the top of the section.
Figure 12. An overview of approximately seven meters of the China Hat tephra deposit with the author for scale. There is substantially less basaltic component to the tephra in this middle portion of the deposit. The pit has been extensively mined since this picture was taken in 2002. Picture by Michael McCurry.
38
Figure 13. Alternating layers of angular palagonitized basalt, both block and lapilli sized rhyolite, some reversely graded and mixed with basalt clasts, and rhyolitic ash in the China Hat gravel pit. Loess tops this section. Wooden staff is 1.6 m long.
Figure 14. Large rhyolite blocks with rounded edges located at the top of the shovel handle. Note the deformed beds around the approximately 30 cm rhyolite clast. This section starts approximately 0.5 m above the top of the section in Figure 11. Wooden staff is 1.6 m long.
39
Near the top of the tephra deposit, there is a layer of block-sized rhyolite at least one
meter thick topped with 0.5 m of very fine ash. On top of this final ash layer is colluvial
talus shed from the dome mixed with loess (Plate 1). While the total section is at least 17
m thick, six meters near the top of the deposit could not be mapped due to overburden.
Since this study, the tephra deposit has been extensively mined for aggregate.
The aerial extent of this tephra deposit is unknown. Mansfield (1927) reports
rhyolitic ash at least eight meters thick along the bank of the Blackfoot River
approximately 3.25 km due east of the China Cap summit. This horizontally bedded fine
grained rhyolitic ash contains a ½ m thick bed of basaltic “ash” about four meters down.
Mansfield indicates that that this deposit might also be the result of erosion and transport
from local sources and that, “local (erosional) conditions varied considerably within short
distances.” Current day local drainages from China Hat flow towards this location (to the
northeast). A hand-dug well, approximately 100 m northwest of China Cap, contained
one meter of rhyolitic ash under approximately 7.5 m of soil (Mansfield, 1927). Also, a
thin bed of ash, of unknown thickness, attributed to the CDF was discovered in mud
cores taken from Gray’s Lake, located 35 km to the north-northeast of the CDF (Ken
Pierce, personal communication, 2003). It is unclear if this is an air fall deposit or fluvial
deposit but the current day smaller scale drainage pattern is to the northwest from the
CDF, not towards Gray’s Lake.
Rhyolite within the CDF shows little variation except in their degree of
vesiculation. The rocks are very light in color, vitric, and crystal poor with phenocrysts
of quartz, euhedral plagioclase, sanidine, biotite and hornblende. The biotite shows
40
variable degrees of weathering indicated by its luster. Fresh, non-weathered crystals are
shiny, mirror-like and black while those that have been weathered are purple, iridescent
or brown. Radiating sprays of post-emplacement topaz have been documented
(Dayvault, 1984) up to two millimeters in length however I looked for them and was
unable to find any. Both spherulitic and lithophysal textures are evident in the field but
are not common with lithophysae generally less than one centimeters where present.
Dense, black obsidian is common and in flowbands and sometimes associated with
devitrification. Unlike the NDF, no brown colored obsidian was found. Two populations
of basaltic enclaves, up to two centimeters, are common but not abundant and were most
easily found on China Cap. One type is typically ovoid and shows crenulated boundaries
with the rhyolite in thin section while the other type has angular margins. These enclaves
will be further detailed in the petrography section.
Other features in the CDF are noteworthy. The three domes of the CDF align
along a N33oE trend and each dome is slightly oblong in shape along a north or northeast
trend. The slopes of the dome margins are consistently steep with 35 to 45 percent
grades with three exceptions. A coulee nearly encircles China Hat resulting in a
noticeable inflection in the slope approximately halfway up the dome (Figure 15). The
southwestern side of the North Cone is the steepest area within the CDF with a shear cliff
over 10 m high while the northeastern side of the dome is manifested by hummocky
terrain and is the least steep area in the CDF (Figure 16, Plate 1). As one approaches the
tops of all three domes, the terrain becomes less steep. All three domes contain spines
with nearly vertical flow banding and block fields. While spines and pressure ridges
exist on all three domes they are particularly common on China Cap (Figure 17). The
41
Figure 15. A coulee nearly encircles China Hat creating a noticeable inflection in the slope of the dome surface. The dome is approximately 300 m tall. Picture taken from Highway 34 looking west.
Figure 16. Three domes (from left to right, China Hat, China Cap and North Cone) showing relative sizes. Note the hummocky terrain on the northeast side of North Cone (right-hand side of the picture) is in stark contrast to the exceptionally steep area on the opposite side of the dome. Picture taken from the Blackfoot River Road looking west.
collapse of pre-existing spines during or soon after their formation may produce the block
fields (Macdonald, 1972), most common on China Hat. Some of block fields have areas
over 0.05 km2 and contain blocks over three meters in size.
Along the same N33oE trend of the domes, two large pits, now containing
Burchett and Gronewall Lakes, exist between China Cap and North Cone. The Burchett
Lake depression (Figure 18) is slightly elongate to the northeast and has a depth of at
42
Figure 17. Many prominent spines exist on China Cap as well as on the other domes. Some of the spines can be 10 m above the local dome topography. Photo taken from China Cap Road looking northwest.
Figure 18. Burchett Lake partially fills a crater between China Cap and North Cone and lies within the linear trend created by the CDF. Note fields on either side sloping away from the crater rim. Photo taken from China Cap looking northeast.
43
least 34 m but this could be considerably greater depending upon the thickness of the
sediments in the bottom of the lake. The Gronewall Lake depression is curvilinear, has a
depth of at least 34 m and forms a moat around the southwestern side of North Cone
(Plate 1). The area between China Cap and North Cone, excepting the craters, is
generally higher in elevation over the surrounding valley by up to 60 m and slopes
downwards along a trend perpendicular to that of the dome alignment (Plate 1) at dips of
five to seven degrees. Basalt flows exposed on the north edge of the Burchett Lake crater
dip a similar direction and amount.
Access to the CDF for all vehicles on paved roads is available along both China
Cap Road and Dike Lake Road and there are other access points that require high
clearance vehicles. The tops of China Hat and China Cap are both public land but
permission for extended periods of access should be secured from David Hubbard of
Soda Springs, ID (China Hat), Lonnie Cellan of Inkom, ID (China Cap), and Joe Elsmore
of Grace, ID (North Cone). Short stops, especially into the gravel pits, should not present
a problem unless the pits are being actively mined.
3.2 Petrography:
A total of twenty-six thin sections were examined using a petrographic
microscope including seventeen from the CDF, eight from the NDF and one from the
SIR. Samples from respective dome sets are all very similar in phenocryst type,
abundance and size as well as textures and generally vary only in the quantity of post
emplacement features such as vapor-phase crystallization, devitrification, spherulitic
aggregate formation, flow banding and vesicle concentration. Additionally, four slides
44
contain enclaves that will be discussed later in this section.
3.2.1 Petrography of the CDF:
The CDF silicic rocks are best described as hololeucocratic hyalo-rhyolites (using
the IUGS classification of LeMaitre et al., 2002) in that they have a color index (M’) of
under ten and contain over eighty percent glass. Common phases include quartz >
sanidine > plagioclase > hornblende ≈ biotite > Fe-Ti oxides (magnetite > ilmenite) and
comprise approximately eight percent of the rock. Accessory phases include zircon >>
apatite > thorite ≈ allanite, with the last two only positively identified in a thin section
made from a heavy mineral separate. As mentioned in the field methods section,
typically dense samples from the glassy carapace were collected and thus the samples
studied in thin section likely under represent the overall crystal content of the domes.
Vesicle concentration ranged from zero to five percent in the thin sections studied,
pumiceous samples should contain a higher proportion of vesicles.
Phenocryst textures:
Quartz: There are two distinct textural types of quartz crystals 1. euhedrally shaped
crystals and 2. strongly embayed crystals. Both types are up to 1 millimeter in size,
occasionally fractured and commonly contain melt inclusions, some of which have
daughter crystals. Many quartz crystals contain small, hourglass shaped embayments
that may be the result of ruptured melt inclusions. Crystals that show embayment-like
features often have some crystal faces that are very straight and otherwise unaltered
(Figure 19). Quartz crystals comprise approximately five percent of the rock on a
45
Figure 19. Strongly embayed quartz likely due to chemical dissolution. Note the unaltered sides retain the very straight edges of the originally euhedral crystal. 10X, crossed polars, long dimension = 0.7 mm.
Figure 20. Square-shaped, skeletal sanidine, indicative of rapid growth, armoring plagioclase. 10X, crossed polars, sanidine is 1 mm on a side.
Figure 21. Blocky magnetite crystal with sharp edges indicating equilibrium with the melt. Note the euhedral quartz along the bottom edge. 10X, plain polarized light, long dimension of the magnetite is 0.4 mm.
46
vesicle-free basis. Euhedral crystals were in textural equilibrium with the melt while
embayed crystals have been resorbed at some point after their original formation.
Sanidine: Sanidine crystals comprise about two percent of the rock on a vesicle-free
basis and the crystals are up to two millimeters on a side. Some crystals show evidence
of inclusion that have leaked along cleavage fracture plains. Euhedral in form, most
crystals are equant and square in cross section while some crystals exhibit evidence of
skeletal growth produced by rapid crystallization. This rapid crystallization can lead to
interesting phenomena where two or three “separate” crystals are in optical continuity
with each other but less than two millimeters apart. It is likely that this is a single
skeletal crystal that was connected in the third dimension. Sanidine crystals occasionally
mantle plagioclase (Figure 20) and were also in textural equilibrium with the melt.
Plagioclase: There are two texturally distinct types of plagioclase in the CDF rhyolites.
The larger-sized crystals, up to two millimeters long, commonly exhibit both Albite and
Carlsbad twinning and comprise approximately one percent of the rhyolite on a vesicle-
free basis. Crystals can be long and slender, mantled by sanidine (Figure 20), broken or
bent and in some cases and show evidence of skeletal growth but all appear to be in
textural equilibrium with the melt. The smaller sized plagioclase crystals are part of the
groundmass and will be described below with the matrix.
Mafic minerals: Hornblende, biotite and Fe-Ti oxide phenocrysts make up far less than
one percent of the CDF rhyolite. Hornblendes often occur as euhedral pleochroic crystals
that exhibit excellent cleavage and have a maximum length of one millimeter. Black,
shiny, euhedral biotite crystals over one millimeter are uncommon but obvious in hand
sample but the biotites preserved in thin section are generally less than 1/2 mm in size.
47
Both magnetite and ilminite are present with the latter always less than 1/5 mm in length.
Magnetite crystals are blocky (Figure 21) and are up to one millimeter in size. All four of
these phases appear to be in textural equilibrium with the melt with no resorption textures
evident.
Accessory minerals: The dominant accessory phase is zircon. Some zircons occur as
inclusions in hornblende while others occur as an intratelluric phase. Apatite is rare and
was only observed as small, less than 1/5 mm acicular crystals in two thin sections. The
existence of thorite and allanite was confirmed with EDS scans from an electron
microprobe (EMP) on a heavy mineral separate and later observed petrographically when
the grain mount for the probe was cut to thin section thickness. Thorite is a tetragonal
mineral with optical properties similar to zircon except that is has a faint apple green
color and is weakly pleochroic.
Matrix: The matrix consists primarily of light brown to colorless glass and vesicles,
some of which appear very long and flat or stretched and varying degrees of spherulitic
aggregate. It also contains microlites of plagioclase crystals, many with swallow tail
forms. These small crystals only occur in long stringers, similar to flow banding that
often appear to flow around larger crystals and sometimes pond around them, perhaps in
previous void spaces. The plagioclase crystals in these stringers are loosely oriented with
the long axis parallel to the length of the stringer (or presumed flow). These
concentrations of microlitic plagioclase can extend across the entire length of the slide
and are found with a very fine grained, less than 0.05 mm, dark, acicular, pleochroic
mineral, likely hornblende, lenticular shaped stretched vessicles and possibly very fine
grained sanidine.
48
Figure 22. Glomerocryst containing quartz, with a melt inclusion, sanidine, and a 1.4 mm long opaque (ilmenite?) mineral. Note this photo shows some of the post emplacement spherulitic aggregate (SA) formation (grey). 10X , plain polarized light.
Figure 23. Same as above but with crossed polars. Some of the crystals are in optical continuity indicating that they may exhibit skeletal growth or are connected in the third dimension. Other cognate pieces contained plagioclase as well.
49
Glomerocrysts: Glomerocrysts containing quartz, plagioclase, sanidine, hornblende and
Fe-Ti oxides make up approximately two percent of the rock. These wall-rock pieces
display intergranular texture (Figure 22, 23) and are in textural equilibrium with the melt
as there are no resorption textures. It appears that the glomerocrysts are similar to the
overall rhyolite in phenocryst assemblage and are predominantly quartz.
3.2.2 Petrography of the NDF:
The NDF rhyolites are also hololeucocratic hyalo-rhyolites and contain the same
major phases as the CDF in nearly the same proportions including quartz (~4.5%),
sanidine (~3%), plagioclase (~1.5%), and hornblende, biotite and Fe-Ti oxides, all far
less than one percent with crystals making up about ten percent of the rock by volume.
Delineation of whether a phase was part of a silicic xenolith or an individual intratelluric
crystal was not made for the NDF rhyolites but they do contain at least some of these
glomerocryst which appear to be in textural equilibrium with the surrounding glass. The
heavy mineral separates EMP mount is dominated with zircon and although allanite or
thorite were not observed, these phases likely exist.
Textures for the two dome sets are very similar for all minerals except for the
following. There were no zones of microlitic plagioclase with swallow tail features and
hornblende found in the NDF rhyolites. Also, unlike the CDF, the NDF did contain some
examples of granophyric texture (the coexistence of graphic and vermicular textures,
Hibbard, 1995) but only in glomerocrysts, not individual phenocrysts (Figure 24).
50
Figure 24. Granophyric texture of glomerocryst in the NDF rhyolites. This texture was not seen in the CDF rhyolite. 20X, crossed polars, field of view ~0.75 mm.
Figure 25. Oscillatory zoned plagioclase within a basaltic magmatic enclave. Note the centers of the crystals exhibit skeletal growth. 10X crossed polars, center crystal is 0.5 mm.
Figure 26. Boxy cellular plagioclase within the basaltic magmatic enclave with rim of later euhedral growth surrounding the quickly formed original crystal. 10X, crossed polars, crystal 1 mm.
51
3.2.3 Petrography of the SIR:
The SIR thin section contains very few crystals, most are smaller than 0.5 mm,
and has substantial post emplacement spherulitic aggregate formation. The thin section
shows flow banding with alternate white and peach colored stripes crossing the slide.
Two textures of plagioclase are present, 1. large euhedral crystals up to 0.5 mm and 2. a
small quantity of microlitic plagioclase. Microlitic sanidine is also present in
small quantities. The slide contains some quartz crystals, at levels less than one percent
by volume.
3.2.4 Mafic Enclaves of the BVF:
While basaltic xenoliths undoubtedly occur within the rhyolite and are ubiquitous,
especially in the lower sections of the tephra deposit, the rhyolite also contains basaltic
enclaves of magmatic origin. A crenulated margin and fine-grained chill texture easily
delineate magmatic enclaves from non-magmatic xenoliths and both the NDF and CDF
contain magmatic enclaves. Some of the margins between the rhyolitic and basaltic
magmas appear to be hybridized.
The mafic enclaves are texturally similar to olivine tholeiites. They contain
olivine crystals that are either unaltered or show evidence of dissolution with crenulated
edges or eroded centers and at least four textures of plagioclase including 1. very fine
grained swallowtail forms, 2. unaltered larger phenocrysts, 3. oscillatory zoned crystals
(Figure 25) and 4. boxy cellular or skeletal crystals rimmed with euhedral plagioclase
growth (Figure 26). In at least one petrographic slide, there is evidence of hybridization
where the basaltic and rhyolitic magmas have begun to mix (Figure 27). This hybridized
52
Figure 27. Zone of hybridization between the rhyolitic magma and basaltic magma. Note small white quartz crystal in the mixed zone (Figure 28). 2.5X, plain polarized light, width of mixed zone ~1.4 mm.
Figure 28. Quartz crystal within the hybridized zone (Figure 27) showing pyroxene- mantled texture with a small gap between the quartz and hornblende. 20X, crossed polars, crystal is 0.25 mm.
Figure 29. Spongy cellular or fingerprint texture in a sanidine crystal approximately 2 cm away from the basaltic magmatic enclave. 40X, crossed polars, width of partially melted zone = 0.05 mm.
53
zone occasionally contains quartz crystals that appear resorbed but does not contain
sanidine (from the rhyolite) or olivine (from the basalt). Quartz within this zone is
pyroxene-mantled (Figure 28). Within the rhyolite, near the basaltic magmatic enclaves,
sanidine crystals exhibit spongy cellular or fingerprint textures (Figure 29). This texture
was not noted in slides that did not also contain a basaltic magmatic enclave.
3.3 Geochemistry of the BVF:
This section first examines the major- and trace-element composition of the BVF
rhyolites, comparing and contrasting the geochemistry of the NDF and CDF to each
other. The BVF rhyolites are then compared to rhyolites in other local provinces
including Quaternary Eastern Snake River Plain (QESRP) rhyolites (e.g. Big Southern
Butte and East Butte), Yellowstone and Tertiary Snake River Plain (YandTSRP)
rhyolites, Cascade rhyolites, Basin and Range rhyolites and topaz rhyolites. Unless
otherwise noted, the major-element data are normalized to 100% on an anhydrous basis
with all iron represented as FeO(t). As noted in the methods section, the thirty-five BVF
rhyolite samples were collected using non-specific criteria that included the degree of
weathering and density/glassiness. Great care was used to remove any weathering rind,
alteration, enclaves or xenoliths and some samples were run multiple times to check data.
