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thesis entitled
PETROCHEMISTRY OF THE QUINN CANYON SILICIC

INTRUSIVES, NYE COUNTY, NEVADA: A COMPARISON
WITH CLIMAX-TYPE PORPHYRY MO SYSTEMS

presented by

James S. Guentert

has been accepted towards fulfillment
of the requirements for

 

 

Masters degree in Geologx

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PETROCHEMISTRY OF THE QUINN CANYON SILICIC
INTRUSIVES, NYE COUNTY, NEVADA: A COMPARISON
WITH CLIMAX-TYPE PORPHYRY Mo SYSTEMS
By

James S. Guentert

A THESIS

Submitted to
Michigan State University
in partial fulfillment of the requirements
for the degree of

MASTER OF SCIENCE

Department of Geological Sciences

1988

Qua- coau(

ABSTRACT

PETROCHEMISTRY OF THE QUINN CANYON SILICIC
INTRUSIVES, NYE COUNTY, NEVADA: A COMPARISON
WITH CLIMAX-TYPE PORPHYRY Mo SYSTEMS
By

James S. Guentert

The Quinn Canyon Range (QC) has anomalous concentrations
of Mo, W, Sn, Pb and fluorite mineralization associated with
the emplacement of 23 m.y. high silica rhyolite intrusives.
This investigation focuses on the comparison of QC dikes and
plugs with the rhyolitic compositions associated with porphyry
Mo, W, Sn, and lithophile element metallization.

The most evolved QC rhyolites have lower Rb, Th, and Nb,
higher Ba and Sr and smaller La/Ybu ratios than the
granitoids at Climax, Urad-Henderson and Mt. Emmons (except
Rb). The QC rhyolites are chemically more similar to the
rhyolites generating Mo mineralization at Pine Grove, Utah.
Based on the elements analyzed (no F or Cl) the QC rhyolites
seem to qualify as topaz rhyolites.

Because of general similarities with Pine Grove, the
possiblity of Mo mineralization at Quinn Canyon cannot be
discounted at this stage of study. Differences in trace
element chemistry and source material with Climax and
Brad-Henderson granitoids suggest that any Mo mineralization

probably would be uneconomic.

ACKNOWLEDGEMENTS

I would like to thank my advisor, John Wilband for all his
help, encouragement, and advise during the preparation of this
thesis. His willingness to take the time (even on weekends)
to answer questions and assist me in any way is greatly
appreciated. I am particularly indebted to him for his help
in the field. I also would like to acknowledge the interest
and assistance of committee members Dr. Tom Vogel and Dr. Bill
Cambray.

Special thanks go to my good friend Dave Westjohn for
suggesting this study, reviewing and critiquing the thesis and
in general offering his encouragement and advise when
necessary.

I would like to express my gratitude to the friends I have
made during my tenure at MSU. Among these are: Don Jessup,
Dave Cook, John Salvino, Steve Young, Bob Cuniff, Soo Meen
Wee, Jim Mills, Jim Call, Paul Carter and especially Bill
Sack.

My most sincere and heartfelt thanks go to my wife and
best friend Linda.

Amoco Minerals Co. supplied exploration geochemical data
for this study. Financial support was provided through a
Chevron Graduate Fellowship and field work award and an MSU

Teaching Assistantship.

ii

TABLE OF CONTENTS

LIST OF FIGURES . . . . . . . . . . . . . . . . . . . . . v

LIST OF TABLES . . . . . . . . . . . . . . . . . . . . . viii

INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . .1
Problem . . . . . . . . . . . . . . . . . . . . . . . . 1
Method .

PREVIOUS WORK .

FIELD WORK .

GEOLOGY OF THE QUINN CANYON RANGE .

COCOQKIU'I

Sedimentary rocks .
Igneous rocks . . . . . . . . . . . . . . . . . . . . .10
Structure . . . . . . . . . . . . . . . . . . . . . . .11
Economic geology . . . . . . . . . . . . . . . . . . . 12
PETROGRAPHY . . . . . . . . . . . . . . . . . . . . . . . .17
Mineralogy and texture . . . . . . . . . . . . . . . . 17
Alteration . . . . . . . . . . . . . . . . . . . . . . 20
PETROCHEMISTRY . . . . . . . . . . . . . . . . . . . . . . 22
Major elements . . . . . . . . . . . . . . . . . . . . 22
Alteration . . . . . . . . . . . . . . . . . . . . . . 22
Trace elements . . . . . . . . . . . . . . . . . . . . 27

Fluorite . . . . . . . . . . . . . . . . . . . . . . . 40

iii

REGIONAL GEOLOGY .
Lineaments .
Cenozoic igneous activity in the western U.S.
Tertiary metallization in the western U.S.
CLIMAX-TYPE PORPHYRY MOLYBDENUM DEPOSITS .
Geology .
Mineralogy and alteration .
Petrochemistry of granitic rocks .
Origin of granitic rocks .
DISCUSSION AND RESULTS .
Quinn Canyon silicic rocks .

Petrochemical comparisons .

Source material comparisons: isotopic evidence .

Economic potential of the Quinn Canyon Range .
CONCLUSIONS .
Future work .
APPENDIX A. CHEMICAL ANALYSES .
APPENDIX B. ANALYTICAL METHODS .
APPENDIX C. LOCATION OF SAMPLES .
LIST OF REFERENCES .

iv

41
41

.43

45
49

.50
.51

52
53
56
56

.62

71
78

.80
.82
.83

91

.93

94

LIST OF FIGURES

Figure 1. Generalized geologic map of the Quinn Canyon
Range (modified from Stewart and Carlson, 1977). Numbered
areas denote location of anomalous element suites
discussed in text.

Figure 2. The Blue Ribbon Lineament and other mineral
belts of eastern Nevada and western Utah. . . .

Figure 3. Major element variation diagrams for the Quinn
Canyon silicic rocks. Highly altered samples denoted by a<b
See figure 5 for group symbols. . . . .

Figure 4. Plots of Sr, Zr, and Eu vs 8102 for the Quinn
Canyon rocks. Highly altered samples denoted by a . See
figure 5 for group symbols. . . . . . . . . . . .

Figure 5. Plots of La and Zr vs Y for the Quinn Canyon
sample groups. . . . . . . . . . . . . .

Figure 6. Arthimetic mean of selected trace elements of
the Quinn Canyon sample groups, normalized to U.S.G.S.
standard RGM-l (rhyolite). . . . . . . . . .

Figure 7. Chondrite-normalized REE patterns for the
rhyodacites and low silica rhyolites. . . . . . .

Figure 8. Chondrite-normalized REE patterns for the
rhyolites and incompatible element-rich rhyolites.

Figure 9. Chondrite-normalized REE patterns for the
muscovite-bearing rhyolites group 1 and 2. . .

Figure 10. Plots of Mo and Sn vs Pb for the incompatible
element-rich rhyolites. Analyses by Bondar and Clegg, Ltd.

24

26

.31

.33

34

.35

.36

.39

Figure 11. Distribution of topaz rhyolites of the
western U.S. relative to the location of the Quinn Canyon
Range, the Pine Grove porphyry Mo system and the
Urad-Henderson, Climax and Mt. Emmons porphyry Mo
deposits of the Colorado Mineral Belt (COMB).

Figure 12. Plot of Rb/Y vs Rb indicating possible partial
melting (solid lines) and fractional crystallization
(dashed lines) trends for the Quinn Canyon silicic rocks

Figure 13. Plot of molecular A1203 vs molecular
Na20+K20+CaO for the incompatible element-rich

rhyolites and muscovite bearing rhyolites (group 1).
Sample QC-13 has not been included because of alteration
affecting NaO.

Figure 14. Plots of Ba and Eu vs Nb, La and Zr vs Rb,
and La and Zr vs Y for the incompatible element-rich
rhyolites and muscovite-bearing rhyolites (group 1).

Figure 15. Enrichment factor diagrams of the Quinn
Canyon incompatible element-rich rhyolites normalized to
the mean of selected trace elements of rhyolites and
granites form: a) Pine Grove (PG; this study), b) Pine
Grove (PG; Keith, 1982), c) Mt. Emmons (ME) and

d) Urad-Henderson (UH). The ranges of values for the
Quinn Canyon rhyolites are shown by vertical bars.

Figure 16. Chondrite-normalized REE patterns for
Mt. Emmons and Urad-Henderson samples. . . . .

Figure 17. Chondrite-normalized REE patterns for
rhyolites of Pine Grove (this study) and Pine Grove
(Keith, 1982). . . . . . . . . . . .

Figure 18. Range of values of selected trace elements
for topaz rhyolites of the western U.S. compared to the
Quinn Canyon incompatible element-rich rhyolites. All
values normalized to U.S.G.S. standard RGM-l. Vertical
bars represent ranges of concentrations for the Quinn
Canyon incompatible element-rich rhyolites (after
Christiansen et al., 1986).

vi

Figure 19. Map illustrating the intrepreted western

limit of Precambrian basement (line 1) and western limit

of a granulite grade lower crust (line 2) (Farmer and

Depaolo, 1983). Lead isotopic provinces from Zartman

(1974) are defined by stipled bands and identified by

Roman numerals. See text for further explanation

(modified from Farmer and Depaolo, 1983). . . . . . . . . . 74

Figure 20. Lead isotope data for Pine Grove (PG), Quinn
Canyon (QC), Urad-Henderson (UH), Mount Emmons (ME) and
Climax (CM). The lead growth curves are from Stacey and
Kramer (1975). Data from Zartman (1974), Stacey et al.,

1968 and Stein (1985) (after Stein, 1985). .77

vii

LIST OF TABLES

Table 1. Petrography of the Quinn Canyon silicic rocks.

Table 2. Arthimetic mean of selected trace elements for
the Quinn Canyon silicic rocks.

Table 3. Arthimetic mean of selected trace elements for
the Climax- type rhyolites and the Quinn Canyon
incompatible element- rich rhyolites. .

Table 4. Chemical analyses of the Quinn Canyon silicic
rocks.

Table 5. Chemical analyses of fluorite, highly altered
samples and other samples from Quinn Canyon.

Table 6. Commercial analyses done by Bondar and Clegg,
Ltd.

Table 7. Chemical analyses of rhyolites and granites
from Mt. Emmons, Urad-Henderson, and Pine Grove.

Table 8. Precision and accuracy.

viii

.18

29

.64

.83

.87

.88

.89

92

INTRODUCTION

Problem

The Quinn Canyon Volcanic complex of southeast central
Nevada consists of Tertiary eruptive sequences, flows and
hypabyssal dikes, plugs and sills, which overlie and intrude
Paleozoic sedimentary rocks (Sainsbury and Kleinhampl, 1969).
This mineralized volcanic center lies along the trend of the
Pioche belt and the Blue Ribbon Lineament as defined by Shaw
and Stewart (1976) and Rowly et al. (1978), a mineral belt
characterized by the occurrence of fluorine, tungsten, gold,
copper, molybdenum, lead-zinc-silver, manganese, uranium and
berylium minerals (Figures 1 and 2).

The most evolved Quinn Canyon rhyolites are part of a
chemically distinct group of high silica, fluorine-rich
rhyolites located in the western United States. Typically,
the hypabyssal intrusive and flow rocks are 10-30 m.y. old and
their emplacement histOry generally coincides with late
Cenozoic extensional faulting, and the onset of transitional
to bimodal magmatism marking the end of Cenozoic subduction
(Christiansen and Lipman, 1972; Lipman et al., 1976; Westra
and Keith, 1981; Christiansen et al., 1983).

High silica, fluorine-rich rhyolites are frequently
associated with epithermal base and precious metal veins (Ag,

Au\

, Pb, Sn, W), fluorite-rich replacement and fracture

1

 

      

 

 

 

Canyon Range
1977). Numbered areas
tes discussed in text.

alized geologic map of the Quinn
lous element sui

of anoma

1. Gener
(modified from Stewart and Carlson,

ote locat1on

Figure
den

 

 

 

  

120

 

160 KILOMETERS

 

Figure 2. The Blue Ribbon Lineament and other mineral belts of
eastern Nevada and western Utah (modified from Rowley et a1,
1968) .

4

controlled deposits and volcanogenic Be, U and Li deposits.
Many workers recognize that fluorine-rich rhyolites are
associated with, and may be indicators of, subsurface
metallization, such as the economically important Climax-type
stockwork molybdenum deposits in the western U.S. (Lamarre and
Hodder, 1978; Westra and Keith, 1981; White et al., 1981; Burt
et al., 1982; Christiansen et al., 1986; Keith et al.. 1986).

Keith (1980 and 1982) and Keith et al. (1986) suggest that
molybdenum mineralization at Pine Grove, Utah, is cogenetic
with spatially associated extrusive high-silica rhyolite
tuffs. The Pine Grove deposit is unique among major porhyry
molybdenum deposits in having a direct correlation between
eruptive and ore-bearing intrusive units. Because of its
location and characteristics the Pine Grove mineralized system
provides a possible "link" between the Climax-type stocks and
ore bodies of the Colorado Mineral Belt and mineralization
related to the emplacement of high level lithoPhile
element-rich rhyolites in other areas of the western U.S.,
such as in the Quinn Canyon volcanic complex.

