AN OXYGEN ISOTOPE, FLUID INCLUSION, AND MINERALOGY STUDY OF
THE ANCIENT HYDROTHERMAL ALTERATION
IN THE GRAND CANYON OF THE YELLOWSTONE RIVER,
YELLOWSTONE NATIONAL PARK, WYOMING
By
Allison R. Phillips
A thesis submitted in partial fulfillment of the requirements for the degree of
MASTER OF SCIENCE IN GEOLOGY
WASHINGTON STATE UNIVERSITY School of Earth and Environmental Science
MAY 2010
ii
To the Faculty of Washington State University:
The members of the Committee appointed to examine the thesis of ALLISON R.
PHILLIPS find it satisfactory and recommend that it be accepted.
________________________________________ Dr. Peter B. Larson, Chair ________________________________________ Dr. David A. John ________________________________________ Dr. Franklin F. Foit
iii
ACKNOWLEDGMENT
This research was funded by: NSF Grant #EAR-0609475 The Roger V. LeClerc II Memorial Fellowship from Washington State University A Society of Economic Geologists Foundation Grant A very special thank you to those who helped in the field: Brian Pauley Allen Andersen Chad Pritchard Jennifer Manion Dr. David Cole Dr. Mike Cosca Dr. Todd Feeley
And thank you to the following for laboratory assistance: Dr. Jean Cline Haraldo Lledo Dr. Charles Knaack The GeoAnalytical Lab staff at Washington State University
iv
AN OXYGEN ISOTOPE, FLUID INCLUSION, AND MINERALOGY STUDY OF
THE ANCIENT HYDROTHERMAL ALTERATION
IN THE GRAND CANYON OF THE YELLOWSTONE RIVER,
YELLOWSTONE NATIONAL PARK, WYOMING
Abstract
By Allison R. Phillips, M.S. Washington State University
May 2010
Chair: Dr. Peter B. Larson
The Grand Canyon of the Yellowstone River displays regions of pervasive
hydrothermal alteration, formed in the shallow parts of an ancient hydrothermal system. The
altered protolith, the 480 ka post-collapse Tuff of Sulphur Creek, is a high silica, low δ18O
rhyolite tuff. The localized alteration is controlled by an underlying caldera ring fault from the
640 ka caldera collapse. Incision of the canyon exposed 350 vertical meters of altered rock in
the Sevenmile Hole vicinity. The alteration shows evidence of both acid-sulfate and neutral-pH
fluid chemistry. There is a kaolinite to illite transition at a depth of ~100 m below the current
canyon rim elevation. Boiling of groundwater at depth creates a neutral-pH environment (illite is
precipitated), and condensation of rising sulfuric vapors above 100 m creates an acid-sulfate
environment (kaolinite is precipitated). At ~50 m, the silica phase precipitated changes from
quartz at depth to opal at shallow elevations. Mineralization in the canyon can be divided into
seven alteration assemblages based on vertical transitions and proximity to a high heat and fluid
v
upwelling zone. Fluid inclusion homogenization temperatures in quartz samples range from
160° to 350°C. Homogenization temperatures generally increase with depth and are higher than
reference boiling point curve temperatures. Additional hydrostatic pressure from a glacial ice
sheet with a thickness of <460 m, could account for anomalously high temperatures at shallow
depths. Freezing of inclusions yield salinities of 0.35-0.71 wt % NaCl eq. δ18O values of
magmatic and hydrothermal quartz were measured for 50 samples using laser fluorination
techniques. Values ranged from -5.7‰ to 1.3‰, all of which are more depleted than magmatic
quartz in the fresh TSC at 1.7‰. Low salinities and negative δ18O values indicate a dominantly
meteoric water source for ancient hydrothermal fluids. The δ18O values of quartz are controlled
by the intensity of alteration and the water-rock ratio. The high heat and fluid upwelling zone
had the most evolved water-rock ratios and lowest δ18O values, whereas less altered zones had
higher δ18O values. Calculated fractionation between quartz and water yields δ18O values
ranging from -10.8 to -19.6 ‰ for ancient hydrothermal waters.
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TABLE OF CONTENTS
Page
ACKNOWLEDGEMENTS…………………………….……………………………………….iii
ABSTRACT…………………………………………….……………………………………….iv
LIST OF FIGURES…………………………………….…………..……………….….………..ix
LIST OF TABLES…………………………………………………………………...….……….x
CHAPTER 1: INTRODUCTION...…………………….………………….………….…………1
1.1 Regional Geologic Setting………….……………….…………………….…………1
1.2 The Yellowstone Caldera….……………………..………………….…….…………5
1.3 The Tuff of Sulphur Creek…….………………..………………………..….……….8
1.4 Glacial History of the Yellowstone Plateau since TSC……………….…………….10
1.5 Hydrothermal Alteration………………………..…………………….……………..11
1.6 Stable Isotopes Ratios: Oxygen…………………...………………….….………….14
1.7 Importance of Research…………………………………..………….………….…..16
CHAPTER 2: SAMPLING AND ANALYTICAL METHODS.………………….…………….19
2.1 Sampling…………….……………………..………………………….…………….19
2.2 PIMA………………….…………………………...………………………………..19
2.3 XRD…………………….…………………..……………………………….………19
2.4 SEM……………………….……………………..……………………….…………20
2.5 Radiocarbon Dating…….………………..……………………………….…………20
2.6 Fluid Inclusions…….…………………..…………………………………...………20
2.7 Oxygen Isotopes……………………..………………………………….…………..21
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CHAPTER 3: STRUCTURE AND HYDROTHERMAL MINERALOGY…………….…..….22
3.1 Caldera Ring Fault Control on Hydrothermal Activity……………………….…….22
3.2 Surficial Expression of Hydrothermal Activity……………………………….…….25
3.3 Alteration Mineralogy and Textures at Depth………………………………….…...32
3.4 Alteration Mineral Assemblages…………………………………………….………36
3.5 Model of Hydrothermal Zoning at Sevenmile Hole………………………...………50
CHAPTER 4: FLUID INCLUSIONS………………….….…………………………….……...54
4.1 Homogenization Temperatures……………………………………………….……..54
4.2 Freezing Point Depressions/Salinity………………………………………………...59
4.3 Homogenization Temperatures Compared to the Boiling Point Curve......................59
CHAPTER 5: STABLE ISOTOPES RATIOS: OXYGEN……………………...……….….….62
5.1 Fractionation and Exchange of Oxygen……………………………………………..62
5.2 δ18O Values of Magmatic and Hydrothermal Quartz……………………...………..63
5.3 δ18O Values of Hydrothermal Quartz Habits………………………………………..63
5.4 Calculation of δ18O Values for the Ancient Hydrothermal Altering Fluid………….67
CHAPTER 6: TIMING AND PARAGENESIS ON THE HYDROTHERMAL SYSTEM...….73
6.1 Incision of the Grand Canyon of the Yellowstone River..………………………….73
6.2 Age of Alteration……………………………...…………………………………….74
6.3 Higher than Reference Boiling Point Curve Temperatures…..…….……………….74
6.4 Cause of the Additional Hydrostatic Head………………………………………….75
6.5 Paragenesis: Structural Cause and Timing of the Alteration..……………………...78
CHAPTER 7: SUMMARY OF CONCLUSIONS…………….………………………..………84
REFERENCES…………………………………….……………….……………………………88
viii
APPENDICES
A. FIELD SITE AND SAMPLE LOCATIONS MAP…..…………….…….…..……..98
B. FIELD SITE LOCATION DESCRIPTIONS..………………………………..…...100
C. SAMPLING LOCATIONS AND ALTERATION MINERALOGY TABLE.....…109
D. X-RAY DIFFRACTION PATTERNS AND PIMA SPECTRA..…...……………..114
E. STABLE OXYGEN ISOTOPE RATIOS TABLE..……………..………….….….121
F. FLUID INCLUSION ASSEMBLAGES TABLES…...………...……….…………125
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LIST OF FIGURES
Page
1. Map of Western US and Volcanic Centers of the Snake River Plain………….……..…..….2
2. Map of Three Nested Yellowstone Calderas………………………………….….………….6
3. Map of Yellowstone Caldera and Hydrothermal Areas………………………….………….7
4. δ18O of Yellowstone Rhyolites……………………………………………………..….…….9
5. Map of the Grand Canyon of the Yellowstone River…………………………….……..….12
6. Meteoric Water Line and δ18O Shift for Waters in Yellowstone National Park…….…...…17
7. Map of Hydrothermal Areas Along Caldera Margin and Grand Canyon…………….…….23
8. Photograph of Sevenmile Hole Field Site………………………………………..…………24
9. Google Earth Image of Sevenmile Hole Field Site….…………………………..………….26
10. Description and Photos of Sinter Field Locations…………………………….…..……….27
11. Radiocarbon Analysis of Wood Sample in Sinter……………………………….. ……….28
12. Schematic of Grand Canyon of the Yellowstone River Incision……………..………..…..31
13. Photographs of Pervasively Altered Textures………………………………..…..…..…….34
14. Photographs of Less Intense Alteration Textures……………………………..…...………35
15. Photographs of Hydrothermal Sulfides.……………………………………...…………….35
16. Cross Section of Grand Canyon at Sevenmile Hole……………………………...………...37
17. Photographs of Two Crystal Morphologies of Alunite…….………………….……….…..42
18. Photographs of Neutral-pH Alteration Mineralogy……………………………..….………44
19. Photographs of Chalcedony Veins….…………...………………………………………….46
20. Photographs of Silicified Ridges from Sevenmile Hole……………………………………48
21. Alteration Assemblage Map……………………….………………………...……………..49
x
22. Theoretical Cross Section of Hydrothermal System…..……………….………..….………53
23. Photograph of Fluid Inclusion Assemblage………………………………………..……….55
24. Frequency Histogram for Primary Fluid Inclusions……………….………...….………57-58
25. Haas Boiling Point Curve and and Homogenization Temperature Plot of FIA…….…....…61
26. δ18O Values for Unaltered and Altered Magmatic, and Hydrothermal Quartz………….….64
27. δ18O Values for Different Quartz Habits……………………………………………………66
28. δ18O Values Compared to Alteration Assemblages…………………………………………68
29. Predicted and Actual Water δD and δ18O Values for Yellowstone…………………………70
30. Water-Rock Ratio Evolution…………………………………………………………….….71
31. Effect of Glaciation on Boiling Point Curve Temperature………………………….………77
32. Paragenesis Cross Section of Hydrothermal System in Sevenmile Hole………...…………83
LIST OF TABLES
Page
1. Hydrothermal Mineral Description………………..…………….…………….…………….33
2. Hydrothermal Mineral Assemblages……………………………….…………….………37-38
3. Sampling Locations and Alteration Mineralogy…………………………………………….110
4. Stable Oxygen Isotope Ratios……………………………………………………………….122
5. Fluid Inclusion Assemblages………………………………………………………………..126
1
Chapter 1: INTRODUTION AND GEOLOGIC HISTORY
The Yellowstone Caldera hosts a vigorous hydrothermal system manifest by the
remarkable surficial thermal features that include hot springs, geysers, mud pots, and fumaroles,
and attracts multitudes of people to Yellowstone National Park. These features are indicators of
the vast and complex circulation of fluids at depth below the surface of the Yellowstone Caldera
(Fournier, 1989). The Yellowstone volcanic center cycled through three periods of explosive
caldera collapsing eruptions, and intermittent minor eruptions and hydrothermal activity
occurred during the lulls in major volcanic activity (Smith and Bailey, 1968). Heat from shallow
magmatism and faulting related to volcanic activity has provided the right environment for an
extensive hydrothermal system. The focus of this study is the hydrothermal alteration of a post-
third collapse rhyolitic unit, the Tuff of Sulphur Creek, at Sevenmile Hole in the Grand Canyon
of the Yellowstone River. Alteration mineral zoning at Sevenmile Hole has been mapped and
temperature and fluid chemistry of the hydrothermal system have been estimated using mineral
phases and assemblages, fluid inclusion analyses, and oxygen isotope ratios. These data can be
compared to active systems in Yellowstone National Park and other hydrothermal systems
around the world.
1.1 Regional Geologic Setting
The Snake River Plain (SRP)/Yellowstone chain of time transgressive volcanic centers
has a long history of volcanism beginning in northern Nevada and southeastern Oregon about
16.7 Ma (Rytuba and McKee, 1984; Pierce and Morgan, 1992). The track stretches across
southern Idaho, and into northwestern Wyoming, creating a topographic low referred to as the
SRP volcanic province (Christiansen and Yeats, 1992) (Figure 1). Conflicting models for the
development of SRP volcanism include southwest migration of the North American Plate over a
2
Figure 1: (a) Map of western US. Red box outlines location of Snake River Plain. (b) Map of Snake River Plain/Yellowstone volcanic province showing major volcanic centers and their ages.
(a)
(b)
3
mantle plume or hot spot at a rate of approximately 2 cm/year (Morgan, 1972; Christiansen,
2001; Yuan and Dueker, 2005), volcanism following a pre-existing lithospheric structure in the
continental craton (Iyer and Healy, 1972), or volcanism beginning concordantly with the onset of
Basin-Range extension that propagated eastward along the SRP and westward forming the High
Lava Plains in Oregon (Jordan et al., 2004; Christiansen et al., 2002), along the regional tectonic
boundary between the Basin-Range province to the south and older intrusive complexes to the
north that include the Idaho and Boulder Batholiths (Christiansen and McKee, 1978;
Christiansen and Yeats, 1992). Either the head of a mantle plume is responsible for the
stretching and thinning of the continental cratonic lithosphere, or the shearing between two
terrains created the proper environment for the weakening and thinning of the continental
lithosphere (Christiansen and McKee, 1978). This extension allowed basaltic magma from the
mantle to rise into shallow levels of the crust and produced large amounts of rhyolitic magma
from crustal partial melting (Leeman et al., 2009). Many major eruptions along the SRP path
have been large; volumes for individual ignimbrite units of the SRP generally vary from 5 to 500
km3 (Boroughs et al., 2005).
The progression of volcanic centers begins with the western-most McDermott caldera at
16.7 Ma, and continues to the east with the Owyhee-Humboldt center at 14.5 Ma, the Bruneau-
Jarbidge center at 12.5 Ma, followed by the Twin Falls center at 10.8 Ma, the Picabo center at
10.2 Ma, the Heise caldera at 6.6 Ma, and finally the currently active Yellowstone caldera
beginning at 2.1 Ma (Figure 1). Each center of volcanic activity has an approximate lifespan of
2 to 3 Ma (Pierce and Morgan, 1992). The youngest and eastern-most Heise and Yellowstone
volcanic centers exhibit nested caldera features (Bindeman et al., 2007). Yellowstone is a classic
nested caldera, composed of three superimposed calderas (Christiansen, 2001), while the older
4
eruptions along the SRP were more spatially sporadic, and vents and/or well defined calderas are
hidden due to burial, hence they are referred to as volcanic centers and not calderas.
The SRP volcanism is unique in many respects such that it is perhaps the type locality for
a new classification of intracontintental bimodal volcanism called Snake River (SR) type
volcanism by Branney et al. (2008). The eruptive units associated with this volcanic province
are generally bimodal basalt and high temperature, water-poor rhyolite pyroclastics and flows;
the rhyolites erupted from calderas, and the basalts erupted contemporaneously in extra-caldera
settings or postdate major volcanic activity.
Another characteristic feature of this volcanic province is the low δ18O nature of the
rhyolitic lavas and ignimbrites. The typical range for “normal” igneous rocks was found by
Taylor (1968) to be 5 to 10‰, with typical siliceous or felsic (granitic and rhyolite) magmatic
values ranging from ~7 to 10‰. Many of the SRP rhyolites are significantly depleted in δ18O.
The Cougar Point Tuff III from the Bruneau-Jarbidge eruptive center has the lowest magmatic
δ18O values at 3.8‰ (Boroughs et al., 2005). The evolution of the centers generally progress
from higher δ18O values in the earlier eruptions to later lower δ18O values the longer a volcanic
center is active in one location (Bindemann et al., 2007). Although the cause of the δ18O
depletion in these units is not fully understood, the most likely process involves the melting and
assimilation of meteoric-hydrothermally altered continental crust (Larson and Taylor, 1986).
Thus, (1) the generation of low δ18O magma must involve meteoric water, the only large
negative δ18O reservoir on Earth. This presents a problem because a magma under lithostatic
pressure cannot simply absorb meteoric water under hydrostatic pressure because of the huge
pressure difference from hydrostatic to much greater lithostatic. (2) There is also a mass balance
difficulty; the amount of water needed to exchange with magma in order to significantly lower
5
δ18O is several times the saturation level of water capable of being dissolved in a magma
(Hildreth et al., 1984). (3) In order to melt meteoric-hydrothermally altered rocks that can only
be generated at shallow crustal levels, there must be a mechanism to get that altered material to a
depth where melting is possible. (4) A large δ18O depleted volume of altered material is required
to produce the large erupted volumes (5 to >2,000 km3) of low δ18O rhyolites of the SRP and
Yellowstone.
1.2 The Yellowstone Caldera
There have been three major caldera forming eruptions at Yellowstone in the past 2.1 Ma.
The oldest and largest of the Yellowstone eruptions is the Huckleberry Ridge Tuff which erupted
at 2.1 Ma. The tuff has an estimated volume of 2,450 km3. The next major eruption occurred at
1.3 Ma, the Mesa Falls Tuff. This was the smallest of the three Yellowstone caldera forming
eruptions with a volume of 280 km3. The latest eruption produced the Lava Creek Tuff at 0.64
Ma. The Lava Creek Tuff has a volume of 1,000 km3 and established the current Yellowstone
caldera (Christiansen, 2001) (Figure 2).
Since the Lava Creek eruption, there has been an extensive hydrothermal system
associated with the Yellowstone caldera (Fournier, 1989; Christiansen, 2001). Active and very
recently active hydrothermal features are shown in Figure 3. Most of this hydrothermal activity
is found either around the 0.64 Ma caldera margin, around the two resurgent domes found within
the caldera (the Mallard Lake to the southwest and the Sour Creek to the northeast), or along a
north-striking fault system referred to as the Mammoth-Norris corridor (Christiansen, 1975;
Fournier, 1989).