Twenty-one samples were analyzed from the CDF, thirteen from the NDF and
one from the SIR while five samples were analyzed from Big Southern Butte (and
included in the QESRP dataset). Major- and trace-element compositions and sample
locations are given in Table 2.
54
3.3.1 Major-Element Geochemistry:
Samples from the both the CDF and NDF are classified as a high-K rhyolite on a
Le Bas, et al. (1986) total alkali silica (TAS) diagram and are slightly metaluminous with
molecular (CaO + Na2O + K2O) > Al2O3 > (Na2O + K2O). Sheep Island is also a high-K
rhyolite but is very slightly peraluminous. Note that there is only one analysis of this
location and that a decrease in Al2O3 content by less than two relative percent (0.3 weight
percent) or increasing the Na2O or K2O by five relative percent, both within analytical
errors for these elements, would shift the SIR into the metaluminous classification.
A major-element variation diagram (Figure 30) gives a graphical representation of
the above data. Assuming a normal distribution about the mean for the two populations,
z-tests at the 99% confidence level indicate that the CDF and NDF differ in their
concentrations of some major elements. The CDF has greater Al2O3, MnO, Na2O while
the NDF has more FeO(t), and K2O with a significantly higher K2O/ Na2O ratio (1.30 vs.
1.07). The remaining major elements are not distinguishable between the two
populations at the 90% confidence interval.
Of special note are the five samples that are above 0.8 weight percent CaO.
These samples are clearly higher than the means of either the CDF or NDF analysis.
Four of the five samples tested positive for carbonate with a 10% HCl solution, likely
from caliche deposition. The one sample with high CaO from the CDF (7160108) did not
react with acid application. Eight other samples that are near the average CaO content
were chosen at random and tested with acid and none showed any reaction, indicating a
lack of caliche.
55
Table 2: Major- and Trace-Element Geochemistry: Central Dome Field; China Cap and China Hat. Major elements expressed as wt % oxide, trace elements are in ppm. UTM points NAD 27, zone 12.
Sample 7290101 8020102 7160105 8020107 7280106 8020101 7290103 7160101Dome CC CC CC CC CC CC CC CHUTM east 451581 451748 451072 451530 451478 451676 451282 450686
UTM north 4741206 4740825 4740919 4740786 4741129 4740872 4741224 4739647
SiO2 76.4 76.5 76.9 76.9 76.8 76.8 77.1 76.6TiO2 0.09 0.08 0.08 0.07 0.06 0.06 0.07 0.08Al2O3 12.6 12.6 12.6 12.5 12.5 12.5 12.5 12.7FeO(t) 0.93 0.90 0.90 0.94 0.86 0.91 0.95 0.94MnO 0.06 0.06 0.05 0.05 0.05 0.06 0.05 0.05MgO 0.08 0.04 0.03 0.02 0.03 0.01 0.03 0.05CaO 0.68 0.56 0.51 0.57 0.59 0.53 0.56 0.55Na2O 4.39 4.48 4.57 4.46 4.20 4.29 4.18 4.47K2O 4.79 4.73 4.34 4.36 4.88 4.72 4.45 4.55P2O5 0.02 0.02 0.01 0.06 0.03 0.05 0.03 0.04
Original total 96.8 97.3 98.2 99.3 97.6 96.1 97.7 97.7
Sc 2.19 2.22 2.18 2.25 2.19 2.28 1.99 2.20Rb 486.5 490.5 517.5 504 500 520 495 479.5Sr 3.59Cs 15.6 15.8 17 16.5 16.2 15.7 17 16.1Ba 152 159 161 179 139 144 161 152La 28.4 18.5 19.3 19.4 24.1 17.1 17.3 19.5Ce 70.6 62.6 69 66.8 77.3 61.6 64.5 69.4Nd 30.1 20.7 29 27.5 24.9 13.8 17.1 22.4Sm 12.4 11.4 12.1 11.8 11.9 11.1 10.8 10.7Eu 0.08 0.07 Tb 3.81 3.55 3.48 3.39 3.94 3.47 2.90Yb 15.5 15.4 16.3 15.7 14.4 14.6 14.4 14.8Lu 2.40 2.27 2.52 2.28 2.42 2.47 2.39 2.30Zr 128 108 126 129 137 131Hf 6.41 6.57 7.04 6.62 6.74 7.02 7.06 6.58Ta 9.81 9.80 10.4 10.1 10.0 10.6 10.1 9.96Th 55.1 52.0 54.3 52.7 53.5 51.8 47.3 50.6U 29.8 31.1 31.7 32.0 30.4 30.5 30.1 29.3
56
Table 2 Continued: Central Dome Field; China Hat. Major elements expressed as wt % oxide, trace elements are in ppm. UTM points NAD 27, zone 12.
Sample 7210103 7280102 7160103 7210102 7160104 7210104 7210101 G1Dome CH CH CH CH CH CH CH CHUTM
easting 450269 451205 450167 450114 450167 450412 449989 451124
UTM northing 4739033 4739494 4739907 4739195 4739822 4738943 4739329 4740456
SiO2 76.4 76.5 77.0 76.6 76.7 76.5 77.0 77.0TiO2 0.10 0.09 0.08 0.06 0.06 0.08 0.07 0.06Al2O3 12.7 12.5 12.6 12.6 12.6 12.6 12.5 12.7FeO(t) 0.99 0.93 0.91 0.90 0.88 0.92 0.88 0.83MnO 0.05 0.05 0.05 0.05 0.05 0.05 0.05 0.05MgO 0.08 0.04 0.03 0.05 0.02 0.05 0.02 0.07CaO 0.59 0.51 0.51 0.55 0.53 0.55 0.52 0.55Na2O 4.39 4.40 4.44 4.45 4.40 4.43 4.09 4.05K2O 4.73 4.78 4.34 4.74 4.69 4.85 4.87 4.66P2O5 0.02 0.12 0.00 0.03 0.03 0.02 0.03 0.03
Original total 97.2 96.6 98.4 96.9 99.4 97.2 95.2 100.7
Sc 2.22 2.15 2.25 2.23 2.24 2.35 2.23Rb 500 487 492 485 491 492 505Sr 2.22 1.58 2.28 1.97 3.27 1.84 2.26Cs 16.1 15.6 15.9 15.7 16.6 15.8 16.4Ba 163 178 171 156 162 185 172La 18.2 21.7 21.6 23.3 20.8 23.5 21.6Ce 69.8 71.7 70.3 79.2 73 77.1 72.6Nd 27.3 25.0 24.6 25.0 27.7 27.9 25.7Sm 11.3 11.6 11.7 11.9 11.6 12.0 11.7Eu 0.05 0.07 0.08 0.06Tb 3.37 3.54 3.51 3.47 3.48 3.25 3.49Yb 15.8 15.3 15.5 15.6 15.6 15.2 14.8Lu 2.40 2.26 2.24 2.37 2.35 2.27 2.36Zr 129 128 131 141Hf 6.76 6.46 6.49 6.55 6.48 6.57 6.78Ta 10.0 9.62 9.85 9.66 9.96 9.92 10.0Th 55.1 58.6 53.6 56.5 52.7 53.2 51.7U 31.3 32.1 32.0 30.8 30.5 33.8 30.6
57
Table 2 Continued: Central Dome Field; China Hat and North Cone. Major elements expressed as wt % oxide, trace elements are in ppm. UTM points NAD 27, zone 12.
Sample G2 G3 8220101 8220104 8220103 7160108 8220102Dome CH CH NC NC NC NC NCUTM
easting 451124 451124 452659 453139 452460 453304 452482
UTM northing 4740456 4740456 4742710 4742959 4742324 4742702 4742443
SiO2 76.9 76.9 77.0 76.6 76.6 76.5 76.7TiO2 0.06 0.06 0.09 0.08 0.09 0.07 0.06Al2O3 12.5 12.6 12.7 12.6 12.7 12.4 12.7FeO(t) 0.91 0.86 0.92 0.90 0.92 0.97 0.89MnO 0.05 0.05 0.05 0.05 0.05 0.05 0.05MgO 0.04 0.02 0.03 0.04 0.03 0.06 0.03CaO 0.64 0.52 0.52 0.55 0.50 1.09 0.55Na2O 4.19 4.37 4.25 4.32 4.52 4.27 4.33K2O 4.67 4.55 4.48 4.86 4.61 4.60 4.71P2O5 0.01 0.03 0.02 0.04 0.04 0.03 0.03
Original total 99.5 96.9 96.4 97.6 98.1 98.6 95.3
Sc 2.20 2.16 2.18 2.17 2.28 2.30 2.22Rb 500 495 485 476 514 471 500Sr 2.41 2.02 2.71Cs 16.2 16.1 12.0 15.3 16.5 15.6 16.1Ba 159 155 112 156 180 173 157La 18.4 19.1 19.3 22.2 25.2 24.8 18.8Ce 68.7 66.1 65.7 72.2 80.7 81.3 64.6Nd 21.3 24.1 21.0 28.8 33.1 29.9 25.9Sm 11.4 11.2 11.0 11.9 12.7 12.1 11.4Eu 0.08 0.06 0.05 0.06 0.08Tb 3.49 3.45 3.40 3.43 3.28Yb 14.8 14.5 16.4 15.1 16.2 15.1 14.8Lu 2.38 2.33 2.42 2.24 2.36 2.26 2.36Zr 134 133 122 131Hf 6.95 6.79 6.38 6.44 6.76 6.31 6.98Ta 10.0 10.1 13.0 9.62 10.2 9.18 10.1Th 51.2 49.3 50.2 53.8 56.5 63.2 50.8U 31.3 31.0 31.2 30.6 33.3 31.4 30.8
58
Table 2 Continued: Northern Dome Field; eastern most and middle dome. Major elements expressed as wt % oxide, trace elements are in ppm. UTM points NAD 27, zone 12.
Sample 6280204 6280202 6280201 6280205 6110201 9010103 9010105 9010104Dome NDE NDE NDE NDE NDM NDM NDM NDMUTM
easting 457137 456891 456891 457297 456031 455344 455382 455330
UTM northing 4764127 4764467 4764469 4764312 4764626 4764875 4763640 4764655
SiO2 77.0 76.9 77.2 77.0 77.3 76.3 76.7TiO2 0.07 0.07 0.07 0.11 0.08 0.11 0.09Al2O3 12.4 12.6 12.5 12.3 12.5 12.4 12.3FeO(t) 1.03 1.02 1.05 1.19 1.07 1.14 1.05MnO 0.04 0.04 0.04 0.04 0.03 0.04 0.04MgO 0.06 0.03 0.07 0.13 0.05 0.12 0.06CaO 0.64 0.51 0.59 0.63 0.47 0.90 0.54Na2O 3.82 3.83 3.68 3.59 3.72 3.98 4.11K2O 4.74 4.98 4.84 4.95 4.77 4.93 5.04P2O5 0.18 0.03 0.03 0.04 0.03 0.03 0.03
Original total 96.3 97.7 99.8 98.4 102.1 95.8 96.1
Sc 1.78 1.76 1.81 1.71 2.05 1.91 1.80 1.84Rb 345 355 345 375 290 345 306 348Sr 6.30 3.90 9.34 6.79 4.86Cs 9.71 9.41 9.79 10.1 8.70 8.79 8.47 9.56Ba 62.8 78.0 96.4 75.6 71.9 65.3 79.0 80.6La 29.1 32.8 29.1 27.1 28.8 33.1 32.5 31.9Ce 76.3 83.1 78.8 71.5 75.2 83.5 85.3 79.4Nd 23.9 26.4 26.3 23.6 23.5 26.1 26.4 26.4Sm 8.6 9.1 8.7 8.6 8.1 9.3 8.6 8.9Eu 0.10 0.05 0.12 0.15 0.10 0.10Tb 3.28 3.36 3.31 3.20 2.97 3.49 2.05 2.37Yb 10.4 10.0 10.6 10.2 9.0 10.1 9.1 10.3Lu 1.52 1.50 1.56 1.54 1.37 1.51 1.36 1.48Zr 121 141 115 108Hf 4.86 4.99 4.95 4.94 5.60 5.11 4.50 4.89Ta 4.35 4.51 4.43 4.37 3.83 4.47 3.68 4.26Th 52.2 51.0 52.3 49.1 46.2 51.0 49.6 52.6U 13.9 13.4 14.2 13.6 12.8 14.0 12.5 14.6
59
Table 2 Continued: Northern Dome Field; western most dome and Sheep Island. Major elements expressed as wt % oxide, trace elements are in ppm. UTM points NAD 27, zone 12.
Sample 7170201 7110201 7110202 7170203 10070001 7130201Dome NDW NDW NDW NDW NDW SIRUTM
easting 453411 453927 453036 452892 453399 449942
UTM northing 4766554 4766941 4766910 4765950 4767036 4750957
SiO2 77.3 76.9 76.1 76.3 76.7 76.6TiO2 0.07 0.07 0.06 0.06 0.06 0.04Al2O3 12.5 12.3 12.3 12.3 12.4 13.1FeO(t) 1.01 1.01 0.98 1.00 1.00 1.0MnO 0.04 0.04 0.04 0.04 0.04 0.04MgO 0.03 0.04 0.06 0.04 0.02 CaO 0.50 0.51 1.97 1.20 0.51 1.16Na2O 3.58 3.86 3.74 3.73 4.15 3.67K2O 4.90 5.27 4.69 5.34 5.11 4.34P2O5 0.02 0.03 0.03 0.03 0.04 0.03
Original total 98.0 94.6 95.0 96.0 95.3 96.3
Sc 1.75 1.75 1.67 1.72 1.75 1.77Rb 370 335 365 290 345 335Sr 4.30 3.21 7.08Cs 9.52 9.35 8.56 9.02 9.39 8.54Ba 115 77.5 69.7 88.3 96.5 87.0La 28.6 30.6 29.3 28.6 32.4 27.5Ce 74.0 77.9 75.1 73.1 84.4 72.7Nd 23.9 25.6 19.9 23.8 28.4 23.9Sm 8.6 9.1 8.4 8.5 8.9 8.5Eu 0.10 0.12 0.09 0.10 0.08 0.07Tb 3.31 3.20 2.92 2.90 2.35 2.29Yb 9.91 9.97 9.86 10.2 9.00 10.4Lu 1.51 1.56 1.44 1.45 1.50 1.52Zr 122 115 113 Hf 5.13 5.08 4.75 5.10 4.75 4.85Ta 4.37 4.33 4.22 4.16 4.39 4.35Th 50.6 49.5 48.2 49.0 52.8 52.2U 13.7 14.4 13.4 13.3 14.3 13.9
60
Table 2 Concluded: Big Southern Butte Samples. Major elements expressed as wt % oxide, trace elements are in ppm. UTM points NAD 27, zone 12.
Sample 11080301 11080301 11080302 11080303 11080304Dome BSB BSB BSB BSB BSBUTM
easting 335862 335862 Top 334788 334640
UTM northing 4806574 4806574 4807258 4807616
SiO2 76.01 75.71 76.18 76.55 76.15TiO2 0.09 0.09 0.09 0.09 0.09Al2O3 12.61 12.72 12.50 12.32 12.40FeO(t) 1.60 1.63 1.62 1.64 1.60MnO 0.04 0.04 0.03 0.04 0.04MgO 0.04 0.04 0.01 0.00 0.01CaO 0.39 0.38 0.45 0.23 0.19Na2O 4.39 4.58 4.52 4.57 4.78K2O 4.82 4.80 4.59 4.54 4.72P2O5 0.03 0.02 0.01 0.02 0.01
Original total 98.6 96.5 97.9 98.1 97.0
Sc 0.14 0.14 0.18 0.19Cr 19.4 19.2 17.2 20.5Rb 301 294 302 291Sr 3.47 2.95 0.06 0.81 0.96Cs 2.78 2.41 2.71 2.14Ba 64.7 55.4 53.0 52.9La 58.1 64.6 50.7 77.1Ce 128 126 113 149Nd 38.6 45.8 41.1 57.0Sm 12.4 14.0 12.5 17.3Eu 0.25 0.25 0.27 0.35Tb 3.1 3.46 3.14 4.43Yb 17.6 18.9 16.3 20.8Lu 2.35 2.52 2.08 2.60Zr 297 306 316 304 308Hf 15.8 16.2 15.5 16.6Ta 18.6 18.8 19.0 18.6Th 20.1 19.1 19.1 20.0U 13.0 13.8 13.1 13.4
Table 2 Notes:
Major elements reported are normalized to 100% anhydrous. P2O5, where italicized, was estimated at 0.03% for normaliza-tion purposes only. All major elements determined using ICP-AES and trace elements determined with INAA except FeO and Na2O measured with both techniques with disputed measurements redone and Sr, Ba, and Zr determined with ICP-AES. Abbreviations as follows: Central Dome Field: CH = China Hat CC = China Cap NC = North Dome Northern Dome Field: NDE = eastern most dome NDM = middle dome NDW = western most dome Other rhyolites SIR = Sheep Island BSB = Big Southern Butte No data indicates the analyte was not measured for that sample or was below detection limits (MgO and Eu only).
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The contribution of calcium from caliche was not calculated. It is likely that most
of the “extra” calcium above 0.6 weight percent can be attributed to caliche formation in
the NDF rhyolites. These five samples also have lower SiO2 concentrations as a result of
both increased CaO concentrations and the loss of CO2 in the dissolution process. The
high CaO in the one sample from the CDF remains unexplained. The data from the ICP
analysis is robust and a non-quantitative examination of the hand sample and thin section
give no indication of significant differences from other samples analyzed (i.e. increased
plagioclase percentage or the existence of other calcium-bearing phenocrysts). The SIR
also has a high CaO percentage and contains at least some caliche but it is unclear how
much calcium came from the caliche. There are no apparent systematic variations of
other common caliche constituents such as Mg and Sr in the five samples with high CaO
concentrations.