The Quinn Canyon volcanic complex exhibits many of the
characteristics of economically significant mineralized
systems, including some of the fundamental features of
Climax-type districts. These features are: 1) the presence of
high silica, apparently fluorine-rich peraluminous rhyolites,
2) anomalous concentrations of Mo, Pb, Zn, Sn, W, Au, and Ag
dispersed in halos spatially and temporally associated with
the rhyolite porphyry dikes and plugs, 3) fluorite and quartz

veining associated with metal anomalies, 4) silicic, argillic,

5

potassic and propyllitic hydrothermal alteration, surrounding

intrusive centers, 5) base-metal vein deposits, 6) precious

metal bearing quartz veins, and 7) emplacement of high silica
rhyolites coincident with the transition from compressional to
extensional tectonics.

Although there are similarities between the Climax-type
mineral districts and the Quinn Canyon Range, drill testing of
one Mo-prospect in this area, has yielded generally negative
results (Westjohn, personal communication).

The purpose of this investigation is to:

1) Characterize the petrochemistry of the Quinn Canyon
hypabyssal silicic rocks and evaluate possible genetic
relationships among the rock types.

2) Evaluate the economic potential of the Quinn Canyon Range,
by comparing the petrochemistry and source material of the
Quinn Canyon rhyolites to that of the granitic and
rhyolitic plutons that generated Climax-type porphyry
molybdenum deposits and topaz rhyolites that are
associated with volcanogenic Li, Be, U, and Sn

mineralization.

flatbed

The following methods were used to address the problem:
1) The chemical composition of the Quinn Canyon igneous
samples were determined by XRF (X-ray fluorescence) for major
and selected trace elements and INAA (instrumental neurtron

activation analysis) for rare earth elements (REE).

6
Petrochemical characterization of the suite of samples was
accomplished by a series of major element variation diagrams,
chondrite-normalized REE plots, and enrichment factor
diagrams. Trace element composition and petrography were used
to "group" the samples and to examine the relationships

between the rocks of this suite.

2) Petrochemical comparisons of Climax-type granitioids with
the Quinn Canyon rhyolite intrusives, were made possible by
data generated at Michigan State University (Westjohn, 1985).
A suite of samples from Urad-Henderson, Pine Grove, and Mt.
Emmons, representing surface and subsurface samples, were
analyzed for major and trace elements (including REE). The
elements that appear to charaterized Climax-type plutons are
Rb, Ba, Sr, Nb, Y, and Th (Westra and Keith, 1981). These
elements, in combination with REE data, were used to make
regional geochemical comparisons.

Pb isotope data from Zartman (1974), Stein (1985) and
Stacey et al. (1968), and Nd-Sm isotope data published by
Depaolo and Farmer (1984) provide a means of indirectly and
directly comparing the source material of the igneous rocks
and metallization associated with Quinn Canyon volcanic

complex and Climax-type mineral districts.

PREVIOUS WORK

Sainsbury and Kleinhampl (1969) published the only written
report on the geology of the Quinn Canyon Range with emphasis
on the fluorite deposits. The authors describe the general
geologic setting of the range including structure, plutonic
and hypabyssal rocks, and sedimentary and extrusive volcanic
rocks. Most importantly, Sainsbury and Kleinhampl produced a
geologic map of the Quinn Canyon Range based on field work
done between 1956-1965.

Kral (1951) describes in considerable detail the economic
history of the range including locations of mines, types of
ores, production, tonnage, and assays. '

Lead isotope data from the Quinn Canyon Range is published
in the appendices of Zartman’s (1974) publication on the
geologic significance of lead isotope provinces in the Western
U.S.A. In addition, the Quinn Canyon Range is mentioned in
the Rowly et al (1978) publication on the Blue Ribbon
Lineament.

Economic evaluations were conducted by Amoco Minerals Co.

in the early 1980’s and will be discussed in a later section.

FIELD WORK

Field work was conducted over a 4 day period in September
1985. The focus of the field work was the sampling of
hypabyssal and intrusive rocks, with emphasis on collecting
fresh samples of rhyolite porphyry dike, plug, and flow
rocks. The geological description of the study area, that
follows, is based almost entirely on the work of Sainsbury and
Kleinhampl (1969), with information added by D.B. Westjohn

(1985). Sample locations are available in Appendix C.

GEOLOGY OF THE QUINN CANYON RANGE

The Quinn Canyon Range is located in southeastern Nevada,
approximately 90 miles south and slightly west of Ely, Nevada
(Figure 1). The Quinn Canyon Range is a distinctive unit of
the Grant Range, separated from it by the Cherry Creek Canyon
to the north (Kral, 1951; Sainsbury and Kleinhampl, 1969).

Tertiary extrusive volcanics and Paleozoic sedimentary
sequences are the main rock-types that outcrop in the Quinn
Canyon Range (Figure 1). Intruding these are a centrally
located hypabyssal quartz-latite pluton (7) and a series of
.chemically variable dacitic-rhyolite porphyry dikes, plugs,

and sills and andesite dikes (Sainsbury and Kleinhampl, 1969).
We

The sedimentary strata is a typical eastern assemblage
miogeosynclinal sequence (Sainsbury and Kleinhampl, 1969;
Stewart and Poole, 1974; and Stewart, 1977). The exposed
rocks are limestone, dolomite, shaly limestone, shale and
quartzite of Early to Middle Paleozoic age, congolmerate and
fresh water limestone of early to middle Tertiary age and
tuffaceous sedimentary rocks (intercalated with volcanics) of

middle to late Tertiary age (Sainsbury and Kleinhampl, 1969).

10
Wk:

The extrusive units are primarily rhyolite and rhyolitic
to quartz latitic welded tuffs of Oligocene and early Miocene
age. Two of the units have tentatively been assigned to the
Shingle Pass Tuff and the Needles Range Formation. The
uppermost member of the Needles Range Formation has yielded a
K-Ar date of 27.2 m.y., in the nearby Grant Range (Armstrong,
1970). The basal member, in other localities of Utah and
Nevada has given consistent age values of 29.7 m.y.
(Armstrong, 1970). Sainsbury and Kleinhampl suggest that the
source area of the Needles Range Formation is to the east and
the Shingle Pass Tuff to the northwest. A third undivided
volcanic unit and a rhyolite tuff unit were mapped by
Sainsbury and Kleinhampl (1969). Andesitic and basaltic
breccias and flows are found at the base of the volcanic
pile. The volcanic stratigraphy is poorly understood.

The dacitic to rhyolitic dikes range from several inches
to greater than 300 feet in width, and are emplaced along east
northeast trending fractures. They cut the Paleozoic
sedimentary package and most of the Tertiary volcanic sequence
(Sainsbury and Kleinhampl, 1969). A K-Ar date of 23 m.y. has
been obtained for alteration associated with one rhyolite
intrusive center (Westjohn, personal communication, 1985).
Sainsbury and Kleinhampl assume that the quartz latite
hypabyssal "stock" is genetically related to the various
rhyolitic dikes and plugs. This assumption will be examined

in later chapters. A granite stock, located in the northwest

11
portion of the study area, is believed to be older than both
the rhyolite porphyry dikes and the quartz latite stock
(Sainsbury and Kleinhampl, 1969; Westjohn, personal

communication, 1985).

51:22:21.:

The Quinn Canyon Range is composed of a series of blocks
that are bounded by Basin and Range faults. Paleozoic rocks
outcrop in blocks to the north and east in the study area.
Both high angle and thrust faults occur, with the thrust
faults appearing to pre-date volcanic activity.

‘ High angle faults appear to be more dense and complex in
the Paleozoic section, then those observed in the Tertiary
sequence. Sainsbury and Kleinhampl (1969) suggest that faults
in the Tertiary volcanic units are difficult to identify,
implying that high angle faulting in the Tertiary may be more
prevalent than their maps suggest.

The contact between the Paleozoic and Tertiary rocks is
engimatic. Sainsbury and Kleinhampl report that in some cases
the sharp contact does not appear to be fault related, but is
depositional in nature. They cite the area north of Adaven as
an example of where Tertiary volcanics appear to lie on the
prevolcanic surface. There is evidence of fault contacts near
the High grade, Spar, and Mammoth properties. Faults in this
area seem to be a major structural feature of the range

(Sainsbury and Kleinhampl, 1969).

12
W

The Quinn Canyon Range has been the subject of sporadic
mining activity and exploration interest since 1971 (Kral,
1951). Although there has been production from numerous small
mines, the return from these workings has been relatively
insignificant with no major deposit of economic significance
known to date. Kral (1951) estimates that the total
production amount at that time was less than $100,000.

Three main types of deposits have been recognized by
previous workers. They are: 1) fluorite replacements and
veins; 2) base metal deposits; and 3) gold-bearing quartz
veins (Kral, 1951; Sainsbury and Kleinhampl, 1969).

Native gold occurs in quartz veins, frequently with
arsenopyrite. The majority of the gold deposits are located
in limestone near Willow creek, and appear to be associated
with a quartz monzonite stock (Kral, 1951; Sainsbury and
Kleinhampl, 1969). Kral also describes a gold vein adjacent to
an andesite dike in this area.

Although little description is available for the base
metal deposits, Kral states that the minerals are typically
arsenopyrite, sphalerite, pyrite, galena, and occasionally
tetrahedrite, which occur in veins and at least in one
instance are asSociated with a large rhyolite dike.

The fluorite deposits have been studied in considerable
detail by Sainsbury and Kleinhampl (1969). The authors
describe eight fluorite deposits and mention that numerous

prospects are scattered throughout the range. The deposits

13

are grouped into two types; those that occur as veins and
tabular replacement bodies in carbonates, and a less
economically important and less diverse group occuring as
fissure veins and tabular bodies in volcanic rocks. The
fluorite is spatially associated with rhyolite dikes as well
as deposited along faults. In the volcanic rocks, fluorite
veins are found within altered zones. Sainsbury and
Kleinhampl state that the alteration zones in the Cottonwood
area are up to 4,000 feet long and 500 feet wide and locally
parallel fault zones. The alteration zones as a group define
a northwest trending band that may be aligned along a major
Pre-Tertiary fracture zone (Sainsbury and Kleinhampl, 1969).

The Quinn Canyon Range, was the target of exploration
activity in the late 1970’s and early 1980’s. Amoco Minerals
Co. initiated a reconnaissance exploration program in 1980,
with a more detailed study in 1981-1982, which included
geologic mapping, geochemical sampling and geophysical
surveying (Westjohn, personal communication, 1985).

Amoco has provided geochemical data and maps for this
study, which will be integrated into a brief summary of the
key geochemical anomalies in the range. Rock chip and stream
sediment samples were analyzed, by Amoco for at least some,
and in many instances almost all of the following elements;
Au, Ag, As, Hg, Sb, Sn, Pb, W, Zn, Ba, Mo, F, and Mn.

Eight different hydrothermally altered centers that are
spatially associated with rhyolite intrusives have been
recognized by Amoco Minerals in the Quinn Canyon Range

(Westjohn, personnel communication, 1985). Surrounding these

14

intrusive centers are specific and somewhat exclusive
metallogenic halos. Four of these areas have been studied in
detail by Amoco and display distincly different enrichment
trends. The anomalous elemental suites are: 1) Mo and F, 2)
Mo, Sn, Pb, Zn, W, Ag, and F, 3) Mo, Pb, Zn, Ag, and F, 4) Mo,
Ag, Au, and F (Westjohn, personal communication, 1985). There
is a change in the elemental distribution pattern from one
area to another in the Quinn Canyon Range. This change is
subtle, but in the case of the second and fourth elemental
halos, the presence of Sn and Au is distinctive. These
differences may be due to variations in source igneous rock
chemistry, and/or type of country rock (Paleozoic sedimentary
vs. Tertiary volcanic).

The intrusive center marked by presence of anomalous Sn
(up to >1000 ppm) is located in the extreme eastern portion of
the range and is composed of several rhyolite porphyry dikes
and a large sill (area 1 of Figure 1) . In addition to the
elements mentioned previously, somewhat anomalous W (mostly
<10 ppm, one 100 ppm value), and Mn (to > 10,000 ppm at the
Roadside Mine) are characteristic of this suite. Mo ranges
from 7-53 ppm, Pb values consistently over 4000 ppm
(particularly in the southern portion of this area), F) 10,000
ppm and Ag reaching concentrations greater than 20 ppm.

The region with consistently high Au and associated
element (As, Ag, and Hg) values is in Water Canyon, in the
western portion of the study area (area 2 of Figure 1). In
this location Mo occurs in isolated concentrations greater

than 400 ppm,-predominately in quartz veins. Au values of

15

over 1 ppm and Ag values as high as 50 ppm are recorded,
although the majority of samples have much lower
concentrations. The highest concentrations of the previously
mentioned elements were for the most part from rock chip
samples of country rock, with quartz veins, although intrusive
rock samples were moderately to highly enriched in some of the
elements of economic interest. Field evidence suggests that
the metals were carried in fluids exsolved from the
crystallizing silicic porphyry intrusives and deposited in the
country rock.