All of the major caldera forming eruptions and the smaller intermittent intracaldera
rhyolite eruptions (with the exception of the first caldera collapse rhyolite, the Huckleberry
6
Figure 2: Map of Yellowstone National Park showing locations of three nested calderas that are collectively referred to as the Yellowstone Calderas. Modified from Christiansen (2001).
7
Figure 3: Map of the 640 ka Yellowstone caldera showing post-third collapse rhyolites, fault trends, and hydrothermal areas. The thick black lines represent the topographic rim of the youngest caldera. The resurgent domes, Mallard Lake to the southwest and Sour Creek to the northeast, are circled with associated faults. The north-trending Mammoth-Norris corridor fault zone is located to the north of the caldera. Areas of active or recently active hydrothermal alteration are shown in red and post-collapse rhyolites are shown in gray. Box outlines area of Figure 5, which includes the Grand Canyon of the Yellowstone River and Eocene Mt. Washburn to the north of the Sour Creek dome. Modified from Christiansen (2001), Morgan et al. (2007), and Larson et al., (2009).
8
Ridge Tuff) have low primary δ18O values (Hildreth et al., 1984). The temporal pattern in low
δ18O values of the Yellowstone rhyolites is somewhat comparable to the SRP rhyolites from each
volcanic center (Bindemann and Valley, 2001). In Yellowstone, there is a trend of minimal δ18O
depletion for the major collapse/eruptive units followed by more δ18O depleted post-collapse
rhyolite units of much smaller volume (Hildreth et al., 1984) (Figure 4). The best explanation
for this phenomenon is that caldera collapse allowed hydrothermally altered, δ18O depleted crust
to descend to melting depths; this explanation also makes mixing more plausible due to the
smaller volumes (Taylor, 1980; Bindeman and Valley, 2001). In Yellowstone, the large-volume
Lava Creek Tuff at 0.64 Ma has magmatic δ18O values of 5.5‰, and the resulting third caldera
collapse was followed by a period of smaller intracaldera rhyolite units, which have some of the
lowest δ18O values of the entire SRP province (as low as 0‰) (Bindeman et al., 2007). These
intracaldera rhyolitic units are referred to as the Plateau Rhyolites (Christiansen and Blank,
1975) and are divided into three members: the older Upper Basin Member, the Mallard Lake
Member, and the younger Central Plateau Member, which is the youngest erupted material in the
park with volcanism ending about 70 ka. The Upper Basin Member rhyolites, include the 486 ±
42 ka Dunraven Road Flow, the 484 ± 15 ka Canyon Flow, the 479 ± 10 ka Tuff of Sulphur
Creek (which is the sole protolith of this alteration study), and the similar aged Tuff of Uncle
Tom’s Trail which pre-dates other UBM eruptions (Gansecki et al., 1996; Christiansen, 2001).
1.3 The Tuff of Sulphur Creek
The Tuff of Sulphur Creek is located in the northeast part of the caldera and erupted after
a period of resurgence that created the Sour Creek dome (Christiansen, 2001). This unit is a
fallout tuff that grades into a densely welded section approximately 300 meters thick. Vertical
exposure of the entire sequence is present in the Grand Canyon of the Yellowstone River due
9
Figure 4: Graph displaying δ18O values of primary quartz phenocrysts from major caldera forming eruptions and subsequent post-collapse units. Notice post-collapse units are significantly lower in δ18O values, especially the focus of this study, the Tuff of Sulphur Creek of the Upper Basin Member Rhyolites. δ18O values are reported as per mil VSMOW. (Modified from Hildreth et al., 1984).
10
to erosion by the river. The thickest exposure of this tuff is to the east, and it is postulated that
the vent for this unit and a petrologically similar overlying rhyolite, the Canyon Flow, is located
there (Christiansen, 2001). The unaltered tuff is vitrophyric and has a phenocryst mineral
assemblage consisting of approximately 3% modal plagioclase and 2% quartz, with <1%
sanidine, and lesser Fe-Ti oxides (Girard and Stix, 2009). Welding has flattened pumice lithics
into fiamme. The tuff and other Upper Basin Member rhyolites have the lowest δ18O values for
eruptive units in the park; the Tuff of Sulphur Creek has magmatic quartz δ18O values of 1.7‰
(Hildreth et al., 1984).
1.4 Glacial History of the Yellowstone Plateau since TSC
Yellowstone National Park lies mostly between 2,100 and 2,500 meters elevation above
sea level and has experienced two major periods of glaciation since emplacement of the Tuff of
Sulphur Creek, at 479 ka. The youngest is the Pinedale Glaciation. This period reached its
maxima somewhere around 18.8 to 16.5 ka on the Yellowstone Plateau but has an ice limit age
of 14.2 ka in the northern plateau/Tower Falls/Grand Canyon of the Yellowstone River area
(Licciardi and Pierce, 2008). Pierce et al., (1976) found the youngest recessional deposits ranged
from 10 to 15 ka. Effects of this glacial period spanned from 11 to 75 ka (Pierce et al., 1976).
The older glacial period was the Bull Lake Glaciation which has an ice limit age of
approximately 136 ka from 10Be exposure ages of terminal moraine boulders (Licciardi and
Pierce, 2008). Other techniques, such as obsidian hydration dating from glacial abrasion,
suggest this period of glaciation reached its terminus sometime around 130 ka (Pierce et al.,
1976). Both ranges correspond to the global glacial maximum of ~140 to 150 ka (Martinson et
al., 1987) with glacial times spanning from ~127 to 170 ka in Yellowstone (Pierce et al., 1976).
11
1.5 Hydrothermal Alteration
Pervasive hydrothermal alteration of the Tuff of Sulphur Creek around the Lower Falls,
Inspiration Point, and Artist Point scenic areas has produced the drastic beauty and coloring, and
resulting extensive erosion of the Grand Canyon of the Yellowstone River. The Tuff also is
pervasively altered in the canyon about six kilometers downstream (to the northeast) in an area
called Sevenmile Hole (Figure 5). Here, the canyon is widened most likely due to the
preferential erosion of altered material. Incision of the canyon has exposed more than 300
vertical meters of hydrothermally altered tuff and provides an exceptional cross section of the
hydrothermal system.
The hydrothermal alteration seen in the TSC is typical of shallow epithermal systems
driven by heat from deeper magma (Cooke and Simmons, 2000; John, 2001; Simmons et al.,
2005). Similar systems provide a source of geothermal energy and also form economically
significant Au-Ag epithermal deposits (Henley and Ellis, 1983). Epithermal deposits such as
these have been classified extensively in the literature (e.g., Hayba et al., 1985; Heald et al.,
1987; White and Hedenquist, 1990; Cooke and Simmons, 2000; Hedenquist et al., 2000; John,
2001; Simmons et al., 2005). Most classification schemes, however, generally put alteration
types into two major categories based on fluid chemistry and related alteration mineral
assemblages; these variables are predominantly dependent on pH and redox state of the
hydrothermal fluids (John, 2001; Simmons et al., 2005).
The first environment has been referred to as high-sulfidation (White and Hedenquist,
1990), advanced argillic, magmatic-hydrothermal, and quartz-alunite, but will here be referred to
as acid-sulfate (Heald et al., 1987). This altering fluid type is more acidic with a pH generally
less than 4 and high oxygen and sulfur fugacities. The fluids do not precipitate silica sinter on
12
Figure 5: Geologic map of the Grand Canyon of the Yellowstone River. Eocene Absaroka Volcanics are located to the north just outside of the caldera margin, shown in plus pattern. Upper Basin Member Rhyolites that form the walls of the canyon and are hosts to the hydrothermal alteration are shown in orange. Pervasively altered areas are shown by the diagonal lines. Box outlines Sevenmile Hole field area, the focus of this thesis. Modified from Christiansen and Blank (1975), Prostka et al. (1975), Christiansen (2001), and Larson et al. (2009).
13
the surface, however, leached vugs are characteristic and these are commonly later filled with
secondary silica, and massive veining and silicification occur at depth. Alteration minerals
include quartz, alunite, pyrophyllite, kaolinite, and dickite. Pyrite and other sulfide minerals are
commonly disseminated or massive replacement in this environment and can contain economic
Au, Ag, and sometimes Cu (Hedenquist et al., 2000).
The other environment has been called low sulfidation (White and Hedenquist, 1990;
John et al., 1999; Hedenquist et al., 2000), alkali-chloride, or adularia-sericite (Heald et al.,
1987) but will be referred to here as neutral-pH. The waters in this environment are more
alkaline, ranging from weakly acidic to neutral in pH (generally pH = 5 to 8) and have lower
oxygen and sulfur fugacities than the acid-sulfate fluids. Opaline sinter can actively precipitate
from fluid discharge at the surface. Associated alteration minerals include quartz, adularia, illite
(sericite), and often carbonate minerals. Pyrite and marcasite are common along with other
sulfides containing Au, Ag, and significantly more Hg than acid-sulfate systems. The Ag:Au
ratio is generally higher in these neutral pH systems relative to acid-sulfate systems (John, 2001).
A hydrothermal system produces interaction between rock and heated water that results
in the exchange of mass and energy. Water is present in most upper crustal rocks, and by
intragranular flow through rocks and molecular diffusion/exchange it is able to alter both its own
composition and the composition of the host rock (Norton, 1984). In near surface or shallow
crustal environments, permeable rocks can be infiltrated by groundwater. As these shallow
fluids are heated (in the case of Yellowstone the fluids are driven by the heat released from
underlying magma) they become less dense or boil and rise adiabatically along permeable zones
while reacting with the host rock (White et al., 1971; Fournier, 1989). The fluids circulating in
these rocks travel along a steep thermal gradient where the temperature is maintained at the
14
boiling temperature of the water under the resulting pressure from depth (Haas, 1971). As these
heated, lower density fluids rise upward, cold groundwater replaces them, and thus heat is
transferred convectively by the circulating fluids (Fournier et al., 1976; Morgan et al., 1977;
White, 1978).
The degree of fluid and wall rock interaction is denoted by the water/rock ratio (W/R)
(Taylor, 1979). If a hydrothermal system is able to remain active for a long enough period of
time, the exchange of mass and energy between the water and rock should reach equilibrium.
The path to equilibrium can be modeled using the equation:
[W/R]open = ln[(δ18OW,f + ΔR-W - δ18OR,i)/(δ18OW,i – (δ18OR,f – ΔR-W)]
Where W/R is the water to rock ratio in the system, δ18OW,f is the final value of the water, ΔR-W is
the fractionation between rock and water, δ18OR,i is the initial value of the rock, δ18OW,i is the
initial value of the water, δ18OR,f is the final value for the rock.
1.6 Stable Isotope Ratios: Oxygen
Stable isotopes are a key to understanding shallow hydrothermal deposits for many
reasons as outlined by Campbell and Larson (1989). They provide information on (1) the
temperatures of systems and mineral deposition, (2) the source of the hydrothermal fluids, and
(3) the degree of water-host rock interactions.
Stable isotope ratios are reported as delta values, δ, by taking the moles of the less
abundant isotope and dividing them by moles of the more abundant isotope to obtain a ratio R.
In the case of oxygen this R is 18O/16O. The R for a sample is then normalized to a known
standard (in the case of oxygen, the standard used is Vienna Standard Mean Oceanic Water
(VSMOW)) and is reported as per mil (‰) in the familiar δ notation. The equation for δ18O is:
δ18O (‰) = (18O/16O)sample _ 1 * 103 = Rsample _ 1 * 103 (18O/16O)standard Rstandard
15
The differential partitioning of light stable isotopes into different phases (between
mineral and fluid, mineral and mineral, or fluid and fluid) is referred to as fractionation. The
magnitude of light stable isotope fractionation is a function of temperature. For quartz, if
temperatures are known, water δ18O values can be calculated from measured quartz δ18O values,
assuming equilibrium and using the equation:
Δ18Oquartz-water = δ18Oquartz – δ18Owater
If the two phases quartz and water are assumed to be in equilibrium (this can be used to calculate
the δ18O of the altering fluid and therefore suggest potential sources), the Δ18Oquartz-water can be
assumed to be equal to the fractionation factor, 103lnαquartz-water, and is represented:
Δ18Oquartz-water ≈ 103lnαquartz-water
This fractionation factor is a temperature dependent function and can be written as:
103lnαquartz-water = A*(106/T2) + B
Where A and B are constants relating to the specific phases (in this case quartz and water).
Numerous calibrations of the quartz-water equilibrium fractionation have been published from
observed, experimental, and theoretical calculations (Shiro and Sakai, 1972; Matsuhisa et al.,
1979; Zhang et al., 1989; Sharp and Kirschner, 1994), however, the equation by Clayton et al.
(1972), in which 103lnαquartz-water = 3.38*(106/T2) – 2.90, is used in this study because it is most
applicable to the typical temperature range in epithermal systems of 200 to 500˚C.
As in most near surface hydrothermal systems the predominant fluid responsible for the
alteration is of meteoric origin (Norton, 1984; Campbell and Larson, 1989; Fournier, 1989).
Meteoric water is the only large δ18O reservoir on the Earth with negative values that range from
0‰ (representing the value of sea water and by definition equal to the δ18O of VSMOW) to
16
-55 ‰. The meteoric water line determined by the relationship between δD and δ18O values of
precipitation and resulting groundwater on Earth was discovered by Craig (1961) to be
δD = 8 (δ18O) + 10 as a general fit for nearly all locations on Earth. Local variations in the
equation of the line occur due to the effects of temperature, latitude, and altitude. For
Yellowstone National Park the present day meteoric water line is δD = 8.2 (δ18O) + 14.7
(Kharaka et al., 2002). Present day cold meteoric waters in the Park have δ18O = -15 to -20 ‰
(Parry and Bowman, 1990; Ball et al., 2002; Kharaka et al., 2002). Modern day thermal waters
have values ranging from -3 to -24 ‰ (Parry and Bowman, 1990; Ball et al., 2002), which
indicates there is some process at work responsible for the shift and fractionation of δ18O values.
Thermal water values show a δ18O shift away from the meteoric water line (Figure 6), which is
common for hot springs. This is produced by the water-rock interaction in the hydrothermal
system. The δ18O depleted meteoric waters exchange oxygen with the rock which are
significantly higher in δ18O and progressively impart higher δ18O values in the water. Water
interaction with rock, however, does not drastically change the δD of the water because rocks do
not contain a significant amount of hydrogen. The δD shift in acid-sulfate fluids is therefore a
result of boiling and evaporation where the light stable isotope is preferentially lost. Figure 6
plots δ18O shifts for both types of hydrothermal fluids, neutral pH and acid-sulfate, from the
local meteoric water line.
1.7 Importance of Research
Much of the information obtained on the active hydrothermal systems in Yellowstone
National Park is from drill cores. However, it is very difficult to study these active systems at
depth, because they are very fragile, erode easily, and are relatively short lived. In 1929-30, two
drill holes (Fenner, 1936), and in 1967-68, thirteen drill holes (White et al., 1975) were drilled
17
Figure 6: Graph showing meteoric water line from Yellowstone National Park (Kharaka et al., 2002). Also displayed is the theoretical δ18O shift, plotted with δ18O values of present day thermal waters (Kharaka et al., 2002), and calculated ancient thermal water δ18O values from hydrothermal silica phases. Acid-sulfate type thermal water chemistry is shown by the red arrow. Neutral-pH thermal water chemistry is shown in yellow.
18
into various active geothermal areas in the park by the USGS. These holes, in addition to being
very difficult to drill into active hydrothermal systems, disrupted the hydrothermal systems by
creating new permeability and channels for fluid flow, causing the discharge of fluids from the
drill holes and making some adjacent geysers extinct (White et al., 1975). In addition, the
stipulation that aesthetics and hydrothermal features are preserved limits the access scientists
have to the deeper portions of thermal areas in Yellowstone. Therefore, this study is significant
because the exposure of an ancient hydrothermal system, from the paleo-surface to
approximately 350 meters depth due to erosion of the Grand Canyon of the Yellowstone River,
allows sampling of the system without tampering, changing, or interfering with park policies and
visitor experiences. It is a window to the inner workings of the hydrothermal system, allows
estimates of temperature and mass transfer of the system, and provides information about how
these systems fit into the geothermal, magmatic, and structural history of the Yellowstone
Caldera. Because there have been virtually no previous studies of hydrothermal alteration of the
Tuff of Sulphur Creek, the initial objective of this project is to map part of the Sevenmile Hole
area and characterize the extent of alteration and its mineralogy. To accomplish this goal, both
hydrothermal minerals and assemblages and other hydrothermal features were mapped.
Mineralogy was characterized using a PIMA (portable infra-red mineral analyzer), standard
powder X-ray diffraction (XRD), and by optical and electron microprobe analysis of thin
sections. Oxygen stable isotope ratio analyses were used to provide information on
hydrothermal system and altering fluids such as sources, water rock interaction rates, diffusion of
ions, and precipitation of hydrothermal minerals. Fluid inclusion homogenization temperatures
provide information on temperatures, salinities, and often state of precipitating fluids in the
hydrothermal system.
19
Chapter 2: SAMPLING AND ANALYTICAL METHODS
2.1 Sampling
Samples of the TSC were collected from the altered area at Sevenmile Hole. The goal for
collecting was to obtain a diverse set of alteration types with a focus on hydrothermal quartz or
other hydrothermally precipitated minerals. Samples were collected over an area of ~3km2 to
evaluate the distribution of alteration phases. The field area spanned a vertical distance 350 m,
from the rim of the canyon to the base of the TSC that was exposed near river elevation.
2.2 PIMA
Mineral identification was determined using a Portable Infrared Mineral Analyzer
(PIMATM) field spectrometer, which measures percent reflectance in the short wave infrared
range (1.3 to 2.5 μm) of the electromagnetic spectrum, to identify characteristic absorption
features primarily of hydrous minerals. Scans were performed shortly after collection in the field
(courtesy of Dr. David John, USGS) by placing the hand sample in front of 1 cm diamater sensor
window for 30 seconds. Many samples were analyzed when still damp which possibly yielded
strong water absorption peaks that may not represent structurally bound water, and may therefore
have obscured clay diagnosis.