3.3.2 Comparison of Major-Element Geochemistry of the BVF to Other Rhyolites:
The CDF, NDF and SIR are similar but differentiable in their major-element
constituents (Figure 30 – note scales). Their differences are indistinguishable when
plotted on diagrams that contain a wide span of rhyolite compositions and thus are
generically referred to as Blackfoot Volcanic Field (BVF) rhyolites for comparison to
rhyolites from other tectonic settings. As the major-element variation diagrams in Figure
31 show, the rhyolites of the BVF lie within the boundary of or near the edge of the area
defined by Yellowstone and Tertiary Snake River Plain (YandTSRP) rhyolites, except
that they are lower in FeO(t). Likewise, data from the Basin and Range province (which
include Long Valley, Death Valley and the Latir Volcanic Field) often overlap with the
63
BVF data. This relationship, where the major-element compositions of the BVF rhyolites
are similar in nature to both YandTSRP and Basin and Range rhyolites, mimics the
results seen in the BVF basalts (Pickett, 2004).
It is important to again clarify the distinction between the YandTSRP and
Quaternary Eastern Snake River Plain (QESRP) rhyolites. While the two groups of
rhyolites occupy the same physiographic province, they are distinct in many of their filed,
compositional and mineralogical characteristics. The YandTSRP (“hot-spot track”)
rhyolites are produced by massive, caldera forming eruptions, generally emplaced by
ignimbrites and contain significant Proterozoic or Archean crustal component, up to 35
percent (McCurry et al., 2002) while the QESRP rhyolites are emplaced as low volume
domes (and cryptodomes in the case of East Butte, Fairy Butte, Buckskin Dome, etc.) and
are formed by extreme fractionation of a mafic parent with very little crustal component,
as little as one percent (McCurry et al., 1999). Several papers (e.g. Spear, 1979; Hildreth
et al., 1991; Leeman, 1982a; Leeman, 1982b; Pierce and Morgan, 1992; McCurry et al.,
1999; McCurry et al., 2002; McCurry and Ganske, 2005) cover the formation of these
two highly distinguishable rhyolite suites.
From Figure 31 note that both QESRP and topaz rhyolites have similar major-
element compositions to the BVF rhyolites. On some variation diagrams (e.g. FeO(t) and
TiO2), two groups of QESRP data points can be delineated. These represent data points
for the two largest domes exposed on the surface, East Butte (typically higher iron) and
Big Southern Butte, with points from other QESRP (e.g. Unnamed Butte, INL Core Hole
1 and Cedar Butte) scattered and more variable. In comparison to the BFV rhyolites, the
QESRP rhyolites are somewhat higher in FeO(t) and correspondingly lower in SiO2 and
64
Table 3. References for data represented for major-element variation diagrams (Figures 31) and trace-element variation diagrams (Figure 33). Quaternary Eastern Bretches, 1984 Snake River Plain Ganske, in preparation Rhyolites Hayden, 1992 (QESRP) Leeman, 1982a Morse, 2002 Spear, 1979 This study Tertiary Snake River * Binderman and Valley, 2001 Plain and Yellow- Eckren et al., 1982 stone rhyolites Hildreth et al., 1991 (YandTSRP) Honjo, 1990 * Kellogg and Marvin, 1988 Morgan et al., 1984 Parker, 1996 Watkins, 1998 Wright, 1998 Basin and Range * Asmerom, 1994 Rhyolites * Binderman and Valley, 2002 * Heumann and Davis, 1997 * Johnson and Lipman, 1988 Cascade rhyolites * Bacon, 1985 * Brophy et al., 1996 * Drophy and Dreher, 2000 * Bullen and Clynne, 1990 * Condie and Hayslip, 1975 * Gerlach and Grove, 1982 * Grove and Donnelly-Nolan, 1986 Topaz rhyolites Christiansen et al., 1986
* From the PLUTO Geochemical database (Baedecker et al., 1998)
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67
East Butte is higher in TiO2. The topaz rhyolite points come from eight temporally and
spatially separated sets of domes resulting in more scatter in the data than seen with the
QESRP of BVF data sets. References for Figure 31 are given in Table 3.
3.3.3 Trace-Element Geochemistry:
While the CDF and NDF rhyolites are similar, but distinguishable, in their major-
element compositions, more substantial differences occur in some of their trace-element
characteristics. These differences are distinguished on a primordial mantle normalized
multi-element spider diagram (Figure 32). The average CDF rhyolite contains
significantly (at least 25%) more Cs, Rb, Ba, U, Ta, Hf and Sm and they are moderately
higher (10 to 25% more) in Zr and Tb than the NDF. The NDF rhyolite only has
significantly higher La and Sr with moderately higher Ce. Th and Nd are similar (less
than a 10% difference). All of the major elements have less than a ten percent difference.
The multi-element spider diagram does not show all of the elements from Table 2 and the
remainder of the trace elements are represented on a chondrite-normalized REE diagram
(Figure 34). Yb and Lu are significantly higher in the CDF while Sc is moderately higher
while Eu appears to be higher in the NDF (Table 4).
Four elements warrant special consideration. P2O5, as stated in the major-element
chemistry section, is estimated at 0.03 for a number of compositions for normalization
purposes. Also, there were a few less samples analyzed for both Zr and Sr calculations
(as indicated in Table 2). Finally, the Eu concentration in both dome fields was near
detection limits and thus there is also less data for this analyte. None of the above
68
conditions altered the overall trends of the data represented in the spider diagram (Figure
32). Data comparing the two BVF dome sets is summarized in Table 4.
Table 4. Summary of the comparison of trace element abundance between the average CDF and NDF of the Blackfoot Volcanic Field. Specific sample values can be found in Table 2 and averaged values are listed in Table 5. Graphical representations are provided, along with additional physiographic province data in Figures 33 and 34. The CDF rhyolite contains significantly to moderately higher concentrations (greater than ten percent more) of nearly all the trace elements. >25% more
in CDF 10-25% more in CDF
>25% more in NDF
10-25% more in NDF
Nearly the same in both
Sc X
Rb X
Sr X
Cs X
Ba X
La X
Ce X
Nd X
Sm X
Eu X
Tb X
Yb X
Lu X
Zr X
Hf X
Ta X
Th X
U X
69
Figure 32. Primordial mantle-normalized multi-element spider diagram of all samples from the CDF (23 samples – black lines) and NDF (13 samples – grey lines). This figure illustrates some of the obvious differences in the trace-element composition of the two dome sets. Normalization values are from Taylor and McLennan (1985), except for P from Sun (1980), and are consistent with those values used by Pickett (2004) for the basalts of the BVF.
0.1
1.0
10.0
100.0
1000.0
10000.0
Cs Rb Ba Th U K Ta La Ce Sr Nd P Hf Zr Sm Ti Tb
70
3.3.4 Comparison of Trace-Element Geochemistry of the BVF to Other Rhyolites:
While the BVF rhyolites are similar to a number of other rhyolites from various
tectonic and physiographic settings (e.g. YandTSRP, Basin and Range, QESRP and topaz
rhyolites) in their major elements, the concentrations of many of the trace elements lie
outside the ranges of most other groups of rhyolites, except for topaz rhyolites. These
include U, Th, Cs and Rb which are present in concentrations that are significantly higher
than rhyolites from other physiographic provinces examined in this study (Table 3). The
BVF and QESRP rhyolites both have higher concentrations, as compared to other
rhyolites, in Tb, Yb, Lu, and Ta. Additionally, both the BVF and QESRP rhyolites are
lower than other rhyolites in Sr, Eu and Ba. Eu and Sr concentrations in the BVF,
especially the CDF, merit special mention as both show extreme depletion with only Big
Southern Butte of the QERSP having less Sr. Table 5 summarizes the average values for
the various tectonic and physiographic provinces for trace elements. Some trace-element
variation diagrams are shown in Figure 33 and help to visualize the data in Table 5.
Both Dayvault et al. (1984) and Christiansen et al. (1986) have classified the CDF
as a topaz rhyolite based upon both the chemical makeup of the rocks and on the presence
of post-emplacement topaz formation. Both dome sets in the BVF, despite their trace
element differences (Table 4), are more similar to topaz rhyolites than to rhyolites from
the afore mentioned suites (Table 5). Further similarities between the topaz rhyolites of
the western United States and the BVF rhyolites will be presented with chondrite
normalized REE data next and in the Discussion section.
71
Table 5. Average concentrations, in ppm, of trace elements from various physiographic regions with the BVF delineated into the CDF and NDF, topaz = topaz rhyolites of the Western United States, B and R = Basin and Range rhyolites, and the QESRP delineated into EB = East Butte and BSB = Big Southern Butte. Sources are the same as in Table 3 except East Butte is Ganske (2005, in preparation) and Big Southern Butte is this study and Spear (1979). Sr for Big Southern Butte, marked with the asterisk, shows an average of eights sample. Dr. Michael McCurry has further refined the value to be ~0.6 ppm (personal communication, 2001). Ba values for topaz rhyolites are widely variable and REE element data was only represented only in chondrite normalized graphs by Christiansen et al. (1986).
CDF NDF topaz B and R Cascades YandTSRP EB BSB
Sc 2.2 1.8 2.2 2.6 5.2 5.6 0.65 0.16
Rb 495 335 600 135 100 175 210 295
Sr 2.3 5.7 12.6 100 225 98 11 1.8*
Cs 16 9.2 30 4.7 5.7 3.3 1.7 2.5
Ba 160 82 625 815 970 115 64
La 21 30 37 22 80 110 63
Ce 71 78 66 42 150 220 140
Nd 25 25 20 19 64 89 51
Sm 12 8.7 3.2 3.4 12 23 15
Eu 0.06 0.10 0.52 0.68 1.61 1.3 .26
Tb 3.5 2.9 0.49 0.55 1.8 4.4 3.5
Yb 15 9.9 1.9 1.9 6.1 18.3 20
Lu 2.3 1.5 0.27 0.35 0.85 2.1 2.4
Zr 130 120 120 170 175 460 380 185
Hf 6.7 5.0 6.8 4.2 4.7 14 16 16
Ta 10 4.3 14 1.7 0.83 3.3 12 19
Th 53 50 52 16 10 27 23 20
U 31 14 18 5.1 3.2 6.6 7.1 13
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75
The chondrite-normalized (using Taylor and McLennan, 1985 values) rare earth
elements (REE) patterns of the BVF are unique, along with the topaz rhyolites as
compared to other rhyolites (Figure 34). Both the CDF and NDF show relative depletion
of LREE and enrichment of HREE with the CDF having a nearly 1:1 ratio of LaN to YbN
and the NDF a 2:1 ratio with those for topaz rhyolites commonly falling between 1:1 and
3:1 (Christiansen et al., 1986). Rhyolites of the BVF have an exceptionally deep Eu
anomaly with Eu/Eu* = 0.014 for the CDF and 0.026 for the NDF, again, within the
range of topaz rhyolites. The REE trend for Big Southern Butte (QESRP) is nearly
identical to the NDF with normalized values approximately twice those for the NDF. In
contrast, the YandTSRP rhyolites have a La/YbN ration of 10 and a Eu/Eu* approaching
0.5 with an overall REE pattern that closely approximates those of A-type granites
(Collins, 1982). The Cascades are generally depleted in all REE, with only a slight Eu
anomaly. Basalts from the Blackfoot Field are similar to Snake River Plain basalts
(Pickett, 2004) and have a Eu/Eu* slightly above 1 on average (Figure 34). Sources for
the aforementioned data are given in Table 4 except as listed in Figure 34 and Big
Southern Butte (this study and Spear, 1979).
As presented earlier, the BVF rhyolites have many similarities to the QESRP
rhyolites in their trace-element geochemistry. Some similarities are also present in the
REE data. While normalized LREE values vary, all are relatively enriched in HREE and
all show strong Eu anomalies. Christiansen (1986) indicates that La/LuN and Eu/Eu*
values are inversely proportional to uranium content in topaz rhyolites. This holds true
for both the BVF and QESRP rhyolites, including Big Southern Butte, East Butte and
Unnamed Butte.
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77
3.4 Electron Microprobe Results on Phenocrysts of the Central Dome Field:
Thirty-five EMP analyses on sample G1 of the CDF were completed at the IRC in
Idaho Falls, Idaho. These included core and rim analysis on a plagioclase, three sanidine,
two biotite, two amphibole (“hornblende”), two magnetite and two ilmenite crystals as
well as four melt inclusions and two matrix glass areas. Complete anaylses are given in
Table 6 and the specific EMP methods are given in Appendix 1.
Feldspars: Only the plagioclase crystal showed evidence of compositional zoning. The
plagioclase crystal is anothite14 with an orthoclase component of seven percent (on
average) with a slightly more sodic and less calcic rim (normal zonation). On average,
sanidine phenocrysts are orthoclase65 with anorthite contents less than one percent
(Figure 35).
Fe-Ti oxides: Magnetite phenocrysts have an average of ulvospinel26 and contain just
over one weight percent MnO and Ilmenite averages hematite10 with 2.5 weight percent
MnO (Figure 35).
Amphibole: The average amphibole is classified as a ferro-edenite.
Mica: The mica is a biotite with approximately annite60 and phlogopite40 with molar
aluminum content slightly higher than these two end members. The extra molar
aluminum may exist in the octahedral sight resulting in up to 15 percent siderophyllite
component.
Melt inclusions: The four melt inclusions are somewhat give variable results in the
original analyses but consistent when normalized to 100 percent anhydrous. Three of the
analyzed melt inclusions are located in quartz crystals while one is in a magnetite crystal
(sample MI 5). The analysis of the melt inclusion in magnetite appears to over-represent
78
concentrations of FeO and TiO2 indicating potential contamination from the host crystal.
This is also potential host crystal contamination in sample MI 1, hosted in quartz.
Sample MI 3 has a low total of oxides attributed to one of two phenomena; either the
inclusion are tight (or sealed) and the majority of the missing mass, approximately six
percent, is water or the microprobe beam spot also incorporated some epoxy used in the
slide preparation or some other foreign substance during the analysis. Analysis of sample
MI 4 clearly clipped the edge of a bubble, possibly filled with epoxy, and had a total of
only 83.4 percent. The average melt inclusion analysis is similar to the bulk rock
analyses except that Na2O is low (Figure 30).
Matrix glass: Two matrix glass measurements are similar to each other and give original
totals near 100 percent. The matrix glass analyses are similar to the bulk rock chemistry
for nearly all major elements, except Na2O. Complete analyses are given in Table 6 and
melt inclusion and matrix glass average values are plotted with bulk chemistry on Figure
30.
Figure 35. Ternary diagrams for the microprobe analyses of sanidine, plagioclase, magnetite and ilmenite. The Fe-Ti oxides are connected by a dotted tie line.
79
Table 6. Normalized, to 100% anhydrous unless otherwise indicated, individual phenocryst chemistries from EMP for both rims and cores (shown with original totals listed). Abbreviations as follows Sa = sanidine, MI = melt inclusions, Mt = magnetite, Ilm = ilmenite. MI 5 is in a magnetite crystal while other melt inclusions are located in quartz. Average values for MI and MG are plotted with bulk chemistry on Figure 30. Mineral standards and integration times are given in Appendix 1.
Sample Sa 1 Sa 1 Sa 2 Sa 2 Sa 3 MI 1 MI 3 MI 4 MI 5 core/rim core rim rim core na na na na
SiO2 65.80 66.13 66.01 66.04 66.13 80.57 78.09 77.96 78.95 Al2O3 18.43 18.57 18.69 18.57 18.51 10.81 13.09 12.44 12.28 TiO2 0.02 0.01 0.03 0.08 FeO 0.08 0.14 0.12 0.08 0.12 0.75 0.82 0.99 1.91
Cr2O3 MgO 0.00 0.00 0.03 0.00 MnO 0.06 0.07 0.01 0.00 NiO ZnO CaO 0.15 0.17 0.20 0.19 0.15 0.44 0.47 0.49 0.47 Na2O 4.26 3.88 3.95 4.09 3.99 3.14 2.65 3.15 1.91 K2O 11.28 11.11 11.03 11.03 11.10 3.98 4.46 4.64 4.11 SrO P2O5 0.01 0.00 0.00 0.00
F 0.16 0.24 0.13 0.24 Cl 0.16 0.25 0.24 0.18
SO3 0.00 0.00 0.01 0.01 orig. total 99.35 100.60 100.80 100.64 100.58 98.96 93.28 83.43 99.19
Sample Mt 1 Mt 1 Mt 2 Mt 2 Ilm 1 Ilm 2 Ilm 2 Ilm 2 core/rim rim core rim core na na na na
SiO2 0.04 0.09 0.07 0.08 1.02 0.02 0.01 0.00 Al2O3 0.74 0.76 0.80 0.76 0.26 0.03 0.05 0.03 TiO2 9.50 9.88 9.95 10.61 47.35 48.07 48.35 47.67 FeO 88.16 87.77 87.67 87.03 48.51 49.02 48.78 49.28
Cr2O3 0.00 0.01 0.03 0.00 0.01 0.01 0.03 0.00 MgO 0.13 0.11 0.14 0.16 0.32 0.31 0.30 0.33 MnO 1.14 1.22 1.13 1.21 2.52 2.45 2.37 2.60 NiO 0.00 0.02 0.00 0.02 ZnO 0.29 0.17 0.21 0.14 0.00 0.07 0.09 0.06 CaO 0.00 0.00 0.00 0.00 Na2O K2O SrO P2O5
F Cl
SO3 orig. total 97.25 97.07 96.69 97.61 96.70 98.69 99.49 97.79
80
Table 6 Concluded. Abbreviations as follows: Plag = plagioclase, MG = matrix glass, Bt = biotite (not normalized), Hlb = hornblende (not normalized). Analytes for which there is no data were not tested for in respective phases (e.g. TiO2 in plagioclase).