A total of seven rotary drill holes were drilled in Water
Canyon by Amoco Minerals 00., during 1982. Six of these
exploration holes were relatively shallow (200’-500’), and
bottomed in Tertiary rhyolite tuff. The drilling consistently
indicated stockwork quartz veining, weak-moderate
silicification and pyritization, localized zones of
weak-moderate argillic alteration and trace MoSz.

The seventh drill hole was completed to a depth of 1024 ft
and cored from 440 ft to the bottom. Silicified and locally
argillized Tertiary rhyolite tuff, with associated
quartz-latite lithic fragments was encountered in core from
474 ft to 860 ft. In this interval, stockwork
quartz-pyrite-fluorite veins were common with pyrite exceeding
1% in most core intervals, and locally exceeding 17%. MoSz
(Jordisite) was identified in quartz veins. Propyllitic
alteration, characterized by chlorite and quarz-pyrite-calcite
veins was observed from a depth of 790 ft-864 ft argillically

altered quartzelatite occured at a depth of 864 ft-880 ft,

16

with intense quartz-calcite-pyrite-fluorite veins from 874
ft-880 ft. Chlorite altered rhyolite intrusive breccia
extended from a depth of 880 ft-914 ft. A lithic-poor
fractured rhyolite was encountered from 914 ft-934 ft,
followed by a rhyolite porphyry intrusive (7), featuring
resorbed bipyrammidal quartz phenocrysts in a glassy matrix.

Geochemical analysis of core samples, revealed anomalously

high Mo values at 516 ft (112 ppm) and 791 ft (885 ppm).

PETROGRAPHY

The following petrographic summary will focus on the
silicic hypabyssal, extrusive dome and flow rocks collected in
the Quinn Canyon Range. Three groups can be crudely
delineated, based on petrography. For ease of description

these groups are named:

1) rhyolites
2) less evolved or low silica rhyolites

3) rhyodacites

Within the rhyolite group two smaller populations can be
delineated according to major and particularly trace mineral
phases. See Table 1 for sample numbers and brief descriptions

of each group.

NW

The rhyolites range from distinctly porphyritic to aphyric
and predominately have euhedral to subhedral phenocrysts of
quartz, potassium feldspar (orthoclase) and subordinate
amounts of plagioclase (albite-oligoclase). These samples
either have a matrix composed of entirely recrystalized
spherulitic alkali feldspar and quartz intergrowths, or a
combination of this texture occuring as "patchs” within a

17

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19
cryptocrystalline or microcrystalline matrix. These
spherulites often occur as halos radiating out from and
surrounding phenocrysts. Micrographic intergrowths of alkali
feldspar and quartz are also pervasive in the matrix of many
of the samples. Resorbed quartz grains are common.
Plagioclase phenocrysts exhibit albite twinning, are generally
unzoned, and in some instances have alkali feldspar
overgrowths. Accessory minerals include: zircon, sphene,
apatite, opaques, and biotite.

A distinct series of rhyolite samples contain primary(?)
euhedral to subhedral muscovite in significant concentrations
in the groundmass and lack spene and zircon.

Another subgroup appear to be the most "evolved” of the
rhyolite population and are characterized by a large portion
of euhedral-subhedral quartz phenocrysts relative to potassium
feldspar and albite. Some of these phenocrysts exhibit a
graphic intergrowth of quartz and alkali feldspar. Embayed
quartz phenocrysts are prevalent. This group of rhyolites has
a microcystalline matrix of albite, quartz, and potassium
feldspar with samples Qc-19c and QcQ-B having the spherulitic
groundmass as seen in many of the other rhyolites. Qc-19 is
the only rock of the group to have significant amounts of
sanidine and trace amounts of relatively unaltered biotite.
Biotite in Qc-19, and fluorite in Qc-IQC and QcQ-B. appear as
accessory mineral phases. Muscovite, with the exception of a

small amount in sample Qc-ZO, zircon, and sphene are absent.

20

The less evolved rhyolite dike samples were all collected
in the same general area. They differ from the rest of the
silicic rocks by having a large amount of plagioclase
phenocrysts (oligocene-andesine), and potassium feldspar
phenocrysts, few, if any quartz phenocrysts and a significant
amount of altered biotite and amphibole(?). Accessory
minerals include zircon, apatite and Opaques.

The three rhyodacite samples taken from the quartz latite
plug exhibit marked extrusive, and in one instance flow-like
texture. Many of the phenocrysts are broken and angular, and
in sample Qc9-1 are crudely aligned. The phenocrysts phases
present include; large, embayed, and in some cases broken

quartz grains, highly altered plagioclase (andesine),

potassium feldspar, biotite and amphibole.

The degree and type of hydrothermal alteration varies from
sample to sample, but all of the silicic rocks collected at
Quinn Canyon are altered to some degree. The principal types
of hydrothermal alteration are silicification, sericitization,
the addition of and replacement of mineral phases by calcite,
and to a lessor degree, argillization and chloritization
(Table 1).

Typical alterations are: plagioclase to sericite and
calcite, alkali feldspar to clays and chlorite, biotite and
amphibole to chlorite (some penninite), amphibole to epidote,

the silicification of major mineral phases, as phenocrysts and

21

in the groundmass (primarily plagioclase and alkali feldspar)
and quartz veining. Generally, the highly evolved rhyolites
(Qc-19, Qc-ZO, etc.) are primarily Silicified, while the other
rhyolites and rhyodacites, in addition to being Silicified,
have calcite, sericite and chlorite as alteration products.

With several exceptions, this rock suite shows a weak to
moderate amount of alteration, and as such, most of the trace
elements and many of the major elements should reflect

approximate magmatic concentrations.

PETROCHEMISTRY

mm

The chemical data for all rocks are presented in Appendix
A, by methods described in Appendix B. The silicic rocks
collected from the Quinn Canyon Range are peraluminous
rhyolites and rhyodacites. The "freshest” most evolved
rhyolites have K20 values which range from 4.32%-5%, NazO
values from 2.25%-3.4% and are depleted in CaO (generally less
than .2%), MgO (<.03%, frequently not detected), P205
(.01%) and TiOz (<.04%). K20/Na20 ratios range from
1.2-2.4 for relatively unaltered rocks of the entire suite.
Silica values are variable (72% to >80%) depending on the
degree of alteration and differentiation. Major element
variation diagrams are shown in Figure 3. Group divisions are
based primarily on trace element chemistry as outlined in a

following section.

mm

The majority of the scatter and extreme deviation of the
major element oxides can be easily attributed to alteration.
The affects of alteration on major element chemistry are
twofold: first, the addition of major cations through
alteration and secondly the addition of non-measured ions such

22

23

Figure 3. Major element variation diagrams for the Quinn
Canyon silicic rocks. Highly altered samples are denoted by
a Q . See figure 5 for group symbols.

24

 

 

 

 

 

 

 

 

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25
as C (in CaCOa), F (in Can) and structurally bound
H20. The addition of significant amounts of 820, CaCOa
and/or Can is recognized by either anomalously high CaO
values and/or anomalously low measured major element oxide
totals.

Silicification and argillization are common alteration
types that produce in the first instance an increase in
SiOz, and in the second case, an increase in K20 and a
decrease in NazO. The addition of calcite and the
alteration of feldspars to sericite and chlorite also cause
some scatter in the plots of CaO, K20, MgO and NazO.

Samples that consistently deviate from the "trend" in
major and selected trace element variation diagrams, such as
Zr, Eu and Sr vs SiOz, and/or exhibit extensive alteration
in thin section, were not included in group comparisons
(Figures 3 and 4). The small grouping of samples with the
lowest silica values are not highly altered, but rather are
significantly less "evolved".

Samples that show minimal amounts of alteration, such as
the small amount of silicification seen in some of the more
evolved rhyolites were used in trace element comparisons.

Most workers believe that under many alteration conditions
(including hydrothermal alteration), the elements Nb, Y, Sc,
Hf, Zr, Ti and REE, particularly heavy rare earth elements
(BREE) are relatively immobile (Winchester and Floyd, 1977;
Floyd and Winchester, 1978; Campbell et al., 1984; Palacious
et al., 1986; and Maclean and Kranidiotis, 1987). Other

workers maintain that under extreme hydrothermal alteration,

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Canyon rocks. Highly altered samples are denoted by a<b. See
figure 5 for group symbols.

27
Y, Sc, Nb, and REE can be significantly mobile (Alderton et
al., 1980; Finlow-Bates and Stumpfl, 1981). They suggest that
only Zr and Ti are immobile and can be used to identify the
degree of magmatic differentiation in a hydrothermally altered
rock . These authors have described and studied the results
of extreme alteration. Except for a small group of highly
altered samples, the general weak to moderate alteration
exhibited by the Quinn Canyon rocks should have little

influence on trace element chemistry, particularly Y, Nb, Zr.

Hf, Ti and HREE.

W

The Quinn Canyon rhyolitic rocks can be ”grouped”
according to trace element chemistry, particularly REE
chemistry. The arthimetic mean of selected trace element
concentrations for each group are presented in Table 2. The
differences in trace element composition between groups can be
attributed to variability in source material and/or degree of
igneous processes, such as partial melting and crystal
fractionation. With the exception of one group, samples
within each group are assumed to be genetically similar.
Different degrees of alteration, partial melting and crystal
fractionation are probably the cause of variation that exists
in some groups.

Six groups can be roughly defined and named for ease of
description and comparison. The "justification" for these

groupings was evaluated by R-Mode factor analysis. Initially,

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29

 

 

 

 

 

 

 

 

 

 

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runs using all chemical variables, grouped the samples in
similar clusters. Later runs, in which chemical variables
which had similar loadings were deleted, left essentially the
same clusters. These clusters were grouped according to Rb,
Sm, La, Sr, and SiOz as the dominant variables. The six
groups are:

1) rhyodacites

2) low-silica rhyolites

3) rhyolites

4) muscovite-bearing rhyolite group 1

5) muscovite-bearing rhyolite group 2

6) the most "evolved" or incompatible element-rich

rhyolites.
Sample numbers and trace element compositions are presented in
Appendix A. This grouping of samples is evident in many trace
element plots, and is well illustrated in plots of La vs Y,
and Zr vs Y (Figure 5). The identification of samples as
rhyodacites and rhyolites is loosely based on a classification
scheme according to the position of samples on a Zr/Ti vs Nb/Y
plot (Winchester and Floyd, 1977).

The composition of several chemically variable, generally
silicified rhyolites (1102b, 1001, 1107, and 1109) taken from
the western portion of the Quinn Canyon volcanic complex
(Water Canyon) are presented in Appendix A, but will not be
included in the following comparisons.

An enrichment factor diagram of the mean of the groups,

normalized to U.S.G.S. standard RGM-l, was used for displaying

element distribution patterns and for group comparisons

3 l

 

 

 

 

 

 

 

 

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+ LOH SILICA RHYOLITES
+ X RHYOLITES
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0 50 100 150 200
Y

Figure 5. Plots of La and Zr vs Y the Quinn Canyon sample
groups.

32
(Figure 8). RGM-l is a rhyolite obsidian from Medicine Lake
Volcano, California and was used by Christiansen et a1 (1986)
for characterizing topaz rhyolite trace element chemistry.
The elements of this plot are arranged with increasing field
strength to the right, after Pierce et al. (1984) and
Christiansen et al. (1986).

The three rhyodacite samples comprising the least evolved
group were collected from the large centrally located "quartz
latite” dome in the Quinn Canyon Range. Compared to the other
silicic rocks of this suite, these samples are depleted in Rb,
Y, Nb, and Th and enriched in Eu, and Sr (Figure 6). The REE
distribution pattern roughly parallels that of the group of
low silica rhyolite dikes, having a small negative europium
anomaly, La/Ybu of 9-11, but with overall lower
chondrite-normalized abundances (Figure 7).

The series of low silica rhyolite dike samples display a
consistent trace element chemistry, having subequal
proportions of Rb and Sr, being depleted in Nb, Y, Th and
enriched in Eu, Ba, REE elements, particularly light rare
earth elements (LREE), Zr and Hf, relative to the most evolved
rhyolites of this suite (Figure 6). This group has a straight
line REE pattern, with a small negative Eu anomaly. La/YbN
ratio averages about 13 (Figure 7).

The rhyolite group shows the most chemical variation. It
is characterized by a La/YbN ratio of 4-12, Sm/Yb ratio of
2-3, and an enrichment trend which, in general, lies between
the low silica rhyolites and the highly evolved rhyolites

(Figures 6 and 8).

33

GROUP MEANS

 

 

 

 

 

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Figure 6. Arthimetic mean of selected trace elements of the
Quinn Canyon sample groups, normalized to U.S.G.S. standard
RGM-l (rhyolite).

34

 

 

 

 

 

 

 

 

RHYODACITES
__ I T I I I I I I I I T I I I I
r- —4
100 E- 2:.
E E
m _ -1
B L s
E
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LOW SILICA RHYOLITES
I I I I I I l I I I I I I r I _
I'- d
1— a-J
100 L— 1
E 3
In - ..
a:
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H
:‘5
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l l l l l l J_ l l l l l L l 1
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Figure 7. Chondrite-normalized REE patterns for the
rhyodacites and low silica rhyolites.