2.3 XRD
Standard powder X-ray diffraction (XRD) analyses were performed using an automated
Siemens D-500 X-ray diffractometer at Washington State University. Whole rock and mineral
separates were powdered for analysis. Clays were separated using float separation techniques by
grinding sample, mixing in a vile with water, allowing powder to settle, and then removing the
finest fraction off the top with a pipette. The separate was air dried. Powders were then placed
on glass slides for X-ray scans. The diffractometer uses CuKα radiation (λ=1.5418) and was
20
operated at 30 mA and 35kV. The whole rock powder patterns were collected from 2o to 60o 2θ
with a step width of 0.03° 2θ and a step count time of one second. The clay separates were
collected from 2o to 30o 2θ with a step width of 0.02° 2θ and a count time also of one second.
2.4 SEM
In addition to thin sections, back-scattered images and semi-quantitative energy
dispersive (EDS) analyses were obtained from a LEO 982 digital field emission SEM with an
Oxford EDS spectrometer courtesy of Dr. David John, USGS.
2.5 Radiocarbon Dating
Two samples of wood that were embedded in a surficial sinter deposit were dated by
radiocarbon methods at the NSF/University of Arizona Accelerator Mass Spectrometry
Laboratory with a NEC 3MV Pelletron AMS Machine. The error for the age is approximately ±
34 years. The ages were calibrated using the methods developed by Stuiver (1998).
2.6 Fluid Inclusions
Fluid inclusion analyses were performed on hydrothermal quartz that grew in prismatic
vugs. Homogenization temperatures were measured to obtain minimum temperatures for the
hydrothermal quartz precipitating fluids. A few freezing point depressions were measured for
one sample to obtain salinity estimates. To prepare samples prismatic quartz crystals were
picked out of vugs from 8 samples from different elevations. Crystals were cut and doubly
polished to a thickness ranging from 100µm to 2mm and super glued onto a glass slide. Crystals
were examined closely under a microscope to identify fluid inclusion assemblages (FIA) of
primary origin that would provide confident temperature data based on the criteria developed by
Goldstein and Reynolds (1994). FIA’s contained 10 to 40 individual fluid inclusions and 2 to 7
separate FIA’s were studied in different quartz crystals from one sampling location. Heating and
21
freezing temperature measurements of fluid inclusions were collected on a Linkam THMSG 600
stage attached to a liquid nitrogen cooling pump at the University of Nevada, Las Vegas, in Dr.
Jean Cline’s laboratory. Heating phase changes are accurate to a temperature range of 10oC, and
to 0.1oC for freezing temperatures.
2.7 Oxygen Isotope Ratios
Oxygen isotope analyses of quartz were performed using a laser fluorination line with a
direct gas feed to a FinniganTM gas source isotope ratio mass spectrometer at Washington State
University’s GeoAnalytical Lab. Hand samples were crushed and various habits of quartz
(including veins, vugs, replaced feldspar phenocrysts, massive matrix silicification, other open
space filling such as fiamme, and primary magmatic quartz phenocrysts) were picked by hand.
Approximately 2.5 mg of quartz were placed into the sample holder and into a sample chamber.
Samples were pre-fluorinated for three brief periods of approximately three minutes in order to
remove any water or other sources of oxygen that may contaminate the samples. Samples were
then heated slowly with a 20W CO2 laser. Oxygen in the quartz samples was liberated by
reaction with the oxidizing reagent BrF5 (Clayton and Mayeda, 1963; Sharp, 1990). The
released oxygen gas was then passed through a vacuum line and cleaned with cold traps and KBr
as outlined in Sharp (1990). The δ18O values of each sample were measured with a FinniganTM
Delta S Isotope Ratio Mass Spectrometer using the ISODAT NT software system. Oxygen
isotope ratios are expressed in δ-notation, which represents the difference in isotope values
between the sample and the standard Vienna Mean Standard Ocean Water (VSMOW). Results
are then reported in parts per thousand or per mil (‰). The δ18O values of the samples were
corrected using standard bracketing with UWG-2 garnet standard which has a δ18O value of
5.8‰ (Valley et al., 1995).
22
Chapter 3: STRUCTURE AND HYDROTHERMAL MINERALOGY
3.1 Caldera Ring Fault Control on Hydrothermal Activity
Most of the hydrothermal activity in the Yellowstone Caldera is concentrated around
permeable structures that allow upflow of high temperature hydrothermal fluids. The most
apparent structures on a large scale are the ring faults associated with the youngest caldera
collapse (Lava Creek Tuff (LCT)) and the Mammoth-Norris corridor of north-south-striking
faults (Figure 7a). Some of the most pervasively altered rocks, representing the highest
temperature and lowest pH alteration, are found at Norris Geyser Basin where these two fault
zones intersect (Christiansen, 2001). The Sevenmile Hole field site is located in the Grand
Canyon of the Yellowstone River which crosses the northeastern edge of the youngest caldera
margin (Figure 7b). The topographic wall of the caldera in the Grand Canyon vicinity is mostly
the pre-caldera 52-55 Ma Absaroka Volcanics (Feeley et al., 2002) with minor overlying post-
third cycle collapse Upper Basin Member (UBM) units including the Tuff of Sulphur Creek
(TSC). During the 0.64 Ma LCT eruption and resulting Yellowstone caldera collapse, part of the
south side of the Eocene Mt. Washburn volcanic edifice foundered into the caldera. The
remaining part of the edifice forms the present day caldera margin. Later, when the 479 ka TSC
erupted, it slumped into the caldera depression forming slump folds at the caldera’s edge (Figure
8), now exposed in the canyon walls across the river from Sevenmile Hole (Christiansen, 2001).
The TSC likely buried the caldera ring faults and resulting slump blocks (Larson et al, 2009).
These faults provide conduits that allow hydrothermal fluids to rise and circulate into the tuff and
thereby produce the localized alteration in this vicinity.
23
Figure 7: (a) Map of Yellowstone 640 ka LCT caldera (black line) with hydrothermal areas and post collapse rhyolites. Orange band is probable buried caldera ring fault along which hydrothermal alteration is focused. (b) Enlarged map of the Grand Canyon of the Yellowstone River showing pervasively altered areas. Box outlines Sevenmile Hole field site.
(a)
(b)
(b)
24
Figure 8: Photograph of Sevenmile Hole field site looking to the northeast and down-river. The inferred caldera margin, shown in pink, marks the approximate contact between the Eocene Absaroka Volcanics and younger post-third caldera collapse, Tuff of Sulphur Creek (TSC) at 480 ka. The slump fold in the TSC, outlined in orange, is due to slumping of TSC over the edge of the Absaroka Volcanics into the 640 ka caldera depression.
Absaroka (Eocene volcanics)
Caldera margin
slump fold
Tuff of Sulphur Creek (post-LCT caldera collapse)
25
3.2 Surficial Expression of Hydrothermal Activity:
Past and present hydrothermal activity at depth is expressed on the surface by hot spring
activity and/or sinter precipitation. The hydrothermal fluids at depth interact with the rock at
high temperatures and silica is dissolved from the rock; as the reaction progresses the fluid is
driven toward silica saturation. In the Sevenmile Hole field area, this is displayed in the
alignment of extinct, intermittently active, and currently active sinter fields, roughly parallel to
the caldera margin (Figure 9). Erosion of the Sulphur Creek canyon, to the north of Sevenmile
Hole, exposes the contact between pre-caldera Absaroka Volcanics and post-caldera TSC and
forms the caldera margin (Christiansen, 2001). The alignment of the three sinter fields is likely
the result of a buried/underlying structure such as the caldera ring fault that channels the rise of
hydrothermal fluids. Boiling of these thermal fluids further concentrates silica in the liquid
phase by removal of water vapor creating a silica over-saturated fluid and precipitating quartz at
depth. If fluids rise rapidly enough to maintain high temperatures to shallow levels of the
system, enough silica can remain dissolved in solution to precipitate as amorphous silica on the
surface when the over-saturated fluid discharges and cools (Fournier and Rowe, 1966).
Three locations of silica sinter at Sevenmile Hole are described here to constrain the
evolution of the hydrothermal system (Figures 9 and 10). Sinter field location #1 is located at
the present-day rim of the canyon at an elevation of approximately 8060 ft (2460 m). This site is
extinct with respect to water discharge and sinter precipitation, however, there are active
fumaroles within the field which are hydrothermal features characterized by the discharge of gas.
The discharge of sulfur-rich gases is apparent from the precipitation of yellow crystalline native
sulfur around vents (Figure 10, photo 1b and 1c).
26
Figure 9: Google Earth satellite image of the Sevenmile Hole field area. Proposed caldera margin which runs along Sulphur Creek canyon is shown in red. Blue circles indicate areas of currently or recently active sinter precipitation and hot spring activity.
27
Sinter Field Location #1: Sinter field at present day rim elevation of canyon (8060 ft). There is no water discharge or active sinter precipitation, however an active fumarole (1b) is apparent from sulfur crystals (1c). Sinter Field Location #2: Large sinter cone (elevation of 7470ft) north of Ridge 7741 that was thought to be extinct in 2007, however as seen in photograph 2b, thermal waters were discharging in 2008. Photograph 2c shows a trapped piece of wood that was dated by 14C methods (Figure 11). Sinter Field Location #3: Large active sinter field currently discharging east of Ridge 7741. Photo 3b shows an extinct sinter cone. Photo 3c shows the opalized feeder pipe for the cone. Max elevation ~7320 ft. Figure 10: Photographs and description of sinter fields in Sevenmile Hole field area.
1a 1b 1c
2a 2b 2c
3a 3b 3c
28
Sample ID Material δ13C F 14C age BP 10A wood -27.0 0.9753 ± 0.0041 201 ± 34
10B wood -25.8 1.210 ± 0.012 post-bomb
Figure 11: Photograph of wood fragment trapped in sinter deposit from sinter location #2 (the large sinter cone). The table shows data for two fragments of wood that were dated by 14C methods at the University of Arizona/NSF Accelerated Mass Spectrometry Lab. Ages in years before present (BP, present being 1950) are shown in the last column. Post-bomb indicates the wood sample is post-1950.
29
Sinter field #2 is dominated by a large sinter cone located to the north of Ridge 7741 at
an elevation of approximately 7470 ft (2280 m) (Figure 10, photo 2a). During field work in the
summer of 2007, the cone was thought to be extinct. Two samples of wood fragments embedded
in layers of sinter approximately 5 cm from the outer and youngest layer were collected from the
base of the cone (Figure 10, photo 2c). One wood fragment provided a radiocarbon date of 201
± 34 years BP (BP=years before present, present being 1950 AD) and the other fragment was not
datable because it gave an age of post-1950 (Figure 11). During field work the following year
(2008) the apex of the cone contained water obviously of hydrothermal origins due to the higher
than ambient surface water temperature and growth of algae (Figure 10, photo 2b).
Sinter field location #3 has a maximum elevation of approximately 7320 ft (2230 m) and
is the lowest elevation of the three sinter fields in the canyon. This sinter field location
encompasses the largest area, and lies on the east side of Ridge 7741 in the satellite photograph
(Figure 9). There are several hot spring pools near the top of the field and a few extinct sinter
cones at the lower end of the field (Figure 10, 3b). The cone has eroded so that the central feeder
pipe/fluid conduit is visible. The pipe is lined with opal (Figure 10, photo 3c).
There is also active hot spring activity and precipitation of silica sinter within 20 meters
above river elevation. In 2007 and 2008, these hot springs had a visibly larger fluid discharge,
frequency, and clustering density than any sinter producing areas at higher elevation.
From each of these locations it is apparent that a major control on the hydrothermal
activity, surficial mineral precipitation, and fluid discharge, is the groundwater table within a
distance of the surface where it can be discharged. At sinter field #1, thermal waters are not
discharging on the surface, and it appears to be the oldest sinter deposit in the field area. Sinter
is typically precipitated on the periphery of the system where the water table intersects the
30
surface and therefore the rising high temperature silica saturated fluids can be neutralized and
discharge. However, sulfur rich gases are discharging, suggesting that hydrothermal fluid
alteration is still occurring at depth, although liquid cannot reach the surface in this location but
discharges at lower elevations. At sinter field #2 the large cone appeared to be extinct during
the dry summer season of 2007, however, in 2008, a wetter year, the cone was discharging
thermal waters. This suggests the cone is intermittently active and was vigorously active at 201
± 34 years BP, and therefore sits on a threshold of activity based on the availability and depth of
the shallow water, which may vary seasonally and annually. At sinter field #3, hot spring
activity appeared fairly constant during both field seasons, suggesting a near surface
groundwater source. Hot spring activity at lowest elevations in the canyon was more frequent
and discharging larger volumes of liquid, likely due to proximity to the river, where the
groundwater table intersects the land surface. The temporal downward progression of hot spring
water discharge activity and intensity to lower elevations is a result of a declining water table due
to incision of the Grand Canyon of the Yellowstone River. As the canyon is eroded deeper the
water table roughly parallels the topography creating a depression that in effect “starves” the
higher elevation hot springs of a fluid supply and recharge (White et al., 1971). Figure 12 shows
a schematic model of the progression of canyon incision and resulting drop in water table
elevation around the river canyon.
31
Figure 12: Schematic model depicting incision of the Grand Canyon of the Yellowstone River in cross section. Hot spring centers progress to deeper elevations as the water table (dashed blue line) drops due to incision of the canyon, effectively “starving” higher elevations. Sinter field locations are labeled by number. In the bottom cartoon, fumaroles are shown with dotted vertical arrows at rim elevation. The red magma chamber is the heat source and that drives shallow groundwater convection (dashed arrows) (depth and size of magma not drawn to scale).
32
3.3 Alteration Minerals and Textures at Depth
The intensity of alteration in the TSC at Sevenmile Hole can be evaluated using
hydrothermal alteration mineralogy and textures, including the amount of replacement of igneous
minerals. Table 1 presents a summary of the alteration minerals found in the Sevenmile Hole
vicinity. Textures of altered rocks range from weakly leached and partially replaced magmatic
feldspars to intensely and pervasively altered, in which original tuff textures are completely
obliterated and groundmasses are completely replaced by hydrothermal minerals. Silica is
abundant in rhyolitic rocks and is readily dissolved by hydrothermal fluids at epithermal
temperatures of 100-350oC. It is then reprecipitated in various silica minerals and phases
(Fournier and Rowe, 1966).
In the Sevenmile Hole area, there is a zoning of decreasing alteration intensity away from
Ridge 7741, which is inferred to be a center of hydrothermal activity. Original textures of the
tuff are completely destroyed in the most intensely altered rocks. Open space cavities, mostly
formed by leaching of flattened pumice (fiamme) and phenocrysts, are filled with hydrothermal
minerals, predominantly quartz and clay (Figure 13). Veins are larger and more frequent, large
leached vugs are filled with prismatic quartz. Bands and veinlets visible in hand sample connect
fiamme, phenocrysts, and small vugs. These bands alternate between layers of quartz and clay
minerals that give the rock a marbled bacon-like appearance (Figure 13). Where alteration is not
as intense, euhedral phenocryst grain shapes and tuff textures are preserved, and igneous
feldspars are only partially replaced (Figure 14). Hydrothermal sulfide minerals include pyrite
and marcasite, that occur as disseminated fine-grained crystal aggregates and are found in veins
at all depths (Figure 15). However, the sulfides in veins are typically oxidized to Fe oxides.
Jarosite, hematite, and other iron oxide minerals impart a rusty red, orange, and yellow staining
33
due to weathering and exposure of sulfides to atmospheric oxygen (Larson and Taylor, 1987).
Hydrothermal veins and breccia matricies throughout contain oxidized Fe minerals (hematite and
jarosite), which may indicate original precipitation of Fe sulfides (pyrite and marcasite) that have
since weathered.
Table 1: List of alteration minerals, chemical formulas, general occurrence and formation explanations. Alteration Mineral Formula Formation/Explanation kaolinite Al2Si2O5(OH)4 -clay mineral
-low pH alteration product of feldspars -typically formed in acid-sulfate type alteration
dickite Al2Si2O5(OH)4 -high temperature polymorph of kaolinite
nacrite Al2Si2O5(OH)4 -high temperature polymorph of kaolinite
illite (K, H3O)(Al, Fe, Mg)2(Si, Al)4O10(OH)2·(H2O)
-clay mineral -alteration product of feldspars -typically formed in neutral-pH type alteration
muscovite KAl2AlSi3O10(OH)2 -high temperature alteration mineral -likely signifies presence of highly ordered illite
alunite KAl3(SO4)2(OH)3 -low pH alteration product of feldspar -typically formed in acid-sulfate type alteration
walthierite Ba0.5Al3(SO4)2(OH)3 -alunite group mineral, similar genesis as alunite
huangite Ca0.5Al3(SO4)2(OH)3 -alunite group mineral, similar genesis as alunite
sanidine KAlSi3O8
microcline KAlSi3O8
orthoclase KAlSi3O8
-igneous feldspars that do not occur in unaltered TSC -likely the product of varying degrees of alteration and recystallization of sanidine -not stable
adularia KAlSi3O8
hyalophane (K,Ba)[Al(Si,Al)Si2O8]
buddingtonite NH4AlSi3O8
-hydrothermal feldspars -form at low temperatures relative to igneous feldspar
34
E
DC
A B
Figure 13: Textures of pervasively altered rocks from Sevenmile Hole. A, B, and C taken at 100X, camera field of view is 0.64mm by 0.44 mm. (A) Thin section photograph of fiamme replaced with quartz. (B) Thin section photograph of feldspar grain outline replaced with quartz and kaolinite. Matrix shows alternating bands of quartz and kaolinite. (C) Thin section photograph of quartz vein first precipitated prismatic quartz, then later filled with massive quartz. (D) Secondary permeability flow banding with quartz veinlets that connect fiamme and minor vugs. (E) Leached vug filled on outer edges with massive quartz and prismatic crystals in the center. qtz = quartz, kao = kaolinite.
qtz
qtz
qtz
kao
kao
35
Figure 14: Textures of less intensely altered rocks from Sevenmile Hole. (A) Thin section photograph of magmatic feldspar partially replaced by clay. (B) Hand sample that is altered but maintains the original texture and phenocryst outlines of the TSC. (C) Leached vugs partially filled with kaolinite (pink) and opal. sn = sanidine, kao = kaolinite, qtz = quartz. Figure 15: (A) Oxidized Fe sulfide minerals in veins that are apparent from rusty orange color, occur at shallow levels in oxidizing zone. (B and C) Hydrothermal disseminated marcasite grains in quartz matrix. qtz = quartz, mar = marcasite.