Sample Plag 1 Plag 1 Plag 1 Plag 1 Plag 1 Plag 1 MG 1 MG 3 core/rim rim core rim rim rim core na na
SiO2 65.61 65.35 64.80 64.83 65.73 64.69 77.75 77.68 Al2O3 21.02 21.53 21.77 21.62 21.08 21.93 12.32 12.42 TiO2 0.02 0.01 FeO 0.15 0.12 0.16 0.13 0.11 0.17 0.84 0.85
Cr2O3 MgO 0.00 0.00 0.00 0.00 0.00 0.01 0.01 0.01 MnO 0.04 0.00 NiO ZnO CaO 2.42 2.77 3.07 3.01 2.38 3.06 0.53 0.50 Na2O 9.41 9.02 9.13 9.23 9.34 9.18 3.69 3.64 K2O 1.38 1.21 1.08 1.18 1.36 0.97 4.51 4.59 SrO P2O5 0.01 0.00
F 0.23 0.28 Cl 0.17 0.16
SO3 0.01 0.01 orig. total 101.14 100.45 100.89 100.74 100.92 100.49 100.86 101.30
Sample Bt 1 Bt 1 Bt 1 Bt2 Bt2 Hbl 2 Hbl 2 Hbl 3 Hbl 3 core/rim core rim rim rim core core core core rim
SiO2 36.43 36.45 36.21 35.95 36.28 41.64 40.82 42.12 42.12 Al2O3 12.26 12.05 12.04 12.23 12.14 7.81 8.29 7.41 8.04 TiO2 3.49 3.64 3.59 3.52 3.75 1.37 1.57 1.13 1.30 FeO 28.97 29.36 28.57 28.70 29.09 30.07 30.19 27.53 27.71
Cr2O3 MgO 6.47 5.54 6.29 6.82 6.66 3.37 3.64 5.17 4.99 MnO 0.65 0.58 0.58 0.60 0.60 1.15 1.13 1.03 1.01 NiO ZnO CaO 0.04 0.00 0.03 0.01 0.00 10.23 10.37 10.05 10.26 Na2O 0.39 0.45 0.42 0.37 0.41 2.06 2.04 2.02 2.11 K2O 8.78 8.79 8.81 9.18 8.61 1.15 1.27 1.07 1.17 SrO P2O5
F 0.17 0.17 0.19 0.26 0.26 Cl 0.27 0.29 0.27 0.30 0.31 0.33 0.37 0.31 0.30
SO3 orig. total 97.78 97.18 96.86 97.74 97.92 99.10 99.59 97.77 98.95
81
3.5 Isotope Results:
Nd and Sr isotopes on four samples from the CDF and one sample from the NDF
were analyzed by Dr. John Chadwick at the University of Florida Department of
Geosciences and are presented in Table 7 and plotted on Figure 39 (Chapter 4). Also
included in the table are basalts from the BVF (Pickett, 2004) and a rhyolite sample from
China Cap of the CDF (Michael McCurry, personal communication, 2001). The BVF
rhyolites are slightly more evolved (higher 87Sr/86Sr and lower σNd) than the high-silica
rhyolites of the QESRP (McCurry et al., 1999) but are closer to the QESRP rhyolites than
to Basin and Range rhyolites (Gans et al., 1989, not plotted). The isotope ratios of the
basalts of the BVF are more similar to QESRP basalts than to Basin and Range basalts
(Pickett, 2004).
Table 7. Nd and Sr isotopic ratios for rhyolites and basalts from the BVF. The rhyolite data are initial concentrations corrected using the given Sr and Nd concentrations (Table 2) except for the China Cap MM sample for which average Sr and Nd concentrations were used. Time corrections used were 50 ka for the CDF and 1.41 Ma for the NDF. The basalt values listed are the measured values. Plus or minus values are X 10-6. China Cap MM from M. McCurry (personal communication, 2001) and KEP basalts from Pickett (2004).
Sample Rhyolites Sr
(ppm) 87Sr/86Sr
initial +/- Nd
(ppm) 143Nd/144Nd
initial +/- Sigma
Nd 7160101 CDF 2.2 0.710591 6 22.4 0.512080 6 -10.9 8220101 CDF 2.02 0.709523 25 21 0.512083 7 -10.8 G1 CDF 2.23 0.7097306 7 25.7 0.512130 40 -9.9 G2 CDF 2.2 0.709595 8 21.3 0.512074 5 -11 China Cap MM CDF 0.710341 10 0.512075 5 -11 9010104 NDF 4.86 0.709354 12 26.4 0.512023 4 -12 Basalts (measured) (measured) KEP 004 "primitive" 350 0.706646 5 26.8 0.512358 30 -5.5 KEP 027 "evolved" 350 0.708905 4 81 0.512241 17 -7.7 KEP 029 Willow Cr 375 0.707929 5 16.8 0.512374 26 -5.1
82
Chapter 4 DISCUSSION
This chapter focuses on the CDF and starts with a discussion of the physical
volcanology and the emplacement history for the igneous rocks of the BVF followed by
petrography interpretation. I then examine the state of the pre-eruptive system including
the phenocryst assemblages and intensive variables (e.g. pressure, temperature and fO2).
Current models for creating topaz rhyolites with fractional melting and the QESRP
rhyolites with assimilation and fractional crystallization are then evaluated as analogs to
the BVF system. The patterns of rock chemistry in the BVF are evaluated to link the two
members of the bimodal field with fractional crystallization. Because there are no
intermediate rocks in the BVF, MELTS modeling (Ghiorso and Sack, 1994) is attempted
to help establish the paragenesis for these intermediate melt compositions during the
preferred fractional crystallization evolution model for the BVF system. These results
could better constrain the timing of phase initiation within the evolving magma to help
determine the bulk distribution coefficients. Assimilation is first modeled with via single
stage bulk mixing and then with EC-AFC models (Spera and Bohrson, 2001; Bohrson
and Spera, 2001). These models are designed as plausibility studies, can we produce the
CDF rhyolite from basalts in the BVF with reasonable constrains? Finally, I give a brief
discussion on the NDF and on the regional implications of my results.
4.1 Physical Volcanology:
Within the vicinity of the domes in the CDF, the basalts were emplaced and cut
by faults before the domes formed. While small normal faults in the valley center are
83
likely present but hidden by still younger basalt flows, many of the surface basalts of the
Blackfoot Lava Field are cut by numerous north to northwest trending normal faults
(Oriel and Platt, 1980) (Figure 2). Some of these fault traces intersect the rhyolite domes
(Figures 10 and 11) but there is no evidence of offset in either obsidian bands or tephra
units on the domes. Also, there is a small but noticeable lobe on the southwestern side of
China Hat where the rhyolite flowed over the edge of a pre-existing fault scarp (Plate 1).
Additional evidence of rhyolite production after localized basalt flows is the
covering of in-situ tephra with loess on the north side of China Hat at an elevation of
approximately 1875 m (6150 feet) (Figure 13), and not basalt while there is a basaltic
vent less than 0.5 km to the north from the tephra deposit with an elevation exceeding
1896 m (6220 feet) (Plate 1). It is possible that this local basaltic vent predates the dome
formation while other vents in the lava field postdate dome formation. No evidence of
basalts on-lapping the domes is apparent, thus any basaltic eruptions occurring after
dome emplacement must be localized and away from the domes.
Dates of basaltic vents from Luedke and Smith’s (1983) map range from 0.01 to
1.0 Ma while the most recent sanidine Ar-Ar age for the CDF is 0.057 ± .008 Ma
(Heumann, 2004). Valley wide, it is clear from the dates above that basalt flows, faulting
and rhyolite dome formation are coeval. Field work in this study favors rhyolite
formation after localized basalt emplacement and faulting.
A small mafic dike was described by Mansfield, 1927 on the north side of China
Cap that may have been emplaced after the rhyolites. I interpret this feature as a slab of
formerly flat-lying surface basalt that was pushed up or rafted and broken by the rhyolite.
Glassy rhyolite coats some of the large, greater than one meter pieces of basalt in a layer
84
that is brittle easily removed with a hammer. A piece of basalt over 100 cm found in this
vicinity likely weathered out of this feature. Xenoliths of basalt in rhyolite are common
near this dike-like feature and resemble a similar area that is rich in xenoliths on the
southwest side of China Cap just above the gravel pit.
The rhyolite eruption began with tephra production and had a hydrovolcanic
component as evidenced by the palagonitized basalts (Figure 13) in the exposed tephra
cone section on China Hat. The tephra beds are largely made up of both pyroclastic
surge and fallout deposits with some larger bocks of rhyolite, up to 0.5 m, indicative of a
crumble breccia (Figures 12 and 14, Plate 1). The lower sections of the tephra deposit
contain more basalt fragments, from pre-existing surface flows, than do the upper
portions of the deposit and likely represent a vent cleaning and widening phase early in
the eruption. Only one tephra deposit is exposed, on the north side of China Hat while
other tephra deposits, if produced in the formation of the other two domes of the CDF,
are covered by the dome or are present but covered by loess. Another unlikely
alternative, due to the good preservation of the deposit around China Hat, is that they
have been eroded away and no evidence of them is present.
Dome growth began before tephra production had stopped as there are pieces of
crumble breccia within the mid- to upper-portions of the tephra deposit (Figure 14).
Dome growth continued, mainly by endogenous processes and resulted in the formation
of many spines (Figure 17) and breccia (block) fields (Plate 1) created by spine collapse
(Macdonald, 1972). Field evidence indicates only a single, nearly continuous
emplacement of the domes, not multiple eruptions separated by considerable time as
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86
suggested by Mansfield (1927). Geochemical evidence also backs this theory as there is
no discernable chemical variation within the domes or within the whole of the CDF.
Partial dome collapse resulted in the region of hummocky terrain or run out feature on the
northeastern side of the North Cone (Figure 16, Plate 1). A four-part schematic dome
emplacement timetable is given as Figure 36.
Two hypotheses may be invoked to explain the depressions, now partially filled
with Burchett Lake (Figure 18) and Gronewall Lake, between China Cap and North Cone
(Plate 1). The Burchett Lake depression is oblong in the same direction as the trend of
the three domes while the Gronewall Lake depression is reniform in shape. The area
adjacent to these two depressions is 60 m (200 feet) higher than the surrounding valley
floor and slopes down to the valley floor at ~5o along a northwesterly trend,
perpendicular to the dome alignment trend. One possible way to produce these structures
is as explosion craters, an idea promoted by Mansfield (1927) or maars. Evidence to
support this includes local hydrovolcanics (i.e. in the China Hat tephra pit). Evidence
against this interpretation is the lack of palagonitized basalts within the craters or along
the rims, distal, columnar jointed basalts within the craters and the exposed basalt dips
away from the crater center.
Another hypothesis is that the pits are produced by one of two similar phenomena,
either by localized extension above a dike (Scenario 1, Figure 37) or by collapse of the
surface after magma is withdrawn from a dike that fed the three domes (Scenario 2,
Figure 37). In either the above situations or the Mansfield (1927) hypothesis, the local
tumescence between China Cap and the North Cone is caused by the dike that fed the
rhyolitic eruptions. The formation of the depressions is contemporaneous with the dome
87
Figure 37: Hypothetical, diagrammatic cross-section between China Cap and the North Cone of the CDF perpendicular to the N33oE trace of the domes showing dike injection causing tumescence and two related possibilities for the formation of depressions (i.e. Burchett Lake and Gronewall Lake). Not to scale.
formations as the North Cone likely filled a more circular hole that now contains the
moat-like Gronewall Lake (Plate 1).
4.2 Petrography Interpretation:
All phenocrysts, with the exception of quartz and those crystals in or near
basaltic magmatic enclaves, appear to be in textural equilibrium with the surrounding
melt. Some quartz crystals are embayed (Figure 19) while others are euhedral (Figure
21). Since quartz generally has a uniform chemical composition, these different textures
likely represent one set of crystals that has been destabilized, possibly due to
depressurization of the melt (Hibbard, 1995), and one set that formed after this
depressurization incident.
88
Within the basaltic enclaves, the boxy cellular or skeletal plagioclase crystals
rimmed with euhedral plagioclase growth (Figure 26) indicate initial rapid under cooling
due to the transfer of heat to the cooler rhyolitic magma followed by growth of the non-
cellular rim zone as the enclave approached equilibrium (Hibbard, 1995). Swallowtail
forms also indicate rapid crystal growth (due to cooling). Both forms are consistent with
the introduction of hot basaltic magma into a rhyolitic system (Hibbard, 1995).
Oscillatory zoned crystals, perhaps due to localized diffusion-supersaturation interplay
(Vance, 1962 referenced in Hibbard, 1995), also indicate localized cooling with a change
in calcium concentration.
Some of the basaltic enclaves show localized evidence of hybridization with
rhyolitic magma (Figure 27). Quartz within this zone is pyroxene-mantled with the space
between the pyroxene and quartz being caused by minor resorption or dissolution of the
quartz (Figure 28). Some olivine crystals within the enclaves also show evidence of
resorption. These features are compatible with magma mixing (Hibbard, 1995). Within
the rhyolite, near the basaltic magmatic enclaves, there is evidence of some partial
melting of sanidine crystals as they exhibit spongy cellular or fingerprint textures (Figure
29). Heat from the basaltic magma is likely responsible as this texture was not evident in
slides that did not contain a basaltic magmatic enclave.
Alignment of the microlitic swallowtail plagioclase crystals coupled with long,
flattened or stretched vesicles are indicative of shear during and shortly after
emplacement (M. McCurry, personal communication, 2005). The silicic glomerocrysts
contain the same phases as the rhyolite and are in textural equilibrium as they do not
exhibit resorption textures. The Fe-Ti oxides within these glomerocysts can be
89
significantly larger than those seen in the rhyolite (Figure 22). These factors indicate that
the glomerocrysts are magma chamber wall rock pieces that were quarried just prior to
eruption.
Section 4.3 The State of the Pre-Eruptive System:
Four analyses of hornblende crystals (ferro-edenites) by electron microprobe
(Table 6) were input into the program AMPHCAL (Yavuz, 1996) to obtain the
geobarometry of the crystals. Both the Johnson and Rutherford (1989) and the
Hammarstrom and Zen (1991) geobarometers, contained with the AMPHCAL program,
gave similar results for the four analyses with a total range of 3.47 ± 0.41 kbars and an
average of 3.5 kbars. This puts the depth of formation for the hornblende crystals, and
thus the magma source of the rhyolite, at 12.7 to 13.4 km using average crustal densities
of 2.8 g/cc and 2.67 g/cc respectively. This places any restite formed by crystallization at
a depth that is equivalent to the top of the mid-crustal sill located under the ESRP (Peng
and Humphreys, 1998; Figure 3).
Petrographic interpretation of the CDF shows that phenocrysts in the rhyolite are
in textural equilibrium, except where localized conditions are altered by the incorporation
of a basaltic magmatic enclave. As a partial check to this interpretation, the Mg/Mn
partitioning into the Fe-Ti oxides, concentrations of which are obtained from EMP data
(Table 6), are plotted on Bacon and Hirschmann’s (1988) graph of empirical data on
fresh volcanics rocks. The results show that the ilmenite and magnetite are in
equilibrium (Figure 38). This allows for the use of QUILF (Andersen et al., 1993), a
program to determine the temperature, fO2 and ∆FMQ (or DFMQ – the log fO2 relative to
90
Figure 38. Log / log plot of atomic Mg/Mn concentrations in ilmenite and magnetite crystals that appear to be in textural equilibrium (small dots) and the CDF, compositions from microprobe analysis (square, covering the range of analyses). While this does not conclusively prove the two phases are is equilibrium, many rocks for which the two phases are not in equilibrium plot outside the 2-sigma lines. Adapted from Bacon and Hirschmann (1988).
91
the FMQ buffer at a given temperature and pressure) of the melt. QUILF results, using a
pressure of 3.5 kbars, show that the temperature of the melt is 758oC ± 10oC and is
insensitive to reasonable changes in pressure. The hematite component of ilmenite was
varied by plus or minus ten percent and the temperature was varied by ± 20oC to check
the robustness of the oxygen fugacity results. The range for log fO2 is -14.5 ± 0.6 and
0.68 ± 0.14 for the DFMQ (see Table 8) for the variations stated above.
Table 8. Magma chamber properties for the CDF. Uncertainty values for pressure are the total range given by a number of different geobarometers Uncertainties for depth are the maximum and minimum range using 2.8 g/cc and 2.67 g/cc crustal densities and the average pressure. Uncertainties in temperature are calculated by QUILF are an indication as to how much the input parameters would need to change in order to get an exact solution. Ten degrees is indicative of a fully determined system and represents the uncertainty in the model reactions used (Anderson et al., 1993). Uncertainties for the log fO2 are determined by changing input parameters, temperature by ± 20oC and hematite component of ilmenite by ± 10 percent. Pressure Depth at given
pressure Temperature Log fO2 DFMQ
Value 3500 bars 13 km 758 degrees C -14.5 units 0.68 log units
Uncertainty ± 400 bars ± 0.4 km ± 10 degrees ± 0.6 0.14 log units
Section 4.4 Current Models For Nearby Physiographic Provinces:
This section examines two existing models for the formation of the BVF
rhyolites, 1. the rhyolites are a product of partial melting or 2. the rhyolites are a result of
fractional crystallization coupled with limited upper crustal assimilation.