35

 

 

 

 

 

 

 

 

 

 

RHYOLITES
,_ I r I T I I I I I I r I I I I _1
L. .1
100 E— __
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In ‘— d
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c : 1
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ii _ _
x 1— 4
8 I. J
c:
1 E— a
E Z
)— —1
L g 1 1 1 1 l 1 1 1 1 1 l 1 1
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INCOMPATIBLE ELEMENT-RICH RHYOLITES
I I I I I I I I I I I I I I I .1
100 F .1
h .1
.. 1
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In - s
3 1 _
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c 10 E- '7'.
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1 1 1 1 1 l 1 1 1 1_ L 1 1 1 1
La Ce Sm Eu Tb. Yb Lu

Figure 8. Chondrite-normalized REE patterns for the rhyolites
and incompatible element-rich rhyolites.

36

MUSCOVITE-BEARING RHYOLITES GROUP 1

 

 

 

 

 

 

 

 

 

 

__ T I I I I I I I— T I I T I Ifi
L- -1
_ 4
100 h— _=
I j
m "’ —1
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MUSCOVITE-BEARING RHYOLITES GROUP 2
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1 a- 1
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,_ —J
l l l L l l l L J l l l L l 1
La Ce Sm Eu Tb YD Lu

Figure 9. Chondrite-normalized REE patterns for the muscovite
bearing rhyolites group 1 and 2.

37

The first group of ”muscovite-Dearing” rhyolites, marked
by the presence of muscovite in most samples, are distinctly
depleted in Zr, Bf, Eu, Ba and Ti, and slightly enriched in
Rb. With the exception of the rhyodacite group, these
rhyolites show no enrichment in the incompatible elements Nb
and Th relative to other silicic rocks of this suite. This
group also displays anomalously large LREE and middle rare
earth elements (MREE) depletions, with La/YbN ratios of
.5-2 and Sm/Ybu ratios of .5-1.5 (Figures 6 and 9). The low
Ti, Eu, and Ba concentrations, suggest a significantly
differentiated or "evolved" rhyolite, much like the group of
incompatible element-rich rhyolites described below. The
peculiar extreme depletion in LREE and MREE could be a result
of extreme differentiation of light and middle REE-rich
mineral phases, a reflection of the source material, and/or
possibly a result of alteration. Possibly related to this
group are several "less evolved" samples (muscovite-bearing
rhyolite group 2), some with muscovite, which have higher REE
abundances, yet display a similar tendency for MREE
depletion. These rhyolites are more enriched in Ti and Zr
(Figures 6 and 9).

The most evolved or incompatible element-rich rhyolites of
the Quinn Canyon range are distinctive in having a large
negative Eu anomaly, a ”flat" REE distribution pattern
(La/Ybu ratio 1-3) and are extremely enriched BREE, Nb, Y,
Rb, and Th and depleted in Sr, Ba, Bf, Zr, and LREE (Figures 8
and 8). Overall REE abundance is greater than the previously

described muscovite-bearing rhyolite group 1. This type of

38
trace element signature is typical of other highly evolved,
incompatible element-rich rhyolites of the Western U.S. (Burt
et al., 1982; Christiansen et al, 1988). The LREE depletion
is probably due to the fractionation of LREE-rich trace
mineral phases such as allanite, monazite, and chevkinite. A
contributing factor to the low La/Yb ratio could be that BREE
complex more readily with fluorine, and may be preferentially
enriched in the roof of a fluorine-rich magma chamber, during
convection-driven thermogravitational diffusion as suggested
by Bildreth (1979). The low Hf and Zr values are probably due
to their compatibility with fractionating zircon. Likewise,
the Ba depletion may be due to the fractionation of potassium
feldspar.

Mo, Pb, and Sn were analyzed for rhyolites of this group
and display a general negative correlation of Pb, with Mo and
Sn (Figure 10).

Sample Qc-20 is unique in having trace element
characteristics of both the previously mentioned group of
rhyolites and the muscovite-bearing rhyolite group 1.
Specifically, it is enriched in Rb, somewhat enriched in Y and
Th and shows little Nb enrichment. This sample has a low
La/Ybu ratio (1) and an overall low REE abundance (Figures
6 and 8).

39

10L 0

Sn -

 

Mo

 

Figure 10. Plots of Mo and Sn vs Pb for the incompatible
element-rich rhyolites. Analyses done by Bondar & Clegg Ltd.

4O

Fluorite

The trace element composition of a fluorite sample taken
from an open pit just west of the Hi Grade Prospect, is

included in Appendix A.

REGIONAL GEOLOGY

Lineaments

Several east-west trending mineralized belts and
structural linaments have been recognized in Utah and Nevada,
based on aeromagnetic data; fluorine, tungsten, berylium,
uranium and base metal mineralization; hydrothermal
alteration, and the presence of Tertiary eruptive and
hypabyssal centers (Rowly et al., 1978; Stewart et al., 1977;
and Shawe and Stewart, 1976). The best defined belts are the
Deep Creek-Tintic belt of Central Utah, the Qquirrh-Unita belt
of north central Utah and the overlapping Wah-Wah-Tushar belt,
Pioche belt and Blue Ribbon Lineament of south central Utah
and eastern Nevada (Rowley et al., 1978). The Blue Ribbon
Lineament has been defined by Rowly et al. (1978) as a
distinct structural zone within the Pioche belt as defined by
Shawe and Stewart (1976) (Figure 2).

Rowly et al. (1978) and Stewart et al. (1977) concur that
these lineaments may be due to an east-west trending
structural warp or break in the subducting slab. Rowly et a1.
(1978) suggest that the fractures associated with the Blue
Ribbon Lineament (BRL) were important conduits for
mineralizing solutions or, more accurately, acted as zones of
weakness for emplacement of rhyolitic-granitic rocks which

were the source for mineralization. In any event, as the
41

42

authors emphasize, the BRL has important exploration
significance.

The Quinn Canyon Volcanic center is located along trend at
the western end of the BRL and Pioche belts in the Quiet Zone
(Rowly et al., 1978). The Quiet Zone (QZ) is a north-south
oriented area from about 115 to 116 W., where the aeromagnetic
signature, characteristic of the BRL and other mineral belts,
is absent (Figure 2). This zone is puzzling because
relatively large volumes of intermediate calc-alkaline igneous
rocks outcrop here and these types of igneous rocks are
associated with high intensity positive magnetic anomalies in
other areas within the BRL.

Stewart (1976) and others suggest that perhaps few large
intrusive bodies exist in the Quiet Zone, even though large
volumes of extrusives are present (in the Quinn Canyon Range
this may be supported by Sainsbury and Kleinhampl’s conclusion
that a large portion of the extrusive sequence is sourced from
the east and northwest). Also, they suggest that the QZ is
located in and parallel to a thick portion of the
miogeosynclinal sequence as noted by Stewart and Poole
(1974). Rowly et al. (1978) believe that the lineament passes
through the Q2 based on the nature of the surface geology and
the presence of fluorite and base metal deposits in the Quinn
Canyon Range. The Warm Springs lineament is just west of
Quinn Canyon and on trend with the BRL, lending support to the
theory that the lineament passes through the QZ (Rowly et al.,

.1978).

43
C zo' ' c 'v' ' W .S.

The igneous activity associated with the east-west mineral
belts, seems to have migrated southward starting at 40 N.,
43-34 m.y. ago moving to 37 N, 17-6 m.y. ago (Stewart et al.,
1977). Armstrong et al. (1969) present a variation on this
theme, with volcanism beginning about 40 m.y. ago in east
central Nevada, migrating as a front outwards toward the
margin of the Great Basin. One possible explanation proposed
by the authors is that crustal fracturing began in the core
area and slowly propagated outward, resulting in a slowly
migrating volcanic front.

The general pattern of middle to late Cenozoic volcanism
in the western U.S. was the eruption of Oligocene-Miocene
intermediate volcanics (quartz latite-low silica rhyolite)
followed by the eruption and emplacement of basalt and high
silica rhyolite. This change in type of volcanism was
coincident with the change from subduction related
compressional tectonics to an extensional regime at the
termination of subduction (Lipman et al., 1972; Christiansen
and Lipman, 1972; Lipman et al., 1978). Eaton (1984) presents
a more specific scenario in which 3 different periods of magma
genesis are defined. The first is the emplacement of
calc-alkaline andesites, rhyolite and quartz latite magmas
during convergence-related compression. The second is the
emplacement of high silica rhyolites, basaltic andesites and
alkali basalts during intra-arc and back-arc spreading at

rapid rates. The final stage, is marked by the emplacement of

44
tholeitic basaltic magma during reduced extensional rates. In
the Basin and Range province, the change from intermediate to
bimodal volcanism occurred in general between 22 and 24 m.y.
ago.

This sequence and change in rock composition has been
documented in the San Juan volcanic complex, the Thomas Range
(Spor Mountain) and the Wah Wah Mountains (Lipman et al.,
1976; Lipman et al., 1978; Lindsey, 1982; and Keith et al.,
1986).

In these volcanic centers, particularly Spor Mountain and
the San Juan volcanic system, significant mineralization (F,
W, Be, and U, was associated with high silica alkalic
rhyolites emplaced 5-15 m.y. after the end of the caldera
cycle and coincident with the transition from compressional to
extensional tectonics. The mineralization in the western San
Juan mountains occurred in pulses from the most intense at
22.5 m.y. ago, to 15-17 m.y., and a final pulse at 11 m.y. ago
(Lipman et al., 1976).

Both Oligocene and Miocene magmatism is evident in the
southern Wah Wah Mountains, which host the Pine Grove
intrusions and porphyry molybdenum deposit. Oligocene
magmatism was characterized by dacitic ash flows and andesite
lava flows. Miocene igneous activity is marked by a
predominately bimodal assemblage of high silica rhyolite
domes, dikes, ash flow tuffs and trachyandesite lava flows.
The Pine Grove system is believed to have been formed 22-23
m.y. (Keith, 1980; Keith, 1982, and Keith et al., 1986).

45

Ter ' t l' at' ' e We er .8.

Several types of mineralization and metallization appear
to be spatially and/or genetically related to the intrusion or
extrusion of high-silica, fluorine-rich rhyolite magmas (topaz
rhyolites). Among these are: 1) base and precious metal
epithermal vein deposits (Ag, Au, Pb, Zn, W); 2) volcanogenic
deposits of U, Be, F, Li, Sn and Cs; 3) fluorite breccia
pipes, fracture fillings and as replacement bodies; 4)
Climax-type porphyry deposits (Mo, W, Sn) and Mo-rich breccia
pipies; and 5) pegmatites (Burt et al., 1982; Christiansen et
al., 1983; Christiansen et al., 1986).

Many authors recognize that the magma types for these
deposits are similar, possibly locally related, and were
emplaced or extruded in a similar tectonic setting (Figure
11). There is evidence supporting similar magmas and magmatic
processes in the genesis of these deposits (Burt et al., 1982;
Christiansen et al., 1983; Christiansen et al., 1986).
Bildreth (1979) demonstrates that Mo, Sn, W, Pb, Cs, Li, Be, F
and U are enriched in the initial high-silica eruption of the
Bishop Tuff.

From an exploration standpoint, the presence of base metal
veining (particularly W, and Sn) and fluorite mineralization
derived by high silica rhyolite porphyry intrusives may
reflect a porphyry molybdenum deposit at depth (Burt et al.,
1982; Christensen et al., 1986). The close spatial and
temporal relationship between major fluorospar districts and
Climax-type districts is widely recognized and accepted (Van

Alstine, 1968; Lamarre and Hodder, 1978).

46

 

 

 

 

 

 

 

 

 

 

 

 

 

 

' TOPAZ RHYOLITES

O URAD-HENDERSON
O CLIMAX

0 MT. EMMONS

O PINE GROVE

U

QUINN CANYON

Figure 11. Distribution of topaz rhyolites of the western U.S.
relative to the location of the Quinn Canyon Range, the Pine
Grove porphyry Mo system and the Grad-Henderson, Climax and Mt.
Emmons porphyry Mo deposits of the Colorado Mineral Belt (COMB)
(modified from Christiansen et al, 1986).

47

The genetic relationship between Au metallization and
rhyolite porphyry intrusion is tenuous. It is possible that
this magma type is not the main source for this metal, in many
mineralized volcanic systems. This is supported by evidence
that Au-quartz veins are sometimes found adjacent to
quartz-latitic, andesitic, and rhyodacitic intrusives, and
that Au, has been shown to be depleted in high silica magma
chambers (Sainsbury and Kleinhampl, 1969; Bildreth, 1979; and
Keith, 1980).

Based on limited Sr isotope data, Christiansen et al.
(1986) believe that topaz rhyolites have a lower crustal
source. Assuming approximate crustal province age,
Christiansen et al. conclude that the Sr isotope ratios
correspond to crustal source areas with relatively low Rb/Sr
ratios. Citing the work of Pettingillet et al. (1984),
Christiansen et al, suggest that a low Rb/Sr ratio is
consistent with a granulite grade source. The broad
distribution of topaz rhyolites in the western U.S. suggests
that there probably is significant variability in source
material (Figure 11).