A B
C
A
B
C
sn kao
qtz
qtz
mar mar
kao
qtz
500 μm
36
3.4 Alteration Mineral Assemblages
The distribution and association of hydrothermal minerals and textures found in the
hydrothermal system at Sevenmile Hole can be divided into seven distinct mineral assemblage
groups that vary laterally and vertically in the system. These assemblages are summarized in
Table 2. Each assemblage can be broadly characterized based on the dominant clay mineral
present, either kaolinite or illite. There is a distinct vertical transition in clay mineralogy in the
walls of the canyon at Sevenmile Hole. Kaolinite is the dominant clay mineral at depths
shallower than 100 m below the present rim, and illite is the dominant clay mineral at depths
deeper than 100 m below the rim (Figure 16). X-ray diffraction patterns and PIMA spectra for
ten representative samples of kaolinite and illite are shown in Appendix D.
A similar transition in clay mineralogy with depth below the surface is observed in the
Yellowstone drill holes in active hydrothermal systems and occurs at temperatures of 150o-200oC
(Keith and Muffler, 1978; Bargar and Muffler, 1982; Barger and Beeson, 1984). Kaolinite is
stable in a more acidic environment than illite (White et al., 1971), and forms where H2S vapor
that is boiled off deep groundwater, rises to near the surface, oxidizes and condenses, forming
sulfuric acid (H2SO4). The sulfuric acid reacts with the near-surface rocks, resulting in leaching
and precipitation of kaolinite, creating an acid-sulfate alteration environment (Heald et al., 1987;
Hedenquist et al., 1994). In the boiling zone at depth, H2S and CO2 gas can separate from the
liquid phase and rise to higher elevations in the system, the loss of CO2 from this environment
increases the pH (Giggenbach, 1997). This zone may also become silica saturated because vapor
contains low silica content and most silica remains in the residual liquid (Fournier and Rowe,
1966). The combination of these events may produces a neutral-pH alteration environment,
where illite is the stable clay mineral and quartz precipitates (Figure 16).
37
Table 2: Description of alteration mineral assemblages in Sevenmile Hole. Temperatures and pH are from Reyes (1990), Hedenquist et al. (2000), Bethke et al. (2005), Simmons et al. (2005). Dominant clay type
Alteration Mineral Assemblage Occurance/Explanation
1. kaolinite + tridymite/cristobalite/ opal C/opal CT
-shallow alteration -some original feldspars not altered -vugs only partially filled with kaolinite and opal (contain empty space) - above 50 m (~paleo-water table depth) Interpretation: -acid-sulfate environment -temperatures <120oC -original opal may have dewatered and ordered into more crystalline phases (tridymite and cristobalite) -indicates older steam-heated environment or boiling of a paleo-water table (pre-canyon incision) -opal at deeper elevations may indicate fresh overprinting, post-canyon incision
2. kaolinite + quartz
-similar to Assemblage 1 (shallow alteration), however quartz is the dominant silica phase instead of opal -few original feldspars not altered -quartz veinlets connect and replace fiamme and small vugs Interpretation: -acid-sulfate environment -above 100 m in system, but below paleo-water table (50 m) -locally controlled by structure (commonly composes linear and vertical silicified ridges at highest elevations) -temperatures 150-200oC -pre-canyon incision
kaol
inite
3. kaolinite + dickite/nacrite + tridymite ± quartz
-shallow levels, typically above 50 m -tridymite and quartz (instead of opal) always found occurring with dickite -original feldspars lost/completely altered with dickite Interpretation: -temperatures 100-280oC -localized higher heat and fluid upflow zones (possibly structurally controlled) -fluid rise above paleo-water table -pre-canyon incision
38
4. kaolinite + alunite + opal
-shallow alteration similar to Assemblage 1 -open space/vug fill relatively empty, (partly filled with kaolinite, alunite, and opal) -lacks relict magmatic feldspars -pseudocubic alunite crystals fill open space Interpretation: -acid-sulfate, steam-heated environment -oxidation of H2S from boiling of groundwater table -temperatures <120oC -pH 2-3 -found on canyon rim (pre-canyon incision) and also as overprinting at deeper elevation (post-canyon incision)
kaol
inite
5. kaolinite + alunite/walthierite/ huangite + quartz ± dickite
-localized zones of intense alteration, below paleowater table but within oxidizing zone -original feldspars completely altered -tabular alunite group minerals replace matrix Interpretation: -acid-sulfate, magmatic-hydrothermal -magmatic gas (SO2) condenses in groundwater -temperature range 200-350oC -pH < 2 -pre-canyon incision
6. illite + quartz ± various hydrothermal feldspars (hyalophane, buddingtonite, adularia)
-deep levels in system -formed below 100 m -igneous feldspars altered to illite, adularia, hyalophane, buddingtonite Interpretation: -neutral-pH (>4) -temperatures >200oC -pre-canyon incision -likely core/feeder beneath Assemblage 5
illite
7. illite + chalcedony + muscovite
-deep elevations >100 m, high temperature -unlikely actual muscovite (not high enough temperature) but more ordered illite -original feldspars replaced Interpretation: -neutral-pH (>4) -temperatures >200oC -massive veins of chalcedony are late stage -over printing of chalcedony on deep illite environment
39
Assemblage 1 consists of kaolinite + tridymite/cristobalite/opal C/opal CT. This
assemblage is found at the highest elevations in the system, from the rim of the canyon to a depth
of ~50 m. The elevation of the paleo-water table during alteration likely lies within this zone.
The feldspars range from original magmatic to completely replaced, but grain outlines are
typically preserved (Figure 14A and B). Leached vugs are only partially filled with kaolinite and
opal (the remainder of the vug is empty) (Figure 14C). Opal is precipitated at shallowest
elevations and like this assemblage was only found above 50 m depth in Yellowstone drill core
Y-11 (Bargar and Muffler, 1982). At depths near the water table, cristobalite and tridymite tend
to form where fluid temperature ranges from 100o to 160oC (White et al., 1975; Simmons et al.,
2005). It is also possible that amorphous opal at the highest elevations may have dewatered and
ordered into more crystalline phases (tridymite and cristobalite) producing opal C and opal CT.
This may occur on timescales of thousands of years (Lynne et al., 2005), which indicates that an
older and now extinct shallow hydrothermal environment was present. Fresh amorphous opal
found at deeper levels may indicate younger overprinting on older, deeper alteration
assemblages. This opal precipitates post-canyon incision because opal typically only forms at
near surface depths <50 m in active systems observed in Yellowstone drill cores (Fenner, 1936;
White et al., 1975; Bargar and Muffler, 1982). Temperatures of opal precipitation in this region
can be as low as 20oC, however, in order to precipitate at the surface a temperature of at least
175oC must be achieved at depth to dissolve enough silica in solution (Fournier and Rowe,
1966). This assemblage near and above the paleo-water table, and 50 m below the paleo-surface
tends to the maximum extent of opal precipitation and this assemblage zone. Sulfur in this
assemblage is oxidized in the near surface environment and therefore produces an acid-sulfate
type alteration.
40
Assemblage 2 contains kaolinite + quartz + various feldspars. This alteration is very
similar to Assemblage 1, however it is slightly deeper and quartz forms instead of opal. At high
temperatures (for epithermal systems ~ 350oC (Simmons et al., 2005)), quartz precipitates
readily down to about 150oC (White et al., 1975). The occurrence of quartz at shallow depths in
Sevenmile Hole suggests localized zones of upwelling, high temperature fluids. Cooling of this
high temperature fluid in permeable zones may cause silica-oversaturation of fluids and produce
massive silicification in the shallow portions of the epithermal system (<150 m) (Hedenquist et
al., 2000). This assemblage commonly composes the linear and vertical silicified ridges at
highest elevations in Sevenmile Hole, suggesting this alteration is highly structure controlled.
The quartz replaced and intensely silicified ridges may represent localized conduits of greater
upflow of high temperature silica saturated fluids. Leaching is not as intense, hence some
original feldspars may survive alteration (Figure 14A). This assemblage zone is found below the
paleo-water table and below ~50 m but above the kaolinite-illite transition at 100 m. The 100 m
depth in the system represents the lower extent of a supergene or surficial oxidizing zone,
therefore producing acid-sulfate type alteration and precipitating kaolinite. Like assemblage 1
this likely formed before the canyon incised and water table declined with the topography.
Assemblage 3 is composed of kaolinite + dickite/nacrite + tridymite ± quartz. The
location of this assemblage appears to extend into the amorphous silica precipitating zone which
could possibly be above the paleo-water table. Like assemblages 1 and 2, this zone has acid-
sulfate chemistry, however, dickite and sometimes nacrite are found here. Dickite and nacrite
are high temperature polymorphs of kaolinite. Dickite can form at temperatures up to 280oC,
whereas kaolinite tends not to form above 200oC (e.g., Reyes, 1991). Tridymite and quartz
(instead of opal) are typically found occurring with dickite, which also indicate higher
41
temperatures. Magmatic feldspars are destroyed and/or completely altered. This indicates these
areas are localized higher heat and fluid upflow zones creating more intense alteration than
assemblages 1 and 2, possibly where permeable structures in the tuff allow an acidic, silica
saturated fluid to rise above the surrounding water table. This assemblage likely formed pre-
canyon incision, as evidenced by the high temperatures at shallow depths.
Assemblage 4 is kaolinite + alunite + opal. Like the previous assemblages this forms at
shallow depths, above a paleo-water table (50 m) and above the oxidizing boundary (100 m).
The presence of amorphous opal indicates this is fresh, young, shallow alteration similar to
assemblage 1. However this assemblage contains alunite and lacks original feldspars. Open
space/vug is filled with kaolinite, pseudocubic alunite, and opal, but remains relatively empty.
Pseudo-cubic open space filling alunite (Figure 17) occurs at shallow, near-surface depths in the
hydrothermal system. XRD diffraction patterns and PIMA spectra for ten representative samples
of alunite are shown in Appendix D. Alunite in this form was found at the rim of the present day
canyon and also on the deepest elevation river terrace. This crystal morphology precipitates in a
steam-heated environment, where H2S rich vapor is boiled off the hydrothermal fluids at depth,
ascends, and condenses in shallow oxidizing vadose zone (Rye et al., 1992; John, 2001). The
pseudo-cubic alunite found on the deep river terrace is likely recent alteration that formed after
incision of the canyon when the hydrothermal activity had migrated downward. It is precipitated
along with kaolinite and opal which suggest formation at shallow depths and at relatively low
temperatures; steam-heated alunite tends to form at temperatures below 120oC (Dill, 2001;
Simmons et al., 2005). The sulfur responsible for the H2S vapor and precipitation of alunite is
most likely sulfur recycled from hydrothermally precipitated sulfides such as pyrite or marcasite
that are disseminated throughout the altered tuff and not a direct magmatic source (John, 2001).
42
A and B: pseudocubic steam-heated alunite (Assemblage 4)
C and D: tabular magmatic-hydrothermal alunite (Assemblage 5)
Figure 17: Contrasting morphologies of alunite. (A and B) Pseudocubic, steam-heated alunite formed above the water table. Pseudocubic alunite precipitates as open space vug fill from oxidation of sulfur-rich gas to sulfuric acid in shallow near surface environments (Assemblage 4). (C and D) Tabular alunite formed below the water table in an acid-sulfate, magmatic-hydrothermal environment (assemblage 5) from magmatic SO2. (C) Massive tabular matrix replacing alunite surrounding a magmatic quartz phenocryst. (D) Tabular vug-filling alunite.
A
C
B
D
Empty vug space
Empty vug space
al al
al al
qtz
qtz
43
Assemblage 5 consists of kaolinite + alunite/walthierite/huangite + quartz ± dickite. This
assemblage occurs at shallow levels (above the 100 m kaolinite-illite transition but below the
paleo-water table and 50 m), in localized zones of pervasive alteration. Original feldspars are
obliterated and replaced by alunite and clay. Tabular alunite group minerals massively replace
matrices of the host tuff, or are intergrown with quartz and kaolinite in vugs (Figure 17).
Tabular matrix-replacing alunite is characteristic of magmatic-hydrothermal alunite and forms at
higher temperatures ranging from 200o to >350oC (Dill, 2001). A sample of tabular alunite from
this assemblage was 40Ar-39Ar dated. The fluids that produce magmatic-hydrothermal alunite are
formed by disproportionation of SO2 released from a degassing magma (Rye et al., 1992). This
acid-sulfate environment typically has pH <2 due to the greater influence of H2S, sulfuric acid,
and a high heat flow likely due to convective vapor transport of energy and mass, which formed
before canyon incision. Dill (2001) noted that the precipitation of alunite from acidic thermal
fluids produces an increase in silica and pH in the fluid by the reaction: 2 kaolinite + 2 K+ + 6H+
+ 4SO42- 2 alunite + 6 H4SiO4(aq), which may drive quartz precipitation and contribute to a
change in fluid chemistry from acidic to slightly more neutral. The base of this zone, like
assemblage 2, occurs at the lower limit of the oxidizing zone.
Assemblage 6 is illite + quartz + various alteration feldspars (hyalophane and
buddingtonite) ± adularia. Illite occurs in hydrothermal veins and flow bands, occurs intergrown
with quartz, and may be relatively coarse-grained and tabular in vug filling (Figure 18). Original
feldspars are altered and replaced with hydrothermal feldspars such as hyalophane,
buddingtonite, and adularia. Adularia also occurs as fine-grained crystals in veins or alteration
flow bands intergrown with quartz and illite (Figure 18). Illite dominated clay mineralogy is
found at deep elevations, below 100 m. It forms at deeper and at higher temperatures (>200oC)
44
Figure 18: SEM Photographs of neutral-pH alteration minerals from assemblage 6. (A) Feldspar phenocryst replaced by coarse tabular illite, with quartz and minor pyrite/marcasite (white). (B) Feldspar phenocryst replaced with quartz and rhombic-shaped adularia (light gray). (C) Veins or flow bands of fine grained quartz, adularia, and illite. qtz = quartz, il = illite, ad = adularia.
C
B
A
il
il
ad
ad
qtz
qtz
qtz
45
than kaolinite in many epithermal systems (Reyes, 1990; Simmons et al., 2005). This mineral
assemblage, quartz + illite ± adularia, is evidence for neutral-pH hydrothermal fluid chemistry
which results from the loss of H2S due to boiling. The H2S in this reducing environment is not
oxidized to sulfuric acid and therefore more neutral assemblages of minerals can precipitate
(illite and adularia). Considering that this assemblage is characteristic of relatively deep
epithermal environments and high temperatures, it must have formed before canyon incision.
This is likely the core or feeder assemblage that lies directly beneath assemblage 4.
Assemblage 7 consists of illite + chalcedony + muscovite. Original feldspars are
completely replaced. The muscovite may not actually be hydrothermal muscovite, since that
typically occurs at much deeper and hotter porphyry type temperatures in excess of 350oC
(Lowell and Guilbert, 1970). Instead it is probably more ordered and crystalline illite (which has
a very similar XRD diffraction pattern). This assemblage is the product of intense alteration by
high temperature neutral-pH fluids (>200oC) below 100 m, similar to assemblage 6.
Chalcedony is common in massive veins that generally increase in diameter with depth (Figure
19), and cross cut the illite alteration suggesting this is late stage overprinting. Chalcedony is
more soluble than quartz and, therefore, requires a higher silica concentration in the fluid than
quartz (Fournier, 1989). Silica over-saturation of a fluid at depth likely formed due to the boiling
loss of water vapor, which concentrates silica in the residual fluid may have been achieved in
this deep illite producing zone below 100 m (pre-canyon incision). When the canyon began to
incise, cooler temperatures were forced deeper into the system, which may have caused the rapid
cooling of this silica over-saturated fluid, hence the late stage chalcedony veins. It is suggested
in Simmons et al. (2005) that chalcedony may occur on the periphery of the epithermal
environment, at or near the elevation of the water table. This assemblage may therefore mark the
46
location of a later and lower elevation paleo-water table resulting from canyon incision.
Chalcedony veins are sporatically found throughout Sevenmile Hole alteration area, however not
with the frequency, prevalence, and size in this deep illite-dominant zone composing assemblage
7. In most cases these veins appear to be the latest stages based on cross cutting relationships
and may indicate interaction and cooling of a silica saturated hydrothermal fluid with the water
table interface.
Figure 19: Massive chalcedony veins of assemblage 7.
47
Assemblages 2, 3, 4, and 6 were often found composing prominent silicified ridges
throughout the field site that have resisted erosion as the canyon incised. Conversely, areas that
comprise the scree fields and are easily eroded tend to have higher clay content within their
matrix based on visual estimation. The silicified ridges have a laterally linear or vertical pipe-
like geometry (Figures 20 and 21). These structures are likely permeable zones of focused
upwelling of a silica saturated hydrothermal fluid. Altered rock along these structures contain as
much as 92 weight percent silica (preliminary XRF analyses), which is an increase in silica by
more than 15% from the high silica rhyolite protolith that contained about 76 weight percent.
The linear distribution of these features suggests that they are controlled pre-alteration structures
in the TSC that provide permeable conduits for the flow of hydrothermal fluids. These conduits
may allow temperatures higher than predicted by the reference boiling point curve to shallow
elevations.
A map of the distribution of the seven mineral assemblages of the Sevenmile Hole field
site is shown in Figure 21.