There are a number of models put forth for topaz rhyolite formation in the Great
Basin and Basin and crustal extension areas of Mexico. Christiansen et al. (1986) has
92
four models of which only one refers to topaz rhyolite domes with the absence of caldera
structures. This model suggests partial melting of felsic granulites in the lower crust,
caused by basaltic magma ponding at the crust – mantle boundary, followed by either
extensive crystallization as the partial melts pass through colder crust or extensive
crystallization in a shallow magma chamber. They did not have access to Nd isotopes for
their 1986 study.
Orozco-Esquivel et al. (2002) report on “topaz rhyolites” from the Mesa Central
(MC) of the Sierra Madre Occidental Volcanic Province in Mexico. The rhyolites they
studied span a small range of SiO2 compositions (76 to 78.2 %) and their analysis show a
trend opposite to what would be produced by fractional crystallization with Rb, Nb, Ta,
and Th decreasing and Sr and Ba slightly increasing with increased SiO2 concentration.
Their model for the formation of these rhyolites involves both chemical and isotopic
disequilibrium partial melting of granulitic lower crust. This involves dehydration
melting of hydrous phases and buffering of the melt by residual feldspars.
Unlike some of the topaz rhyolites in the Great Basin, the rhyolites described by
Orozco-Esquivel et al. (2002) do not occur in a basalt – rhyolite bimodal field. There are
two sequences of rhyolites, a lower sequence (NOT a topaz style rhyolite) formed by
fractional crystallization of a mafic parent with some crustal assimilation and the upper
sequence, modeled above. The Nd isotopes, at εNd values less than -3 for all their
samples, are more radiogenic than those for the BVF basalts and the crustal rocks are of
Grenville age and thus the model described by Orozco-Esquivel et al. (2002) may not be
analogous to the BVF system with its much older crust. Additionally, they assert that the
geochemical trends they see are counter to fractional crystallization mechanisms. Their
93
range in SiO2 concentration is nearly within analytical error at 2-sigma if reported as
RSD (error reported only as ± 2%) on the eight topaz rhyolite samples analyzed. It is
also unclear that if the feldspars act as residual, buffer phases and biotite is the primary
phase melted, how SiO2 concentrations reach rhyolitic levels. They further state that
isotopic disequilibrium melting of the lower crust is “ …still a matter of discussion” and
that little is known about the mineral composition and physical conditions in the lower
crustal rocks under the MC.
Halliday et al. (1991) modeled the petrogenesis of high Rb/Sr silicic magmas
although not topaz rhyolites specifically. They favor fractional crystallization over
crustal partial melting to achieve high Rb and low Sr concentrations. Using a lower crust
bulk distribution coefficient (D) for Sr of 0.05 (as suggested by Bohrson and Spera,
2001), and a starting Sr content of 60 ppm (one of the lowest of the SRP xenoliths, from
Leeman et al., 1985), the Sr increases with batch melting. Even using a very high DSr (to
represent feldspar buffering) of 3, the Sr content in the CDF rhyolites can not be modeled
with a single stage of melting. A minimum of three partial melting episodes of less than
10 percent are required with the subsequent melt coming from the previous melted and
segregated crust. Some Leeman et al. xenoliths would require up to six of these partial
melting episodes and the high DSr values also rule out melting of upper crust.. Rb values
are equally unrealistic with reasonable DRb values.
Another reason crustal melts could not be the source of the BVF rhyolites is based
on isotopic signatures. The lower crust could not have been partially melted to produce
the BVF rhyolites because the Sr isotopes ratios are too low in the Wyoming age rocks.
Any combination of lower crustal melts and assimilated upper crust would produce Nd
94
isotopes that are too evolved (εNd < -25 based on the Leeman et al., 1985 xenoliths). The
partial melting of upper crust for the BVF rhyolites can be ruled out based upon the work
of Halliday et al (1991). Additionally, Christiansen (personal communication, 2004),
with new research (and given the ~2.2 Ga age of the crust in the eastern Great Basin) is
currently working on new models involving extensive fractionation of a mafic parent
coupled with minor upper-crustal assimilation, not partial melting, for topaz rhyolites.
Halliday et al. (1991) favor fractional crystallization with or without minor crustal
contamination and do not rule out rhyolite magma formation by fractional crystallization
of basalts. As Dsr increases with increasing SiO2 content, the percentage of fractional
crystallization decreases. This is the model favored by McCurry et al., (1999) for the
formation of the QESRP rhyolite domes (e.g. Big Southern Butte, Cedar Butte and East
Butte). The BVF and QESRP rhyolites have more similarities in their geochemistry than
compared to rhyolites from other tectonic or physiographic province, excepting topaz
rhyolites. Pickett (2004) showed a genetic link between the basalts of the two provinces
and Peng and Humphreys (1998) indicate a zone of partial melt running beneath both
areas (Figure 3) indicating the mafic parent material is similar for both regions. The
QESRP rhyolites have Sr and Nd isotopic ratios that are closer to the parent magma than
the BVF rhyolites and since parent magma is similar, the QESRP rhyolites must have less
crustal assimilation (Figure 40; McCurry et al., 1999). Hanna Nekvasil (personal
communication, 2004) has also produced rhyolites from the sequential crystallization of
SRP olivine tholeiite experimentally, showing that this range in chemistry is possible
utilizing only fractional crystallization.
95
Section 4.5 Patterns of Rock Chemistry in the BVF:
The relationships between the trace-element chemistry of the basalts and rhyolites
of the BVF can be examined to see if it is possible to link the two types of rocks using
fractional crystallization models. Figure 39 is spider diagram illustrating evolved basalt
(KEP-027) and the BFV rhyolite compositions, normalized to the most primitive BVF
basalt (KEP-004). This diagram, coupled with estimated partition coefficient (Kd’s) for
common igneous minerals can delineate plausible fractionating phases and possibly
timing.
Figure 39. Spider diagram of the CDF (23 samples – black lines), NDF (13 samples – grey lines) and KEP-027, the most evolved BVF basalt (triangles) using the most primitive basalt (KEP-004) for normalization. There were no uranium analyses of the basalts of the BVF however the CDF rhyolite is 20 to 60 times enriched in U as compared to the most primitive SRP basalts (Hughes et al., 2002a), similar to Th.
0.001
0.01
0.1
1
10
100
Cs Rb Ba Th K Ta La Ce Sr Nd P Hf Zr Sm Ti Tb
Roc
k / K
EP-0
04
96
The evolved basalt (KEP-027) is enriched over the most primitive basalt as the
phenocryst phases present in the basalt, olivine and plagioclase, are incompatible with
most of the elements listed. For REE and high field strength (HFS) elements, this
enrichment is uniformly about 3 to 1, except for Ti. This indicates that ilmenite was
produced (Ti Kdilmenite ≈ 200) at some point between these two end member
compositions. Likewise, the amount of plagioclase produced must be enough to raise the
bulk distribution coefficient (D) to unity in order to keep Sr concentrations equal for both
samples. The production of plagioclase may be enough to slightly increase bulk D for the
large ion lithophile (LIL) elements as plagioclase has higher Kd values for these elements
than for the REE or HFS elements listed. The Kdolivine is less than 0.05 for all the listed
elements. MELTS modeling (Pickett, 2004) indicate that cpx might fractionate at
moderate to high pressures (middle crust or greater) and fractionation of this phase could
explain the anomalous Sc pattern observed in the basalts. CPX however is not observed
as a phenocryst in the basalts.
While 50 to 60 percent of the phenocrysts in the CDF rhyolite are quartz (Kd
essentially zero for all listed elements), it also contains ~25 percent sanidine, ~15 percent
plagioclase and at least one percent hornblende and biotite and some magnetite and
ilmenite. The accessory phases (zircon >> apatite > thorite ≈ allanite) while minor in
abundance, have very large Kd values for some elements.
The rhyolites of the BVF show greater enrichment over the evolved basalt in most
of the LIL elements and slight depletion of most of the REE and HFS elements. Ba
shows marked depletion amongst the LIL elements and is indicative of alkali feldspar
fractionation. This fractionation is also responsible for the depletion of K, Sr and the
97
large Eu anomaly (Figure 34). The light REE are only slightly depleted compared to the
evolved basalt and the trace phases, especially allanite which has massive Kd’s over 2000
for the light REE, are likely the cause. The middle REE (Sm and Tb) plot near the
evolved basalt while the heavy REE (not shown) are enriched over the evolved basalt.
Another line of evidence for allanite fractionation is the 1:1 La to Yb ratio observed
(Figure 34) as allanite preferentially incorporates the light REE. Phosphorous depletion
may be a result of apatite fractionation and while apatite is an accessory phenocryst phase
in the rhyolite, it is possible that more significant apatite fractionation took place when
the magma was more mafic (and MELTS modeling, described later, support this
scenario). The HFS elements are slightly depleted as compared to the evolved basalt and
those depletions are attributed to zircon fractionation. The depletion of Ti is again mostly
due to ilmenite but hornblende and magnetite might also reduce Ti. Th is still
substantially enriched despite the fact that thorite, which has Th as a stoichiometric
constituent is an accessory phase. This phase may have nucleated late in the evolution of
the magma chamber and thus only small quantities were able to fractionate out before the
eruption. Fractionation of allanite would also remove Th from the melt. Typical allanite
analyses show 10 to 20 fold enrichments of La over Th (Deer et al., 1986) and thus large
quantities of the light REE elements are removed with allanite fractionation while
substantially small amounts of Th are removed. Approximate Kd’s used above were
taken from the Geochemical Earth Reference Model (GERM) (http://earthref.org/) and
Rollinson (1998).
98
Section 4.6 MELTS Modeling To Determine Crystallization Sequences and Timing:
In this section, MELTS modeling (Ghiorso and Sack, 1994) is used to help
constrain the timing of phase formation so that bulk distribution coefficients for the
magma can be estimated. I failed to mimic the phases or bulk chemistry of the BFV
rhyolites with numerous runs invoking both equilibrium and fractional crystallization
with reasonable variation in other parameters (e.g. participating phases and intensive
variables). Resultant chemistries however were rhyolitic. Christiansen et al. (1983)
indicate that high fHF stabilizes biotite over fayalite in topaz rhyolites but it is unclear if
MELTS accounts for these unusual magmatic conditions.
Pickett (2004) had poor results modeling the basalts of the BVF with MELTS.
Ghiorso and Sack (1994) state their model is only applicable above 900oC and I could not
model the rhyolites from the evolved basalts using MELTS. MELTS is unable to model
accessory phases (e.g. allanite, thorite) which play substantial roles in determining the Kd
values for many trace phases and the magma chemistry of this system (Figure 39 and
Patterns of Rock Chemistry in the BVF discussion). I was unable to reproduce the
phenocryst assemblage of the rhyolites resulting in less robust control of calculated
magma distribution coefficients used in EC-AFC modeling (next section). This
represents possible limitations within the MELTS program, due at least in part to the
unusual chemistries of topaz rhyolites and extensive fractionation to temperatures below
900oC, and does not indicate problems with the liquid line of decent (fractional
crystallization) hypothesis that links the basalts and rhyolites of the BVF.
99
Section 4.7 Isotope Modeling With Bulk Assimilation and EC-AFC:
After describing the upper crustal rocks under the ESRP, two approaches are used
to model the isotopic evolution of the CDF magma 1. single stage bulk mixing models
and 2. EC-AFC (Energy-Constrained Assimilation and Fractional Crystallization - Spera
and Bohrson, 2001; Bohrson and Spera, 2001). Both of these models have a given set of
assumptions, some of which are poorly constrained, and thus results shown in this section
are intended to only show some consistency between petrographic, geochemical and
model results. This is designed as a plausibility study; can we produce CDF rhyolites
from basalts in the BVF with reasonable constraints?
The crust beneath the BVF consists of Wyoming Terrane cratonic rocks (Figure
5) (O’Brien et al., 1995). The age of these rocks, approximately 2.7 Ga, results in upper
crustal isotope signatures that are in stark contrast to those of the mantle reservoirs
(Faure, 2001; Zindler and Hart, 1986). Actual isotope measurements from crustal
xenoliths from the northern margin of the Snake River Plain (SRP), brought to the
surface by Quaternary volcanism, and the Archean Albion Range (along the southern
margin of the SRP) are documented by Leeman et al., (1985) and given in Table 9.
Leeman et al. (1985) states that the “least modified” xenoliths equilibrated at ~5 kbars
which corresponds to a depth of 18.5 ± 1 km, right at the upper crust – lower crust
boundary (Sparlin et al., 1982) or just a few km above it (Peng and Humphreys, 1998;
Figure 3). Matty (1984) calculated pressures greater than 5 kbars for most xenoliths.
The assertion by Leeman et al., (1985) and Matty (1984) that most of his xenoliths are
representative of the deep crust is in contrast to the elevated Sr isotopic signatures of
100
these 2.7 Ga or older rocks which would place the xenoliths in the upper crust (Faure,
2001). While Leeman et al., (1985) does investigate some xenoliths from
the lower crust, based upon 87Sr/86Sr ratios of less than 0.706, none of those were
included in this study as they have lower 87Sr/86Sr values than those of the rhyolites and
indeed the basalts of the BVF. Assimilation of lower crustal xenoliths could not produce
an increase in the 87Sr/86Sr ratios and in fact would make the rhyolites appear more
primitive, isotopically, than the basalts. A working assumption is that the crustal rocks
that underlie the BVF are similar to those under the SRP as volcanics from both areas
must pass through the Wyoming Terrane rocks.
Table 9. 87Sr/86Sr and 143Nd/144Nd isotopic ratios for selected crustal xenoliths from the Snake River Plain and Albion Range (south-central Idaho) with abbreviations as follows: SK = Spencer-Kilgore; COM = Craters of the Moon; SM = Square Mountain and ALB = Albion Range. Errors for Sr isotope data are ± 0.00004 or better while errors for Nd isotopes are ± 0.000030 or better and average ± 0.000018. As stated in the text, all of these xenoliths are proposed to be from the upper crust based on their age and evolved Sr isotopic signatures. All data from Leeman et al., 1985.
Location Sample Sr
(ppm) 87Sr/86Sr Nd
(ppm) 143Nd/144Nd sigma Nd SK 73-68X 175 0.72550 22.3 0.510895 -34.0 COM CKI-1 211 0.73359 13.5 0.510499 -41.8 COM 70-40 324 0.71525 9.8 0.510454 -42.6 COM SI-1 70 0.71795 85.0 0.510590 -40.0 COM COM-1 97 0.81728 9.2 0.510228 -47.1 SM SM-2G 559 0.72092 7.1 0.509973 -52.0 SM DM-103 310 0.73878 25.7 0.510853 -34.9 ALB YAG-799 115 0.8060 19.0 0.511278 -26.6 ALB YAG-800 181 0.89260 59.0 0.510675 -38.3
101
Simple single-stage mixing hyperbolas between the most primitive SRP
compositions (Hughes et al., 2002) and selected Leeman et al., (1985) xenoliths and
Albion range rocks are graphed in Figure 40. The hyperbolas show that in order to
achieve the isotope ratios in CDF rhyolites, between 30 and 45 percent bulk assimilation
of crustal component is required if the most primitive SRP magma is used as a parent. If
the most primitive BVF basalt is used as the parent, 20 to 35% crustal component is
required (Ford and McCurry, 2003). The mixing hyperbolas indicate that more crustal
component must be added to the magma to form the isotopic signature of the parent of
the BVF rhyolites than the parent of the QESRP domes (prior to the bulk of fractional
crystallization).
Figure 40 (overleaf) Simple one-stage mixing models using isotope ratios of the most primitive basalt from the SRP (Hughes et al., 2002) and crustal xenoliths from Leeman et al., 1985 (see Table 9). Each tick mark represents an additional five percent bulk assimilation of crustal rocks to the hypothetical basaltic parent magma. For example, the first tick mark along mixing hyperbola COM SI-1 (located on the left hand side of the graph) represents five percent bulk addition of xenolith COM SI-1 to the parent magma, the second tick mark represents ten percent xenolith assimilation and so on. Pickett, 2004, established an isotopic and chemical link between the BVF basalts and those of the SRP and indeed, the most primitive BVF sample (Table 7) plots in the SRP field. Additional evidence that links the two basaltic fields is a contiguous zone of inferred partial melt at the crust – mantle boundary (Peng and Humphreys, 1998 – Figure 3). Quaternary rhyolites of the Eastern Snake River Plain (e.g. Big Southern Butte, East Butte and the silicic compositions of Cedar Butte) plot in or very near the SRP basalt field and contain less than five percent crustal component, perhaps much less depending on the starting parent basalt isotopic ratios (McCurry et al., 1999). The CDF data points (one point with large Nd isotope errors is not represented here – see Table 7) lie along lines represented by crustal xenoliths COM 70-40 and SK-73-68X. Between 30 and 45 percent assimilation of custal rocks with these isotope signatures into the most primitive SRP basalt is required to produce the isotope signatures in the CDF rhyolites while only 20 to 35 percent crustal assimilation is needed if the parent magma has the isotopic composition of the most primitive BVF basalt. The one isotope point from the NDF rhyolites (Table 7) requires approximately five percent more crustal assimilation than the CDF rhyolites.
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103
The curves also indicate that the chemically evolved basalts in the BVF need to
have assimilated some crust to achieve their isotopic signatures. To produce the isotopic
composition of most evolved basalts (KEP-027) from the most primitive (KEP-004)
requires 6 to 18 percent bulk assimilation, depending on the xenolith (Ford et al., 2004).
While this model can reproduce the isotopic signature of the BVF rhyolites, the resultant
hybrid composition would be andesitic and requires fractional crystallization to produce
the chemical signature of the BFV rhyolites. This bulk assimilation model will be
contrasted against the EC-AFC model.