Stewart et al (1977) define two major metallogenic
provinces in the Western U.S. The western province,
characterized by precious metals, antimony, arsenic, and
mercury, is delineated by a north-south trending line,
parallel to and 100km west of the western boundary of the
Quiet Zone. This line is roughly coincident with the western
edge of the Precambrian craton as positioned by Farmer and

Depaolo (1983) and the appearance of transitional assemblage

48

sedimentary rocks. The eastern metallogenic province is
characterized by base-metal and lithophile element

mineralization.

CLIMAX-TYPE PORPBYRY MOLYBDENUM DEPOSITS

The Climax-type porphyry molybdenum deposits are a
specific variety of stockwork molybdenum deposits, which are
confined in terms of geologic setting and time of
emplacement. These deposits have produced over half of the
worlds production of molybdenum, with the mine at Climax,
Colorado, being the worlds single largest producer of
molybdenum (Westra and Keith, 1981; and White et al., 1981).

Climax-type porphyry molybdenum deposits are characterized
by the style and intensity of mineralization and alteration,
age and petrochemistry of the associated granitic plutons.
Deposits discovered to data include Climax, Urad-Benderson,
Mt. Emmons-Redwell Basin, and Pine Grove (Wallace et al.,
1978; White et al., 1981; and Westra and Keith, 1981). With
the exception of Pine Grove (Utah), the Climax type deposits
are confined to the Colorado Mineral Belt (Figure 11), a
northeast-trending, mineralized shear zone cutting across
western Colorado (Tweto and Sims, 1963). Pine Grove has been
classified as a Climax-type deposit, although some authors
have expressed a degree of uncertainty (Keith et a1, 1986).

Volumes of material have been published on the Climax-type
systems. Westra and Keith (1981) propose a classification
scheme for stockwork porphyry molybdenum deposits, based on
their magma series chemistry, focusing on the characteristics

and origin of the Climax-type. Mutschler et al (1981) present
49

50

a different classificiation system for stockwork molybdenum
deposits based on pluton chemistry and age. A detailed
description of Climax and Urad-Benderson is presented by
Wallace et a1. (1968) and Wallace et al (1978), respectively.
White et al. (1981) describe the important features of
Climax-type deposits and the possible origin of magmas and
metals. These publications represent fairly exhaustive
studies of the ore-forming systems, but are a small fraction
of the published and unpublished studies.

The following section is a brief overview of the
Climax-type deposits, specifically the petrochemical signature
of the source rocks, mineralization, alteration, emplacement

history, and the origin of the granitic intrusives and metals.

Eagles:

Molybdenum mineralization at Climax, Urad-Benderson, and
Mount Emmons-Redwell is generated by granitic and rhyolitic
porphyry stocks. These stocks range in size from 1600 ft-4500
ft in diameter and were emplaced between 33 and 17 m.y. ago,
during crustal relaxation, coincident with and following
initiation of the Rio Grande Rift (White et al., 1981; Westra
and Keith, 1981; Stein, 1985 and Christiansen et al., 1986).

A key feature of these districts is the multiple intrusion
of magmas. At Climax, this results in stacked,
umbrella-shaped ore shells, which cap and correspond to
hydrothermal activity associated with successive pulses of

granitic magma. The ore is distributed in

51
quartz - molybdenite veinlets in the ore zone and the ore
forming solutions appear to have exsolved directly from the
crystallizing magma (Wallace et al., 1968; Wallace et al.,
1978; Mutschler et al., 1981; White et al., 1981; and Westra
and Keith., 1981). Ore grades range from .2% to 1.0%, with
averages between .3% and .45% (White et al., 1981; Edwards and
Atkinson, 1986).

Unlike Pine Grove, the deposits at Climax and
Urad-Benderson appear to have no associated comagmatic
eruptive volcanic units. It should be noted, however, that
they do occur in deeply eroded terranes, in part caused by
large-scale post-ore movement along major regional faults. LBy
reconstructing the Paleozoic section above the deposit, the
upper ore body at Climax was formed between 6,600 and 9,600 ft
below the surface. Molybdenum metallization at
Urad-Benderson, occurred at depths of 2,500-3,400 ft and
5,700-7,500 ft respectively (White et al, 1981). The Pine
Grove vent is believed to be eroded 1.7 km below the
pre-eruption surface (Keith, 1980; Keith, 1982; Keith et al.,

1986).

WWII

Four main types of alteration zones are found above and
peripheral to the ore bodies and plutons. In order from inner
to outer, they are; the potassic zone, the
quartz-sericite-pyrite zone (phyllic zone), an upper and lower

argillic zone and a propyllitic zone (dominated by chlorite,

'm'mm 1mm .1.

“1771.

52
clay, epidote, and calcite) Wallace et al., 1968; Lowell and
Guilbert, 1970, Mutschler et al., 1981; White et al.. 1981:
and Westra and Keith, 1981).

Tin, tungsten and base metal halos frequently develop
above, or in the case of tungsten, can coincide with the
molybdenite ore body. Buebnerite, wolframite and less
commonly scheelite are the principal tungsten-bearing
minerals. Cassiterite has been recognized by Wallace et a1
(1968), with its crystallization synchronous with tungsten
mineralization. Fluorite, and to a lesser degree, topaz, are
found in veins in the phyllic zone and the potassic zone. The
highest fluorine concentrations are found directly above the
ore zone. Paragenetically late lead, zinc, silver and to a
lesser degree copper, are deposited in zones well above the
ore bodies and adjacent to late rhyolite dikes (Westra and
Keith, 1981; White et al., 1981). Copper minerals are rare,

with a high Mo/Cu being a prominent feature of the deposits.

E ! o I I ! E '!° 1

The igneous rocks associated with Climax-type Mo deposits
are highly differentiated peraluminous~metaluminous granites
and rhyolites. They are characterized by high Si02, and
K20 (with K20>Na20) and low Mg0, Ti02, CaO and Fe.

The trace element signature is defined by an enrichment of Rb,
Nb, F, Mo, W, Th, Sn, and U, and a depletion in Ba and Sr
(Westra and Keith, 1981; and White et al., 1981). Although

the granitic rocks display these general characteristics,

53

there is a variation in the range of trace element values
between districts. The typical trace element ranges, as
defined by Westra and Keith (1981), are; 200-800 ppm Rb (with
isolated values exceeding 1000 ppm), 25 to greater than 250
ppm Nb, generally less than 100 ppm Sr (frequently less than
20 ppm), less than 150 ppm Ba, and less than .2% Ti02. Values
for the Climax district and nearby Urad-Benderson tend toward
the extremes of the ranges, with values of less than 10 ppm Sr
and less than 80 ppm Ba common. I
Published REE data is virtually nonexistent, but Stein
(1985) states that REE patterns display relatively large
negative europium anomalies. Analyses from this study, also
show significant Eu anomalies and a La/Ybu ratio of about
6. Detailed petrochemical comparisons with the Quinn Canyon

high silica rhyolites follows in a later chapter.

Wrecks

Two main theories have been proposed for the origin of
Climax-type granites and rhyolites. The first, supported by
Westra and Keith (1981), Lamarre and Bodder (1978) and Van
Alstine (1976), is a mantle source for the magmas and
molybdenum. The second theory precludes a mantle source and
suggests a middle and/or lower crustal source for the magmas.
This theory, based on isotopic evidence, has found wide
support among various workers, such as White et al. (1981),
Farmer and DePaolo (1984), Lipman et al. (1978), Zartman

(1974) and Stein (1985).

Flt.

Westra and Keith (1981) propose that melting of the upper
mantle above a subducting slab produced alkali-calcic magmas
enriched in volatiles, such as F, and incompatible elements.
The release of F-bearing hydrous fluids was caused by the
dehydration of phologopite in the subducted slab at depths
greater than 250 km. This hydrous fluid would initiate
partial melting of mantle peridotite. The resultant melt
would be highly enriched in K, F and other incompatible
elements. Mo and Sn, the authors suggest, would be added as
the melt was introduced into the upper mantle.

Work by Zartman (1974) and more recently Farmer and
DePaolo (1984) and Stein (1985) utilizes Pb, Nd, and Sr
isotopes in defining a probable source for Climax-type
deposits, particularly in the Colorado Mineral Belt.

Molybdenum deposits of the Colorado Mineral Belt fall into
Area I of Zartman’s (1974) lead isotope classification system
for the Western U.S. In this province, igneous rocks and ore
samples displayed a considerable range in lead isotope
composition, but were generally indicative of a lower crustal
source or possibly upper mantle.

Several conclusions were reached by Farmer and DePaolo
(1984), among these were: 1) Cu-mineralized granites in
general have a mantle source, 2) Climax-type Mo-mineralized
granites have a crustal source, with a mantle component being
highly unlikely, and 3) the granites at Henderson appear to

have a Rb-rich middle crustal source, because of their high

wk “4. ._._

55

6 Sr . The authors propose that this high 6 Sr may be a
result of contamination by radiogenic Sr derived during
hydrothermal alteration of the nearby Silver Plume Granite.

Stein (1985) suggests based on Pb and Sr isotope data,

that the Climax-type granites have a 1.4 b.y. lower crustal
source, which had previously undergone granulite facies
metamorphism. This source material probably had a low U/Pb
ratio, but a close to average Th/Pb ratio. Stein states that
Climax-type granites cannot be simple mantle melts. Likewise,
a mantle source contaminated by upper crustal material or a
purely upper crustal source seems implausible. Stein
concludes that Climax-type granites are not differentiates of
or genetically related to the source material producing the
older intermediate intrusives in the Colorado Mineral Belt.
The author further proposes that this apparent difference in
source material may be due to "localized" or "specialized"
partial melting of a lower crust material during the change

from compressional to extensional tectonics (p. 271).

DISCUSSION AND RESULTS

9 . C 5.]: . E 1

Several groups have been delineated according to trace
element chemistry and to a lessor extent, petrography.
Although a quantitative and unequivocable determination of the
relationship among these groups or individual rocks is not
attempted in this study, some meaningful conclusions have been
reached.

Although not definitive, some variation diagrams,
including a process identification diagram (see Minster and
Allegre (1978) and Allegre and Minster (1978) for explanation)
of Rb/Y vs Rb, suggests that the rhyodacite dome samples are
not cogenetic with the other rhyolitic hypabyssal samples
(Figure 12).

The group of rhyolites displaying the most extreme
enrichment in the incompatible elements (Nb, Y, Th and Rb)
appear to have a different source than the muscovite-bearing
rhyolite samples (group 1). The strength of this comparison
is that both "groups" have at least some characteristics of
being highly differentiated, particularly in terms of the
depletion in Ba, Zr, La, Ti, and Eu. The lines of evidence
supporting a different source are:

1) The muscovite-bearing rhyolites are more peraluminous

(Figure 13).

57

10 1 r I I

 

Rb/Y

   
 

FRACTIONAL
CRYSTALLIZATION
— - — — — fi —. .— ‘

it if

 

 

L

l l

 

100 200 300 400 500 600
Rb

RHYODACITES

LOH SILICA RHYOLITES

RHYOLITES

MUSCOVITE-BEARING RHYOLITES 2
MUSCOVITE-BEARING RHYOLITES 1
INCOMP. ELEMENT-RICH RHYOLITES

*l3 DIX-+(3

Figure 12. Plot of Rb/Y vs Rb indicating possible partial
melting (solid lines) and fractional crystallization (dashed
lines) trends for the Quinn Canyon silicic rocks.

Ln
0)

 

 

 

 

-
-
h

1 1
A .5 .6 .7 .8 .9 10

N020 + K20+ CaO(m°l%)

* INCOMPATIBLE ELEMENT-RICH RHYOLITES

C] MUSCOVITE-BEARING RHYOLITES GROUP 1

Figure 13. Plot of molecular A1203 vs molecular

NazO+K20+CaO for the incompatible element-rich rhyolites

and muscovite-bearing rhyolites (group 1). Sample QC-13 has
not been included because of alteration affecting NaO.

£3

2) The muscovite-bearing rhyolites appear highly
differentiated, being equally or more depleted in Eu, Zr, Ti,
Ba, and LREE, yet are only minimally enriched in Th, Nb, Y and
Rb, when compared to samples Qc-19, Qc-ch, Qc-13 and Qc9-3
(Figure 14).

It may be possible to explain the trace element
composition of the muscovite-bearing rhyolites by invoking a
similar source which has undergone a previous partial melting
event. It is more likely, however, considering one group is
more peraluminous, that they are not cogenetic.

Fluorine-rich rhyolites, such as topaz rhyolites, are
generally classified as A-type or anorogenic granitiods, and
are believed to be produced by the partial melting of a lower
crustal source (Burt et al., 1982; Christiansen et al., 1983;
Christiansen et al., 1986). The muscovite-bearing rhyolite
group 1, relative to the incompatible element-rich rhyolites
at Quinn Canyon have a trace element signature more similar to
that of S-type granites. These characteristics are a
depletion in Ba, Rb, Zr, Y, Ce and Nb and an enrichment in Sr,
compared to A-type granites (White and Chappell, 1983).
Although the muscovite-bearing rhyolites may not be the
high-level equivalent of ”pure" S-type granites, they seem to
have many of the trace element compositional features.