48
Figure 20: Photographs of erosion resistant highly silicified ridges from Sevenmile Hole. (A) Vertical pipe-like structures. (B) Silicified linear fracture which provided the permeability to channel silica-rich fluids upward. (C) Ridge that runs down Ridge 7741 in the background and the scree field of more clay rich alteration that is not as resistant to erosion. (D) Along the top of Ridge 7741, where the tuff is pervasively leached and vuggy silica that is found in the central highly altered core of the Sevenmile Hole hydrothermal system.
B
A C
D
49
Figure 21: Alteration assemblage map of the Sevenmile Hole altered field area showing location of the seven mineral assemblages. Approximate cross section of Figure 33 is also shown.
50
3.5 Model of the Hydrothermal Zoning at Sevenmile Hole
There are vertical and lateral patterns in the distribution of the seven alteration
assemblages. The exposed system seems to consist of a central pervasively altered and leached,
vuggy silica zone. The diverse alteration mineralogy (alunite group minerals and alteration
feldspars) and heavy leaching indicate acidic vapor, hydrothermal cation, and silica saturated
fluid upflow. This alteration destroys original volcanic texture. The high permeability of this
region allows for recharge likely form surrounding less altered areas, fast enough so that a silica
saturated fluid can reach the surface and deposit sinter. This central zone is where the largest
and most frequent veining, hydrothermal brecciation, heaviest silicification, and overall most
pervasive alteration occurs. This zone is composed of mineral assemblages 5 and 6 below the
elevation of the water table during alteration. The depth of the paleo-water table is assumed to
be ~50 m due to a transition from quartz to opal at the shallowest depths (Figure 22). At depths
above the water table, boiling of groundwater and/or hydrothermal fluid creates a steam heated
environment where assemblages 1 and 4 may form. Assemblage 3 which consists of localized
dickite/kaolinite and quartz; that may represent the rise of high temperature hydrothermal liquid
above the water table due to permeable fractures in the host tuff.
Moving out laterally away from this pervasively altered core to regions farther from the
focus of upflow the alteration intensity weakens. Original tuff textures can be preserved and
varying degrees of replacement occur as described in assemblages 1 and 2. The transition from 1
to 2 is identified by the transition from quartz in the groundwater saturated zone to opal in a
shallower steam-heated zone. This depth of approximately 50 m may represent the depth of an
original paleo-water table before the canyon incised. This zone may be locally heavily silicified
and alterated to clays, however alteration minerals such as alunite group minerals and alteration
51
feldspars such as adularia, buddingtonite, and hyalophane are not present. Instead feldspars
identified in XRD diffraction patterns are sanidine (magmatic), microcline, orthoclase, and
anorthoclase. Besides sanidine, these do not occur in the unaltered TSC and tend not to form in
low temperature epithermal environments, so it is probable that these feldspars signifiy the
disordering and alteration/replacement of original feldspars due to alteration. Localized zones of
assemblage 3 (dickite and quartz) occur along structures that allow high temperature liquid
above water table elevation. Steam-heated alunite (assemblage 4) does occur at the shallowest
elevations locally, the sulfur source in this case may be from recycled hydrothermal sulfides, as
opposed to a massive magmatic upflow in the central conduit. The kaolinite-illite transition at
100 m is also apparent in this zone, represented by the contact between assemblage 2 and 6.
Farther out in the lowest intensity alteration textures of the tuff are well preserved
although may still experience weak alteration. There is much less hydrothermal silicification,
the rock seems to resist erosion due to it original welded nature.
It is apparent that both alteration types (acid-sulfate and neutral-pH) are found within the
Sevenmile Hole altered area. There is a clear transition at depth from kaolinite in the shallow
portion of the system to illite in the deeper portion. It is possible and likely that the system has
fluctuated in water chemistry between acid-sulfate and neutral-pH throughout time, however the
sharp vertical transition between clay types implies that both alteration types are actually
components of one coeval hydrothermal system. This vertical transition is also observed in the
Yellowstone drill cores (Fenner, 1936; White et al., 1975). This kaolinite-illite transition may be
caused by increasing temperature at depth. At temperatures <~150-200oC kaolinite is the stable
clay. At temperatures >200oC illite is the stable clay. This clay transition may also mark the
52
lower boundary or extent of influence of atmospheric oxidation. Groundwater at depth in the
system boils off vapors and gases such as H2S. When the H2S gas reaches the oxidizing zone it
condenses to form sulfuric acid and results in more acidic alteration, hence heavy leaching, and
kaolinite, quartz, alunite precipitation (assemblage 5). The loss of H2S and water vapor from the
residual liquid at depth results in neutral fluid chemistry and an increase in silica concentration,
hence illite, quartz, adularia precipitation (assemblage 6).
Chalcedony veins are present sporadically throughout however at deep elevations
(occurring with assemblage 6) the veins are much more prevalent. They are late stage and cross
cut earlier alteration. In similar systems chalcedony was found to occur on the periphery of the
intensely altered zone at depths near the water table. Assuming paragenesis with the rest of the
assemblages this zone is well below the paleo-water table depth. Therefore, the massive size and
frequency of the veins at deep elevation suggest that a silica over-saturated fluid was once the
fluid responsible for assemblage 6, but then was cooled relatively rapidly to precipitate
chalcedony near a later stage, deeper water table. The drop in temperature and water table can
be explained by the incision of the canyon.
53
Figure 22: Theoretical cross section of the hydrothermal system at Sevenmile Hole showing distribution of mineral assemblages and relation to paleowater tables, paleosurface, the kaolinite-illite transition at 100 m depth, and structures. The vertical fingers of assemblages 2 and 7 are to indicate the structural control that may cause these zones to extend beyond the predicted temperature at depth ranges in system.
54
Chapter 4: FLUID INCLUSIONS
Estimating temperature and fluid composition are important parts of understanding the
hydrothermal system in the Sevenmile Hole vicinity or any other epithermal system in the world.
One method for estimating temperature is the analysis of fluid inclusions. In this study, fluid
inclusion analyses were performed on hydrothermal quartz. Ideally, only primary inclusions are
analyzed, meaning that the fluid is entrapped during growth of the crystal it is enclosed in, and
therefore records the conditions when the crystal grew. Secondary inclusions are inclusions
trapped after crystal growth usually in fractures, and these can provide useful temperature and
fluid composition data. Primary inclusions were recognized by characteristics such as alignment
in planar arrays that are parallel to crystal faces, alignment along growth planes, or
crystallographical orientation (Goldstein and Reynolds, 1994). A group of inclusions that are
petrographically related and assumed to have formed at the same time and under the same
conditions and from a single homogeneous fluid is referred to as a Fluid Inclusion Assemblage
(FIA). Figure 23 is a photograph of a typical FIA analyzed for this study. FIA were selected
from different prismatic crystals from within a single vug or often from different vugs at the
same sampling location to obtain an average temperature for that elevation of mineralization.
All of the analyzed fluid inclusions contain two phases: a liquid phase and a vapor bubble. The
vapor bubble typically ranged from 1 to 10 volume percent of the inclusion based on visual
estimates. All of the inclusions homogenized to liquid when heated.
4.1 Homogenization Temperatures
The FIA analyzed from the eight samples of prismatic vug quartz crystals (including
primary, secondary and unknown origin inclusions) gave homogenization temperatures ranging
from 150o to 350oC. Primary FIA temperatures ranged for
55
Figure 23: Photograph of a doubly polished cross section of a prismatic quartz crystal from sample YS8-19 showing an assemblage of primary fluid inclusions. Inclusions are aligned nicely along crystallographic axes and all inclusions are in one plane of view, which are ideal characteristics of primary inclusions. Freezing point depressions were measured on this FIA and gave salinities of 0.35 to 0.71 % wt NaCl eq.
YS8-19
10μm
56
precipitation of prismatic vug quartz averaged 180o to 280oC (Figure 24). This range is typical
for shallow epithermal hydrothermal systems (Haas, 1971; Hedenquist et al., 2000; Simmons et
al., 2005). The more variable and larger range of temperatures in the secondary and unknown
origin inclusions suggest a fluctuating system and multiple phases of quartz precipitation (which
is also evident from two stages of quartz precipitation in veins observed in thin section (Figure
13C)). The secondary and unknown origin inclusions produced homogenization temperatures
typically higher than the primary inclusions. A silica saturated fluid will precipitate minerals
upon cooling, which lowers the solubility of the silica phase (quartz in this case) in solution. The
primary homogenization temperatures, therefore, represent the minimum temperature of mineral
(quartz) formation (Goldstein and Reynolds, 1994; Kesler, 2005).
Figure 24 shows samples in order of depth with YS-07-75 being the shallowest in the
system (highest elevation), and YS8-24 the deepest. There is a crude trend of increasing
temperature with depth in the system. These temperatures are comparable to those found in the
drill cores drilled into active systems in the park and other epithermal deposits around the world
(White et al., 1975; Hedenquist et al., 2000; Simmons et al., 2005).
In a few crystals there is evidence of boiling in the FIA, which is common in
hydrothermal systems where temperatures remain at the boiling point temperature for that
hydrostatic pressure over the entire (epithermal) depth in the system (Haas, 1971). Evidence of
boiling is preserved when both liquid and vapor phases are simultaneously trapped within the
inclusion because both phases were present during crystal growth. The inclusions in the
prismatic crystals that show evidence of boiling have variable liquid/vapor ratios, or some of the
inclusions in the assemblage may appear dark because they are completely filled with vapor.
57
Figure 24: Frequency histograms of fluid inclusion homogenization temperatures from 8 samples of prismatic vug quartz from varying depths in the Sevenmile Hole system. Only primary fluid inclusions are shown. Samples are arranged from shallowest depth of formation at the top (YS-07-75) to deepest at the bottom (YS8-24). Inclusions typically have 1 to 10% volume gas. Generally the inclusion assemblages increase in homogenization temperatures with depth. Freezing point depressions were performed on sample YS8-19 (the second to lowest sample in elevation) and gave salinities of 0.35 to 0.71 % wt NaCl eq.
YS-07-75
YS-07-84
YS8-1
YS-07-88
58
YS8-17
YS8-18
YS8-19
YS8-24
59
4.2 Freezing Point Depressions/Salinity
Freezing point depressions were measured on one sample (YS8-19) (Figure 23), in order
to estimate bulk composition of the trapped fluid in weight percent NaCl equivalent. The
temperature at which the frozen liquid (ice) within these inclusions melted ranged from -0.2o to -
0.4oC, and the majority of them melted at -0.3oC. The salinities that correspond to these freezing
point depressions are 0.35 to 0.71 weight percent NaCl equivalent (Bodnar, 2003; Goldstein and
Reynolds, 1994). There was no CO2 detected in any of the inclusions. A low salinity is very
typical of epithermal systems and of present day thermal waters in the park and indicates the
thermal fluid is dominantly meteoric water (Campbell and Larson, 1989; Cooke and Simmons,
2000; Simmons et al., 2005). For calculations using Haas’s (1971) boiling point curve the
salinity of the hydrothermal fluids is assumed to be 0 wt % NaCl or a pure water system because
0.71 weight percent salinity has a negligible effect on boiling point.
4.3 Homogenization Temperatures Compared to the Boiling Point Curve
Figure 25 plots average homogenization temperatures for primary FIA versus their
associated sample elevation within the hydrothermal system. The boiling point curves for 0 wt
% NaCl hydrothermal fluids at depth from Haas (1971) are also plotted. The red curve is
assuming the current canyon rim elevation of 2470 m (8100 ft) is the paleo-surface during this
ancient alteration and therefore the top of the effective hydrothermal system during time of
quartz growth. Most of the FIAs plot above this reference boiling point curve (in red) and do not
all align to a single boiling point curve, but instead produce a range of temperatures greater than
would be expected from precipitation under the current canyon rim elevation. A possible
explanation is the boiling point curve was shifted up due to additional overhead hydrostatic
pressure. The yellow curve assumes additional hydrostatic pressure that would result from the
60
top of the system being located at an elevation of 2560 m (an additional 90 vertical meters of
hydrostatic pressure above current rim elevation), green of 2650 m (additional 180 m), blue of
2745 m (additional 275 m), and orange of 2930 m (additional 460 m). This upward shift in the
boiling point curve from the addition of hydrostatic pressure could either be due to a previously
higher elevation land surface that has been since reduced by erosion, or the additional hydrostatic
pressure due to a glacial ice sheet, assuming the alteration formed in a glacial period. The
elevation of a glacial ice sheet may produce more variable changes in hydrostatic head over a
shorter time period due to climatic accumulation and ablation. Various periods of prismatic
quartz growth may have been controlled by this flux in hydrostatic head above the site of
precipitation, and may be a cause of the temperature variation measured in samples at the same
depth.
61
2000
2100
2200
2300
2400
2500
150 170 190 210 230 250 270 290 310
Average Homogenization temperature (oC)
Elev
atio
n (m
eter
s ab
ove
SL)
Figure 25: Elevation of samples selected for fluid inclusion assemblages (FIA) plotted versus the average homogenization temperature for primary FIA for each sample. There is a weak trend showing increase in temperature with depth of sample. Also plotted are boiling point curves for 0 wt % NaCl hydrothermal fluids at depth from Haas (1971). Red curve is assuming current canyon rim elevation of 2470 m is the paleosurface during time of quartz growth. Yellow curve assumes maximum limits of surface elevation and resulting pressure of 2560 m, green of 2650 m, blue of 2745 m, and orange of 2930 m.
62
Chapter 5: STABLE ISOTOPE RATIOS: OXYGEN
5.1 Fractionation and Exchange of Oxygen
The Tuff of Sulphur Creek (TSC), like many rhyolitic tuffs, is host to the circulation of
hydrothermal fluids, although the welded character is not favorable for permeability. The high
silica composition of the TSC (76% weight SiO2 (Hildreth et al., 1984; Christiansen, 2001))
provides abundant silica for producing a silica-saturated hydrothermal fluid. Fluid-rock
interaction is responsible for the exchange and fractionation of oxygen isotopes between a fluid
and hydrothermal minerals such as quartz. Original magmatic quartz in the TSC is already low
δ18O with values averaging approximately 1.7‰ (Hildreth et al., 1984; Bindeman and Valley,
2001), which is likely a result of magmatic recycling of previously meteoric-hydrothermally
altered crust.
Meteoric water is the dominant fluid involved in the alteration and hydrothermal
processes in the Yellowstone hydrothermal system (Truesdell and Fournier, 1976; Truesdell et
al., 1977) and this is the case for most epithermal environments elsewhere (Taylor, 1974 and
1979). The δ18O values of the altered host rock are shifted to significantly more negative values
(Campbell and Larson, 1998), therefore, the rocks must be exchanging oxygen with the only
large reservoir of negative δ18O values on the planet, meteoric water. Present day pristine
meteoric waters in Yellowstone National Park have δ18O values of -15 to -20‰ (Parry and
Bowman, 1990; Ball et al., 2002; Kharaka et al., 2002). This meteoric water permeates into the
ground and is circulated through the host rock (Criss et al., 1987). In addition to exchanging
oxygen, the water continually dissolves more silica from the host rock to become saturated, and
as boiling and minor changes in temperature occur, the fluid may precipitate hydrothermal silica
minerals (Sturchio et al., 1990).
63
5.2 δ18O Values of Magmatic and Hydrothermal Quartz
The δ18O values of hydrothermal quartz were measured in order to calculate the δ18O
values of the fossil hydrothermal fluid responsible for the alteration in Sevenmile Hole.
Temperatures for this system are estimated from previous studies on epithermal systems (White
et al., 1975; Cooke and Simmons, 2000; Simmons et al., 2005), drill core in situ temperatures in
Yellowstone (Keither and Muffler, 1978; Bargar and Muffler, 1982; Barger and Beeson, 1984),
mineral assemblages, and experimentally with fluid inclusion homogenization temperatures
(Chapter 4).
Magmatic quartz phenocrysts in the altered tuff, which started out at approximately
1.7‰, show a slight depletion in δ18O values of approximately 1‰, averaging 0.6‰ as opposed
to the 1.7‰ in the fresh tuff (Figure 27). This suggests a minor degree of oxygen exchange with
altering fluids. In thin section it is apparent that many of the magmatic phenocrysts are
fractured. This may facilitate the diffusion of oxygen into the crystal by providing more surface
area for contact with the fluid. Also apparent in a few samples are hydrothermal rim growths of
quartz. If these were inadvertently sampled along with magmatic quartz it may have lowered the
measured δ18O values. Quartz is relatively resistant to isotope exchange at temperatures below
250oC (Clayton et al., 1972; Criss and Taylor, 1983). However, fluid inclusion analyses of
secondary inclusions in quartz suggest that temperatures may have reached up to 350oC, and may
have therefore increased the oxygen diffusion into relict phenocrysts.
5.3 δ18O Values of Hydrothermal Quartz Habits
Hydrothermal quartz is assumed to form in equilibrium with the fluid from which it
precipitates. In the case of this system, the fluid is a near boiling, silica saturated water of
meteoric origins. The δ18O values of the resulting quartz precipitates ranges from 1.3‰ to as
64
2050
2100
2150
2200
2250
2300
2350
2400
2450
2500
-7.0 -6.0 -5.0 -4.0 -3.0 -2.0 -1.0 0.0 1.0 2.0 3.0 4.0
Elev
atio
n (m
)
magmatic quartz phenocrystshydrothermal quartz
δ18O (‰) VSMOW
Figure 26: Graph comparing unaltered magmatic quartz phenocryst values for fresh Tuff of Sulphur Creek (gray box) (from Hildreth et al., 1984; Bindeman and Valley, 2001), to measured magmatic quartz phenocrysts (blue data points) in altered sections and all hydrothermal quartz habits (pink data points) in altered tuff.
65
low as -5.7‰ (Figure 27). These are all more negative than magmatic quartz in the fresh TSC.
The hydrothermal quartz has variability in δ18O values of as much as 7‰. There is no pattern in
δ18O values with depth in the system which is assumed to increase temperature. Variable
temperatures corresponding to depth in the system would ideally produce predictable
fractionation of δ18O values. The lack of a trend suggests that temperature controlled by depth in
the system is not solely responsible for the variable δ18O values.