EC-AFC modeling (Spera and Bohrson, 2001; Bohrson and Spera, 2001) allows
for simultaneous assimilation and fractional crystallization of a magma and incorporates
fractional melting of country rock, the latent heats of crystallization and heat transfer to
country rocks using an EXCEL spreadsheet form. The theoretical basis for an extended
model incorporating magma recharge (EC-RAFC - Spera and Bohrson, 2002; Bohrson
and Spera, 2003) has been published, but was not yet included in the EXCEL macro,
therefore I used the program without magma recharge (EC-AFC).
Magma recharge does not affect the late-stage evolution of the magma chamber.
Petrographic evidence does indicate the addition of olivine tholeiite-like basaltic magma
to the rhyolitic chamber (i.e. magmatic recharge) with at least some localized
hybridization (Figure 27) just prior to eruption. Field work indicates that such enclaves
are rare and trace element calculations show that the rhyolite to basalt ratio must be at
least 100:1 (i.e. less than one percent basalt composition magma added and mixed with
the rhyolitic magma). The best example of this is Sr. Even if the rhyolitic magma started
104
with absolutely no Sr, an unlikely scenario, an addition of one percent typical BVF basalt
(Pickett, 2004 at 350 ppm) would raise the Sr content in the magma to 3.5 ppm, higher
than the average for the CDF rhyolites. Magmatic recharge may or may not have
happened at earlier stages in the magma chamber evolution but the recharge that certainly
did occur shortly before the eruption, as evidenced by the enclaves, was very minor and
can be ignored for modeling purposes.
A number of parameters for the EC-AFC model are briefly detailed below. Each
parameter is adjusted to best represent the local magma chamber conditions. For
example, the assimilant trace element distribution coefficient (Da) for Sr can be given
either typical upper crustal or lower crustal values, depending upon where the user thinks
assimilation is occurring.
Thermal parameters:
- (Tlm) magma liquidus temperature: can be estimated using MELTS runs or by some
other means. This parameter helps to constrain the depth of the system.
- (Tm0) magma initial temperature: often the same as Tlm.
- (Tla) assimilant liquidus temperature: estimated based on type of assimilant.
- (Ta0) assimilant initial temperature: Based on a number of factors including depth of
assimilation, the geotherm of the area and local conditions (i.e. the possibility that
recent basaltic intrusions have increased the local country rock temperature).
- (Ts) assimilant solidus temperature: an estimate that must be hotter than the Ta0.
- (Teq) equilibration temperature: With the above thermal input, the user has the option
of choosing an equilibration temperature from a pull-down menu. Spera and
Bohrson (2001) suggest this temperature be set below the eruptive temperature of
105
the most evolved rocks in a volcanic system that contains a spectrum of
compositions. This increases mass of country rock affected by the heat from the
magma body. If Teq < Ts, no anatexis will take place.
- (∆T) normalized temperature change, set to -0.005 for all runs in this study.
Geochemical parameters:
This includes magma elemental concentrations (bulk rock) of Sr and Nd as well as the
87Sr/86Sr and 143Nd/144Nd isotopic ratios and the trace element (i) bulk distribution
coefficients for the magma (Dmi) and assimilant (Dai). Maximum and minimum bulk
distribution coefficients for the basaltic magmas of the BVF were calculated using the
equation: Ai
Ai KdWD *∑=
where W is the weight fraction of phase A, and i is the element of interest. Resultant Dmi
values are given in Table 10 and were calculated using phenocryst crystallization
percentages calculated by Pickett (2004) and approximate Kd’s from the Geochemical
Earth Reference Model (GERM) (http://earthref.org/) for olivine and plagioclase in
basaltic rocks and specific gravities for these phases.
Table 10: General parameters used in EC-AFC modeling for the basalts of the BVF. The Tlm corresponds to a depth of approximately 22 km and the Ta0 is higher than the geothermal gradient to indicate local country rock heating by basalts flooding the region over time. The Di’s for the magma were figured using the method stated above. Starting Sr and Nd isotopic and trace element data for KEP-004 are given in Table 7. Linear melting functions were used for all runs. Tlm Tm0 Tla Ta0 Ts Teq
value DmSr min.
DmSr max
DmNd min
DmNd max
1300 1300 1100 950 1000 minimum 1.07 1.45 0.029 0.036
Other parameters:
106
Other options include assigning different crystallization or fusion enthalpies,
changing specific heats of the magma and assimilant, assigning linear or user-defined
non-linear melting functions and including additional isotope and trace elements in the
modeling. None of these parameters were altered for these runs. There are also various
ways to change the output format.
The first step in using EC-AFC model was to try to produce the most evolved
basalt in the BVF (KEP-027) from the most primitive basalt in the BVF (KEP-004).
Bohrson and Spera (2001) suggest using DaSr values of 1.5 if in the upper crust, 0.05 for
Sr in the lower crust, and 0.25 for Nd in either the upper or lower crust. Runs using
lower crustal values of Dai are not reported as all crustal xenoliths used in modeling have
upper crustal isotopic signatures. Of the possible assimilants given in Table 9, if upper
crust Dai values are used, only xenolith SM-2G has a chance of elevating the Sr content
back to 350 ppm in the evolved basalts after some fractional crystallization of plagioclase
from the most primitive basaltic composition (Section 4.5 and Pickett, 2004). While Nd
and Sr composition data can not be simultaneously modeled with any possible assimilant
or Dmi values in the range given in Table 10 and upper crustal Dai values, the isotopic
data can be modeled and indicates crustal assimilants in the 5 to 12 percent range (Figure
41), slightly less but similar to the bulk assimilation model discussed earlier.
Modeled Nd concentrations can reach expected levels for those assimilants with
high concentrations of Nd (e.g. SK-73-68X, SM-DM-103, YAG-800) however the Sr
concentrations in these assimilants is too low and thus as soon as assimilation begins, the
modeled Sr content drops rapidly. Using the minimum DmSr (1.07), the Sr content just
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109
during the fractional crystallization, before assimilation begins, drops about 25 ppm.
This is represented by approximately 58 percent of the magma chamber being
crystallized, before the country rock begins to assimilate, at the thermal parameters giver
in Table 10. Thus, if the starting value for the most primitive magma was increased to
375 ppm, a value that is certainly plausible, using compositions from Pickett (2004) and
with the expected low percentages of assimilation, only the assimilants with the lowest Sr
contents and/or the least evolved Sr-isotope values (e.g. COM-SI-1, SK-73-68X) fail to
model the expected Sr content. A similar statement is true for Nd concentrations. If the
stating Nd is increased by three to six ppm, the model can more accurately predict the
expected Nd concentrations and assimilation percentages range from 4 to 12 percent
(Figure 42).
Adjustments to the poorly constrained Ta0 (assimilant initial temperature) can
change the Sr and Nd concentrations. A lower value of Ta0 (colder crust) results in
assimilation being retarded and allows for more fractional crystallization before
assimilation. Higher values of Ta0 result in assimilation happening earlier in the
evolving magma chamber. A number of runs indicate that a value of 930oC (slightly less
than was used in the above work) yields the best results using minimum Sr and Nd Dmi
values with xenolith SM-2G (the high Sr xenolith). While Sr concentrations modeled
with this run are still slightly low (330 ppm), the isotope (87Sr/86Sr = 0.70871 and
143Nd/144Nd = 0.51224) Nd results (80 ppm – compare to KEP-027, Table 7) nearly
mimic the evolved BVF basalt values with 6.2 percent assimilation and approximately 66
percent crystallization. Pickett (2004) suggests that between 64 and 72 percent
crystallization is required to model the La and Th values in the basalts and MELTS runs
110
also indicate 70 percent crystallization is required to obtain the proper MgO
concentration. Neither of Pickett’s calculations incorporated assimilant geochemistry and
Hughes et al. (2002b) suggest that this level of crystallization results in rocks that are
more differentiated that what is seen in the BVF basalts.
The overall sensitivity of the model for the BVF basalts is more dependent on the
concentrations of the analytes and Ta0 values than on the expected variations in Dmi
values. For example, a change in the DmNd value within the limits given above barely
changed the resultant Nd concentrations. A minimum DmNd value of 0.001 was tried and
resulted in slightly higher Nd concentrations at similar Sr concentrations but again, the
change was not enough to correctly model the Nd concentrations (using the temperature
constraints from Table 10). Using the maximum value for DmSr (Table 10) only widened
the gap between the observed data and the modeled data and given that there is no trend
between SiO2 and Sr in the basalts (i.e. no decline in Sr as the basalt increases in SiO2 as
a result of assimilation and fractional crystallization), a value closer to 1 is more
appropriate for DmSr.
Bohrson and Spera (2001) give upper crustal values of DaSr = 1.5 and DaNd =
0.25. Changes in DaSr to values less than one could increase the Sr content of the magma
chamber after assimilation begins for those assimilants with low Sr concentrations (i.e.
less than the starting magma). Such low bulk distribution coefficients for upper crustal
assimilant are unwarranted given the modal analysis of Snake River Plain xenoliths
(Matty, 1984; Leeman et al., 1985). All of their xenoliths have modal plagioclase plus
alkali feldspar plus interstitial glass greater than 50 percent. Many of the xenoliths have
more than 40 percent modal plagioclase plus alkali feldspar which would yield DaSr
111
values greater than 1.2 (for Sr Kdplagioclase = 3 and Kdalkali feldspar = 4). Additional evidence
that DaSr values should be greater than unity is electron microprobe analysis of the
interstitial glass in xenolith DM-103 reported by Matty (1984). The bulk rock Sr values
are 310 ppm while the interstitial glass has 85 ppm indicating Sr prefers the crystalline
phases over the melt. Other typical “feldspar compatible” elements such as Ba and Eu
also have lower concentrations in the glass than in the bulk rock.
Trying to model the rhyolites from the most evolved basalt (KEP-027) is
somewhat more difficult for a number of reasons. First, the basalt sample in question has
no phenocrysts, as does another sample with similar chemistry (Pickett, 2004) and thus it
is not clear what phases are precipitating in the evolved basalts. Pickett’s MELTS runs
indicate that two types of clino-pyroxene and ortho-pyroxene, in addition to spinel,
olivine and plagioclase fractionate at 5 kbar runs. There is no petrographic evidence in
any of the BVF basalts for pyroxene phenocryst development, and pyroxene is very rare
in the similar SRP rocks as well (Hughes et al., 2002b). With no phenocrysts present in
the rocks and inconclusive MELTS results indicating fractionating assemblages that do
not exist in the rocks, the Dmi values can not be accurately calculated.
A second problem is the great span in chemistry between the starting composition
(KEP-027 at 51.5 % SiO2) and the BVF rhyolites (76.7 % SiO2) with no surface rock
samples from this composition gap. It is unclear when many phases, such as alkali
feldspars, will start to form and significantly change the Dmi values. While Cedar Butte
from the SRP spans this compositional gap, it is a poor analog because it’s silicic rocks
do not contain the same phase assemblage as the BVF rhyolites and the most chemically
112
evolved rocks (highest SiO2) have isotopes that are more primitive than KEP-027 from
the BVF (McCurry et al., 1999). Additionally, most topaz rhyolites come from basalt-
rhyolite bimodal fields as well (e.g. Thomas Range, Spor Mountain, Wah Wah
Mountains, Sheep Creek Mountains, and others) and while topaz rhyolites have similar
phases (Christiansen et al., 1986), they also lack the intermediate rocks required to show
the timing for phase development.
The problems associated with the calculation of the magma bulk distribution
coefficients are accentuated by the high volatile (fluorine and chlorine) percentages in the
BVF rhyolites. Topaz rhyolites with similar bulk, trace, and volatile chemistries are the
only rhyolites in a suite of compositions that do not fit models designed to calculate Kd’s
for a number of elements (Ren et al., 2003) in plagioclase. Ren et al. (2003) go on to
state that high F and Cl values, “…might inhibit Sr from entering plagioclase.” His study
indicates plagioclase KdSr values of between 2 and 3 for topaz rhyolites with similar
plagioclase An contents, slightly more CaO and slightly less SiO2 as compared to BVF
rhyolites.
Ren’s (2004) experiments with sanidine Kdi values also indicate that topaz
rhyolites are harder to model than other rhyolites but indicate that KdSr is inversely
proportional to Al2O3 content and increases slightly with SiO2. The BVF rhyolites have
slightly higher SiO2 and lower Al2O3 than the topaz rhyolites in Ren’s (2004) study
indicating that sanidine KdSr is about 4 for the BVF rhyolites, slightly higher than the
topaz rhyolites sited Ren’s work.
One final issue is the low percentage of phenocrysts in the BVF rhyolites.
Calculating relative proportions of the fractionating phases where phenocrysts make up
113
less than ten percent of the rock introduces large errors in calculating Dmi values
especially when trying to account for uncommon trace phases with very high Kd’s (i.e.
Nd Kdallanite ≈ 1600). The unknown starting assemblage in the BVF evolved basalts,
unknown timing of the beginning phase fractionation (e.g. sanidine, zircon, allanite,
hornblende, biotite), changes in Kdsan and Kdplag with changes in bulk chemistry as
fractionation continues (Ren et al., 2003; Ren, 2004), and poor constraints on the relative
proportions of trace phases with high Kd’s for Nd all conspire to make calculation of Dmi
values difficult. Even with all the uncertainty described above, it is unlikely that the
DmSr value would be greater than three in the BVF magma.
Assimilation of crustal rocks must occur earlier in the magma evolution rater than
later because if any assimilation occurs after large amounts of fractional crystallization,
the small volume of resultant magma is changed enormously by the additions of even
small volumes of partially melted assimilant. For example, using evolved basalt
composition KEP-027 and an estimated DmSr of 3 (indicative of adding alkali feldspars at
some point in the fractionation process) to drop the Sr concentration in the magma below
5 ppm requires 89% crystallization of the magma body. Addition of even one percent
COM-SI-1 assimilant, one of the most primitive in Sr isotopic ratios and lowest
concentration of Sr of any of the assimilants, results in magma isotopic ratios that are too
high using upper crustal DaSr values. Other assimilants that are higher in Sr
concentration also increase magma Sr contents to unreasonable levels. Increasing DmSr
above three requires less crystallization of the magma chamber to achieve the same
magma Sr concentration and thus decreases the affects of small amounts of assimilant on
114
the overall isotopic ratios of the magma. Values of DmSr above 3 seem unwarranted
based on the work by Ren et al. (2003) and Ren (2004) which place the Sr Kdsan at
approximately 4 and the final phenocryst assemblage is only about 25 percent (volume)
sanidine based on petrography, while intermediate (un-erupted) compositions may
contain even less sanidine.
If assimilation occurs early in the evolution of the magma chamber, then the total
assimilation required using EC-AFC modeling to get from the most primitive basalt to
the rhyolite isotopic signatures is between 9 and 18 percent, less than the 20 to 35 percent
suggested by bulk assimilation (Figure 40). The percentage of assimilation to produce
the isotopic signature in the rhyolites is only slightly more than required to produce the
evolved basalts. To model the Sr and Nd concentrations, the most primitive magma must
undergo 60 to 70 percent fractional crystallization and the evolved basalt must undergo
an additional 80 to 90 percent crystallization (a total of 92 to 97 percent total fractional
crystallization). Production of high (~50 ppm) concentrations of Th in the rhyolite
requires at least 98 percent fractional crystallization of the most primitive parent and
greater than 95 percent from the more evolved basalts. These fractionation percentages
are higher than those modeled for Sr and Nd concentrations.
In summary, I was unable to produce the rhyolite phase assemblage and chemistry
from the basalts with MELTS models (Ghiorso and Sack, 1994), perhaps because the
model is not well calibrated for the unusual topaz rhyolite chemistry or fractionation at
such low temperatures. Bulk assimilation is a good initial isotopic model but over
estimates the amount of assimilant required. EC-AFC models (Spera and Bohrson, 2001;
Bohrson and Spera, 2001) indicate that assimilation must have occurred early in the
115
evolution of the magma chamber and totaled between 9 and 18 percent. Slight changes to
the original starting composition for Nd and Sr and minor changes in Ta0 can accurately
model the resultant Nd and Sr concentration and isotopes in the basaltic rocks, not
reasonable changes in Dmi values. The model also indicates that the most primitive
basalt must be crystallized between 60 and 70 percent to produce the evolved basalt,
similar to results discussed by Pickett, 2004 to model some trace elements. This evolved
basalt then requires an additional 80 to 90 percent crystallization to reach proper rhyolite
Sr concentrations. Recharge to the system by basaltic magma was not modeled and may
take place early in the evolution of the magma chamber but the recharge that did happen
just prior to eruption (as evidenced by the mafic magmatic enclaves) was very minor.
Section 4.8 NDF Petrogenesis and Regional Implications:
This section first examines some of the similarities and possible differences
between the petrogenesis of the CDF and the NDF within the BVF. I then briefly discuss
the implications and possible causes for the differences between the QESRP and BVF
rhyolites.
All three domes from the CDF are chemically and petrographically identical as
are all three domes from the NFD. The two dome sets are similar in many respects (i.e.
major element chemistry, phenocryst assemblage, the presence of basaltic magmatic
enclaves, emplacement in a basalt-rhyolite bimodal field) but do contain the following
differences: 1. the NDF is ~1.4 million years older than the CDF, 2. the NDF has
evidence of granophyric texture (Figure 24), 3. possible accessory phase differences, 4.
differing enrichments in trace element compositions (Table 2, Table 5, Figure 39) and 5.
116
slightly different isotope ratios (Table 7). Some of these differences may result in slight
changes to the petrogenetic model developed for the CDF.
Granophyic texture is present in silicic glomerocrysts in the NDF but not the CDF
rhyolite. This texture forms as a result of rapid crystallization of alkali feldspar and
quartz due to a rapid loss of volatiles or depressurization (change in PH2O) and occurs at
depths less than ~10 km (Mason, 1985 referenced in Wilson, 1989). Granophyric texture
is also found in East Butte of the SRP in plagioclase crystals, not in glomerocrysts (R.