Farmer and Depaolo (1984) indicate that peraluminous
granites (such as S-type) may be derived from a midcrustal
source (p 10,151). Farmer and Depaolo (1983) state that the
compositional differences seen in miogeoclinal granites of the

Eastern Great Basin may be due to the assimilation of various

Zr

Zr

Eu

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

60
200 T T T T T T T T T T T T T T T T T T T T T T T T T T T
I I r I q I Ix I 35
1- * -l L .1
I- 4
In: S
I- * a- b * ‘- T.
1- -1 _.I
L. n . " an
1— E _ III
F a d '- "
. m . m ,,
50 h L J 1 l l l I I L 1 l l l L l 1 l L L l L L l l l l l l l 1 l L l l o
20 200 20 200
Y Y
200 T T T l T r T I 35
X " *
: '9! ‘ - 1
L m L
1- -I
1- * .1 '- * ‘ ID
. - * —l
I- * q
_ m m _ m as
- ETD - ' El ‘
. m "‘ - *
1- -I up!!!
" El
50 L l l L l_ l l l o
150 300 150 500
an RD
0 I 5 h T T f I T T T T T T T T T T T T T 400
m
1- .1 ,. .1
fi
1. -4 *
.. .1. 8
I. X
K -I
U [I] fl i
.. El -
3* 1* ' I .11
fi
0 . o J g m l l l i L l l L L l l L L l l l o
0 100 0 100
ND ND

El WSCOVITE-BEARING RHYCLITES 1
X INCOHP. ELEMENT-RICH RHYGJTES

Figure 14. Plots of Ba and Eu vs Nb, La and Zr vs Rb, and La
and Zr vs Y for the incompatible element-rich rhyolites and

muscovite-bearing rhyolites (group 1).

61

amounts of pelitic sedimentary material, with the magma body.
This suggests that the muscovite-bearing rhyolites of Quinn
Canyon may have a source higher in the crust than that assumed
for the incompatible element-rich rhyolites of this suite, or
may be the result of the mixing of upper crustal pelitic
sedimentary material with the parent magma.

Sr, Nd and/or Pb isotope values are needed to accurately
constrain the source material of the silicic rocks, both
locally and for regional comparisons.

Both groups of rhyolites are probably a result of some
degree of crystal fractionation from a parent formed by small
amounts of partial melting. Samples Qc-20, Qc-IQC, Qc-13, and
Qc9-3 (incompatible element-rich rhyolite group) are the
result of large degrees of fractionation from a more
"primitive" rhyolite as suggested by Christiansen et al.
(1986). The less evolved, "rhyolite group" probably is
genetically related to and possibly the "parent" of one or
many of the incompatible element-rich rhyolites. This
conclusion is supported by some process identification
diagrams, including a plot of Rb/Y vs Rb (Figure 12).

The extreme enrichment in incompatible elements evident in
sample Qc-19, may indicate that it has been the most
fractionated, and/or possibly is not genetically related to
other members of its "group".

Qc-20 has enrichment trends which "fall“ between the two
groups, but is distinctly peraluminous, indicating that it may
have been derived from a source similar to that which

generated the muscovite-bearing rhyolites (group 1), and/or

62

perhaps is directly related by fractional crystalization.

P o ' C ' o s

The focus of this comparison is on trace element
composition. Specifically, the enrichment in the incompatible
elements that are characteristic of the granitic and rhyolitic
rocks which generate porphyry Mo, Sn, and W, and
lithophile-type (F, Be, Li, and U) mineralization. This
petrochemical signature is marked by an enrichment in Nb, Y,
Th, and Rb and depletion in Sr, Ba, Zr, and Eu. Major element
chemistry is similar for most F-rich rhyolites and granites,
and alone, is not adequate for comparison of magma chemistry,
relative to the potential for generating the metallization and
mineralization mentioned above.

Analyses of samples from Urad-Benderson (8), Mount Emmons
(1) and Pine Grove (3) form the basis of this comparison,
supplemented by data taken from published and unpublished
studies. The arthimetic mean of selected trace element
concentrations are presented in Table 3. Major and trace
element compositions of all samples are available in Appendix
A.

Most of the Quinn Canyon (QC) rhyolites are much less
enriched in the key incompatible elements, even comparing
samples with equivalent Si02. Because the source rocks of
Climax-type deposits and other F-rich rhyolites represent
extreme differentiates, only the high-silica incompatible

element-rich and relatively unaltered samples are considered

63

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65
in detail. The members of this group are not necessarily
genetically related.

"Enrichment factor diagrams” of this group (QC—13, Qc-19,
Qc-19c, Qc-20, Qc9-3) divided by the mean of the Pine Grove,
Brad-Henderson and Mt. Emmons samples illustrates the relative
enrichment trends (Figure 15).

Overall, the trace element composition of the QC rhyolites
is more similar to the Pine Grove rhyolites, than to the
Urad-Benderson and Mount Emmons granitic rocks (traditional
"Climax-type" source rocks). This is particularly evident
when examining Rb values and REE distribution. Compared to
Urad-Benderson, the QC rhyolites are less enriched in Rb, Nb,
and Th and less depleted in Sr and Ba. They are, however,
more enriched in BREE and depleted in LREE, resulting in a
much more "flat" REE distribution pattern and display a
greater negative Eu anomaly (La/Ybn average of 6 for
Urad-Benderson and 1.5 for QC) (Figure 16). Y, behaving in a
similar manner as the BREE, is enriched in the QC rhyolites,
and Ti depleted (Figure 15).

Published data from the Climax Mine, Henderson and Urad,
support the trace element differences noted previously.
Samples from the Climax Bartlett and Wallace Stock and Late
Rhyolite dikes have Rb>600 ppm, Sr<40 ppm, Nb>100 ppm, and Y
values from 38-42 ppm (Stein, 1985).

Examination of the one Mount Emmons analysis from this
study shows a lower Rb and Y value than Urad-Benderson or
Climax, but a similar REE distribution pattern (Figure 16).

Similar Rb, Y, REE and other key element values for the Mount

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Figure 16. Chondrite-normalized REE patterns for Mt. Emmons and
Brad-Henderson samples.

69

 

 

 

 

 

 

 

 

PINE GROVE
_ I I I I I I I I I I I I I I I 4
L— .J
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l_ 1 1 L L 1 L 1 L 1 L 1 1 1 1
La Ce Sm Eu Tb Yb Lu

Figure 17. Chondrite-normalized REE patterns for rhyolites of
Pine Grove (this study) and Pine Grove (Keith, 1982).

7O

Emmons Keystone Porphyry and Redwell Pipe, are listed by Stein
(1985).

Two sets of Pine Grove (PG) data are used for comparison.
The first set is from a group of undivided surface samples and
samples taken at depth, analyzed at Michigan State
University. The second group consists of the average of
analyses of one sample each of the Pine Grove Porphyry, Phase
Five Porphyry, and Phase Four Porphyry (Keith, 1982).

The QC rhyolites are moderately enriched in Nb, Th, Zr, Y,
somewhat depleted in Rb and have similar La/YbN ratios
(1.5-2.0), to those of Pine Grove (PG). Eu and Sr values
differ considerably between the two sets of PG data, with the
analyses from Keith (1982) displaying larger Eu depletions and
higher Sr concentrations (Figures 15 and 17).

The intrusive Pine Grove Porphyry has a Mo concentration
of 4 ppm and Sn concentration of 8 ppm (Keith, 1982), compared
to 3 ppm and 15 ppm respectively, for the most enriched QC
rhyolite dikes. The mineralized Phase Four and Five Porphyry
at Pine Grove have Mo concentrations ranging from 65-150 ppm.

The low La/Ybn displayed by the most evolved QC and PG
rhyolites are common among F-rich rhyolites such as topaz
rhyolites. Topax rhyolites as defined by Christiansen et al.
(1986), typically have La/YbN ratios of 1 to 3, are enriched
in the incompatible lithophile elements (Rb, U, Th, Ta, Nb, Y,
Be, Li, and Cs) and depleted in Sr, Eu, Ba, Ti, Zr, and Hf.
Fluorine concentrations range from less than .2 to more than
1.0 wt%. F/C1 ratios are high (Christiansen et al., 1986).

Although topaz has not been observed in thin section, fluorite

71
is present in some of the QC rhyolites and many key elements
for most samples fall within the range published by
Christiansen et a1 (1986) for topaz rhyolites of the western
U.S. (Figure 18). It must be emphasized that the mean values
of the elements Rb, Ba, Sr, and Th, barely lie within topaz
element ranges.

Pb, Mo, and Sn data for topaz rhyolites is limited, but in
general, these elements occur in higher concentrations than
the few analyzed samples from QC. It is important to note
that the analysis of these elements and F (Christiansen et
al., 1986) are from vitrophyres. Devitrification and/or the
release of magmatic fluids from the QC intrusives would
probably result in a lowering of these element levels,
certainly at least volatiles such as F.

Christiansen et al state that although the Pine Grove
rhyolite tuff lies within the range of all topaz rhyolite
compositions, relative to topaz rhyolites in SW Utah, it is.
enriched in Ba, Sr, Sc, and A1 and depleted in F, Zr, Rb, Nb,
Ta, Th, U, Yb, Y and M0. The authors imply that the 22-23
m.y. rhyolites at Pine Grove are not considered "true" topaz
rhyolites, because of their only marginally qualifying

geochemical signature.

5 l . 1 . = . I . v'i

The Quinn Canyon Range is not only geographically close to
the Pine Grove volcanic system but has a similar age of

rhyolite emplacement (23 m.y. for Quinn Canyon, and 22-23 m.y.

72

 

100 I I I I I I I I I I I

UIUh

I r TUI'IU' I
l l 1 hill. 1 L L LLJJI

JL

1 1 114111

RDCK/HGM-i

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T T t UUTUT'
L L 1141111

 

 

O 01 1 L 1 1 1 1 1 1 1 1 1
no Ba Sr Eu La 5. Yb Y Th Zr ND

 

fl MEAN 0F PCPULATION

Figure 18. Range of values of selected trace elements for topaz
rhyolites of the western U.S. compared to the Quinn Canyon
incompatible element-rich rhyolites. All values normalized to
U.S.G.S. standard RGM-I. Vertical bars represent ranges of
concentrations for the Quinn Canyon incompatible element-rich
rhyolites (after Christiansen et al., 1986).

73
for Pine Grove; Keith, 1982; Keith et al., 1986), a similar
general geologic setting and is located along the same
mineralized lineament. Based on these similarities, it is not
surprising that the Miocene rhyolite flows and intrusives of
the Pine Grove area are compositionally closer to the Quinn
Canyon rhyolites, than the granitoid porphyries associated
with more traditional Climax-type deposits located in the
Colorado Mineral Belt.

Major differences in rhyolite trace element composition
and associated metalization may be attributed to a difference
in source material. As mentioned earlier, Climax-type
granitic source rocks are believed to have a lower crustal
source, which has undergone granulite facies metamorphism
(Zartman, 1974; White et al., 1981; Farmer and Depaolo, 1984;
Stein, 1985; and Christiansen et al, 1986).

Sr and Nd isotope data are not presently available for the
Pine Grove deposit. Pine Grove is located at the boundary of
the Rb-depleted lower crust (granulite facies metamorphism) of
the western carton, as delineated by Farmer and Depaolo (1983)
(Figure 19). This limit is based on a discontinuity in
granite Sr isotope composition, obtained from samples taken
across the northern Great Basin and from Tertiary granites of
Colorado. Low 68: values are interpreted as being derived
from a granulite-grade lower crust, while higher 63:- values
of the miogeocline to the west are suggestive of a basement
with no "depleted" lower crust. A second discontinuity in
Elk! and Es:- to the east marks the basement edge. The

Quinn Canyon Range is located significantly west of the

74

 

 

 

 

 

 

Figure 19. Map showing the intrepreted western limit of
Precambrian basement (line 1) and western limit of a granulite
grade lower crust (line 2) (Farmer and Depaolo, 1983). Pb
isotopic provinces from Zartman (1974) are defined by stipled
bands and identified by Roman numerals. See text for further
explanation (modified from Farmer and Depaolo, 1983).

75

granulite facies line (Figure 19), in the area of
"non-depleted" lower basement.