In all cases hydrothermal quartz habits such as veins, veinlets, quartz replaced feldspars
and fiamme, and minor vugs <1 cm, showed higher (less negative) δ18O values compared to their
silicified matrix component from the same rock sample. And conversely matrix values in each
case display more negative δ18O signatures than their vein or vug component.
Hydrothermal quartz veins and minor vug filling quartz have the most variable range of
δ18O values. Veins yield δ18O values ranging from 0.9 to -5.7 ‰. Minor vugs and pumice
veinlets have values ranging from 1.3 to -4.0 ‰. Matrix silicification and large prismatic vug
quartz generally produce the most negative and most tightly constrained range of δ18O values for
hydrothermal phases. Prismatic vug crystals have values of -3.3 to -4.8 ‰. Massive matrix
silica replacement ranges from -3.1 to -4.3 ‰ (Figure 28).
Factors that could account for the wide spread of δ18O values for veins and minor vugs
versus the tightly constrained matrix and prismatic δ18O values are (1) temperatures of
formation, (2) variability in the δ18O of the hydrothermal fluid due to recharge, (3) surface area
or volumes of rock in contact with fluid, or evolution of the water-rock ratio, (4) variable
permeability in pore space in the host rock (in the case of this system: fractures and highly
permeable zones versus the rock matrix in between these fractures) (Evans and Nicholson, 1987;
DePaolo, 2006).
66
Figure 27: Graph comparing δ18O signatures for the different types of hydrothermal quartz habits or mineralization. Veins and minor vugs seem to span the largest range, whereas prismatic vugs and matrix silicification are confined to narrow more negative ranges.
67
All of these possible factors controlling the δ18O values can be generally assumed to be
affected by the intensity of the alteration. Fluid upwelling zones in the system produce higher
than expected temperatures at shallow depths as indicated by fluid inclusions and mineral
assemblages discussed in Chapters 3 and 4. The δ18O values of quartz in the TSC were then
graphed according to their mineral assemblage number (Figure 29). Generally the assemblages
increase in relative intensity from 1 to 7. Intensity was estimated using temperature, depth,
amount of leaching and replacement, and destruction or original tuff textures. There appears to
be two populations of hydrothermal quartz δ18O values. The lower intensity assemblages 1, 2, 3,
and 4, had typically higher δ18O. The higher intensity assemblages 5, 6, and 7, tended to have
more negative δ18O values. Magmatic quartz phenocrysts show a slight depletion in δ18O values
as intensity increases. This likely indicates that some degree of diffusion of oxygen into relict
crystals is in fact occurring.
5.4 Calculation of δ18O Values for the Ancient Hydrothermal Altering Fluid
These minerals are assumed to form in equilibrium with the fluid and therefore have a
predictable δ18O fractionation factor between fluid and mineral that is dependent on temperature,
and can be calculated using an equation by Clayton et al. (1972):
103lnαquartz-water = 3.38*(106/T2) + -3.40
Temperatures for each measured hydrothermal quartz habit were estimated using fluid inclusion
homogenization temperatures, mineral assemblages and temperature stabilities, and comparisons
to similar epithermal systems. The estimated temperatures range from 100 to 275oC gives
fractionation factors between quartz and water ranging from 20.87 to 7.85‰. Using the
measured δ18O values of quartz habits in the system and the fractionation factor, the δ18O values
of the altering fluid responsible for the precipitation of the hydrothermal quartz
68
1
2
3
4
5
6
7
-7.0 -6.0 -5.0 -4.0 -3.0 -2.0 -1.0 0.0 1.0 2.0 3.0 4.0
δ18O (‰) VSMOWA
ltera
tion
Ass
embl
age
#
magmatic quartz phenocrystshydrothermal quartz from assemblages 1-4hydrothermal quartz from assemblages 5-7
Figure 28: Graph of δ18O values of quartz versus the mineral assemblage the sample occurred in. Assemblages generally increase in intensity from 1 to 7. There are two apparent populations of hydrothermal quartz δ18O values. Higher intensity, low δ18O quartz are shown in orange field, lower intensity, higher δ18O quartz are shown in pink field. Blue arrow indicates increasing depletion in relict magmatic quartz phenocrysts as relative intensity increases.
69
was calculated:
Δ18Oquartz-water = δ18Oquartz – δ18Owater
The total range of ancient thermal waters was -10.1 to -20.9‰ (Figure 30). These are
very typical values when compared to present day thermal waters discharging in Yellowstone
National Park, which have ranges from approximately -3 to -24‰ (Parry and Bowman, 1990;
Ball et al., 2002; Kharaka et al., 2002). The ancient hydrothermal fluids tend to separate into two
distinct fluid ranges, -17.4 ± 2.2 ‰ (n = 30), and -12.2 ± 1.4 ‰ (n = 43). These ranges
correspond to the two populations of quartz values. The high intensity, low δ18O quartz give
ancient altering fluids in the -12.2 ± 1.4 ‰ range. The low intensity, high δ18O quartz give
calculated ancient fluid values within the other range -17.4 ± 2.2 ‰.
Using these calculated ancient thermal water δ18O values and the measured hydrothermal
quartz values, the water-rock ratio was calculated for the Sevenmile Hole area (Figure 31). The
evolution of altering fluid begins with average present day meteoric δ18O values of -18‰ and is
shifted to higher δ18O values due to exchange with host tuff. The average whole rock δ18O
values are shifted to lower δ18O values due to exchange with the altering fluid. The δ18O values
for hydrothermal quartz that would precipitate in equilibrium with the evolving altering
hydrothermal fluid, shows an evolution towards higher δ18O values due to the increasing altering
fluid δ18O values. The range of measured hydrothermal quartz δ18O values from -5.7 to 1.3‰
suggest that the system evolved along with this water-rock ratio as fluid and rock interacted.
Combinations of (1) temperature fluctuation possibly due to depth, boiling, or recharge, (2)
variability in the δ18O values of the hydrothermal fluid due to recharge or boiling, (3) surface
area and volume of rock available for oxygen diffusion or evolution of the water-rock ratio, (4)
70
Figure 29: Graph showing meteoric water line from Yellowstone National Park (Kharaka et al., 2002). Also showing various ranges of measured and calculated waters in the park. The rock and arrows are to indicate the δ18O shift in rocks that occurs during hydrothermal alteration as fluid and host rock interact. δ18O values reported in VSMOW.
71
Figure 30: Graph of water-rock ratios for Sevenmile Hole hydrothermal system. Evolution of altering fluid (blue line) starts out at average meteoric δ18O values of -18 ‰ and is shifted to higher δ18O values due to oxygen exchange with host tuff. In brown the average whole rock δ18O values are shown shifting to lower δ18O values due to exchange with altering fluid. Also shown in red is the line for hydrothermal quartz that would precipitate in equilibrium with the evolving altering hydrothermal fluid. The two calculated ancient hydrothermal fluid ranges are shown in pink and orange.
-12.2 ± 1.4 ‰
-17.4 ± 2.2 ‰
72
and permeability of different modes of transport of fluids are likely responsible for this variable
span. Overall these conditions can be characterized as intensity of alteration on a relative scale.
The alteration assemblages exert a greater control over δ18O values in the system, because factors
such as localized heat and fluid upflow zones add a lateral component to the alteration intensity.
Assemblages 5 and 6 are more pervasively altered and are assumed to form within the fluid
upflow zone and therefore give higher calculated altering fluid δ18O values which indicates more
evolved water-rock ratios of 1 to 1.5. In shallow and outer zones of alteration such as
assemblages 1, 2, 3, and 4 where alteration is not as pervasive δ18O values tend to be closer to
present day fresh meteoric water values (-15 to -20 ‰) which likely represents a less evolved
lower water-rock ratio because less fluid has circulated in contact with the tuff in these less
altered zones.
73
Chapter 6: TIMING AND PARAGENESIS OF THE SEVENMILE HOLE
HYDROTHERMAL SYSTEM
The timing of the ancient alteration in Sevenmile Hole can be constrained initially by the
emplacement of the host, the TSC at 479 ka, and incision of the canyon which ended the
alteration found in the upper elevations of the present day canyon walls.
6.1 Incision of the Grand Canyon of the Yellowstone River
Incision of the canyon is thought to have occurred before the youngest period of
glaciation (Pinedale Glaciation from 11-75 ka) due to the presence of glacial gravel deposits on
lower river terraces within the canyon (Pierce, 1974). Additional evidence is the presence of
glacial gravel deposits, hot spring features, and hydrothermal cement along the rim of the
canyon, stratigraphically above the TSC, that have not been removed by erosion during the most
recent glacial period (Christiansen, 2001; Licciardi and Pierce, 2008). It is probable that
outwash from the end of the older Bull Lake glaciation (~130 ka) initiated incision of the
canyon. As the canyon incised the groundwater table dropped along with the topography,
effectively ending the alteration in the hydrothermal system at high elevations and causing the
progressive migration of sinter precipitation and alteration to lower elevations. The erosion of
rock material would also cause a drop in pressure and temperatures in the hydrothermal system
at depth, which may trigger the late stage chalcedony veining from the cooling of a silica
saturated fluid or brecciation from newly over-pressurized fluids (Muffler et al., 1971). Once the
canyon incised the alteration became extinct, hence the designation “ancient” hydrothermal
alteration.
74
6.2 Age of Alteration
One 40Ar/39Ar age of alteration of 0.154 ± 0.016 Ma from a sample of tabular alunite was
obtained from the Sevenmile Hole field site (Larson et al., 2009) and is bracketed within the Bull
Lake glaciation (127-170 ka).
6.3 Higher than Reference Boiling Point Curve Temperatures
Fluid inclusion homogenization temperatures and hydrothermal mineral assemblages in
the canyon walls at Sevenmile Hole indicate the alteration had to have formed prior to canyon
incision. In order for high temperature to reach shallow levels in the hydrothermal system, there
must have been additional mass above the system contributing to the hydrostatic pressure.
Evidence for this is provided by theoretical temperature depth relationships according to Haas’s
boiling point curve. The present day elevation of the canyon rim is used to approximate the
elevation of the paleosurface during ancient alteration. (1) The dated sample of alunite and other
tabular matrix-replacing alunite was found at a depth of 70 m below the surface. At this depth
according to the reference boiling point curve temperatures should be 160o to 170oC (Haas,
1971). Matrix replacing alunite in this tabular form also suggests higher temperatures of
formation ranging from approximately 200o-300oC from research on other acid-sulfate systems
(John et al., 2005). (2) Fluid inclusion homogenization temperatures from prismatic quartz vugs
sampled from the same outcrop as the dated alunite average 230oC. Other measured fluid
inclusion homogenization temperatures indicate many of the primary inclusions within vug
quartz were trapped at temperatures as much as 70oC higher than predicted by the reference
boiling point curve (Figure 25). Higher than reference boiling point curve temperatures from
fluid inclusions in hydrothermal minerals are apparent in many of the Yellowstone drill cores
(White et al., 1975; Keither and Muffler, 1978; Bargar and Muffler, 1982; Barger and Beeson,
75
1985; Bargar and Fournier, 1988). There must be addition hydrostatic pressure from a water
column above the location of mineralization to account for the suggested temperature of
formation. A water column height of 460 m above the present day canyon rim is necessary to
cause an upward shift in the boiling point curve that would constrain the fluid inclusion and
suggested mineral assemblage temperatures.
6.4 Cause of the Additional Hydrostatic Head
It is improbable, although possible, that erosion has removed 460 m of rock material.
The current rim elevation is likely approximately the elevation of the paleosurface in the region
since the TSC is a relatively young unit and is overlain by the co-eruptive Canyon Flow (CF) and
the younger Dunraven Road Flow (DRF). No younger units in the area have been recognized.
There is a contact between TSC and the DRF where the CF is missing that lies at an elevation of
~2490 m (8150 ft) which is approximately the elevation of the present day canyon rim. Unless a
substantial thickness of DRF or CF was removed from above the TSC at Sevenmile Hole, the
elevation of the present day rim likely approximates the paleosurface. In Sulphur Creek canyon
massive opaline sinter deposits that would have been deposited at or just below the paleosurface
lie at an elevation just above the present day rim of the canyon. The sinter is cut by the erosion
of Sulphur Creek which likely incised around the same time as the main Grand Canyon. This
alteration therefore must have occurred before the canyon incised, during the main phase of
sulfide and quartz mineralization and brecciation in Sulphur Creek canyon, and therefore
suggests this elevation was the paleosurface during alteration. The additional hydrostatic
pressure may therefore be due to a massive glacial ice sheet. During glacial periods the weight
of ice above the hydrothermal system contributes to an increase in hydrostatic head pressure felt
by the fluids in the system at depth. This will shift the reference boiling point curve to a
76
shallower depth (Bargar and Fournier, 1988) and allow higher temperature conditions at
shallower levels in the hydrothermal system, and therefore, affect the alteration mineralogy.
The temperatures of hydrothermal quartz growth from fluid inclusion data is modeled
using the build up of glacial ice as the cause of additional hydrostatic pressure to the system
overhead (Figure 31). For the sake of calculations the density of both water at or near boiling in
the epithermal system and the ice of the glacier was assumed to be 0.99 g/cm3. Although these
phases have different densities in the thousandth decimal place, the difference is negligible for
these calculations. Most homogenization temperatures tend to fit between the boiling point
curves generated assuming present day canyon elevation at 2470 m (8100 ft) as the paleosurface,
and a reference curve assuming the addition of 460 m of hydrostatic head due to a glacial ice
sheet and possibly the additional height of rock material (elevation of which is approximately
2930 m). The 154 ka alunite alteration and all assemblages associated with this likely formed
during the Bull Lake glacial period which spanned from ~127-170 ka (Liccardi and Pierce,
2008). Evidence suggests that the Yellowstone plateau ice cap may have reached an elevation of
3300 m during the youngest Pinedale glacial period (Pierce et al., 2007). Bargar and Fournier
(1988) also found that glacial ice may have reached a thickness of up to 730 m, so an ice sheet of
460 m above the surface which would reach an elevation of 2930 m is plausible for the Bull Lake
glacial period.
77
Figure 31: Model of the effect of glaciation on the hydrothermal system at Sevenmile Hole. Green bars represent ranges of fluid inclusion homogenization temperatures from primary quartz FIA. Two Haas (1971) reference boiling point curves are plotted in pink, the lower one assuming 2470 m (present day canyon rim elevation) is the upper limit in the hydrothermal system, the upper curve assumes the addition of 460 m of glacial ice.
78
6.5 Paragenesis: Structural Cause and Timing of the Alteration The hydrothermal alteration at Sevenmile Hole is likely the result of a caldera ring fault
or slump block fault channeling deeper hydrothermal fluids at depths >350 m below the basal
vitrophyre of TSC. The near surface alteration (<350 m) in the TSC appears to be initially
controlled by the contact of the TSC with the pre-caldera Absaroka Volcanics. Although the
LCT may underlie the TSC within the caldera stratigraphically, the TSC slumps into the caldera
creating a more vertical contact with older Absaroka Volcanics in this region, as indicated by the
slump fold in the canyon walls. The contact may provide a permeable conduit for the ascent of
hydrothermal fluids. The near surface (within 350 meters depth) expression of this alteration in
the region with the most intense alteration lies in the Sulphur Creek canyon stream bed.
Evidence that this area may have experienced more intense alteration is indicated by large vein
size, more hydrothermal brecciation, and higher sulfide deposition in the region. Erosion of the
Sulphur Creek canyon follows the contact between the Absaroka Volcanics in the canyon walls
on the north side and the TSC composing the walls to the south side. Although the alteration has
not been dated in this region it was likely was the initial zone of hydrothermal fluid upflow due
to its higher permeability: the contact (Figure 33). A high initial permeability is also indicated
by the cap of massive sinter at the highest elevations, which was precipitated at or just below the
paleosurface. Massive surficial silicification indicates relatively rapid fluid upflow rates since a
high temperature must be maintained to precipitate silica at shallow levels (White et al., 1971).
The base of the TSC that makes contact with the Absarokas is vitrophyric and is exposed along
the caldera rim on the north side of Sulphur Creek (Figure 33). Vitrophyric amorphous glass
reacts and dissolves fairly easily, which may cause a silica over-saturation in the hydrothermal
79
fluid (hence the massive sinter deposition at the paleosurface), and may also increase
permeability along the contact.
The alteration in the TSC becomes progressively weaker farther south from the contact.
Although pervasive in some areas, veins are generally much smaller in diameter, less frequent,
and less brecciated, hydrothermal breccias are not as prevalent, and sulfide content is less. The
first zone in Sevenmile Hole moving south from the contact is composed of kaolinite + quartz +
alunite (alteration assemblages 5) above 100m, and illite + quartz ± hydrothermal feldspars
(assemblage 6 and 7) below 100 m. This zone contains the most intense alteration in the
Sevenmile Hole field area suggested by higher temperatures based on mineral assemblages and
fluid inclusion, obliteration of original tuff textures, more negative δ18O values of hydrothermal
quartz, and higher calculated δ18O values of ancient altering fluids (-12.2 ± 1.4‰), which
indicate more evolved water-rock ratios.
Moving further from the contact the alteration intensity continues to decrease, the
mineralogy in the next zone is mainly silica phases (quartz and opal) and clay (assemblages 1, 2,
and 3). Localized zones of silicification are prominent as erosion resistant ridges. These
assemblages suggest lower intensity because igneous feldspars are still present, as well as
original tuff textures, δ18O values of hydrothermal quartz are higher, and calculated δ18O values
of altering fluids (-17.4 ± 2.2‰) are closer to present day meteoric waters, which indicate less
evolved water-rock ratios.
The outermost zone of the alteration that grades into unaltered sections in Sevenmile
Hole has been affected by river erosion. This zone appears to be predominantly altered to clay
from visual estimates. Due to the preferential erosion and high clay content this zone is not well
sampled, it mostly composes scree fields of loose material.