Ganske, personal communication, 2005) again indicating magma chamber residence at
less than 10 km. This indicates that the NDF magma chamber may have resided at more
shallow depths than the CDF. Conversely, the presence of this texture may indicate
higher volatile partial pressures in the NDF magma chamber. The exact consequence of
finding this texture in a glomerocrysts is unknown.
The heavy mineral separates EMP mount for the NDF is dominated with zircon
and although allanite or thorite were not evident, it is possible these phases exist (no EDS
scans were performed on this sample nor has the probe mount been made into a thin
section). Although there is two to three times as much heavy separates produced from a
nearly identical amount of rock processed, it is unclear if the NDF sample contained a
higher percentage of phenocryst. While zircons were not represented in the point count
data on the thin sections examined, they appear to be more common in the NDF rhyolite
slides. This may indicate greater amounts of zircon fractionation resulting in lower
concentrations of Zr, Hf, Ta, and U (Figure 39, Figure 32; Table 5), all with high Kdzircon
values. More work is needed to assess the affects of other accessory phases on the trace
element compositions (i.e. the slightly higher La to Yb ratio for the NDF). While there
117
are some differences in the trace elements, it is important to point out that both dome sets
have highly elevated “incompatible” elements (i.e. Rb and Cs) as compared to other types
of rhyolite, other than topaz rhyolites, and the formation of both requires large amounts
of fractional crystallization.
The NDF has slightly different isotope ratios than the CDF indicating
approximately 5 percent more crustal assimilation in the NDF (Figure 40). One possible
hypothesis for this additional assimilation of crustal component in the BVF is that the
middle crust under the NDF was less refractory and therefore easier to melt than the
middle crust under the CDF because the crust had not been receiving basaltic injections
for as long (i.e. lower the Tla or Ts values). Another hypothesis is that the parent basalt
had a different isotopic ratio although the one isotope point from the Willow Creek Lava
Field (Table 7) is similar to that of the most primitive BVF sample (KEP-004) used in the
assimilant calculations. More complete isotope work on the basalts of the BVF, which is
forthcoming, will elucidate this question.
The ESRP crust has received more voluminous basalt injections over a longer
period of time. Additionally, the voluminous Tertiary rhyolites have stripped the crust
under the ESRP of mobile LIL elements. The well defined and refractory mid-crustal sill
(Peng and Humphreys, 1998; Figure 3) is the position of restite material from the partial
melts that produced the Tertiary rhyolites and products of fractional crystallization of the
more recent basalts (McCurry et al., 2002). Thus, the crust under the SRP maybe more
difficult to assimilate and contain lower concentrations of the LIL elements as compared
to the crust under the BVF.
118
Chapter 5 CONCLUSIONS
Both sets of spatially separated rhyolite domes in the BVF are chemically and
petrographically homogeneous. This, coupled with the alignment of the dome sets and
associated features (i.e. craters in the CDF) indicates that the each set is likely connected
to an intrusive dike system. The domes in the CDF were produced after local basalt
emplacement and faulting although other faults or basalt flows located away from the
domes might have occurred after dome formation. The CDF eruptions began with vent-
clearing, hydrovolcanic tephra production before the onset of simultaneous tephra
production and largely endogenous dome growth. The higher elevations between China
Cap and North Cone in the CDF are caused by tumescence above the intruding dike with
the craters (e.g. Gronewall and Burchett Lakes) forming as a result of extensional faulting
above the dike, collapse after magma withdrawal from the dike, or a combination of
these. I was not able to find any post emplacement vapor phase topaz mineralization
(Dayvault et al., 1984).
The amphibole geobarometry indicates the CDF magma equilibrated at ~3.5 kbars
which corresponds to depths or 13 ± 0.4 km, depending on the density of the upper crust.
This depth roughly corresponds to the position of the mid-crustal sill under the adjacent
ESRP (Peng and Humphreys, 1998; Figure 3). The temperature of the magma is tightly
constrained at 758o ± 10o C and the log fO2 is -14.5 units, calculated by QUILF (Andersen
et al., 1993). All phenocryst phases, except for quartz, in the CDF are in textural
equilibrium with the magma. The strongly embayed quartz crystals are likely the result
of a drop in volatile partial pressures which can destabilize quartz (Hibbard, 1995).
119
Previous models for the formation of topaz rhyolites invoke the partial melting of
lower crust (Christiansen et al., 1986, Orozco-Esquivel et al., 2002). These models are
rejected on the basis that unlikely small portions of multiple, sequential partial melts are
required to produce the low Sr concentrations (less than 3 ppm) and correspondingly high
Rb/Sr ratios (over 200), even with incongruent, non-modal melting. Additionally, low
εNd values in the Archean crust (Zindler and Hart, 1986, Faure, 2001) under the BVF
preclude partial melting of either the upper or lower crust or a combination of the two as
a principle mechanism for forming the CDF rhyolites.
Assimilation and fractional crystallization (AFC) evolution of the basalts is the
preferred model for the formation of the CDF rhyolites. This model has been used in the
adjacent ESRP province to model the development of Quaternary rhyolite domes (e.g.
Big Southern Butte, Cedar Butte, East Butte) from basalt by McCurry et al., (1999) and
rhyolite compositions have been produced experimentally by fractionally crystallizing
SRP olivine tholeiites (H. Nekvasil, personal communication, 2004). E. Christiansen
(personal communication, 2004) is also working on a new model utilizing AFC for topaz
rhyolites in the Great Basin.
The BVF and ESRP are linked by a geophysically anomalous zone at the crust –
mantle boundary inferred to be partial melt (Peng and Humphreys, 1998; Figure 3).
Also, the basalts, or parent material, of the bimodal volcanic ESPR are geochemically
and isotopically similar to the basalts of the BVF (Pickett, 2004). Fractionation of the
rhyolite phenocryst phases can produce the trace element patterns of the rhyolites from
the basalts although required percentages were not calculated. Even the low
concentrations of Sr and Eu in the BVF rhyolites can be modeled with extensive
120
fractional crystallization (Halliday, et al., 1991). High concentrations of Th require 95
percent fractional crystallization of the evolved basaltic parent. I was unable to model
the phenocryst assemblages in either the evolved BVF basalt or the rhyolite with MELTS
models (Ghirso and Sach, 1994) and thus the timing of phase development or loss in the
magma is unknown. MELTS models did produce a rhyolitic magma but not of the
compositions seen in the BVF.
About 20 to 30 percent bulk assimilation of upper crust is required to form the
isotopic signatures in the BVF rhyolites when starting with the most primitive BVF basalt
although additional magma processing (e.g. fractional crystallization) is required to create
a “rhyolitic” composition. When Energy-Constrained Assimilation Fractional
Crystallization models are used (Spera and Bohrson, 2001, Bohrson and Spera, 2001), 4
to 12 percent upper crustal assimilation coupled with 60 to 70 percent fractional
crystallization is required to form the evolved basalts. A total 9 to 18 percent upper
crustal assimilation with an additional 80 to 90 percent fractional crystallization is
required to form the rhyolites from the most primitive BVF basalt. The percentage of
assimilant is largely dependent upon the stating xenolith composition and how close the
country rocks (assimilants) are to their solidus temperature. Slight and acceptable
changes in the model parameters can model the Sr and Nd elemental concentrations with
little effect on the total amount of assimilant required to model the isotopes. All of the
assimilation occurred early in the evolution of the magma.
Magma recharge into the system was not modeled and may present a problem for
future workers. The recharge event that did occur just prior to the eruption, as evidenced
by basaltic magmatic enclaves, was very minor and did not affect significantly affect the
121
element concentrations or isotopic ratios in the rhyolite. Although there is no evidence to
suggest more substantial magma recharge earlier in the evolution of the magma chamber,
it is possible.
122
REFERENCES
Andersen, D. J., Lindsley, D. H. and Davidson, P. M., 1993, QUILF: A program to assess equilibria among Fe-Mg-Ti oxides, pyroxenes, olivine, and quartz: Computers in Geoscience, v. 19, p. 1333-1350 Armstrong, F. C., 1969, Geologic map of the Soda Springs quadrangle, southeastern Idaho: U. S. Geological Survey Miscellaneous Investigations Map I-557. Armstrong, F. C. and Oriel, S. S., 1965, Tectonic development of the Idaho-Wyoming
Thrust Belt: American Association of petroleum Geologist Bulletin, v. 49, no. 11, p. 1847-1866.
Armstrong, R. L., Leeman, W. P., Malde, H. E., 1975, K-Ar dating, Quaternary and
Neogene volcanic rocks of the Snake River Plain, Idaho: American Journal of Science, v. 275, p. 225-251.
Asmerom, Y., Jacobsen, S. B. and Wernicke, B. P., 1994, Variations in magma source regions during large-scale continental extension, Death Valley Region, Western Uninted States: Earth Planetary Science Letters, v. 125, p. 235-254 Bacon, C. R., 1985, Magmatic inclusions in silicic and intermediate volcanic rocks: J. of Geophysical Research, v. 91, No. B91, p. 6091-6112. Bacon, C. R. and Hirschmann, M. M., 1988, Mg/Mn partitioning as a test for equilibrium between coexisting Fe-Ti oxides: American Mineralogist, v. 73, p. 57-61. Baedecker, P. A., Grossman, J. N. and Buttleman, K. P., 1998, National geochemical data base: PLUTO geochemical data base for the United States: U. S. Geological Survey digital data series, DDS-47, one optical computer disk. Bindeman, I. N. and Valley, J. W., 2001, Low D18O rhyolites from Yellowstone: magmatic evolution based on analyses of zircons and individual phenocrysts: J. of Petrology, v. 42, p. 1491-1517. Bindeman, I. N. and Valley, J. W., 2002, Oxygen isotope study of the Long Valley magma system, California: isotope thermometry and convection in large silicic magma bodies: Contributions to Mineralogy and Petrology, v. 144, p. 185-205. Blackstone, D., 1977, The Overthrust Belt salient of the Cordilleran fold belt – western Wyoming, southeastern Idaho, northwestern Utah: Wyoming Geological Association 29th Annual Field Conference Guidebook, p. 367-384.
123
Bohrson, W. A. and Spera, F. J., 2001, Energy-constrained open-system magmatic processes II: Application of energy-constrained assimilation-fractional crystallization (EC-AFC) model to magmatic systems: J. of Petrology, v. 42, no.5, p. 1019-1041. Bohrson, W. A. and Spera, F. J., 2003, Energy-constrained open-system magmatic processes IV: Geochemical, thermal and mass consequences of energy- constrained recharge, assimilation and fractional crystallization (EC-RAFC): Geochemisty, Geophysics, Geosystems (electronic journal), v. 4, no. 2, 25p. Bretches, J. E., 1984, A geologic study of East Butte, a rhyolitic volcanic dome on the Eastern Snake River Plain, Idaho: State University of New York at Buffalo Masters Thesis 159 p. Brophy, J. G. and Dreher, S. T., 2000, The origin of compositional gaps at South Sister volcano, central Oregon: implications for fractional crystallization processes beneath active calc-alkaline volcanoes: J. of Volcanology and Geothermal Research, v. 102, p. 287-307. Brophy, J. G., Dorais, J. M., Donnelly-Nolan J. M., and Singer, B. S., 1996, Plagioclase zonation styles in hornblende gabbro inclusions from Little Grass Mountain, Medicine Lake Volcano, California: implications for fractionation mechanisms and the formation of composition gaps: Contributions to Mineralogy and Petrology, v. 126, p. 121-136. Bullen, T. D. and Clynne, M. A., 1990, Trace element and isotopic constraints on magmatic evolution of Lassen volcanic center: J. of Geophysical Research, V. 95, No. B95, p. 19671-19691. Carney, E., 1998, Historic Soda Springs: oasis on the Oregon Trail: Traildust Publishing Company, Wayan, Idaho, 393 p. Christiansen, E. H., Burt, D. M., Sheridan, M. F., and Wilson, R. T., 1983, The petrogenesis of topaz rhyolites from the western United States: Contributions to Mineralogy and Petrology, v. 83, p. 16-30. Christiansen, E. H., Sheridan, M. F., and Burt, D. M., 1986, The geology and geochemistry of Cenozoic topaz rhyolites from the western United States: U. S. Geological Society Special Paper 205, 82 p. Christiansen, R. L. and Lipman, P. W., 1972, Cenozoic volcanism and plate-tectonic evolution of the Western United States. II late Cenozoic: Royal Society of London Philosophical Transactions A, v. 271, p. 249-284.
124
Collins, W. J., Beams, S. D., White, A. J. R., and Chappell, B. W., 1982, Nature and origin of A-type granites with particular reference to southeastern Australia: Contributions to Mineralogy and Petrology, v. 80, p. 180-200. Condie, K. D. and Hayslip, D. L., 1975, Young bimodal volcanism at Medicine Lake volcanic center, northern California: Geochimica Cosmocimica Acta, v. 39, p. 1165-1178. Dayvault, R. D., Rush, S. M., and Ludlum, J. R., 1984, Evaluation of uranium potential in a topaz-bearing rhyolite, China Hat dome, southeastern Idaho: in Reports on field investigations of uranium anomalies: Bendix Field Engineering Corporation open file report GJBX-1(84), p. II-1-II26. DeCelles, P. G., 2004, Late Jurassic to Eocene evolution of the Cordilleran Thrust belt and foreland basin system, western U.S.A.: American Journal of Science, v. 304, p. 105-168. Dorr, J. A., Spearing, D. R., Steidtmann, J. R., Wiltschko, D. V., and Craddock, J. P.,
1987, Hoback River Canyon, central western Wyoming, in Beus, S. S., ed., Rocky Mountain Section of the Geological Society of America: Centennial Field Guide, v. 2, p. 197-200.
Eckren, E. B., McIntyre, D. H., Bennett, E. H., and Marvin, R. F., 1982, Cenozoic stratigraphy of western Owyhee County, Idaho, in Bonnichsen, W. and Brekenridge, R. M., eds., Cenozoic Geology of Idaho: Bureau of Mines and Geology Bulletin 26, p. 215-235. Fauve, G., 2001, Origin of igneous rocks – the isotopic evidence: Springer-Verlay, New York, 496 p. Fiesinger, D. W., Perkins, W. D., and Puchy, B. J., 1982, Mineralogy and petrology of
Tertiary-Quaternary volcanic rocks in Caribou County, Idaho, in Bonnichsen, W. and Brekenridge, R. M., eds., Cenozoic Geology of Idaho: Bureau of Mines and Geology Bulletin 26, p. 465-488.
Finton, J. G., Dodie, J., and Leeman, W. P., 1991, Basic magmatism associated with Late
Cretaceous Extension in the Western United States: Compositional variations in space and time: J. of Geophysical Research, v. 96, no. B8, p. 13693-13711.
Ford, M. T. and McCurry, M., 2003, Petrology of Quaternary rhyolite domes of the bimodal Blackfoot Volcanic Field, SE Idaho: Geological Society of America Abstracts with Programs Vol. 35, No. 6. Ford, M. T., McCurry, M. and Chadwick, J., 2004, Genesis of Quaternary rhyolite domes of the eastern Snake River Plain and Blackfoot Volcanic Field, SE Idaho: Geological Society of America Abstracts with Programs Vol. 36, No. 4.
125
Gans, P. B., Mahood, G. A., and Schermer, E., 1989, Synextensional magmatism in the Basin and Range Province; A case study from the eastern Great Basin: U. S. Geological Survey Special Paper no. 233, 53p. Ganske, R., in progress, The geology and petrology of the East Butte area, Bingham County, Idaho: Idaho State University Masters Thesis. Gerlach D. C. and Grove, T. L., 1982, Petrology of Medicine Lake Highland volcanics: characterization of the end members of magma mixing: Contributions to Mineralogy and Petrology, v. 80, p. 147-159. Ghiorso, M. S. and Sack, R. O., (1995) Chemical mass transfer in magmatic processes IV: A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid-solid equilibria in magmatic systems at elevated temperatures and pressures: Contributions to Mineralogy and Petrology, v. 119, p. 197-212 Grove, T. L. and Donnelly-Nolan J. M., 1986, The evolution of young silicic lavas at Medicine Lake volcano, California: implications for the origin of compositional gaps in calc-alkaline series lavas: Contributions to Mineralogy and Petrology, v. 92, p. 281-302. Grunder, A. L., 1992, Two-stage contamination during crustal assimilation: isotopic evidence from volcanic rocks in eastern Nevada: Contributions to Mineralogy and Petrology, v. 112, p. 219-229. Hayden, K. P., 1992, The geology and petrology of Cedar Butte, Bingham County, Idaho: Idaho State University Masters Thesis, 104p. Hackett, W. R., and Smith, R. P., 1992, Quaternary volcanism, tectonics, and
sedimentation in the Idaho National Engineering Laboratory Area, in Wilson, J. R., ed., Field Guide to Geologic Excursions in Utah and Adjacent Areas of Nevada, Idaho, and Wyoming: Utah Geological Survey, p. 1-18.
Halliday, A. N., Davidson, J. P., Hildreth, W. and Holden, P., 1991, Modeling the petrogenesis of high Rb/Sr silicic magmas: Chemical Geology, v. 92, p. 107-114. Hammarstrom, J. M. and Zen, E., 1986, Aluminum in hornblende: An empirical igneous geobarometer: American Mineralogist, v. 71, p. 1297-1313. Heumann, A., 2004, Timescales of evolved magma generation at Blackfoot Lava Field, SE Idaho, USA: International Association of Volcanology and Chemistry of the Earth’s Interior (IAVCEI) General Assembly, Volcanism and its impacts on society, Pucon, Chile.