Four Pb isotope provinces (Ia, Ib, II, III) are defined by
Zartman (1974) for the western U.S., with each province
reflecting a somewhat specific source material. The
Climax-type deposits all occur in area Ib, characterized by a
fairly unradiogenic PbZOG/Pb204 ratio of 16.2-18.8, 3
PbZOB/Pb304 of 36.5-39.9, and a basement rock age of 1.8
b.y. The Quinn Canyon Range lies in a transition area between
area Ib and area II, characterized by a more radiogenic Pb
isotope composition (PbZOG/Pb2°4=19.1-19.7,
Pb208/Pb204z38.9-40.3) (Figure 19). According to Zartman,

a PbZOG/Pb204 ratio of 18.8 is the most important factor
separating region Ib and II. The placement of this boundary
relative to the Quinn Canyon Range is based on a Pb isotOpe
values from galena sample taken from a vein adjacent to a
dacite dike (PbZOS/Pb3°4=18.886, Pb207/Pb2°4=15.640,
szoa/Pb3°4=39.077; Zartman, 1984). Pine Grove also lies
within region Ib, but very close to the II boundary (Figure
19). A galena sample taken from the southern Wah Wah
mountains, Utah (probably from the Pine Grove system) has a Pb
isotope signature similar to the Quinn Canyon transitional Pb
isotope signature (PbZOS/Pb2°4=18.921,

Pb207/Pb2°4=15.780, Pb303/Pb20‘=39.988) (Stacey et

al., 1968).

Doe and Zartman (1979) state that Mesozoic and Cenozic
igneous rocks of region II, could be generated from the

partial melting of either upper crustal miogeosynclinal rocks

76

or from a non-cratonized Precambrian source (never
metamorphosed above amphibolite facies). The Pb isotopic
signature of these materials are not easily differentiated and
would exhibit a similar relationship to the Pb isotopic
signature of granulite facies lower crustal source materials.

Stein studied Pb and Sr isotope data from the Colorado
Mineral Belt and concluded that even within a somewhat
restricted region, Pb isotope composition was significantly
different among Laramide-Tertiary granites. These small yet
significant variations Stein attributed to subtle differences
in age and composition of source materials. Stein also
.recognized that granitiods associated with major molybdenum
mineralization had noticeably less radiogenic Pb206 and
Pb207 isotopes, compared to barren granitiods.

Paired Pb-Pb diagrams indicate that Urad-Henderson,
Climax and Mount Emmons have considerably less radiogenic
Pb303 and Pb207 and somewhat similar Pb2°8 isotope
values than those of Quinn Canyon and the southern Wah Wah
mountains (Pine Grove). The data from Pine Grove and Quinn
Canyon plot above the average Pb growth curve of Stacey and
Kramer (1975) for both Pb207 and szo8 vs Pb206, while
those from Climax, Urad-Henderson and Mount Emmons plot below
the Pb3°7 curve and above the szo8 curve (Figure 20).

This relationship may suggest that the source area of the
traditional Climax-type granitiods had a higher Th/U ratio.
The apparent depletion of U relative to Th and Pb, in the
source material of the Climax-type granitiods, has been

compelling evidence for a granulite-grade lower crustal source

208 204
Pb / Pb

 

 

 

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206 204
Pb / Pb

. Pb - K FELDSPAR

A Pb-GALENA

Figure 20. Lead isotope data for Pine Grove (PG), Quinn Canyon
(QC), Urad-Henderson (UH), Mount Emmons (ME) and Climax (CM).
The lead growth curves are from Stacey and Kramer (1975). Data
from Zartman (1974), Stacey et al. (1968) and Stein (1985)
(after Stein, 1985).

78
(see Stein, 1985 for further discussion).

Although detailed Pb and Sr isotope study is needed at
Quinn Canyon and Pine Grove, limited data suggests that both
of these Miocene igneous and mineralized systems may be the
result of partial melting of a non-depleted or non-cratonized
source, or at least an appreciably different source material
than that which gave rise to granitiods genetically related to
metallization at Urad-Henderson, Climax, and Mount Emmons.

The implication of Stein’s work is that a specific source
material is necessary for the development of economic-grade
molybdenum mineralization. There appears to be a direct
correlation between the amount of radiogenic Pb isotopes,
enrichment in incompatibles (particularly Rb) and associated
"economic" potential for molybdenum mineralization.

The Pine Grove system is enigmatic in terms of trace
element chemistry, and tentatively, Pb isotopic signature of
source material. Perhaps a redefinition of the
characteristics of Climax-type porphyry deposits, is
warranted, and a re-evaluation of the importance of a specific
type of lower crustal source material for the development of a
major porphyry molybdenum deposit. Alternatively, perhaps the

Pine Grove deposit should not be considered a "Climax-type".
WOW

Considering the similarities in trace element composition,
location, age of emplacement and apparent isotopic signature

for the Quinn Canyon Volcanic system and the Pine Grove

79
system, the possibility of some form of molybdenum
mineralization at depth in the Quinn Canyon Range cannot be
ruled out. It is unlikely that mineralization would be
"economic" because: 1) the "ore body” or pluton(s) would be
found at considerable depth, due to the lack of "dissection"
evident at Pine Grove and other Climax-type deposits, 2) the
rhyolite samples from QC with the most "economic promise”
represent only a small portion of the rhyolites sampled and 3)
the trace element signature and source material of the QC
rhyolites is significantly different than that of Climax and
Urad-Henderson, the most prolific producers of molybdenum in
the world.

The presence of fluorite, base metal vein deposits,
anomalous concentrations of Mo, W, Pb, and Sn associated with
rhyolite intrusives, and some rhyolite dikes being chemically
similar to topaz rhyolites suggests that the Quinn Canyon
Range has some potential for hosting Be, Li, U and Sn

mineralization.

CONCLUSIONS

The major conclusions of this study are:

1) Rhyolitic hypabyssal and extrusive rocks in the Quinn
Canyon Range exhibit considerable chemical variation. Most of
the intrusive and flow samples collected can be placed into
six groups, based on trace and to a lesser degree, major

element chemistry.

2) Two different groups of "differentiated” high-silica
rhyolites can be delineated; a more peraluminous,
muscovite-bearing group, which is moderately enriched in
incompatible elements and depleted in LREE, MREE, Ti, Ba, Zr,
and Hf and an incompatible-element rich group having
relatively high concentrations of Rb, Y, Nb, Th, and BREE and
depleted in Ti, Ba, Eu, Hf and Zr. Chondrite-normalized REE

patterns for this group are flat with low La/YbN ratios.

3) The differences in trace element chemistry of these groups
could be due to subtle differences in source material, with
the muscovite-bearing rhyolites perhaps being the result of
partial melting at a different crustal level or asssimilation

of upper crustal pelitic rocks.

80

81

4) The most incompatible element-rich rhyolites are probably
the result of fractional crystalization of a magma formed by

small degree of partial melting.

5) The trace element signature of the Quinn Canyon rhyolites
is considerably different than the granitic rocks which source
the mineralization at Urad-Henderson and Climax. These
differences are apparent in the lower enrichment in most key
incompatible elements, an enrichment in HREE, and a depletion

in LREE.

6) The rhyolite intrusives and extrusives associated with
molybdenum mineralization at Pine Grove have in many respects
similar trace element enrichment trends, as those seen in the
most incompatible-element enriched Quinn Canyon rhyolite
dikes. This is especially true of REE concentrations and

ratios.

7) According to trace element chemistry, the most "evolved"
Quinn Canyon rhyolites probably can qualify, with minor

exceptions, as topaz rhyolites (Christiansen et al., 1986).

8) Limited Pb isotope evidence taken from published
literature for the Quinn Canyon and Pine Grove area suggests a
non-depleted or non-cratonized source, unlike the
granulite-facies lower crustal source proposed for Climax-type

granitiods of the Colorado Mineral Belt.

82
9) Based on similarities with the Pine Grove system in
geologic setting, trace element chemistry of rhyolites, age of
emplacement and to a lesser extent isotope data, the
occurrence of molybdenum mineralization at depth in the Quinn
Canyon Range cannot be precluded at this stage of study. It
must be emphasized that the generating pluton and
mineralization, if any, would have a significantly different
source material, than that of the Climax-type deposits of the

Colorado Mineral Belt and would likely be non-economic.

Egtgzg wgrk

The relationship between the Quinn Canyon rhyolites, topaz
rhyolites associated with high level lithophile-type
metalization and the granitiods associated with porphyry Mo,
W, Sn mineralization, warrants further attention. Comparisons
would be greatly enhanced by a detailed Pb, and Sr isotope
study of Quinn Canyon and other mineralized intrusive centers
(such as Pine Grove). This type of detailed study, would help
elucidate the importance of a granulite facies lower crustal
source material, for producing magmas which later generate the
types of metallization mentioned previously.

A exhaustive K-Ar, Pb and Sr isotope study, would be
helpful in determining the genetic relationship between
silicic rocks and the igneous and metallization history of the

Quinn Canyon volcanic center.

APPENDIX A

APPENDIX A

Chemical Analyses

Table 4. Chemical analyses of the Quinn Canyon silicic rocks.

6mm 1 1 1 2 2 2 2 2 2
89111111000311 069.1 0092 0011A QC11B QCJW48 QC1SN 00-10 06—96

Wt 16
5102 71.06 70.54 70.13 71.04 70.63 70.76 72.60 71.14 7122
T10 0.51 0.49 0.45 0.32 0.33 0.28 0.23 0.36 0.33

711,5, 14.65 1466 13.96 14.15 1426 13.60 13.65 14.16 14.15
F60 336 2.80 274 2.63 2.75 2.60 201 2.77 235
1100 0.06 0.05 004 0.05 0.05 0.06 004 0.04 0.04
M90 097 053 036 030 0.45 055 0.40 0.42 0.42
030 3.13 2.44 2.58 1.37 1.45 151 0.96 097 0.96
N620 2.67 2.76 2.67 2.75 2.82 1.66 2.66 257 3.36
190 350 3.45 3.43 5.02 456 469 4.76 495 523
P10, 0.10 0.09 0.09 006 0.06 0.06 0.04 0.09 0.09

138 132 170 162 158 147 155 174 216
119 146 184 225
19 16 20 37
186 180 164 318
14 15 13 24
5.1 5.5 4.0 8.4
14.2 14.1 14.3 19.0
11 4 4 4

205
38
343 367 288 343 329
24
8 7
18.5
5
7.2 5.4 5.1 5.8 6.6 5.8 4 .4 5.8 5.6
1270
75.8
140.6
9 4
1 .96

9.1 7.7 8.0 8.7
19.1 19.2 18.5 19.6

551 806 490 1367 1434 1516 1304 1095

81 .0 75.1 62.7 57.2
138 .6 134.1 122. 4 124.5
10.8 10.8 8.9 10.2
1.37 1.36 1.12 2.01 2.13 1.91 1.81 1.83
0.90 0.87 0.85 1.20 1.13 1.16 120 1.18 1.23
1.83 1.94 1.69 3.31 3 .43 3.32 3.34 3.22 2.99
0.33 0.30 0.27 0.49 0.52 0.51 0.50 0. 42 0.50

33.3 29.9 30. 7 64.3
69.0 65.7 67. 4 134.1

2539591: FQQEE§N<Q§§QE

83

Table 4 (cont.

Group 3
1m 14

5102 7239
no, 022
41,0, 1353
F60 205
14110 0.05
M90 024
030 1.01
N320 3.41
190 4.65
9,0, 0.05
2914

011 0
211 53
Rb 169
s: 130
v 46
z: 276
Mb 37
111 79
711 223
Or 4
Se 4.0
Ba 737
L3 595
04 131.0
8111 9.4
Eu 1.07
T0 122
Yb 4.05
Lu 0.60

gssgso

«figfi

5.3

75.42
0.09
12.82
1.30
0.02
0.1 1
0.43
3.28
5.08
0.01

158
37

244

7.5
20.1

3.0

44.3
97.9
10.8
0.82
1 .32
4.46
0.67

84

3 3
008 009.5
76.71 76.88

0.09 0.1 1
12.66 12.72

0.83 1.18

000 0.02

0.00 0.06

0.18 0.1 1

3.” 2.45

4.89 5.12

0.01 0.03

0 11

9 12

185 220
52 97
48 38
247 187
38 22
8.0 5.0
16.9 28.2
3 3
2.1 4.2
205 612
27.7 45.8
65.3 93.0
7.6 8.5
0.56 0.79
1.03 1.32
3.58 2.93
0.60 0.49

3
00-14

75.48
0.10
13.35
1 .43

0.17
0.22
3.48
4.37
0.01

41
127

51

6.8
19.0

4.8

52.0
108.0
1 1.7
1 .33
1 .26
4.52
0.64

3

3

3

009-400-6001A

7734
1252

27.9

3.7
312
32.8
76.1
5.5

0.31
1 .27

3.25
0.51

197

7.1
24.0

3.4
185

44.2
97.6
10.5
0.29
1 .43
4.07
0.69

74.70
0.13
13.20
1 .47
0.04
0.07
0.35
3.19
5.40
0.02

0

243
125

214

8.5
35.8

6.0

65.8
127.2
8.7
1.22
1.56
3.03
0.51

()1

Table 4 (con’t).