80
All of these zones of alteration are assumed to have formed before the Grand Canyon
incised. It seems probable that the initial focus of hydrothermal fluid flow was channeled along
the highly permeable contact of the TSC and the Absarokas Volcanics. As the hydrothermal
system precipitated silica at the surface and at depth it may have begun to self seal along this
zone which perhaps caused the migration of the system from the contact into the less permeable
welded TSC within the present day Sevenmile Hole. The alteration of assemblages 5 and 6 may
be initiated by permeability caused by hydrothermal brecciation at depth which are found in two
locations around Ridge 7741. Hydrothermal breccias are the result of over-pressurization of
heated fluids (well above hydrostatic pressure) beneath partially self-sealed and semi-permeable
hydrothermal cap rocks (Sillitoe, 1985; Hedenquist and Henley, 1985). This over-pressurization
of fluids at depth may also be accomplished by the rapid release of overhead pressure from the
removal of glacial ice (Muffler et al., 1971; Hedenquist and Henley, 1985). Boiling separation
of water vapor and other gases such as H2S would cause the residual fluid to become
oversaturated in silica (silica is not transported in vapor) and increase pH, both of which favor
the precipitation of silica phases (Muffler et al., 1971; Nordstrom et al., 2009). Evidence of this
is apparent in the hydrothermal breccia sample YS-07-77, in which the clasts are all altered TSC
and cemented with a matrix of massive silica (Figure 32). This breccia lacks fresh clasts but
contains all pervasively altered clasts suggesting the hydrothermal system had been active well
before the brecciation event. This breccia displays minimal clast transport due to the jigsaw fit
of the angular clasts which suggests this may be a hydraulic fracture in-situ breccia that formed
in place as described by Davies et al. (2008).
Overpressures at depth to trigger hydrothermal brecciation may also be caused by
trapping of vapor below an impermeable cap (White, 1971). The impermeable cap could be
81
either the densely welded TSC or the silica sealed contact to the north. It is difficult to tell
whether the brecciation inititates the start of more intense alteration in the Sevenmile Hole
vicinity during increased pressure from Bull Lake glaciation and explosion at depth due to self-
sealing, or if the brecciation occurs at the end of the glacial period due to the deglaciation. The
melting of the Bull Lake glacial period and resulting rapid removal of overhead pressure may
also have triggered the hydrothermal brecciation below Ridge 7741. However the latter is
suggested by pervasively altered clasts composing these breccias, indicating the tuff experienced
alteration before brecciation.
The fluid inclusion temperatures and hydrothermal mineral assemblages in the canyon
walls at Sevenmile Hole are alteration that had to have formed prior to canyon incision, when
high pressures and temperature reached shallow levels in the hydrothermal system. In order for
temperatures to occur as high in elevation as they do, there must be addition hydrostatic pressure
from a water column above their location. Evidence indicates that the necessary pressure may be
as much as 460 m to cause the upward shift in the boiling point curve. Either there was
additional thickness of rock since removed by erosion, or the weight of a glacial ice sheet caused
the increase in hydrostatic pressure. The presence of the massive sinter deposits at the current
rim elevation in Sulphur Creek Canyon, assumed to be associated with the pre-canyon incision
alteration, suggests this elevation is the paleosurface during this stage of alteration. The
additional hydrostatic pressure must therefore be due to a massive glacial ice sheet. This stage of
alteration which includes the 0.154 ± 0.016 Ma (Larson et al., 2009) alunite sample occurs
during formed during the Bull Lake glacial period which spanned from ~127-170 ka (Liccardi
and Pierce, 2008), which accounts for the anomalously high temperatures at shallow depths.
82
It is probable that the outwash from the end of the Bull Lake glaciation (~130 ka) is
responsible for the start of incision of the Grand Canyon of the Yellowstone River. Licciardi and
Pierce have shown that the canyon is definitely pre-most recent glaciaion (Pinedale from 11-75
ka). As the canyon incised the groundwater table dropped along with the topography, effectively
ending alteration in the hydrothermal system at high elevations and causing the temporal
downward migration of sinter precipitation to lower elevations. Once the canyon incised the 154
ka alteration was preserved and the system migrated to deeper levels around river elevation.
This younger and currently active hydrothermal alteration (Figure 33). The evidence that this
system is overprinting deeper alteration is (1) pseudocubic alunite that forms in steam-heated
environments at lower temperatures (~100oC), (2) presence of opal/surficial sinter precipitation,
(3) occurance of kaolinite and illite in the same sample (YS-07-31b and 32, see Appendix C),
and (4) massive chalcedony veins of assemblage 7 that cross cut illite alteration in the deepest
elevations in the system, but probably formed during water table drop and cooling.
83
Figure 32: Model diagram of zones of alteration in Sevenmile Hole and associated structure and fluid paths.
84
Chapter 7: Summary of Conclusions
The major conclusions of this work are summarized here:
1. A buried structure such as a caldera ring fault may be responsible for the localized alteration
seen in Sevenmile Hole. This structure provides the conduit for deep hydrothermal fluids below
the TSC below a depth of 350 m. In the shallow <350 m environment the contact provides the
channel. Minor structures or fractures which are upwelling zones can be recognized by the
silicified ridges. The crudely linear alignment of surficial hot spring activity and zones of most
intense alteration may reflect the channelized upflow of hydrothermal fluids above this deeper
structure.
2. The relationship between depth of the groundwater table and topographic land surface is a
controlling factor on alteration mineralogy and surficial discharge. The lowering of the water
table and incision of the canyon has effectively starved higher elevations of the system of an
altering fluid. Surficial sinter at highest elevations is extinct as far as water discharge but is
discharging sulfuric gas as fumaroles. At intermediate depths the surficial water discharge is
intermittent seasonally and annually. At lowest elevations, the hot spring activity is more intense
due to a shallow ground water table which even intersects the land surface (i.e., the river). Sinter
precipitation has progressed down in elevation as the canyon incised.
3. The rocks in Sevenmile Hole show both acid-sulfate and neutral-pH type alteration chemistry
due to the presence of the mineral assemblages of quartz + kaolinite ± alunite ± dickite (resulting
from acid-sulfate fluids) and quartz + illite ± adularia along with sinter deposition at the surface
(resulting from neutral-pH fluids). These two alteration chemistries are apparent by a vertical
distribution of clay minerals in the walls of the canyon. Kaolinite is found at shallow depths
(<100m) in the hydrothermal system and illite is found at deeper levels (>100m). The occurance
85
of acid-sulfate environments overlying neutral-pH environments can be explained as one coeval
hydrothermal system instead of temporal changes in hydrothermal fluid chemistry. The
precipitation of kaolinte above 100m versus illite below is controlled by pH and the evolution
and speciation of vapors (CO2, H2S, SO2). Below 100 m and approximately 200oC where CO2
and carbonates are in equilibrium the system remains neutral in pH. Above 100 m carbonates
are no longer stable and CO2 is lost by degassing, and so the sulfuric vapors condense in
groundwater to produce H2SO4 which reacts with host rock and leaches vugs and precipitates
kaolinite.
4. Decreasing alteration intensity in zones moving away from the caldera margin and contact
between the pre-caldera Absaroka Volcanics and the post-caldera TSC. Next to the contact is an
area
5. Fluid inclusion homogenization temperatures in prismatic vug quartz indicate that most
crystal growth occurred on average between 180o and 280oC. Other inclusions indicate that
temperatures of fluids in the system may have ranged from 160o to 350oC. Homogenization
temperatures generally increase with depth in the system.
6. Homogenization temperatures often lie well above the boiling point curve and temperature
that would be expected at that particular depth. Mineral assemblages also suggest precipitation
at higher than reference boiling point curve temperatures for that depth. This shift to higher
temperatures at shallow levels can be attributed to additional hydrostatic head above the
hydrothermal system either due to land surface that has since eroded away or the mass of a
glacial ice sheet.
7. Meteoric water is the dominant fluid involved in the circulation and alteration of the TSC.
All hydrothermal quartz habits measured have more negative δ18O values (ranging from 1.3 to -
86
5.7‰) than magmatic quartz in the fresh TSC at 1.7‰. These negatively shifted δ18O values
reflect formation in equilibrium with meteoric water, the only large negative reservoir of δ18O on
the planet. The alteration is intense enough to cause magmatic quartz phenocrysts in the host
TSC to show a slight depletion in δ18O values of approximately 1‰ due to diffusion of oxygen
into the crystal or recrystallization and overgrowth.
8. Matrix silicification and prismatic vug quartz have narrow ranges of δ18O values from -3.1 to
-4.8‰. Quartz veins, minor vugs, and quartz replaced pumice fiamme, had δ18O values that span
the entire measured range (1.3 to -5.7‰). This suggests the δ18O values of the system evolved
along with the water-rock ratio.
9. Calculated ancient thermal waters responsible for the precipitation of hydrothermal quartz
ranged from δ18O values of -10.8 to -19.6‰.
10. Two distinct calculated δ18O values of ancient altering fluid are recognized when
considering the alteration intensity of each mineral assemblage. The first is a low δ18O quartz,
high calculated δ18O altering waters (-12.2 ± 1.4‰), and therefore high intensity assemblage,
including assemblages 5, 6, and 7. These are likely hydrothermal fluid upwelling zones with
higher and more evolved water-rock ratios of 1 to 1.5. Hydrothermal quartz occurs as massive
matrix replacement, large prismatic vugs, and veins. The second grouping of calculated ancient
fluid values is lower δ18O of altering waters (-17.4 ± 2.2‰), the quartz δ18O values are higher
and silicification and intensity of alteration is generally weaker. This includes assemblages 1, 2,
3, and 4. The higher, closer to meteoric values, are likely less evolved thermal waters, and have
lower calculated water-rock ratios. Silicification is mostly igneous phenocryst and fiamme
replacement, and minor veinlets to veins, and often textures of the tuff are preserved.
87
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Appendix A:
Field Site Groups and
PIMA, XRD, Fluid Inclusion, and δ18O Sample Locations
98
99
Appendix B:
Field Site Descriptions
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Field Site #1: “Alunite Outcrop” Description and Large Structures: This location is massive outcrop along Sevenmile Hole trail in the upper part of the system. The linear feature that has resisted erosion and is a “rib” that spans downward in elevation for approximately 50m vertically. Mineralogy and Textures: Tabular alunite replaces matrix of rock. An 40Ar-39Ar age date was obtained from a sample of alunite from this location (Larson et al., 2009). Large vugs filled with prismatic quartz are prevalent as shown in photograph to the right above.
101
Field Site #2: “Hill 7741” Description and Large Structures: Hill in the middle of the Sevenmile Hole altered area that is resistant to erosion due to massive silicification. Mineralogy and Textures: All pervasively altered, leached and bleached to white or Fe oxide stained. Dominant clay is illite, with some kaolinite and opal overprinting. Vuggy textures as shown in upper right. Very large quartz vugs with massive fill and prismatic crystals shown in bottom right. A hydrothermal explosion clast-supported breccia with altered clasts and massive silica cement was found on the top of the hill.
102
Field Site #3: “Down by the River” Description and Large Structures: Deepest level of system sampled but well exposed by river erosion. Basal ash fall of Tuff of Sulphur Creek is present and shown in upper right. Active thermal water discharge all along the river, and active precipitation of surficial sinter cementing scree. Mineralogy and Textures: All samples are pervasively altered. Difficult to sample due to float and active hot spring activity. Most hydrothermally active (surface discharge) apparent in presence of opal and sinter cementing scree. Upper right photo shows basal ash fall and breccia of TSC which is highly altered likely due to high permeability. A ferricrete (iron oxide cemented) breccia shown in bottom right, formed in creek bed. All clasts are altered and breccia incorporates some surficial sinter clasts, suggesting formation after canyon incision.
103
Field Site #4: “Surficial Sinter Fields” Description and Large Structures:
Sinter cones, hot springs, sinter fields, and fumaroles. Mineralogy and Textures: Opaline sinter precipitated from thermal fluid discharge. *see detailed description in Chapter 3: Structure and Physical Controls on the Hydrothermal System
4.1
4.2
4.3
104
Field Site #5: “Scree field below Hill 7741” Description and Large Structures:
Large silicified ridges extending off of Hill 7741. Ridges have linear orientation. Large scree field from erosion of weaker and more clay rich altered products. Mineralogy and Textures:
Highly silicified matrices in most samples (bottom right). Iron oxide weathering. There are large veins of quartz and chalcedony up to 7 cm (bottom left). Also large vugs grow to sizes with diameters of up to 15 cm and prismatic crystals up to 2 cm. Variable alteration mineralogy includes alunite and adularia which suggest both acidic and neutral hydrothermal fluid chemistry. The clay present in mostly illite, however opal and kaolinite overprint this alteration.
105
Field Site #6: “Shallow system up Mosquito Creek” Description and Large Structures: Shallowest and highest levels of the hydrothermal system which sits on the edge of the altered zone. This area shows weaker alteration. Various silicified “ribs” running vertically from current rim to deeper in elevation. Mineralogy and Textures: Edge of alteration, the farther eastward, the less pervasive the alteration. Kaolinite is the dominant clay. The tuff has been leached and has a vuggy texture as shown in the photo to the right, however unlike deeper part of the system such as Hill 7741, the vugs contain kaolinite and opal, and tend to have more void space, not filled with hydrothermal minerals. More original rhyolitic textures are preserved here than in more intensely altered areas.
106
Field Site #7: “Deep system below Lookout Point” Description and Large Structures: Linear alignment of large structures and highly silicified vertical silicified upflow “pipes.” Mineralogy and Textures: Highly silicified vertical conduit pipes. Many large vugs with prismatic quartz. The tuff is leached and bleached to a white silica residue (bottom right).
107
Field Site #8: “Upper Scree Field” Description and Large Structures: Top of system, above Alunite Outcrop at point where Sevenmile Hole trail begins to drop into the canyon. Less pervasively silicified compared to deeper outcrops, hence the weathering of the tuff creating a large scree field. Mineralogy and Textures: Area is pervasively altered and leached. The matrix of the tuff contains more clay and less silicification. Kaolinite is the dominant clay, however dickite, a higher temperature polymorph of kaolinite, is also present. Less vuggy than deeper areas of the hydrothermal system. However, there is evidence of silica saturated fluid flow due to the presence of 0.2mm to 2cm veins of chalcedony and opal.
108
Appendix C:
Sample Locations and Alteration Mineralogy
109
Table 3: List of sample, location, elevation, and mineralogy determined by both PIMA and XRD methods. Only minutes are listed for latitude and longitude. The degrees for each sample site in the Sevenmile Hole field area are 44oN latitude and 110oW longitude. Ass. # indicates alteration mineral assebmblage designation used in Table 2. The asterisk next to sample # indicates the PIMA and XRD patterns are shown in Appendix D.