126
Heumann, A. and Davis, G. R., 1997, Isotopic and chemical evolution of the post-caldera rhyolitic system at Long Valley, California: J. of Petrology, v. 38, p. 1661-1678. Hildreth, W., Halliday, A. N. and Christiansen, R. L., 1991, Isotopic and chemical evidence concerning the genesis and contamination of basaltic and rhyolitic magma beneath the Yellowstone Plateau Volcanic Field: J. of Petrology, v. 32, p. 63-138. Honjo, N., 1990, Geology and stratigraphy of the Mount Bennett Hills and the orgin of the west-central Snake River Plain rhyolites: PhD thesis, Rice University, Houston, Tx, 259p. Hughes, S. S. and McCurry, M., 2002, Bulk major and trace element evidence for a time- space evolution of Snake River Plain rhyolites, Idaho, in Bonnochsen, B., White, C. M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 161-176. Hughes, S. S., Smith, R. P., Hackett, W. R., McCurry, M., Anderson, S. R., and Ferdock,
G. C., 1997, Bimodal magmatism, basaltic volcanic styles, tectonics, and geomorphic processes of the eastern Snake River Plain, Idaho, in Link, P. K. and Kowallis, J, eds., Proterozoic to Recent Stratigraphy, Tectonics, and Volcanology, Utah, Nevada, Southern Idaho, and Central Mexico: Brigham Young University Geology Studies, v. 42, part 1, p. 423-457.
Hughes, S. S., McCurry, M. and Geist, D. J., 2002A, Geochemical correlations and implications for the magmatic evolution of basalt flow groups at the Idaho National Engineering and Environmental Laboratory, in Link, P. K. and Mink, L. L., eds., Geology, Hydrogeology, and Environmental Remediation: Idaho National Engineering and Environmental Laboratory, Eastern Snake River Plain, Idaho: Geological Society of America Special Paper 353, p. 151-173. Hughes, S. S., Wetmore, P. H. and Casper, J. l., 2002B, Evolution of Quaternary tholeiitic basalt eruptive centers on the eastern Snake River Plain, Idaho, in Bonnochsen, B., White, C. M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 363-385. Hutchison, C. S., 1974, Laboratory Handbook of Petrographic Techniques: John Wiley and Sons, New York, 527p. Johnson, C. M. and Lipman, P. W., 1988, Origin of metaluminous and alkaline volcanic rocks of the Latir Volcanic Field, northern Rio Grande Rift, New Mexico: Contributions to Mineralogy and Petrology, v. 100, p. 107-128.
127
Johnson, M. C. and Rutherford, M. J., 1989, Experimental calibration of an aluminum-in- hornblende geobarometer applicable to calc-alkaline rocks: Geology, v. 17, p. 837-841. Kellogg, K. S. and Marvin, R. F., 1988, New potassium-argon ages, geochemistry and tectonic setting of upper Cenozoic volcanic rocks near Blackfoot, Idaho: U. S. Geological Survey Bulletin 1806, 19p. Kuntz, M. A. and Dalrymple, G. B., 1979, Geology, geochronology, and potential volcanic hazards in the Lava Ridge – Hell’s Half Acre area, eastern Snake River Plain, Idaho: U. S. Geological Survey Open-file Report 79-1675, 70p. Kuntz, M. A., Covington, H. R., and Schorr, L. J., 1992, An overview of basaltic
volcanism of the eastern Snake River Plain, Idaho, in Link, P. K., Kuntz, M. A., and Platt, L. B., eds., Regional Geology of Eastern Idaho and Western Wyoming: Geological Society of America Memoir 179, p. 227-267.
Le Bas, M. J., Le Maitre, R. W., Streckeisen, A., and Zanettin, B., 1986, A chemical classification of volcanic rocks based on the total alkali-silica diagram: J. of Petrology, v. 27, p. 745-750. Leeman, W. P., 1982a, Evolved and hybrid lavas from the Snake River Plain, Idaho, in
Bonnichsen, W and Brekenridge, R. M., eds., Cenozoic Geology of Idaho: Bureau of Mines and Geology Bulletin 26, p. 193-202. Leeman, W. P., 1982b, Rhyolites of the Snake River Plain-Yellowstone Plateau province, Idaho and Wyoming: a summary of petrogenetic models, in Bonnichsen, W and Brekenridge, R. M., eds., Cenozoic Geology of Idaho: Bureau of Mines and Geology Bulletin 26, p. 203-212. Leeman, W. P. and Gettings, M. E., 1977, Holocene rhyolite in southeast Idaho and
geothermal potential (abs.): American Geophysical Union Transactions – EOS, v. 58, no. 12, p. 1249.
Leeman, W. P., Menzies, M. A., Matty, D. J. and Embree, G. F., 1985, Strontium, neodymium and led isotopic compositions of deep crustal xenoliths from the Snake River Plain: evidence for Archean basement: Earth and Planetary Science Letters, v. 75, p. 354-368. Le Maitre, R. W., Streckeisen, A., Zanettin, B., Le Bas, M. J., Bonin, B., Bateman, P., Belliene, G., Dudek, A., Efremova, S., Keller, J., Lameyre, J., Sabine, P. A., Schmid, R., Sorensen, H., Woolley, A. R., 2002, Igneous rocks a classification and glossary of terms: Cambridge University Press, England, p. 236 Link, P. K., 1982, Idaho-Wyoming Thrust Belt: Regional Geology: Northwest Geology, v. 11, p. 1-12.
128
Luedke, R. G. and Smith, R. L., 1983, Map showing distribution, composition, and age of late Cenozoic volcanic centers in Idaho, western Montana, west-central South Dakota, and northwestern Wyoming: U. S. Geological Survey Miscellaneous Investigations Series Map I-1091-E, scale 1:100,000. Mabey, D. R. and Oriel, S. S., 1970, Gravity and magnetic anomalies in the Soda Springs region, southeastern Idaho: U. S. Geological Survey Professional Paper 646-E, 15p. Macdonald, G. A., 1972, Volcanoes: Prentice-Hall, New Jersey, 510p. Mansfield, G. R., 1927, Geography, geology and mineral resources of part of
southeastern Idaho: U. S. Geological Survey Professional Paper 152, 453 p. Mason, G. H., 1985, The mineralogy and textures of the Costal Batholith, Peru, in Pitcher, W. S., Atherton, M. P., Cobbing, E. J. and Beckinsale, R. D., eds. Magmatism at a plate edge: Blackie, London, p. 156-66 McCurry, M. and Ganske, R., 2005, Genesis of Quaternary high-K “A-type” rhyolites along part of the Yellowstone – Snake River Plain hot spot track, in 15th Annual V. M. Goldschmidt Conference Programme, p. 55. McCurry, M. and Morse, L. H., 2003, Petrology and geochemistry of basaltic andesite to
rhyolite rocks from borehole CH-1: Implications for Quaternary Eastern Snake River Plain-style intermediate-composition volcanism: Geological Society of America Abstracts with Programs, v. 35, no. 6.
McCurry, M., Hackett, W. R., and Hayden, K., 1999, Cedar Butte and cogenetic
Quaternary rhyolite domes of the Eastern Snake River Plain, in Hughes, S. S., and Thackray, G. D., eds., Guidebook to the Geology of Eastern Idaho: Idaho Museum of Natural History, p. 169-179.
McCurry, M., Rodgers, D. W., Hughes, S. S., Price, K., Scarberry, K., Ford, M., 2002, Mantle-Derived Mass Transfer to Continental Crust along the Yellowstone Hotspot Track: Geological Society of America Abstracts with Programs V. 34, No. 5. Mitchell, V. E. and Bennett, E. H., 1979, Geologic map of the Driggs Quadrangle, Idaho: Idaho Bureau of Mines and Geology Map GM-6.
Morgan, L. A., Doherty, D. J. and Leeman, W. P., 1984, Ignimbrites of the eastern Snake River Plain: evidence for major caldera forming eruptions: J. of Geophysical Research, v. 89, No. B10, p. 8665-8678.
129
Morse, L. H., 2002, Basalt alteration and authigenic mineralization near the effective base of the Snake River Plain Aquifer at the Idaho National Engineering and Environmental Laboratory, Idaho: Idaho State University Masters Thesis, 182p. Mueller, P. A., Heatherington, A. L., Kelly, D. M., Wooden, J. L., and Mogk, D. W.,
2002, Paleoproterozoic crust within the Great Falls tectonic zone: Implications for the assembly of southern Laurentia: Geology, v. 30, no. 2, p. 127-130.
O’Brien, H. E., Irving, A. J., McCallum, I. S., Thirlwall, M. F., 1995, Strontium, neodymium, and lead isotope evidence for the interaction of post-subduction asthenoshperic potassic mafic magmas of the Highwood Mountains, Montana, USA, with ancient Wyoming craton lithospheric mantle: Geochimica et Cosmochimica Acta, v. 59, p. 4539-4556. Oriel, S. S., 1968, Preliminary geologicmap of the Bancroft quadrangle, Caribou and Bannock Counties, Idaho: U. S. Geological Survey open file map. Oriel, S. S. and Platt, L. B., 1980, Geologic map of the Preston 1o X 2o quadrangle,
southeastern Idaho and western Wyoming: U. S. Geological Survey Miscellaneous Investigation Series Map I-1127, scale 1:250,000.
Orozco-Esquivel, M. T., Nieto-Samaniego, A. F. and Alaniz-Alvarez, S. A., 2002, Origin of rhyolitic lavas in the Mesa Central, Mexico, by crustal melting related to extension: J. of Volcanology and Geothermal Research, V. 118, p. 37-56. Parker, J. L., 1996, Physical volcanology and geochemistry of the tuff of Wooden Shoe Butte, Cassia Mountains, Idaho: Idaho State University Masters Thesis, 105p. Pickett, K. E., 2004, Physical volcanology, petrography and geochemistry of basalts in the bimodal Blackfoot Volcanic Field, Southeastern Idaho: Idaho State University Masters Thesis, 92 p. Pierce, K. L., and Morgan, L. A., 1992, The track of the Yellowstone hot spot: volcanism, faulting, and uplift, in Link, P. K., Kuntz, M. A., and Platt, L. B., eds., Regional Geology of Eastern Idaho and Western Wyoming: Geological Society of America Memoir 179, p. 1-53. Pierce, K. L., Fosberg, M. A., Scott, W. E., Lewis, G. C., and Colman, S. M., 1982, Loess
deposits of southeastern Idaho: Age and correlation of the upper two loess units, in Bonnichsen, W and Brekenridge, R. M., eds., Cenozoic Geology of Idaho: Bureau of Mines and Geology Bulletin 26, p. 717-725.
Peng, X., and Humphreys, E. D., 1998, Crustal velocity structure across the eastern
Snake River Plain and Yellowstone swell: J. of Geophysical Research, v. 103, no. B4, p. 7171-7186.
130
Ren, M., 2004, Partitioning of Sr, Ba, Rb, Y, and LREE between alkali feldspar and peraluminous silicic magma: American Mineralogist, v. 89, p. 1290-1303. Ren, M., Parker, D. F. and White, J. C., 2003, Partitioning of Sr, Ba, Rb, Y, and LREE between plagioclase and peraluminous silicic magma: American Mineralogist, v. 88, p. 1091-1103. Rollinson, H., 1998, Using geochemical data: Evaluation, presentation, interpretation: Longman Group, Essex, England, 352p. Royse, F. C., Warner, M. A., and Reese, D. L., 1975, Thrust belt structural geometry and related stratigraphic prolems, Wyoming-Idaho-northern Utah, in Bolyand, D. W., ed., Deep Drilling Frontiers of the Central Rocky Mountains: Rocky Mountain Association of Geologists, Symposium, p. 41-43. Scheidegger, K. F., Federman, A. N. and Tallman, A. M., 1982, Compositional Heterogeneity of tephras from the 1980 eruptions of Mount St. Helens: J. of Geophysical Research, v. 87, No. B87, p. 10861-10881. Sparlin, M. A., Braile, L. W. and Smith, R. B., 1982, Crustal structure of the eastern Snake River Plane determined from ray trace modeling of seismic refraction data: J. of Geophysical Research, v. 87, p. 2619-2633. Spear, D. B., 1979, The geology and volcanic history of the Big Southern Butte – East Butte area, eastern Snake River Plain, Idaho: State University of New York at Buffalo PhD Thesis, 136p. Spear, D. B. and King, J. S., 1982, The geology of Big Southern Butte, Idaho, in
Bonnichsen, W and Brekenridge, R. M., eds., Cenozoic Geology of Idaho: Bureau of Mines and Geology Bulletin 26, p. 395-403. Spera, F. J. and Bohrson, W. A., 2001, Energy-constrained open-system magmatic processes I: General model and energy-constrained assimilation and fractional crystallization (EC-AFC) formulation: J. of Petrology, v. 42, no.5, p. 999-1018. Spera, F. J. and Bohrson, W. A., 2002, Energy-constrained open-system magmatic processes 3: Energy-constrained recharge, assimilation and fractional crystallization (EC-RAFC): Geochemisty, Geophysics, Geosystems (electronic journal, v. 3, no. 12, 20p. Sun, S. S., 1980, Lead isotopic study of young volcanic rocks from mid-ocean ridges, ocean islands and island arcs: Philosophical Transaction of the Royal Society of London, v. A 297, p. 409-445. Taylor, S. R. and McLennan, S. M., 1985, The continental crust: Its composition and evolution: Blackwell, Oxford, 312p.
131
Vance, J. A., 1962, Zoning in igneous plagioclase: Normal and oscillatory zoning: American Journal of Science, v. 260, p. 746-760. Watkins, A., 1998, Geochemistry, petrology and stratigraphy of the tuff of Steer Basin, Twin Falls and Cassia Counties, Idaho: Idaho State University Masters Thesis, 177p. Whitehead, R. L., 1992, Geohydrologic framework of the Snake River Plain regional aquifer system, Idaho and Eastern Oregon: U. S. Geological Survey Professional Paper 1408-B, 32p. Wilson, M., 1989, Igneous petrogenesis: A global tectonic approach: Unwin Hyman, Boston, 466p. Wiltschko, D. V. and Dorr, J. A., 1983, Timing of deformation in overthrust belt and
foreland of Idaho, Wyoming, and Utah: American Association of Petroleum Geologists Bulletin, v. 63, issue 8, p. 1304-1322.
Wright, K., 1998, Geochemistry and petrology of the tuff of McMullen Creek, Cassia County, Idaho: Idaho State University Masters Thesis, 172p. Yavuz, F., 1996, AMPHCAL: A QUICKBASIC program for determining the amphibole name from electron microprobe analysis using the IMA rules: Computers in Geoscience, v. 22, p. 101-107. Zindler, A. and Hart, S., 1986, Chemical geodynamics: Annual Reviews of Earth and Planetary Science, v. 14, p. 493-571.
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Appendix 1 Electron Microprobe Methods: Crystals LiF PET TAP LDE1 LDE2 LDEB Specrometer 1: (1) PET (1) TAP (1) LDE1 (1) LDEB Specrometer 2: (2) LiF (2) PET Specrometer 3: (3) TAP (3) LDE2 Specrometer 4: (4) LiF (4) PET Specrometer 5: (5) LDE1 (5) LDE2
Feldspar crystal count time probe diam STD (seconds) (microns) SiO2 Albite (1) TAP 30 10 Al2O3 Kyanite (1) TAP 30 10 FeO* Almandine (2) LiF 30 10 MgO Diopside (3) TAP 30 10 CaO Wollastonite (2) PET 30 10 Na2O Albite (3) TAP 30 10 K2O Orthoclase (4) PET 30 10 SrO Celestite (4) PET 30 10
Melt Inclusions* crystal count time probe diam STD (seconds) (microns) SiO2 Diopside (1) TAP 30 10 TiO2 Rutile (4) LiF 30 10 Al2O3 Almandine (3) TAP 30 10 FeO* Hematite (4) LiF 30 10 MnO Spessartite (2) LiF 30 10 MgO Diopside (3) TAP 30 10 CaO Diopside (1) PET 30 10 Na2O Albite (3) TAP 30 10 K2O Orthoclase (4) PET 40 20 P2O5 Apatite (1) TAP 40 20 S Anhydrite (2) PET 40 20 Cl KCl (2) PET 40 20 F Fluorite (5) LDE1 60 20 * also for biotite and hornblende
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Appendix 1: concluded Olivine crystal count time probe diam STD (seconds) (microns) SiO2 Diopside (1) TAP 30 10 FeO* Almandine (2) LiF 30 10 MnO Spessartite (2) LiF 30 10 MgO Olivine (3) TAP 30 10 CaO Diopside (4) PET 30 10 NiO Ni-metal (4) LiF 30 10
Pyroxene crystal count time probe diam STD (seconds) (microns) SiO2 Diopside (1) TAP 30 10 TiO2 Rutile (2) LiF 30 10 Al2O3 Almandine (1) TAP 30 10 Cr2O3 Cr-metal (2) LiF 30 10 FeO* Almandine (2) LiF 30 10 NiO Ni-metal (4) LiF 30 10 MnO Spessartine (4) LiF 30 10 MgO Diopside (3) TAP 30 10 CaO Diopside (4) PET 30 10 Na2O Albite (3) TAP 30 10
Oxides crystal count time probe diam STD (seconds) (microns) SiO2 Almandine (3) TAP 30 10 TiO2 Rutile (2) LiF 30 10 Al2O3 Almandine (1) TAP 30 10 Cr2O3 Cr-metal (2) LiF 30 10 FeO* Hematite (4) LiF 30 10 MnO Spessartine (4) LiF 30 10 NiO Ni-metal (2) LiF 30 10 ZnO Sphalerite (4) LiF 30 10 MgO Periclase (3) TAP 30 10 CaO Diopside (1) PET 30 10