Mm 3 4 4 4 4 5 5 5 5
swam-12 069-9 2016 2012 2015 2010 2011 2008 005

510, 73.60 7599 7564 75.79 76.72 76.13 7722 76. 72 79.19
1102 0.10 0.12 0.13 0.13 0.13 004 0.04 0.03 0.03
41,0, 13. 14 13.03 13.14 13.66 12.79 1255 12.75 1260 11.65
F60 165 061 094 0.16 0.63 029 0.73 0.61 037
14110 004 000 0.01 0.00 0.00 001 0.01 0.00 0.00
M90 009 0.04 000 000 0.00 000 0.00 000 0.00
060 127 0.19 030 026 0.13 0.16 0.16 0.16 0.11
N330 223 236 299 2.69 227 295 333 3.01 226

190 5.44 5.81 5.38 554 5.17 4.62 4.41 467 469
P20, 000 0.01 0.02 001 000 001 001 001 0.00
0 0 0 0 0 0 0 0 0
25 15 19 9 7 7 7 5
205 233 224 224 262 213 299 297
65 102 116 94 31 45 27 29
61 30 37 25 43 47 46 46
256 175 226 230 91 106 101 93
37 21 22 24 19 24 25 23
63 52 6.5 6.9 39 4.1 4.1 36
272 34.1 205 252 335 212 15.1
13 2 1 4 1 1 1 3

3.2 4.5 4.8
71 1 717 996

5.5 2.8 3.0 3.1 2.9
46 146

54.6 ND 18.7
111.4 34.7 22.4
11 .4 3.9 3.7

10.3 4.7 12.9 5.1 2.9
16.6 4.8 36.4 7.1 6.9
2.1 1.6 4.8 2.0 1.6
0.23 0 .02 0.11 0.03 0.16
0.93 1 .13 1.48 1.03 0.92
2.66 3.55 3.89 3.39 2.94
0.40 0.53 0.60 0.53 0.48

130 1.17 1.40
4.46 274 213
0.47 0.44

6395’95 S’S’QEEEN‘Z’PB‘PE
8
2..

§§§§$§§ $2~E28§aégz
§
2
a:

,—
1:
.°
8

:3 13

Table 4 (con’t).

510, 6001 77.35 7563 7660 75.66 7563 7694
no, 003 0.09 0.05 0.03 005 0.04 0.04
41,0, 1126 12.55 1213 1279 12.37 1220 11.69
F00 091 0.87 104 0.33 1.11 0.66 0.80
14110 000 001 ' 002 001 0.03 0.01 0.06
M00 0.03 001 020 0.03 020 0.62 0.02
020 020 026 069 0.16 126 1.44 026
116,0 206 230 274 293 220 0.14 336

491 492 476 460
0m 0m 0m 0m

:36
Or
86
0‘
88
0‘
SB

0
11 50 45 45
325 541
19 17 20 22
70 118 109 173
111 177 188 130

4.6 8.0 9.1 7.5
24.9 30.0 36.2 51.3

3.0 2.0 1.0 2.7
149 148 263 168

3.6 8.0 15.8 6.7 19.7 32.8 14.7
5.6 13.4 40.2 16.8 39.3 27.7 45.6
1.8 1.7 12.8 5.2 13.5 12.1 11.5
0.09 0.08 0.19 0.09 0.08 0.35 0.04
0.83 1.04 1.84 1.22 1.97 2.08 2.44
2.94 2.49 5.69 4.41 8.12 6.90 8.68
0.51 0.41 1.12 0.74 1.00 1.05 1.58

5639395 Eweasgu<waygi

87

Table 5. Chemical analyses of fluorite, highly altered

samples and other samples from Quinn Canyon.

WM6

531116140017 00-98 0090 1107 1001 1102b 1109 OJWZA QJWZB QJWZC
W196

510, — 71.16 72.05 7650 76.16 6134 79.71 76.40 74.96 73.93
110, — 033 033 0.25 0.17 0.04 0.16 0.10 0.12 0.11
41,0, — 1421 1436 1247 1256 11.43 10.69 13.01 12.71 1337
F90 — 255 2.49 1.40 026 0.73 064 023 153 1.33
M10 - 0.05 0.05 0.03 001 0.00 0.01 0.01 0.03 0.01
M90 — 0.44 032 021 0.00 0.00 0.00 0.63 0.16 1.63
030 — 1.29 0.49 026 022 0.06 0.17 023 1.04 037
112,0 — 3.03 2.76 0.05 2.61 0.07 2.08 0.16 0.07 030
14,0 — 5.04 5.47 4.79 524 267 4.09 6.60 566 556
6,0, — 0 06 0.09 0.05 001 0.01 0.03 0.01 0.01 0.00

0
61 43 15 26 17 16 78 73
226 209 286 133 148 237 196 159

210 83 105 38 166 11 14 15

0
40
196
253
40
339 337 253 150 188 238 264 271 278
25
8.9 8.5 7.0 3.8 9.0 5.8 8.2 7.8 8.4

20 6 20.4 23.9 27.6 34.0 13.5 15.8 18.3 20.7

3
5.4 5.5 3.3 2.7 0.8 2.8 3.2 5.2 3.9
10 1180 1120 1640 215 326 1710 1110 1111 1113

1.3 55.1 54.5 55.1 12.8 4.3 30.9 15.1 58.3 37.8
2.1 128.6 1 18.3 97.8 28.3 15.9 59.5 42.3 130.7 84.1
0.5 10.2 9.8 4.7 2.3 10.9 5.5 3.8 1 1.3 8.4
0.04 1.71 1.75 1.06 0.26 0.04 0.79 0.48 1.34 0.94
0.31 1.18 1.08 1.04 1.11 1.89 0.79 0.71 1.23 1.21
0.38 3.30 258 1.65 1.18 6.55 2.53 4.00 4.01 4.64
0.05 0.48 0.44 0.32 0.22 1 .09 0.33 0.58 0.59 0.69

Eéa‘I-PSQS 999§E§N<Q3§8§

88

Table 6. Commercial analyses by Bondar and Clegg, Ltd.

86 m ple Qc9-3 Qc-13 00.19 (Dc-190 (Dc-20
PPNI
Pb 26 17 11 21 17
Mo 1 1 3 1 2
Sn <5 <5 15 7 10
Ag <5 .5 .8 1<5 (<5

Table 7.

Chemical analyses of rhyolites from Mt.

CD

Urad~Henderson and Pine Grove.

 

 

Prim
94111111314094
Vflfi
510, 76.19
110, 0.12

O 12 60
2450' 065
MM) mm
1190 066
0%) ms
114,0 2.76

0 fl”
Pp, 0.01

49.7
71 .1
6.7
028
1 .97
4.70
0.83

659969: 66yzsw<6999§
2

Lhfl

78.35
0.12
11 .94
0.49
0.12
0.11

2.24
5.65
0.01

Wad—116110615011
0166 um 0124 m
6077 uuos UUPCX 4711140 RMPSI
77.11 79.72 70.42 77.62 7555
0.11 0.13 0.16 0.13 0.13
1326 11.66 12.70 1210 1240
0.32 1.02 0.92 0.57 134
0.13 0.42 0.05 0.05 220
0.03 0.12 0.05 0.04 0.14
0.33 0.35 0.42 0.70 057
026 0.12 029 1.93 0.16
6.40 4.61 5.65 5.54 455
0.00 0.01 0.02 0.01 0.02
11 9 7 9 14
51 276 13 6 506
636 467 476 463 459
4 14 9 36 23
67 77 62 63 73
96 115 146 127 116
113 66 66 94 66
35 42 6.4 4.6 4.0
51.7 572 725 629 59.0
97 101 121 150 146
421 46.1 64.6 424 45.9
59.4 76.6 106.1 68.1 61.1
65 6.6 76 7.0 72
0.16 0.32 0.36 0.29 026
1.82 220 255 2.26 2.16
3.53 5.04 3.64 3.97 3.92
0.62 0.67 0.65 0.69 0.66

 

M5111

75.99
0.23
12.51
1 .35
0.07
0.29
0.70
2.68
5.25
0.04

12
60
321
19
x
209
w
7.9
71.2

100

58.1
81 .3
9.4

0.36
2.45
3.33
0.63

Emmons,

 

 

P63

78.”
0.02
1 1 .39
0.53
0.02
0.04

0.88
8.16
0.01

18
16

21
70

1.2
20.4

1%

12.5
19.4
7.1

0.10
1.27
4.61
0.75

Pine Grove

PG-S

11.8

8.3
0.41
1 .32
3.59
0.58

P66

76.15
0.02
13.48
0.67
0.06
0.17
0.23
0.95
5.37
0.01

42

1 18
57
24.0
1 14
13.8
9.0
0.61
1 43

4.07
0.73

90

Table 7 (con’t).

FWBGmWHNmm1QHY

Phasera ”PineGrova PhaseFour
Pomhm POW Porphvry

 

Sample 152 142 155
88%

$0, 1583 1503 7“”
'W% 003 004 084
A90, 11.67 12.98 12.23
Hq% 087 102 087
MM) 003 042 005
M90 0.17 0.02 0.41
OK) 098 085 144
NE? 102 2J5 1J2
K, 787 an: 708
P,0, 0.02 0.01 0.02
HHI

Cu _. _. ._

21 4 84 89

Rb $5 “2 3w
8' 39 34 33

Y 82 68 78

2' 83 as 99

Nb 41 34 37

1% 3] 35 34

Th 25.5 27.8 27.4
a1 _. _. ._

Sc 78 97 78

89 um 121 12)
La 15.5 18.4 15.8
00 38.4 43.2 41.9
Sm 82 88 85

En 0J0 087 086
Tb 188 193 110
Yb 8“) 8&1 700
I8 083 102 095

* analyses from Keith (1982. Some trace element values

represent the arthimetic mean of replicate analyses.

APPENDIX B

APPENDIX B

Analytical Methods

Representative 1kg. samples were trimmed to remove
weathered rinds, ground to a fine powder (200 mesh), and dried
in an evacuated oven at 100c for 24 hours to remove excess
water.

Major element oxide percentages were determined by XRF
using fused glass wafers. Glass wafers were made by
homogenizing 1.00000 grams of sample, 9.00000 grams of lithium
tetraborate, and .1600 grams of ammonium nitrate, in a
platinum-gold alloy crucible. This mixture was continuously
agitated while being heated at 1100 C for 20 minutes. The
resultant liquid was poured into a heated Pt-Au mold and
slowly cooled to room temperature.

The trace elements Rb, Sr, Nb, Y, Cr, Ni, Cu, Zn, Zr, and
Y were also analyzed by XRF, using pressed powder pellets,
with a sample/celluolose "amercil" binder ration of 4:1.

La, Ce, Sm, Eu, Tb, Yb, Lu, Th, Hf, Sc and Cr were
analyzed by INAA (Instrumental Neutron Activication
Analysis). 1.00000 gram samples were sealed in polyvinyl
vials and irradiated for 18 hours over a 3 day period.

All analyses were performed at the Michigan State
University Department of Geological Sciences XRF and Neutron
Activation Laboratory. Analytical precision and accuracy are
given for replicate analyses of USGS standard W-2 in table 8.

91

Table 8.

Precision and accuracy.

A. Concentration 01 U.S.G.S. standard G-2 determined by X-ray fluorescence (XRF)

W136

$102
AbOa
FeO

F9203

M90
030
NazO
K20
“02
chks
MnO

OQQQOQUIQQQ 3

Mean

69.39
15.41
‘2.39

.76
1.98
4.01
4.48

.48

.13

.03

Std. Dev.

.13
.01
.01
.03
.02
.09
.01
.00
.00
.00

69.22
15.40
1.44
1.07
.75
1.96
4.06
4.46
.48
.13
.03

Reported

(Govlndaralo, 1984)

B. Concentration of U.S.G.S. standard G-2 determined by X-ray fluorescence (XRF).

PPM

Cu
Zn
Rb
Sr
Y
Zr
Nb
Ba

(410011000000 3

11
93
172
481
8
296
13
1921

Std. Dev.

2.52
.56
1.53
2.06
.56
2.00
1.16
66.40

10
84
170
480
11
300
13
1900

Reported

(Govlndaralo, 1984)

C. Concentration of U.S.G.S. standard W-2 determined by instrumental neutron

activation analysis (INAA).

PPM 71 Mean
La 4 11.7
Ce 4 22.7
Sm 4 3.3
Eu 4 1.12
Tb 4 .77
Yb 4 1.50
Lu 4 .36
H1 4 2.7
Th 4 2.2
Cr 4 88

* Fe reported as FeO.

Std. Dev.

Reported
(Govlndaralo,

1964)

APPENDIX C

APPENDIX C

Location of Samples

Sample

QC-lA
QC-BA
QC-5
QC-6
QC-8
QC-QB,C,E
QC-10
QC-11A,B
QC-12
QC-13
QC-14
QC-15
QC-15N
QC-16
QC-17
QC-19
QC-19A,C
QC-20
QCJW2A,JW2B,JW2C
QCJW4B
QCQ-l
QC9-2
QC9-3
QC9-4
QCQ-S
QCQ-Q
2007
2008
2010
2011
2012
2015
2016
2022

93

NE
NE
NW
NW
NW
SE
SE
SE
SE
SE

NW
SE
SW
SW

NW
NW

SW

Location

Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec

SeC'

Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec
Sec

b##khkkhhhbfifihhw

R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R57E
R56E
R57E
R57E
R57E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E
R56E

T2N
T2N
T2N
T2N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N
T3N

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