sample # lat long elev. (ft)
elev. (m) PIMA minerals XRD Minerals Ass.#
YS-07-1* 45.157 25.605 7980 2432 alunite alunite, quartz 5
YS-07-2* 45.157 25.605 7980 2432 alunite, water alunite, quartz 5
YS-07-3 45.157 25.605 7980 2432 montmorillite quartz, kaolinite 5
YS-07-4 45.395 25.635 8090 2466 opal, water 1
YS-07-5 45.395 25.635 8090 2466 water 1
YS-07-6 45.395 25.635 8090 2466 quartz 2
YS-07-7 45.180 25.546 7940 2420 montmorillite chalcedony, sanidine, illite, gypsum 7
YS-07-8* 45.180 25.546 7940 2420 illite/smectite illite, chalcedony/opal 7
YS-07-9 45.180 25.546 7940 2420 kaolinite, water, possible alunite 1/4
YS-07-10 45.144 25.302 7460 2274 montmorillite
YS-07-12b 44.810 25.944 8055 2455 kaolinite 1
YS-07-12c 44.810 25.944 8055 2455 kaolinite 1
YS-07-12d 44.810 25.944 8055 2455 kaolinite 1
YS-07-13 44.810 25.944 8055 2455 kaolinite, water 1
YS-07-14 44.804 25.892 8020 2444 kaolinite, water 1
YS-07-15* 44.932 25.819 7910 2411 kaolinite, water kaolinite, qtz, opal 1/2
YS-07-16* 44.941 25.885 7880 2402 kaolinite, water kaolinite (dickite, nacrite), tridymite 3
YS-07-17 44.941 25.885 7880 2402 kaolinite, water 1
YS-07-18* 44.937 25.916 7915 2412 kaolinite, minor water kaolinite, qtz, tridymite, sanidine, microcline, hyalophane
1/2
YS-07-19* 44.937 25.916 7915 2412 kaolinite, water kaolinite, gypsum 1
YS-07-20* 44.935 25.946 7945 2422 kaolinite, water kaolinite, tridymite 1
YS-07-21 44.935 25.946 7945 2422 montmorillite opal C, barite?, sanidine/ microcline 1
YS-07-22 44.962 26.007 7890 2405 kaolinite, water 1
YS-07-23* 44.959 26.031 7930 2417 kaolinite, water kaolinite, dickite, tridymite, gypsum 3
YS-07-24 44.959 26.031 7930 2417 kaolinite, water 1
110
sample # lat long elev. (ft)
elev. (m) PIMA minerals XRD Minerals Ass.#
YS-07-25* 44.959 26.031 7930 2417 kaolinite, water, minor alunite
opal CT, kaolinite, orthoclase, sanidine 1
YS-07-26a* 44.967 26.060 7905 2409 kaolinite, water tridymite, dickite, kaolinite 3
YS-07-26b 44.967 26.060 7905 2409 kaolinite, water 1
YS-07-27 44.976 26.157 7906 2410 kaolinite, water 1
YS-07-28a 45.024 26.086 8020 2444 kaolinite, water 1
YS-07-28b 45.024 26.086 8020 2444 kaolinite, water 1
YS-07-28c 45.024 26.086 8020 2444 kaolinite, water, minor smectite 1
YS-07-28d 45.024 26.086 8020 2444 kaolinite, water 1
YS-07-29 45.047 24.587 6920 2109 montmorillite, jarosite
YS-07-30 45.158 24.357 6921 2110 water cristobalite, tridymite 1
YS-07-31a 45.047 24.587 6920 2109 kaolinite, alunite tridymite, qtz, sanidine, kaolinite, orthoclase, 3
YS-07-31b 45.047 24.587 6920 2109 montmorillte, possible kaolinite
qtz, orthoclase, microcline, kaolinite, illite 2
YS-07-31c 45.047 24.587 6920 2109 kaolinite 2
YS-07-32 45.023 24.576 6880 2097 illite kaolinite & illite 6
YS-07-33 45.017 24.612 6920 2109 kaolinite, water 1
YS-07-34 45.017 24.612 6920 2109 kaolinite, water, possible alunite alunite, walthierite 4
YS-07-35 45.002 24.620 6920 2109 kaolinite, water 1
YS-07-36a 44.945 24.671 6900 2103 kaolinite, jarosite quartz, orthoclase, sanidine, kaolinite 2
YS-07-36b 44.945 24.671 6900 2103 kaolinite, water, possible smectite
qtz, microcline, kaolinite, sanidine/orthoclase, 2
YS-07-37 44.939 24.707 6910 2106 montmorillite
YS-07-38 44.942 24.771 6945 2117 kaolinite, water 1
YS-07-40 44.977 24.620 6985 2129 illite 6
YS-07-41 44.977 24.620 6985 2129 Opal 1
YS-07-42 45.115 25.252 7560 2304 illite, possible alunite quartz, orthoclase,illite 6
YS-07-43 45.069 25.299 7640 2329 illite/smectite quartz,orthoclase/ microcline, hyalophane, illite
6
YS-07-44* 45.076 25.288 7660 2335 illite/smectite quartz, microcline, illite 6
YS-07-45* 45.096 25.260 7610 2320 illite illite 6
YS-07-46 45.056 25.331 7615 2321 illite, chalcedony, muscovite 7
YS-07-47a 45.049 25.383 7640 2329 illite/smectite, possible alunite
qtz, hyalophane, buddingtonite, orthoclase, illite
6
111
sample # lat long elev. (ft)
elev. (m) PIMA minerals XRD Minerals Ass.#
YS-07-47b 45.049 25.383 7640 2329 illite/smectite, possible alunite
qtz, microcline, hyalophane, illite 6
YS-07-48* 45.034 25.443 7740 2359 illite qtz, hyalophane, buddingtonite, microcline, illite
6
YS-07-49 45.037 25.445 7740 2359 illite/smectite 6
YS-07-50 45.031 25.521 7560 2304 illite/smectite 6
YS-07-50y 45.031 25.521 7560 2304 illite/jarosite 6
YS-07-51 45.025 25.522 7540 2298 opal, water quartz 6
YS-07-52* 44.989 25.472 7535 2297 illite, water qtz, hyalophane, microcline, illite 6
YS-07-53* 44.982 25.443 7590 2313 illite quartz, microcline, illite 6
YS-07-54* 44.988 25.490 7510 2289 illite, possible gypsum qtz, hyalophane, alunite?, illite 5/6
YS-07-55 45.057 26.052 8020 2444 kaolinite, water 1
YS-07-56 45.057 26.052 8020 2444 kaolinite, water, possible smectite 1
YS-07-57 45.057 26.052 8020 2444 kaolinite cristobalite, tridymite, orthoclase, sanidine, kaolinite
1
YS-07-58 45.045 26.038 7970 2429 kaolinite, water, possible smectite 1
YS-07-59 45.047 26.037 7970 2429 kaolinite, water 1
YS-07-60* 45.031 26.006 7875 2400 kaolinite, water, possible smectite quartz, kaolinite 1/2
YS-07-61 45.044 25.949 7855 2394 kaolinite, water, possible jarosite 1
YS-07-62 45.016 25.998 7820 2384 kaolinite, water, possible smectite 1
YS-07-63* 44.990 25.967 7740 2359 kaolinite, water kaolinite (dickite, nacrite), tridymite 3
YS-07-64 44.975 25.958 7780 2371 kaolinite, water 1
YS-07-65a 44.974 25.939 7760 2365 kaolinite, water 1
YS-07-65b 44.974 25.939 7760 2365 kaolinite, water 1
YS-07-66 44.984 25.978 7795 2376 kaolinite, water 1
YS-07-67 44.969 26.048 7960 2426 kaolinite, water 1
YS-07-68* 45.208 25.663 7970 2429 dickite, alunite quartz, dickite, walthierite, huangite, kaolinite 4
YS-07-69* 45.169 25.66 7920 2414 kaolinite, alunite quartz, dickite, kaolinite, alunite, walthierite 5
YS-07-70 45.169 25.66 7920 2414 kaolinite 2
YS-07-71* 45.149 25.659 7895 2406 kaolinite quartz, dickite/kaolinite, walthierite, huangite 5
YS-07-72* 45.156 25.611 7790 2374 dickite, alunite quartz, dickite, alunite, kaolinite 5
YS-07-73* 45.156 25.611 7790 2374 alunite qtz, alunite, (walthierite, huangite), kaolinite 5
112
sample # lat long elev. (ft)
elev. (m) PIMA minerals XRD Minerals Ass.#
YS-07-74* 45.156 25.611 7790 2374 alunite quartz, alunite, walthierite, kaolinite 5
YS-07-75 45.092 25.260 7640 2329 illite/smectite 6
YS-07-76 45.092 25.260 7640 2329 illite/smectite 6
YS-07-77 45.092 25.260 7640 2329 illite/smectite, possible jarosite quartz, hyalophane 6
YS-07-78 45.056 25.335 7610 2320 illite 6
YS-07-81* 45.037 25.445 7720 2353 illite/smectite quartz, jarosite, hyalophane, illite, 6
YS-07-82a 45.152 25.668 7915 2412 alunite 5
YS-07-82b* 45.152 25.668 7915 2412 alunite quartz, alunite, dickite, walthierite, kaolinite 5
YS-07-82c* 45.152 25.668 7915 2412 alunite quartz, dickite, alunite, walthierite, huangite, kaolinite
5
YS-07-83a* 45.154 25.603 7770 2368 alunite, water quartz, alunite/walthierite/ huangite, jarosite, 5
YS-07-83b 45.154 25.603 7770 2368 alunite 5
YS-07-83c 45.154 25.603 7770 2368 alunite 5
YS-07-83x 45.154 25.589 7735 2358 quartz, jarosite, alunite 5
YS-07-84 45.103 25.622 7715 2352 illite, water 6
YS-07-85a 45.053 25.635 7490 2283 illite, water 6
YS-07-85b 45.053 25.635 7490 2283 kaolinite, water 1/2
YS-07-86* 45.026 25.644 7400 2256 illite/smectite illite, chalcedony, muscovite 7
YS-07-87 45.025 25.699 7440 2268 illite, water 6
YS-07-88a 44.998 25.768 7630 2326 illite/smectite 6
YS-07-88b 44.998 25.768 7630 2326 illite, jarosite 6
YS-07-89a* 44.985 25.772 7690 2344 buddingtonite, illite quartz, microcline, hyalophane, illite 6
YS-07-89b 44.985 25.772 7690 2344 illite/smectite 6
YS-07-90 44.987 25.782 7700 2347 illite, jarosite 6
113
Appendix D:
XRD Diffraction Patterns and PIMA spectra
114
X-ray diffraction patterns for ten representative samples of kaolinite. X-axis is 2θ ranging from 0o to 60o. Kaolinite 001 peak is visible at 12.37o 2θ. kao = kaolinite, Q = quartz.
115
PIMA spectra for ten representative samples of kaolinite. Double pronged peak between 1.9 and 1.5 μm is characteristic of kaolinite.
116
X-ray diffraction patterns for ten representative samples of illite. Illite 001 peak is visible at 8.82o 2θ. il = illite, Q = quartz.
117
PIMA spectra for ten representative samples of illite. Single peak between 1.4 and 1.5 μm is characteristic of this clay.
118
X-ray diffraction patterns for ten representative samples of alunite. X-axis is 2θ ranging from 0o to 60o. Alunite Kα peak is visible at 29.97o 2θ, also diagnostic peaks at 15.56o and 17.92o 2θ. Al = alunite, Q = quartz, kao = kaolinite.
119
PIMA spectra for ten representative sample of alunite. Double peak between 1.4 and 1.5 μm is characteristic of alunite.
120
Appendix E:
δ18O Values
121
Table 4: Table of oxygen isotope δ values of various quartz habits from Sevenmile Hole altered area. The δ18O values are corrected based on standard bracketing and reported in per mil (‰) VSMOW. The standard used for correction is UWG-2 garnet. Error calculated based on standard deviation using average of δ18O values of standards. sample quartz habit δ18O 1σ error
YS-07-11 1° magmatic phenocryst 1.3 0.01
YS-07-14-1 1° magmatic phenocryst 1.8 0.08
YS-07-14-2 1° magmatic phenocryst 0.9 0.08
YS-07-14p-1 vein 6.0 0.06
YS-07-14p-2 vein 4.9 0.06
YS-07-15 1° magmatic phenocryst 1.3 0.01
YS-07-17-1 1° magmatic phenocryst 1.5 0.22
YS-07-17-2 1° magmatic phenocryst 1.4 0.22
YS-07-20 1° magmatic phenocryst 0.0 0.19
YS-07-22-1 1° magmatic phenocryst 1.4 0.11
YS-07-22-2 1° magmatic phenocryst 1.4 0.11
YS-07-24 minor vug fill 0.7 0.11
YS-07-24 minor vug fill 0.4 0.12
YS-07-24 1° magmatic phenocryst 1.2 0.07
YS-07-32-1 vein 0.9 0.24
YS-07-32-2 vein 0.7 0.24
YS-07-35 1° magmatic phenocryst 1.0 0.04
YS-07-35-1 minor vug fill 2.7 0.12
YS-07-35-2 minor vug fill 2.4 0.12
YS-07-37 minor vug fill 0.0 0.38
YS-07-37 1° magmatic phenocryst 1.7 0.33
YS-07-37-1 1° magmatic phenocryst 0.2 0.08
YS-07-37-2 1° magmatic phenocryst 0.6 0.07
YS-07-44-1 vein -2.7 0.16
YS-07-44-2 vein -2.6 0.16
122
sample quartz habit δ18O 1σ error
YS-07-47a 1° magmatic phenocryst -0.9 0.08
YS-07-47am-1 matrix silicification -3.3 0.33
YS-07-47am-2 matrix silicification -4.3 0.33
YS-07-48-1 vein -0.6 0.16
YS-07-48-2 vein -1.1 0.12
YS-07-49 prismatic vug -4.6 0.16
YS-07-49-1 1° magmatic phenocryst -0.4 0.08
YS-07-49-2 1° magmatic phenocryst 0.9 0.04
YS-07-50 minor vug fill -3.2 0.07
YS-07-50 1° magmatic phenocryst -0.2 0.22
YS-07-52 1° magmatic phenocryst -0.5 0.33
YS-07-52m-1 matrix silicification -3.6 0.33
YS-07-52m-2 matrix silicification -4.0 0.33
YS-07-54 minor vug fill -3.0 0.19
YS-07-54 minor vug fill -1.0 0.12
YS-07-54-1 1° magmatic phenocryst 0.2 0.08
YS-07-61-1 minor vug fill -0.6 0.25
YS-07-61-2 minor vug fill -0.9 0.25
YS-07-61m-1 matrix silicification -3.4 0.38
YS-07-61m-2 matrix silicification -3.1 0.38
YS-07-63 minor vug fill 0.8 0.11
YS-07-63 minor vug fill 1.2 0.24
YS-07-63-1 matrix silicification 0.5 0.16
YS-07-63-2 matrix silicification 0.6 0.16
YS-07-65-1 minor vug fill 4.0 0.07
YS-07-65-2 minor vug fill 4.4 0.07
YS-07-68-1 matrix silicification -3.1 0.45
YS-07-68-2 matrix silicification -3.4 0.45
123
sample quartz habit δ18O 1σ error
YS-07-69 vein -5.7 0.17
YS-07-7 minor vug fill -3.1 0.19
YS-07-7 minor vug fill -2.6 0.12
YS-07-71 1° magmatic phenocryst 0.6 0.22
YS-07-75m-1 matrix silicification -3.8 0.02
YS-07-75m-2 matrix silicification -3.6 0.02
YS-07-75m-3 matrix silicification -3.9 0.02
YS-07-75p-1 minor vug fill -2.7 0.13
YS-07-75p-2 minor vug fill -2.9 0.13
YS-07-75t prismatic vug -3.3 0.20
YS-07-77 vein -2.8 0.17
YS-07-81it-1 prismatic vug -0.6 0.13
YS-07-81it-2 prismatic vug -0.8 0.13
YS-07-82a-1 1° magmatic phenocryst -0.1 0.17
YS-07-82a-2 1° magmatic phenocryst -0.7 0.17
YS-07-82b vein -4.3 0.24
YS-07-84i prismatic vug -4.8 0.20
YS-07-84m matrix silicification -4.1 0.10
YS-07-84t prismatic vug -3.9 0.33
YS-07-85b 1° magmatic phenocryst 0.6 0.17
YS-07-86 vein -3.2 0.17
YS-07-9 minor vug fill -4.0 0.16
YS-07-90-1 minor vug fill 1.3 0.24
YS-07-90-2 minor vug fill 2.6 0.27
124
Appendix F:
Fluid Inclusion Measurements
125
Table 5: Fluid inclusion data for prismatic hydrothermal quartz samples. Sample number is listed in the upper left corner of each table and fluid inclusion assemblages (FIA) are numbered in the first column. Frequency refers to how many individual fluid inclusions gave homogenization temperatures within the 10oC range listed. Origin symbols are p = primary inclusions, s = secondary inclusion, and u = unknown origins. The last two columns are averages of homogenization temperatures based on frequency and origin. All inclusions began as two phase (liquid and vapor) and the final homogenization modes were all liquid. The photographs of fluid inclusions when available are numbered by sample number and FIA number.
YS-07-75 Average Th oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 170-180 1 p 213 195 180-190 1 p 190-200 5 p 200-210 3 p 240-250 3 u 250-260 2 u 2 200-210 6 p 216 210 210-220 3 p 220-230 1 p 230-240 3 s 3 190-200 2 p 214 214 200-210 2 p 210-220 2 p 220-230 3 p 230-240 1 p 4 170-180 1 u 218 218 210-220 7 p 220-230 1 p 230-240 1 p 250-260 1 u 5 190-200 1 u 231 230 210-220 1 p 220-230 5 p 230-240 5 p 240-250 1 p 250-260 2 u
126
YS-07-84 Average Th oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 180-190 1 p 230 214 200-210 1 p 210-220 2 p 220-230 3 p 240-250 3 u 300-310 1 u 2 200-210 2 p 225 225 210-220 2 p 230-240 6 p 3 170-180 2 u 175 4 170-180 2 u 229 253 180-190 3 u 230-240 1 p 240-250 1 p 250-260 7 p 260-270 1 p
7-84 FIA #2 7-84 FIA #3
7-84 FIA #4a
7-84 FIA #4b 7-84 FIA #4c
127
YS8-1 Average Th oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 180-190 6 p 190 190 190-200 3 p 200-210 1 p 2 160-170 2 p 189 184 170-180 5 p 180-190 7 p 200-210 1 p 210-220 2 p 270-280 1 u
8-1 FIA #1a
8-1 FIA #1b
8-1 FIA #2
128
YS-07-88 Average Th oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 160-170 2 p 194 194 170-180 3 p 180-190 8 p 190-200 2 p 200-210 9 p 210-220 2 p 220-230 1 p 2 190-200 7 p 199 199 200-210 2 p 210-220 1 p 3 170-180 2 p 194 190 180-190 4 p 190-200 3 p 200-210 3 p 210-220 2 p 4 170-180 3 p 188 188 180-190 4 p 190-200 5 p 200-210 1 p
129
YS8-17 Average Th
oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 220-230 2 u 240 230-240 4 u 240-250 2 u 250-260 1 u 260-270 1 u 2 230-240 2 p 245 245 240-250 1 p 250-260 2 p 3 240-250 1 p 269 269 250-260 2 p 260-270 9 p 270-280 16 p 4 220-230 4 u 239 230-240 5 u 240-250 6 u 280-290 1 u
8-17 FIA #2a
8-17 FIA #2b
130
YS8-18 Average Th
oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 200-210 3 u 261 210-220 1 u 230-240 1 u 250-260 4 u 280-290 6 u 310-320 1 u 320-330 1 u 2 210-220 1 u 259 230-240 1 u 250-260 2 u 260-270 1 u 270-280 3 u 280-290 1 u 3 210-220 1 p 240 240 230-240 3 p 240-250 7 p 4 220-230 1 p 251 251 230-240 5 p 240-250 2 p 250-260 5 270-280 4 5 260-270 3 p 282 282 270-280 1 p 280-290 2 p 290-300 2 p 300-310 1 p
8-18 FIA #1
8-18 FIA #3
131
YS8-19 Average Th
oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 220-230 1 p 248 237 230-240 9 p 240-250 3 p 300-310 1 u 340-350 1 u 2 210-220 4 p 220 220 220-230 1 p 230-240 1 p 3 200-210 2 p 224 224 210-220 11 p 220-230 12 p 230-240 3 p 240-250 4 p
8-19 FIA #3a 8-19 FIA #3b 8-19 FIA #3c
8-19 FIA #1
8-19 FIA #2
132
YS8-24 Average Th
oC
FIA
homogenization temperature
(Th)oC frequency origin all FIA primary FIA 1 230-240 8 s 240 240-250 3 s 250-260 2 s 2 240-250 3 p 264 264 250-260 2 p 260-270 2 p 270-280 4 p 290-300 1 p 3 200-210 1 p 219 219 210-220 11 p 220-230 6 p 230-240 1 p 4 190-200 2 u 278 278 260-270 1 p 270-280 6 p 280-290 4 p 310-320 1 u 340-350 2 u 5 210-220 1 u 261 265 260-270 3 p 290-300 1 u 6 230-240 8 s 240 240-250 3 s 250-260 2 s
